NYSGA 1991
NEW YORK STATE GEOLOGICAL ASSOCIATION
63rd Annual Meeting
James R. Ebert, Editor
FIELD TRIP GUIDEBOOK
Department of Earth Sciences
State University of New York
College at Oneonta
Oneonta, New York 13820-4015
Published by the New York State Geological Association. Additional copies may be obtained from the permanent Executive Secretary, James F. Olmsted. Center for Earth and Environmental Science. State University of New York, College at Plattsburgh, Plattsburgh, New York 12901.
Cover: Kingfisher Tower, Otsego lake
PREFACE AND ACKNOWLEDGEMENTS
We are very pleased to welcome you to the 63rd Annual Meeting of the New York State Geological Association and to the State University of New York College at Oneonta. Seven of this year's trip leaders were also contributors in 1977, when this department last hosted NYSGA. In the intervening fourteen years, tremendous advances have been made in our understanding of the geology of central New York.
From precision geochronology of precambrian rocks to processes of deglaciation and interaction of glacial and karst processes, geological investigation in the heart of New York is alive and well. Since 1977, we have witnessed the burgeoning of "sequence stratigraphy", the emergence of hydrogeology as a distinct subdiscipline and the expansion of process-oriented sedimentology. These changes are clearly represented by the spectrum of field trips in this guidebook. Sequence stratigraphy is applied to Lower Devonian Limestones, Middle Devonian clastics, and Quaternary sediments in Finger Lakes' basins. Hydrological studies of caves and glacial sediments are described. Physical processes of sedimentation in Devonian shelf, shore, and fluvial environments are also examined. We hope that you enjoy the field trips and other activities this weekend and that you find them stimulating and enlightening.
The core of an NYSGA meeting is obviously the slate of field trips. The contributions of all the trip leaders (especially those that submitted manuscripts early!) are gratefully acknowledged. Without their efforts, the Association could not exist.
For all endeavors such as this, there is usually an unsung hero who is largely responsible for bringing the event to fruition. For typing the announcements, executing editorial changes, generally greasing the institutional wheels and reassuring a sometimes frazzled editor and frantic president, Moira Beach is our unsung hero.
We would also like to acknowledge the herculean efforts of Wayne Byam and his staff in the college print shop, Dorothy Gill for the cover layout, Bill Harman for supplying the sketch for the cover, Charlie Winters for shooting negatives for some figures and the administration of the college for their support and encouragement.
Finally, we thank the students of the Department of Earth Sciences, especially the Geology Club under the leadership of Heather Finlayson and Kelly Milunich for their logistical and clerical assistance.
James R. Ebert, President New York State Geological Association Department of Earth Sciences State University of New York College at Oneonta Oneonta, NY 13820-4015

SP1: SUBSURFACE GEOLOGY OF THE FINGER LAKES REGION
HENRY T. MULLINS, ROBERT W. WELLNER, JOHN L. PETRUCCIONE, EDWARD J. HINCHEY and STEVEN WANZER
Department of Geology
Syracuse University
Syracuse, New York 13244
INTRODUCTION
The Finger Lakes of central New York State (Fig. 1) have long been recognized as the product of continental glaciation. In 1868 Louis Agassiz spoke of the "glacial heritage" of the region (Coates, 1968) and since then numerous glacial geomorphic investigations have been undertaken (Mullins et al., 1989). However, most previous studies dealt with surficial features and because of this relatively little was known about the subsurface Quaternary geology of these world renown lakes.
That the Finger Lakes are deeply scoured and infilled by thick sediment sequences was known from bathymetric surveys of the lakes completed in the late 1800's (Bloomfield, 1978), a drill record of 1,080' (329m) of unconsolidated sediment at Watkins Glen (Tarr, 1904), and the publication of a line drawing of one seismic reflection profile from Seneca Lake (Woodrow et al., 1969). This paucity of subsurface geologic data has been a major void in our knowledge and understanding of the geologic history of the Finger Lakes. Many questions remained unresolved including the most fundamental --"When and by what processes were the great Finger Lakes troughs eroded?" (Bloom, 1984, p. 61). Based on the one seismic reflection profile from Seneca Lake, Bloom (1984) noted that bedrock beneath the lake floor was more V-shaped like that of a river valley rather than the expected U-shaped glacial trough, and that there had been multiple erosional and depositional events. He concluded that more comprehensive seismic reflection surveys of the Finger Lakes "might reveal a barely suspected missing chapter in the history of the Finger Lakes region" (Bloom, 1984, p. 61).
During the summer of 1984 we initiated a seismic reflection investigation of the Finger Lakes with a pilot study of the northern half of Seneca Lake (Stephens, 1986) supported by a "starter" grant from Syracuse University. Preliminary results were very encouraging and revealed a detailed stratigraphy of a thick sediment fill. These preliminary data formed the basis of a National Science Foundation grant for a comprehensive seismic reflection study of all eleven Finger Lakes which was carried out during the summers of 1986 and 1987 (Mullins and Hinchey, 1989). The scientific objectives of this study were three-fold: (1) the provincial question of the origin and evolution of the Finger Lakes; (2) evaluation of the processes of deglaciation along the southern margin of the Laurentide ice sheet; and (3) to use the thick sediment sequence beneath the lakes as a record of continental environmental (paleolimnology/paleoclimatology) change.
In order to fully realize these scientific objectives it was also necessary to extend our lake-based seismic surveys to adjacent on-land areas and to "ground truth" these geophysical data with subsurface geologic samples. Phase II of our investigation was also supported by NSF, and on-land multi-channel seismic reflection profiles were obtained during the summers of 1988 and 1990. In July 1990 we also recovered a 120 m (400') drill core from the dry valley south of Canandaigua Lake, north of the village of Naples, N.Y. These new drill core samples, which are presently being dated and analyzed, complement 26 piston cores (up to 5 m long) collected from Seneca and Cayuga Lakes. Previous subsurface geologic information on Quaternary deposits beneath the Finger Lakes was limited to stratigraphic descriptions of water wells at Ithaca (Tarr, 1904) and short cores recovered from southern Cayuga Lake (Ludlam, 1967) and northern Seneca Lake (Woodrow et al., 1969).
The purpose of this field trip is to provide a new and unique perspective on the Quaternary geology of the Finger Lakes -- the shallow subsurface. Following a brief tour of surficial features, we will embark on a comprehensive overview of the subsurface Quaternary geology of the eastern Finger Lakes (Skaneateles, Owasco, Cayuga, Seneca, Keuka, and Canandaigua) as well as Tully Valley (a "dry Finger Lake") and Montezuma wetlands (a glacial meltwater system). We have planned 12 stops over a 1.5 day period to examine and discuss seismic reflection, gravity, and drill core data at acquisition sites. This non-traditional field trip has no outcrop stops so leave your shovels but bring your imagination!
SURFICIAL GEOLOGY
The surficial geology of the Finger Lakes region has recently been compiled by Muller and Cadwell (1986; Fig. 2). Much of the upland area in the Finger Lakes region is covered by relatively thin till. However, a series of low-relief, chevroned, recessional till moraines also occur on the uplands in the central and southern regions of the lakes. At the north end of the lakes, recessional till moraines become more continuous in an east-west direction (Fig. 2).
North of the Finger Lakes is an extensive field of drumlins and meltwater channels which extend to Lake Ontario. Glaciolacustrine and swamp deposits selectively infill north-south oriented channels whereas an east-west system of channels is largely filled by outwash.
South of the lakes, and confined to the valleys, are well-developed kame moraines, collectively known as the Valley Heads moraine. These moraines are not connected over the uplands and consist principally of water-laid stratified drift. Based on regional correlations, Fullerton (1986) suggests that the Valley Heads moraines were deposited between 12.95 and 14.1 ka. If correct, Valley Heads deposition correlates with a pulse of rapid melting of the Laurentide ice sheet centered at 13.5 ka as recognized in the oxygen isotope record of North Atlantic and Gulf of Mexico deep-sea sediments (Ruddiman, 1987). To the south of the Valley Heads is a well-developed, arcuate system of outwash-filled valleys that likely served as meltwater channels during deglaciation.
There are relatively few radiocarbon dates available from the Finger Lakes region. Many of those reported (summarized by Muller and Cadwell, 1986) are "dead" dates >35 ka suggesting pre-late Wisconsin glacial deposits. Based on these dates Bloom (1986) has suggested that the valleys around Ithaca, and presumably the other Finger Lakes valleys, were excavated "prior to two ice ages ago" (p. 263), well before the late Wisconsin.

SP1: SUBSURFACE GEOLOGY OF THE FINGER LAKES REGION

East margin of Onondaga Trough
Drive south along east margin of Onondaga Trough toward
Cortland.

Note jointed bedrock along off-ramp.

STOP 1. - dead end of Amidon Road - overview of Tully Valley.
This brief stop at the dead end of Amidon Road provides (weather permitting) an outstanding south-directed view of Tully Valley, a "dry" Finger Lake valley. Clearly visible are the steep glacially-eroded bedrock walls of the valley and its relatively flat, sediment-filled floor which produced an overall U-shape profile. To the south is the Valley Heads moraine at Tully which completely fills the valley but does not extend over adjacent upland divides. The primary purpose of this stop is to give you a view of a "Finger Lake" without the water!
During late Wisconsin deglaciation glacial meltwaters ponded in this valley to form a proglacial lake more than 183 m (600') deep (Hand and Muller, 1972). The valley floor has an elevation today of about 152 m (500') and Bare Mountain on the west side of the valley crests at about 457 m (1500') which means that the valley was filled to about 60% capacity by this preglacial lake or about to the level we are standing (351 m, 1150').
As the ice receded to the north up contiguous Onondaga Trough east-west outlets were progressively exposed and the preglacial meltwaters in Tully Valley began to drain into the Mohawk through the Syracuse meltwater channels (Hand and Muller, 1972). Hand (1978) suggests that the outflow of water from Tully Valley occurred in a series of catastrophic flood events which ultimately led to the "dry Finger Lake" valley we see today. Since the draining of Tully Valley some 12 ka ago, post-glacial alluvium has washed into the valley, with the large alluvial fan at the mouth of the Rattlesnake Gulf being the most prominent.
From the vantage point of this overview stop we will drive down into this "dry Finger Lake" to examine subsurface data from the floor of Tully Valley.

STOP 2: THE FLOOR OF TULLY VALLEY
At this stop along Otisco Road you are now standing at the bottom of what once was a Finger Lake as deep as modern day Seneca Lake. How far below the valley floor does bedrock extend? How thick are the sediments beneath the valley floor and what is their geologic nature? To address these simple
questions we have collected preliminary multichannel seismic reflection and gravity data which have been integrated with available drill hole information.
A gravity transect across Webster Road, 2.5 km north of our current location, suggests a total sediment fill of about 134 m (440') with bedrock eroded to within 30 m (100') of sea level. Sediment thickness of at least 117 m (385') has been confirmed there by a drill hole that did not reach bedrock.
A multichannel (6-fold) weight-drop seismic reflection profile shot across Otisco Road reveals a high-amplitude reflector interpreted as bedrock at a maximum of 0.13 sec. of two-way travel time. Calculation of true depth to bedrock is dependent upon the compressional wave velocity of overlying sediment. Based on wide-angle reflection experiments conducted in the Finger Lakes (Mullins and Hinchey, 1989) as well as seismic refraction data from southern Tully Valley (Faltyn, 1957) unconsolidated sediments beneath the Finger Lake valleys have a P-wave velocity in the range of 1.5-2.2 km/sec. Using this range of velocities, maximum subsurface depth to bedrock beneath Otisco Road is on the order of 98-130 m (321-4262'). If one accepts the higher range of velocities (2.0 km/sec.) there is good agreement between the seismic (130 m) and gravity (134 m) data. Both sets of data show a relatively flat bedrock floor beneath the axial portion of Tully Valley. Total bedrock erosion, as measured from the top of adjacent divides to bedrock beneath Tully Valley, has been on the order of 540 m (1770').
The seismic reflection data also suggest that there are three major stratigraphic sequences beneath Tully Valley (Fig. 5). The youngest sequence is characterized by discontinuous, high-frequency reflections and thickens to the west toward the alluvial fan at the mouth of Rattlesnake Gulf where it is as much as 60-80 m (197-262') thick. We interpret this sequence as post-glacial alluvial in-fill. The middle stratigraphic unit is largely "transparent" or reflection-free indicating massive sediment which we interpret as glacio-lacustrine muds, that are as much as 53-70 m (174-230') thick. The top of this middle unit is undulatory which may be a consequence of loading by the overlying alluvial fill. The lowermost stratigraphic sequence beneath Tully Valley is also transparent and is characterized by a discontinuous, high-amplitude reflection at its top. This oldest sequence is relatively thin (19-25 m; 62-82') and is interpreted as a coarse-grained facies.
Our inferences on depth to bedrock and stratigraphic nature of the sediment fill beneath Tully Valley based on geophysical data are supported by available drill hole data. Getchell (1983) reported the gross stratigraphy of nine drill holes located in the southern end of Tully Valley (Fig. 6). Eight of the nine wells penetrated three major stratigraphic units: (1) an upper gravel/clayey gravel unit up to 24m (80') thick; (2) a middle clay sequence 99-134 m (325-440') thick; and, (3) a basal gravel/sandy gravel unit at least 17 m (55') thick. The one well (WL-6) that did not penetrate the basal gravel unit reached shale bedrock at a subsurface depth of 142 m (465') which is at an elevation of 72 m (235') above sea level.
Combined, the subsurface geophysical and drill hole data reveal a deeply scoured trough beneath Tully Valley with bedrock rising to the south, which is consistent with reflection data from the Finger Lakes (Mullins and Hinchey, 1989). The data also indicate a relatively simple infill stratigraphy of washed subglacial (?) sands and gravel at the base overlain by glacio-lacustrine muds which in turn are capped by coarse-grained alluvial gravels and clayey gravels. Although there is no chronostratigraphic control available, the gross lithostratigraphy beneath Tully Valley argues for a single overall infill sequence.
As we leave Otisco Road we will drive south along Tully Valley and then up the Valley Heads moraine which crests at an elevation of about 366 m (1200') which is some 213 m (700') above the floor of the adjacent valley. Gravel pit exposures here reveal washed sands and gravel that are crudely stratified. The Valley Heads deposits argue for an extended ice marginal environment here with large volumes of coarse-grained debris, both local and exotic, pumped through Tully Valley about 13-14 ka. The thickness of sediment beneath the Valley Heads moraine at Tully and the adjacent outwash plain has yet to be resolved, but is certainly an important question. Durham (1958) suggested that the position of the Valley Heads moraine was controlled by a pre-glacial drainage divide based on a well record at Little York which encountered bedrock at a subsurface depth of only 67 m (220'). However, it is not known whether or not this well, at an elevation of 354 m (1160'), was located along the valley axis. Preliminary gravity data obtained along Little York Road by Steve Wanzer, suggest that bedrock may be considerably deeper than that suggested by Durham (1958). This topic will be part of Steve's master's thesis, and hopefully he will have more definitive results by the time of our field trip.
From the Valley Heads moraine at Tully we will head northwest and drive around the north end of Otisco Lake (the easternmost Finger Lake), and then south to our next stop at the south end of Skaneateles Lake. When leaving the Valley Heads we will be able to see its contact with the bedrock wall of Tully Valley and observe that the moraine does not extend up and over the divide. At the northeast corner of Otisco Lake we will also pass by a hanging delta at Amber some 61 m (200') above the modern lake level of Otisco Lake, attesting to higher glacial lake levels.
Otisco Lake today is relatively shallow (maximum depth of 21 m or 69') and eutrophic which greatly limited our ability to collect seismic reflection data due to gas-saturated sediment (Hinchey, 1986). However, "windows" of seismic data indicate at least 83 m of sediment beneath Otisco Lake (Hinchey, 1986). As we drive around the north end of Otisco Lake we will (weather permitting) have an excellent view down the axis of the lake and Otisco Valley. The steep, U-shaped valley wall on the west side of the valley has a slope of about 35°. From Otisco Lake we will drive over the divide to Skaneateles Lake. Bedrock outcrops along the road attest to the relative paucity of glacial erosion and deposition on the uplands compared to the adjacent valleys.

Contact between Valley Heads moraine and bedrock
Note contact between Valley Heads moraine and bedrock of valley wall as you head west.

Amber hanging delta

STOP 3: SOUTH END OF SKANEATELES LAKE
This stop, high above the lake at the Onondaga-Cortland county line, provides a spectacular north-view of one of the most scenic Finger Lakes -- Skaneateles. We are standing at an elevation of 506 m (1660') which is slightly lower than the surrounding top of the plateau at about 579 m (1900'), but well above lake level which is at 263m (863'). As you look to the north, it is quite apparent that the southern end of Skaneateles Lake is much more deeply incised than it is to the north. Also visible are three "points" along the west shore of the lake which represent post-glacial stream deltas that have built into Skaneateles Lake as higher glacial lake levels receded. The overall morphology of the Skaneateles Lake basin is similar to that observed in Tully Valley except that the valley is partially filled with water.
Maximum water depth in Skaneateles Lake is 90 m (295') which occurs at about the north-south midpoint of the lake. Seismic reflection data indicate that bedrock is relatively shallow beneath the north end of the lake but deepens toward the central portion of the lake where maximum erosion down to about sea level has produced a closed rock basin. Bedrock gradually rises toward the south end of the lake producing an overall spoon-shaped longitudinal profile which is typical of the Finger Lakes (Mullins and Hinchey, 1989). Crossing lines indicate broad, U-shaped bedrock profiles at the north end of the lake which become progressively more incised and V-shaped toward the south, which is also typical of the Finger Lakes (Mullins and Hinchey, 1989). Total erosion of bedrock, from the top of the plateau to the maximum extent of bedrock beneath the lake, has been on the order of 580 m (1900').
Maximum sediment thickness beneath Skaneateles Lake is on the order of 170 m (Mullins and Hinchey, 1989) with most of the sediment in the southern half of the lake basin. The most regionally significant feature of the stratigraphy beneath Skaneateles is onlap at the south end of the lake of transparent to highly reflective sediments onto a southward thickening wedge characterized by a hummocky chaotic seismic facies. The younger package of sediment is interpreted as largely northerly-derived glacio-lacustrine deposits. The older chaotic wedge that thickens to the south, projects to on-land outcrops of Valley Heads. If Fullerton's (1986) correlation of Valley Heads at 13-14 ka is correct, this onlap at the south end of Skaneateles Lake implies that the sediment fill beneath the lake is all less than 14 ka, or of late Wisconsin to Holocene age. If older glacial sequences were deposited in Skaneateles valley they have been effectively removed by late Wisconsin glaciation.
Using this onlap relationship (14 ka) and a maximum sediment thickness of 170 m for Skaneateles Lake, we calculate an average rate of sediment accumulation of 12.1 m/1000 years. However, drilling south of Canandaigua Lake (see Stop #10) suggests that the post-glacial section is quite thin relative to the total thickness of sediment-fill which is consistent with our seismic stratigraphic interpretations. If 90% of the sediment fill beneath Skaneateles Lake was deposited over a 2 ka period between 12 and 14 ka, accumulation rates may have been as high as 77 m/1000 years. If correct, these calculations suggest that large volumes of sediment were rapidly deposited in the Skaneateles Lake basin during deglaciation.
A second stratigraphic feature of significance occurs at about the north-south midpoint of Skaneateles Lake. Here, a series of stacked, transparent wedges, each up to 15 m thick, occur. These wedges pinch-out to the south over a horizontal distance of 3-4 km and are interbedded with highly reflective units interpreted as glacio-lacustrine deposits. Longitudinal (N-S) profiles indicate irregular tops with large diffractions. The seismically transparent (reflection-free) nature of these wedges indicates that they consist of massive sediment.
One interpretation is that these wedges represent "inflow events" deposited rapidly, issuing perhaps from a subglacial conduit; in essence a rapid pulse of massive debris. An alternative interpretation is that these wedges represent "till-tongues" similar to those described by King et al. (1991). Unfortunately, no samples from the wedges are available to resolve which (if either) interpretation is correct. However, Ed Hinchey promises to be on hand during the field trip to defend our view that these wedges represent sudden inflow events.
Regardless of the interpretation of these wedges, they suggest the presence of an ice margin that was either grounded or pinned here, where there is a break in the bedrock slope. North of these wedges total sediment thickness thins appreciably, consisting of a thin, highly-reflective (glacio-lacustrine?) unit overlying a thin, chaotic (subglacial?) sequence. These stratigraphic relationships suggest that the ice margin in Skaneateles valley retreated rapidly (by calving?) to the north from its Valley Heads position to the lake's midpoint where it stabilized and issued large volumes of sediment into a proglacial lake.
From our overview of Skaneateles Lake we will drive down into the valley and around the south end of the lake. This will give you a good opportunity to get a "feel" for how deeply (and steeply) incised the south end of Skaneateles Lake is into bedrock of the surrounding Appalachian Plateau. As we cross the valley floor at the south end of Skaneateles Lake keep in mind that bedrock extends about 150 m (~500') beneath the surface. We will then drive up the west wall of Skaneateles Lake, over the divide to Owasco valley and through the village of Moravia to our next stop overlooking Owasco Lake at Ensnare.

Note bedrock outcrops along west valley wall

STOP 4: OWASCO LAKE
This stop near the intersection of Ensnore Road and State Route 38 provides a good north-oriented overview of Owasco Lake. Elevation here is 323 m (1058') which is about half way between lake level at 217 m (711') and the crest of the upland divide at 402 m (1320') just to our west.
Maximum water depth beneath Owasco Lake is 52 m (170'). Like the other Finger Lakes, it has steep walls and a relatively flat floor. Bedrock beneath the lake extends down to an elevation of ~80 m (262') above sea-level. Total erosion, from the top of the divide here at Ensnare to the bottom of bedrock beneath the lake, has been on the order of 320 m (1050').
Total sediment thickness beneath Owasco Lake is on the order of 100 m (328'). The stratigraphy beneath the lake is particularly well-defined by our seismic reflection data. A basal, seismically chaotic sequence which thickens to the south, is overlain by a highly reflective unit. We infer a single coarse-grained, subglacial sequence overlain by proglacial and post-glacial lacustrine deposits.
In the northern third of Owasco Lake, an axial ridge is buried beneath the highly reflective lacustrine sequence. This ridge, with up to 20 m (66') of relief, is sinuous and can be traced over a lateral distance of at least 5 km. We interpret this buried ridge as an esker which implies the past presence of a large subglacial meltwater conduit along the axis of northern Owasco Lake. Such an ice tunnel may have been responsible for the supply of coarse subglacial debris to the basal chaotic sequence beneath Owasco.
Another significant stratigraphic feature of the sediment-fill beneath Owasco Lake is a remarkably transparent (reflection-free) wedge of sediment in the northern half of the basin. The transparent nature of this deposit indicates massive sediment which is interbedded with highly reflective glacio-lacustrine sequences. Underlying units reveal only slight erosion and the top of this transparent unit is smooth. The wedge is up to 15 m (50') thick in the north and pinches out to zero thickness near the north-south midpoint of the lake. Transverse reflection profiles indicate that the wedge onlaps both the east and west bedrock walls of the lake. Although no subsurface samples are available, we interpret this transparent wedge as a subaqueous outwash fan of massive sand or silt, perhaps analogous to the deposits described by Rust and Romanelli (1973) near Ottawa. If this wedge represents in essence an "inflow event," it further argues in favor of the role of subglacial meltwater in the transport of debris to the southern margin of the Laurentide Ice Sheet, similar to the suggestion by Gustavson and Boothroyd (1987).
From our overlook of Owasco Lake we will head west over the divide toward Cayuga Lake. Just prior to entering the village of Poplar Ridge, we will cross one of the "chevron" till moraines mapped by Muller and Cadwell (1986). Radiocarbon analysis of wood overlying the south end of this till moraine yielded a date of 11,410 ±410 years (Muller and Cadwell, 1986).

Cross "chevron moraine" just east of Poplar Ridge.

STOP 5: CAYUGA LAKE
This stop is located at Long Point State Park at about the midpoint (east shore) of Cayuga Lake, just south of the village of Aurora. Cayuga Lake is one of the two largest Finger Lakes extending north-south for a distance of about 60 km with a maximum width of about 5 km. It has a maximum water depth of 132 m (433') located about two-thirds of the length of the lake form its north end. At this point of maximum water depth, the lake bottom is 15 m (51') below sea-level. Off Long Point maximum water depth is 100m (361'). Lake level is at 116 m (382') which is the lowest elevation of all the Finger Lakes.
We have collected more than 150 km of high-resolution uniboom seismic reflection data from Cayuga Lake which have been correlated with 11 piston cores. Bedrock beneath Cayuga Lake has been eroded as much as 358 m (1174') below lake level, which means that bedrock extends as much as 242 m (794') below sea-level!
A north-south schematic illustrates the overall "spoon-shaped" longitudinal profile of Cayuga Lake. Note that the maximum extent of erosion extends down to the Onondaga Limestone and appears to follow its southward dip before rising up through Hamilton Group shales at the south end of the lake. East-west schematics illustrate the broad, shallow form of the bedrock basin at the north end of the lake which becomes more narrow, and deeply incised to the south.
Maximum sediment fill beneath Cayuga Lake is 226 m (741') which occurs in the southern half of the lake basin. We have divided this fill into six seismically-defined depositional sequences: (I) The oldest sequence occurs only in the axial southern half of the lake basin where it is as much as 100 m (328') thick. This unit is characterized by a chaotic seismic facies and an irregular, hummocky upper surface which projects to Valley Heads outcrops south of the lake; (II) Sequence II is present throughout the lake basin and is up to 135 m (443') thick. It is characterized by a distinct north to south change in acoustic facies from chaotic to transparent. A piston core (#10) located at the crest of an esker-like ridge along the top of sequence II (Fig. 18) recovered less than a meter of fine-grained rhythmites with dropstones before bottoming in clast-supported gravels (Fig. 20); (III) This sequence occurs throughout the lake basin as a transparent (reflection-free) unit up to 59 m (194') thick. At the north end of the lake this sequence is diapiric where it intrudes vertically up to 20 m (66'). Although unsampled, the massive, diapiric nature of this sequence suggests rapidly deposited muds; (IV) Sequence IV is a package of high-frequency, parallel reflections in which individual reflectors can be traced for 10's km along the length of the lake basin. It has a rather uniform thickness of about 30 m (98') except at the north end where it is more than 60 m (197') thick. Piston cores from lake floor outcrops of this sequence, recovered fine-grained, centimeter-scale rhythmites with occasional dropstones that were deposited in a preglacial lake; (V) This sequence records a "turning point" in the history of Cayuga Lake. It is thickest (>30 m; 98') at the south end of the lake basin as well as off major "points" such as Myers Point and Sheldrake Point; and, it is not present at the north end of the lake. This sequence heralds the beginning of northward drainage and lateral sediment input to Cayuga Lake that likely accompanied the lowering of preglacial lake levels and erosion of regional, upland gorges as ice withdrew from Cayuga Valley (Fairchild, 1934a); (VI) Sequence VI is a low-amplitude to transparent seismic unit that is variable in thickness (8-19 m; 26-62') throughout the lake basin. Piston cores have recovered fine-grained organic-rich sediments from this sequence, some of which are "banded" and were perhaps deposited by turbidity currents (Ludlam, 1967).
After lunch at Long Point State Park we will head south to Buttermilk Falls State Park near the south end of Cayuga Lake at Ithaca.

STOP 6: SOUTH END OF CAYUGA LAKE
The purpose of this stop at scenic Buttermilk Falls State Park is to examine and discuss land-based, seismic reflection data (acquired from Cayuga valley across from the Park) that have been integrated with results from 13 wells reported by Tarr (1904). The bedrock reflection extends as much as 0.19 sec. of two-way travel time beneath the valley (Fig. 21). Using a range of interval velocities from 1.5 to 2.0 km/sec., maximum depth to bedrock here is 143 m (467') to 190 m (623'). Tarr (1904) reported a salt well along the western margin of Cayuga valley at Ithaca, that encountered bedrock at 131 m which provides a minimum value for depth to bedrock because of its western location. Maximum sediment thickness beneath Ithaca is probably on the order of 166 m (545').
Our land-based seismic reflection profile (which is of lower resolution than our lake data) also reveals four major stratigraphic units beneath the valley: (l) an upper transparent (reflection-free) unit up to 20 m (66') thick (assuming an interval velocity of 1.5 km/sec.); (2) a high-frequency unit about 19 m (62') thick; (3) a sequence, bound by high-amplitude reflections, up to 28 m (92') thick (assuming a velocity of 1.6 km/sec.), and, (4) a basal unit characterized by discontinuous high-amplitude reflections up to 85 m (279') thick (assuming a velocity of 1.7 km/sec.).
The stratigraphy of Tarr's (1904) 13 wells, all located along the western margin of the valley, is schematically illustrated in Figure 22. The upper layer in these wells is predominantly fine-grained massive clay (up to 18 m or 60' thick) with fragments of mollusks, plants, and wood, including logs. Beneath these organic-rich clays is a series of sand and gravel layers 6-21 m (20-70') thick. According to Tarr (1904), the sands are well-washed and the gravels well-rounded. Plant fragments, mollusks and logs also occur in these coarser beds.
In all wells, these coarser deposits are underlain by "a great thickness of clay" devoid of mollusk fragments (Tarr, 1904). There are also minor coarser units within this clay layer including scratched, angular pebbles. The base of this clay layer is irregular, occurring at subsurface depths between 61 m (200') and 85 m (280') in the 13 wells (Tarr, 1904).
Beneath this clay layer is a heterogenous, coarse-grained unit from which artesian water flowed (up to 300,000 gallons/day!) in many of the wells. The total thickness of this basal coarse-grained unit is not known as 11 of the 13 wells bottomed in it. Tarr (1904) describes this coarse basal unit as consisting of variable sediment ranging from washed sand and gravel, to till, to "quicksand."
Tarr (1904) interpreted all these subsurface sequences as late Wisconsinian --"Neither here nor in the other well that reached rock, nor, in fact, in any of the wells, was any older drift encountered. All the materials are such as might have been brought by the last ice advance, or deposited since the ice-sheet melted away." His view of the heterogenous basal coarse layer was that it is morainic; a subsurface extension of the nearby Valley Heads. We concur with Tarr's interpretation and our seismic data suggest that this basal Valley Heads facies extends to bedrock and is more than 100 m (328') thick. We correlate this unit with seismic sequence I beneath Cayuga Lake.
The middle clay layer with scratched angular stones, is interpreted by both Tarr and us as a proglacial lacustrine sequence, perhaps with dropstones which would suggest the presence of icebergs. We correlate this preglacial clay layer with seismic sequence IV beneath the lake. The overlying washed sands and gravels appear to be a fluvial/alluvial or shallow lacustrine facies, that likely records a drop in lake level and drainage reversal as ice receded from the north end of Cayuga Lake (seismic sequence V; Mullins and Hinchey, 1989). The succeeding organic-rich clays at the top of the section are likely a post-glacial lacustrine facies that records the flooding of the south end of Cayuga valley in response to differential isostatic rebound at the north end of the lake (Tarr, 1904). Since this flooding event, only a thin (8-12 m; 26-39') unit of organic-rich muds has accumulated in Cayuga Lake (seismic sequence VI).
Based on these correlations between our lacustrine seismic reflection data, land-based profile, and Tarr's well data, we have been able to "ground-truth" four of the six seismic stratigraphic sequences beneath Cayuga Lake (sequences I, IV, V, and VI). Sequences II and III pinch-out at the south end of Cayuga Lake and do not extend to Tarr's wells.
From Ithaca we will head west across the upland divide to Watkins Glen at the south end of Seneca Lake.

Note kame deposits in quarry across road.

STOP 7: SENECA LAKE
This stop at Warren Clute Memorial (Lakeside) Park near the south end of Seneca Lake is designed to examine both our lake-based and land-based seismic reflection data coupled with available drill hole data. Seneca Lake is the deepest of all the Finger Lakes having a maximum water depth of 186 m (610') (in the southern half of the lake) which is 50 m (165') below sea level.
The overall bedrock morphology beneath Seneca Lake is similar to the other Finger Lakes: spoon-shaped and deepening to the south in longitudinal (N-S) profile; and, in transverse profile (E-W), broad and shallow in the north becoming more deeply incised to the south (Mullins and Hinchey, 1989). Bedrock extends as much as 434 m (1423') below lake level which is 298 m (978') below sea level! This maximum depth of bedrock erosion occurs about one-third of the distance to the north end of the lake from Watkins Glen. It is only slightly deeper than the 257 m (842') depth below sea level predicted by Fairchild (1934b).
Thickness of the sediment-fill beneath Seneca Lake increases from north to south where it reaches a maximum of 270 m (886'). Seismic stratigraphic sequences beneath Seneca Lake are similar to those beneath Cayuga Lake except that there is a thicker transparent unit (sequence III). However, stratigraphic sequences could not be traced completely to the south end of the lake due to a highly irregular lake floor near the south end.
Similar to Owasco Lake an apparent esker is buried in the subsurface beneath Seneca Lake. A ridge, with up to 35 m (115') of relief, occurs at the base of the axial thalweg of northern Seneca Lake. This ridge can be traced from line to line for several kilometers. We interpret this sinuous ridge as an esker which, if correct, further implies the presence of sub-glacial meltwater conduits during early stages of erosion and infill of the Finger Lakes.
A "half-valley" (east side) multichannel profile collected on-land between Watkins Glen and Montour Falls, further documents large scale erosion and thick sediment fill beneath Seneca valley. The bedrock reflection here is quite distinct and extends at least 0.28 seconds of two-way travel time into the subsurface. Again, using a range of P-wave velocities of 1.5 - 2.0 km/sec., bedrock extends as much as 210-280 m (689-918') beneath the valley floor. The higher end of these possibilities is very compatible with a well drilled at the power station at the south end of Seneca Lake (just to the west of our stop at Lakeside Park) which reached bedrock at 281 m (922') (Wendy McPherson, USGS Ithaca, personal communication, 1991). Thus, bedrock at the south end of Seneca Lake extends only about 145m (476') below sea level compared to 298 m (978') below sea level further north, indicating that bedrock is rising to the south.
Our on-land profile reveals four major stratigraphic units above bedrock: (1) an upper transparent sequence about 20 m (66') thick; (2) a unit characterized by high-frequency reflections about 23 m (76') thick; (3) a largely transparent sequence with discontinuous high-amplitude reflections about 81 m (265') thick; and, (4) a basal, seismically chaotic high-amplitude sequence that is about 108 m (354') thick. By analogy with Tarr's (1904) well data from Ithaca as well as our own drill records south of Canandaigua Lake (stop #10) we interpret the basal chaotic sequence as Valley Heads equivalent; the overlying transparent unit as preglacial lake clays; the overlying high-frequency unit as fluvial/alluvial sands and gravels deposited as glacial lake levels dropped; and, the upper transparent unit as post-glacial lake muds deposited when the south end of the valley was flooded by isostatic rebound to the north. Again, it is our interpretation that there is only a single large infill sequence beneath the Finger Lakes.
From Watkins Glen we will head south to Horseheads passing through one of the largest Valley Heads deposits in the Finger Lakes region. Outcrops indicate that the Valley Heads here consist of water-laid, stratified sands and gravel. Near Horseheads the hummocky topography of the Valley Heads will give way to a more smooth outwash plain. From Horseheads we will drive west-northwest through Corning and then on to Bath where we will spend the evening. If it is still light, this drive will take us along a broad continuous valley along the Southern Tier Expressway which likely served as a channel for meltwater flow from the western Finger Lakes into the Susquehanna drainage system.

Note infilled lake valley to left as you head south from Watkins Glen.

Drive will take you along proglacial meltwater channel.

Drive will take you across Valley Heads and down to dry lake floor.

STOP 8: KEUKA LAKE (DAY TWO)
From our overnight stay in Bath, N.Y. we will head north to Hammondsport and then up along the west shore of Keuka Lake to our next stop at the north end of the west branch of Keuka Lake. Bath is located on an outwash plain just south of Valley Heads fill. Soon after leaving Bath you will be able to see hummocky terrain typical of the Valley Heads. We will then drop down to the dry lake floor on which the village of Hammondsport is located.
Keuka Lake is an anomalous Finger Lake in many regards. Most obvious is its two branches that converge to the south at Bluff Point which strongly suggests a southward directed, preglacial drainage system. Second, there are a number of circular to elliptical closed bathymetric depressions on the floor of both branches of Keuka Lake. Third, there is little north-south variation in the depth to bedrock beneath the northwest branch of the lake. And, fourth, Keuka Lake valley is the only non-through valley of the Finger Lakes.
Maximum water depth beneath Keuka Lake is 58 m (190') which occurs in a closed depression near the junction of the Lake's two branches. Bedrock has been eroded as much as 200 m (656') below lake level which is 18 m (59') above sea level. Total sediment thickness is rather uniform beneath the west branch with a maximum of about 160 m (525').
Our axial (longitudinal, N-S) reflection profiles clearly define the east branch of Keuka Lake as a hanging tributary valley to the west branch. Depth to bedrock beneath the east branch is significantly shallower and there is a distinct bedrock high or "sill" where the east branch joins the more deeply scoured west branch.
Axial profiles from the northwest branch of Keuka Lake indicate that the closed bathymetric depressions on the lake floor are remnants of ice-melt-out features. Reflectors beneath these depressions are displaced downwards suggesting post-depositional collapse and chaotic patterns suggest mass wasting. These melt-out features may be analogous to "dead-ice sinks" described by Fleisher (1986) who noticed that they occur preferentially in non-through valleys. He suggested that stagnant ice conditions develop in non-through valleys because of glacial thinning and detachment at headward divides which are not present in through valleys (Fleisher, 1986). The discovery of ice-melt-out features beneath Keuka Lake suggests a component of down-wasting here during deglaciation unlike the other Finger Lakes where ice retreat appears to have been more rapid, perhaps by calving.
Despite this evidence for down-wasting, the seismic stratigraphy beneath Keuka Lake at its southern end is typical of the other Finger Lakes (Fig. 28). A lower chaotic unit (sequence II) is overlain by a very transparent (reflection-free) sequence (III) which in turn is overlain by a highly reflective unit (IV) derived from the north. Sequence V, which is also highly reflective, thickens to the south and contains southerly-derived mass-movement deposits; and, sequence VI is a relatively low-amplitude facies that probably represents contemporary (post-glacial) lacustrine sedimentation.
From our overlook of the northwest branch of Keuka Lake we will drive west through Italy Valley to the village of Naples passing over Valley Heads fill and then to the south end of Canandaigua Lake.

There will be a brief stop to overlook drill site along Parrish Flat Road in valley.

STOP 9: CANANDAIGUA LAKE
This overlook stop provides a spectacular north-view of Canandaigua Lake -- "the chosen place." Bathymetric data reveal a symmetrical "tub-shaped" lake basin with a maximum water depth of 84 m (276') near its north-south midpoint. Bedrock has been eroded as much as 261 m (856') below lake level which is 51 m (168') below sea-level. Total erosion, from the top of the adjacent divides to the top bedrock beneath the lake, has been on the order of 630 m (2068'). An on-land seismic reflection profile across Valley Heads along Eel Pot Road south of Naples, suggests that although bedrock does rise to the south there may be as much as 250 m (820') of Valley Heads fill.
Maximum sediment thickness beneath Canandaigua Lake is 202m (663') which occurs in the southern half of the lake basin. Similar to the other Finger Lakes we recognize six seismically-defined depositional sequences beneath Canandaigua Lake; however, there are some significant differences. Most notable is sequence III --a transparent (reflection-free) facies --which on our axial line occupies a restricted, "scooped-out" depression near the center of the lake. These massive fine-grained (?) sediments are up to 60 m (197') thick, and are overlain by highly reflective sequences which onlap bedrock valley walls.
Another difference is that collectively, sequences IV-VI (highly reflective facies) thicken to the south implying a largely southerly derived source for these deposits. We have noticed a similar relationship in other Finger Lakes only for sequences V and VI which we have related to a lowering of glacial lake levels and a drainage reversal which accompanied complete ice withdrawal from the Finger Lake valleys. In the case of Canandaigua Lake, these data suggest that sediment input from the south end of the lake began sooner than in the other lakes. This may have been due to the fact that there are multiple valleys, which may have been ice-free at the time, that drain into the south end of Canandaigua Lake.
All seismic stratigraphic sequences beneath Canandaigua Lake, with the exception of sequence III, can be traced to the south end of the lake. These sequences do not pinch out here but appear to continue beneath the wetland and dry lake valley to the south. In fact, this is why we decided to locate our drill core to the south Canandaigua Lake. Figure 29 schematically illustrates the projection of seismic stratigraphic sequences beneath Canandaigua Lake to our drill site ~3 km to the south along Parrish Flat Road.

STOP 10: PARRISH FLAT ROAD (DRILL CORE RESULTS)
Parrish Flat Road, which cuts across the dry valley just south of Canandaigua Lake, provides a unique accessible location to correlate (as directly as possible) our lake-based seismic reflection data with drill core samples. Between the drill site and Canandaigua Lake (3 km to the north) are the High Tor wetlands; directly to the east is a large alluvial fan at the mouth of Conklin Gully; and 3 km to the south, Valley Heads deposits begin to crop out near the village of Naples.
A multichannel reflection profile collected along Parrish Flat Road in 1988 indicated that bedrock extends as much as 0.22 seconds of two-way travel time beneath the valley floor. Using an interval velocity of 1.5 km/sec., there is about 165 m (540') of sediment fill above bedrock here. This profile also suggested the presence of four stratigraphic units: (1) a thin, upper transparent unit about 11 m (37') thick, underlain by (2) a thin (~10 m, 33' thick) high frequency sequence that thickens to east toward the alluvial fan; (3) a thick (~94 m, 308') unit with discontinuous high-amplitude reflections; and, (4) a basal unit, about 50 m (165') thick, beneath an irregular, high-amplitude reflector.
Drill core samples were recovered in July 1990 at two sites along Parrish Flat Road: site 1 extended 120 m (400') into the subsurface; and site 2, located 168 m (550') to the west, extended only 12 m (40') into the subsurface. Although composite core recovery was only ~40%, vertical stratigraphic continuity was extended with the use of downhole geophysical (gamma ray and resistivity) logs. Data from a vertical seismic profile shot at the deeper drill hole indicate that P-wave velocities in the upper 100 m (328') of section are all less than 1.6 km/sec. indicating under-compacted materials.
Based on physical stratigraphy, we recognize five distinct stratigraphic sequences in the Canandaigua drill core: (1) an upper unit (12 m, 39' thick) of cyclic sequences of peat and lake clays; (2) a coarsening-upward, washed sand and gravel unit (containing artesian water) that is 14 m (46') thick; (3) a rhythmically bedded silt and sand sequence plus interbeds of massive clay with dropstones that is 60 m (197') thick; (4) a massive clay unit devoid of dropstones that is 23 m (75') thick; and (5) and basal coarse sand and gravel unit that would not support open-hole drilling. Although bedrock was not reached in this drill core, a gas well located in the center of the valley about 3 km to the south, did encounter bedrock at a subsurface depth of 143 m (470'). The age of these deposits has not yet been determined; however, ten peat samples from the upper 12 m (39') have been sent off for radiocarbon analysis and one sample of massive clay from 109 m (358') has been sent off for thermoluminescence dating. Hopefully results will be in hand by the time of the field trip.
Our stratigraphic results front the Canandaigua drill core are strikingly similar to those reported by Tarr (1904) from Ithaca (Cayuga Lake) as well as available drill core data from Tully Valley. Basically, a basal sand and gravel unit is overlain by a thick sequence of fine-grained, organic-poor sediment which in turn is capped by sands and gravels that grade up into organic-rich lake clays. We interpret this stratigraphy as a single infill sequence resulting from the last glaciation. The basal sands and gravels are interpreted as part of the Valley Heads fill (13-14 ka); the thick, organic-lean clay sequence is interpreted as glacio-lacustrine facies; the upper sands and gravels as prograding alluvial valley fill that formed as glacial lake levels dropped; and, the uppermost organic-rich sequence as post-glacial swamp and lacustrine deposits. The cyclic nature of these youngest sequences suggest that lake levels have fluctuated periodically in Canandaigua Lake, and perhaps the other Finger Lakes, during postglacial time. Much more detailed study of these upper cyclic sequences is currently underway as part of Rob Wellner's doctoral dissertation.
From Parrish Flat Road we will drive north along the western margin of Canandaigua Lake to the city of Canandaigua. Toward the north end of the lake you will begin to notice drumlins which will be very common as we drive east along the New York State Thruway to our next stop at Montezuma wetlands north of Cayuga Lake.

Look to right (east) across Montezuma channel.

STOP 11: MONTEZUMA CHANNELS
Montezuma National Wildlife Refuge is one of the largest freshwater wetlands in New York State. The wetlands actually occupy a southward directed, dendritic system of channels that funnel into the north end of Cayuga Lake, which has the lowest lake level (116 m, 382'). That these channels are relict, is evidenced by the fact that modern outlet drainage from Cayuga Lake is to the north.
The Montezuma channel system is unusually broad (up to 5 km) for having a north-south length of only ~25 km. The channels are quite distinct on satellite images of the Finger Lakes and have been mapped largely as swamp deposits by Muller and Cadwell (1986). The channels occur within a very well developed drumlin field but are neither drumlinized nor do they appear to truncate drumlins. The geologic significance of the Montezuma channels has been largely overlooked in previous regional analysis.
In order to better understand the subsurface geology of the Montezuma channels, we have collected about 12 km of on-land, multichannel seismic reflection profiles that have been integrated with available well data. A profile acquired at our stop along Armitage Road reveals a thalweg along the western edge of this eastern branch of the channel system. Bedrock here extends as much as 40 m (131') beneath the surface of the channel or 78 m (257') above sea level. Maximum depth to bedrock beneath the channels, based on well records, is 55 m (180').
Two well records are available from Armitage Road. One well along the eastern margin of the channel penetrated about 30 m (98') of clay on top of shale before bottoming in limestone at 38 m (125'). The second well, located in the channel's thalweg, bottomed in sand without reaching bedrock. Facies analysis of our reflection data suggest that the sands are restricted to the channel's thalweg.
The postglacial pollen record of the Montezuma channels has been investigated via a 12 m core recovered from Crusoe Lake located about 8 km north of Armitage Road (Cox and Lewis, 1965). This core retrieved about 6 m (20') of organic-rich gytta overlying 5 m (16') of clay (Fig. 36). Two radiocarbon dates from 5.0 m (16') and 1.3 m (4') below the lake floor yielded ages of 6,850 ±150 and 3,200 ±100 years, respectively (Cox and Lewis, 1965). Pollen data indicate that a spruce-pine-fir forest dominated the region prior to 6,850 years ago. The first deciduous forests (dominated by hemlock and oak) developed about 6,500 years ago during a time when the climate here may have been warmer and more moist than today (Cox and Lewis, 1965). Between about 3,000 and 2,000 years ago, beech replaced much of the hemlock suggesting a decrease in available moisture. During the past 2,000 years, the pollen record suggests a decrease in temperature and an increase of available moisture (Cox and Lewis, 1965). These authors also suggest that a large lake, of which Crusoe is a remnant, may have persisted in Montezuma Marsh until about 700 A.D.
We also collected a "sledge-hammer" multichannel seismic reflection profile across the drumlin we are standing on. Results suggest that there are about 18 m (58') of unconsolidated sediment beneath the crest of the drumlin. The bedrock reflection is quite distinct and indicates a "high" with about 6 m (18') of relief. Although these results are from only one of thousands of drumlins in the area, they do suggest that the drumlins here may be controlled by relief on an eroded bedrock surface.
What is the origin of the Montezuma channels? The facts are that it is a southward-directed, dendritic system with a maximum length to width ratio of about 5:1; it is a broad, shallow system filled with up to 55 m (180') of sediment; basal coarse-grained sediments are overlain by lake clays; and, the entire system feeds into the north end of Cayuga Lake despite the fact that contemporary outlet drainage is to the north.
One interpretation is that these channels represent subaerial preglacial drainage from a stabilized ice margin north of the channels. However, there is little direct evidence for a stabilized ice-margin here (Muller and Cadwell, 1986) and the dendritic pattern is inconsistent with preglacial streams which are typically braided. Also, a preglacial stream origin would require subaerial conditions followed by a sub-lacustrine environment.
An alternative hypothesis is that the Montezuma channels were carved by subglacial meltwaters that were funneled into the Cayuga Lake valley. A dendritic pattern has been theoretically predicted for subglacial drainage systems which Boulton and Hindmarsh (1987) argue are required to drain excess water during stable subglacial deformation to form drumlins. A subglacial origin for these channels would also allow for direct evolution of environments from subglacial fluvial to glacio-lacustrine and then lacustrine (without an intervening subaerial stage) as the ice sheet retreated. Muller and Cadwell's (1986) surficial geologic map of the Finger Lakes indicates the presence of a similar, but less well-developed, channel at the north end of Seneca Lake which has an esker-like ridge along its northern axial thalweg.
From our stop on Armitage Road along the western branch of the Montezuma channels, we will head east and drive across the eastern branch of the channels where you can get a good "feel" for the width of these channels. We will then drive south to Route 20 and then east to our last stop at Skaneateles which will take us across the southern edge of the Weedsport drumlin field.

STOP 12: NORTH END OF SKANEATELES LAKE
(WRAP-UP DISCUSSION)
This scenic stop at the north end of Skaneateles Lake is designed to provide us with an opportunity to collectively discuss the subsurface geophysical/geologic data we have examined during the past day and a half, and its implications for the origin and evolution of the Finger Lakes. These data document the large scale erosion (up to 298 m; 978' below sea-level) and infill (up to 270m; 886') of the Finger Lake valleys (Mullins and Hinchey, 1989) including a thick Valley Heads sequence. Seismic stratigraphic relationships indicate that the sediment-fill beneath the Finger Lakes post-dates deposition of Valley Heads at 13-14 ka. Thus, timing of the erosion of the Finger Lakes can only loosely be constrained as sometime between the Devonian and 13-14 ka! If pre-late Wisconsin sediments were deposited beneath the Finger Lake valleys they have been removed by subsequent erosion. This implies that the Finger Lakes were at least deepened by late Wisconsin glaciation. Hughes' (1987) model of Laurentide Ice Sheet deglaciation for 14 ka suggests that ice streamed into the Finger Lakes region from the St. Lawrence Lowlands/Ontario Basin as the ice sheet began to collapse. Based on our subsurface data integrated with Muller and Cadwell's (1986) surficial map, we suggest a highly digitate ice margin in the Finger Lakes region with streams of more rapidly flowing ice moving down the Finger Lake valleys. At this time (13-14 ka) large volumes of coarse-grained debris were being pumped subglacially down through the Finger Lake valleys, perhaps by channelized subglacial meltwaters, and ultimately deposited as the Valley Heads moraines. An alternative interpretation for this stage is that an ice dome developed in the Lake Ontario basin and that rapid ice flow through the Finger Lake valleys was in the form of outlet glaciers which may have also been accompanied by channelized subglacial meltwater (Ridky and Bindshadler, 1990).
As the ice margin began to pull-off its Valley Heads position channelized subglacial meltwaters continued to pump large volumes of finer-grained debris into the Finger Lakes valleys as subaqueous outwash (sequence II; Fig. 41). In this regard we agree with Gustavson and Boothroyd (1987) that subglacial streams (rather than "dirt machines") were the primary source of glacio-fluvial and glacio-lacustrine sediment along the southern margin of the Laurentide Ice Sheet.
Once ice margins had retreated to the north end of the Finger Lake valleys, classical proglacial lacustrine sedimentation continued to fill the valleys which is recorded on our reflection profiles as highly-reflective sequence IV. At this time, with the north ends of the valleys dammed by ice, lake levels were considerably higher than today. However, as soon as ice retreated from the north ends of the valleys, lake levels dropped rapidly and a drainage reversal (from south-to north-directed) occurred (Mullins and Hinchey, 1989). This event is recorded on our profiles by sequence V which thickens to the south in all the lake basins. When lake levels dropped, highstand deltas were left hanging on valley walls and numerous gorges were eroded as local base level was dramatically lowered.
After the ice completely withdrew from the Finger Lakes region ~12 ka (Hughes, 1987), differential isostatic rebound in the north resulted in the flooding of the south ends of the Finger Lakes valleys as evidenced by lake clays overlying fluvial/alluvial sands and gravels in drill cores. Subsequent to this flooding event, post-glacial drainage from the south has partially filled the southern ends of the Finger Lake valleys, while accumulating a thin, transparent "blanket" of sediment beneath the lakes.
The most recent chapter in the natural history of the Finger Lakes is recorded in the cyclic sequence of peats and lake clays recovered in the Canandaigua drill core. There are about six of these cycles that overlay fluvial/alluvial sands and gravels. Do these cycles represent climatically driven lake level fluctuations or post-glacial rebound oscillations? Whatever their cause, the fact that there are about six cycle in the post-glacial (<~12 ka) section suggests that they have a periodicity on the order of 2 ka. Our working hypothesis is that they were climatically driven and may record an environmental record of wet and dry cycles for the Finger Lakes region during the past 12 ka. We continue to work on these cyclic sequences and should have more definitive results in the near future.
We hope that you have not only enjoyed this unique field trip but that you have also gained a better appreciation of the "unseen" geology beneath the Finger Lake valleys. If we are to fully comprehend the Quaternary geologic history of regions such as the Finger Lakes it will be necessary to record subsurface as well as surficial geology. We hope that our integrated subsurface geophysical and geological approach has revealed at least a part of the "barely suspected missing chapter in the istory of the Finger Lakes" alluded to by Bloom (1984)

A1: THE FOUNDERS OF AMERICAN GEOLOGY
THE FOUNDERS OF AMERICAN GEOLOGY: A VISIT TO THEIR TOMBS, LABS, AND THEIR FAVORITE EXPOSURES: THE DEVONIAN LIMESTONES OF THE CAPITAL DISTRICT; A STUDY OF THE SEQUENCE STRATIGRAPHY OF THESE LIMESTONES
GERALD M. FRIEDMAN
Brooklyn College and the Graduate School of CUNY
and
Northeastern Science Foundation affiliated with Brooklyn College-CUNY
15 Third Street, Box 746
INTRODUCTION
Located along the Helderberg Escarpment this classic site is on hallowed ground. Amos Eaton (1776-1842), Ebenezer Emmons (1799-1863), James Hall (1811-1898), William W Mather (1804-1859), Sir Charles Lyell (1797-1875), Benjamin Silliman (1779-1864), James D. Dana (1813-1895), Louis Agassiz (1807-1873), and Sir William E. Logan (1798-1875) have trod here before you, and we will view the memorial plaque erected in their honor. The field sites expose a large variety of Lower and Middle Devonian limestones, including coral reefs, stromatoporoid reefs, storm deposits, stromatolite facies, skeletal-and lime mud facies, solution-collapse features, and karst. This diversity of facies will be studied in terms of transgressive-regressive cycles, known as parasequences, which are defined as conformable successions of genetically-related beds bounded by surfaces of erosion, called parasequence surfaces.
HISTORY OF GEOLOGY
Before viewing the sequence stratigraphy of the Devonian we will pay tribute to the founders of American geology on whose concepts modern geology has built. Among those particular attention will be given to Amos Eaton, James Hall, and Ebenezer Emmons, whose debt we have incurred as builders of our science and at whose graves we will pay our respects. This field trip will begin in Troy, N.Y.

A1: The Founders of American Geology

STOP 1. Amos Eaton grave
Visit the grave of the father of American geology Amos Eaton

STOP 2a. James Hall (Lot 93, Section 18)
James Hall. - James Hall was the single most prominent American geologist of the 19th century. The impressive honors he received are worth listing chronologically. Hall obtained the degrees of A.B. and A.M. from Rensselaer in 1831 and 1832. In 1837 he was elected Member of the Imperial Mineralogical Society of St. Petersburg, Russia. In 1840 he was one of the founders of the American Association of Geologists which became the American Association for the Advancement of Science. In 1843 he was elected correspondent of the Academy of Natural Sciences of Philadelphia. In 1848 he was elected Foreign Member of the Geological Society of London, the number of foreign members being limited to 50. In 1848 he was elected fellow of the American Academy of Arts and Sciences. In 1856 he became President of the American Association for the Advancement of Science and in 1863 he was named, by an Act of Congress, to be one of the 50 original members of the National Academy of Science. The King of Italy conferred on him the title of Commander of the Order dei Santi Maurizzio e Lazzaro. Many other honors too numerous to he cited may he added.

STOP 2b. Ebenezer Emmons (Lot 46, Section 6).
Ebenezer Emmons.--alumnus and Junior Professor at Rensselaer, member of the New York State Geological Survey, founder of the North Carolina Geological Survey, and State Geologist of North Carolina; father of the Taconic System. A graduate of Rensselaer in the first class of 1826, Emmons had been inspired by Eaton. Emmons became Junior Professor at Rensselaer, a position he held for ten years, and a member of the New York State Geological Survey in 1836. Later he was State Geologist of North Carolina, spreading Rensselaer's influence in American geology through his textbooks and advocacy of the Taconic system.· Emmons had noted the presence of a group of rocks between the Potsdam Sandstone, the lowest of the sedimentary formations in New York and what was at the time called the Primitive Rocks of Central Vermont. This interval he proposed to call the Taconic System. Emmons acquainted the public with the Adirondack Region and gave the names to principal mountains. Classics which Emmons published include Manual of Mineralogy and Geology (1826), Report on the Second Geological District of New York (1842), Natural History of New York (1848), American Geology Containing a Statement of Principles of the Science With Full Illustrations of the Characteristic American Fossils (1854), Treatise Upon American Geology (1854), The Swamplands of North Carolina (1860), and Textbook of Geology (1860).
The antagonism of Hall helped chase Emmons out of New York. Yet in death their graves, close to each other, make their old controversy seem remote. Hall never acknowledged Emmons' thought-provoking studies on the Adirondack and Taconic Regions. One of the Adirondack peaks has been named Mount Emmons.

STOP 3. James Hall office and lab
The annex to Sunshine School is the historical laboratory of James Hall. For fifty years this building served as his office and laboratory.

STOP 4. Indian Ladder Trail at the John Boyd Thacher State Park.
Before descending to see the exposures read the sections titled Significance, the Stratigraphic Column, and Parasequences. Study especially figure 5. Then descend.
Study the section as you descend to the base of the trail: identify the facies, parasequence surfaces, contacts, and formations. On your return to the top of the section measure the thickness of various stratigraphic and facies units and plot them on a striplog or in a notebook. Compare your data with those shown on figure 5. After completion of the exercise visit the memorial plaque at the top of the trail near Mine Lot Creek parking lot. It is this plaque which explains why the ground on which you tread at this site is sacred.

A2: GEOLOGY AND GEOCHRONOLOGY OF THE SOUTHERN ADIRONDACKS
James McLelland
Department of Geology
Colgate University
Hamilton, NY 13346
PROLOGUE
Geologic investigations of the Adirondack Region began in the nineteenth century with the early surveys of Ebenezer Emmons and later with important contributions from Kemp, Smyth, and Cushing (see Buddington 1939, for complete bibliographic references of these and other early workers). in the first three decades of the twentieth century a large number of significant publications were forthcoming from Miller, Newland, Alling, and Balk, as well as those cited previously, including Buddington. The essence of these contributions is summarized in Buddington's classic Geological Society of America Memoir (1939) entitled "Adirondack Igneous Rocks and their Metamorphism". The central, dominant theme of this work is that the Adirondacks consist of a vast collection of intrusive igneous complexes ranging in composition from anorthosite to granite. This consensus view was strongly supported by a wealth of field and chemical data and may be regarded as comprising Phase I in the history of Adirondack geology. The culmination of Phase I corresponds with the · publication of U.S. Geological Survey Professional Papers 376 and 377 by Buddington and Leonard (1962) and Leonard and Buddington (1964 ), respectively. These works were the result of government sponsored mineral exploration efforts during World War II. Both reports manifest the classical view of the Adirondacks as an igneous-plutonic domain.
Phase II begins with Engel and Engel (1958, 1964) in the lowlands and Walton and deWaard (1963) and deWaard (1964, 1965) in the highlands. Conceptually, this phase is characterized by paradigms of stratigraphy and stratigraphic correlation. Its modus operandi was the recognition and definition of rock sequences, interpreted as stratigraphic, and correlated over great distances. Depending upon the particular investigators, this stratigraphic framework was cross-bred with varying degrees of granitization and metasomatic transformations.
A late stage of Phase II is represented by the investigations of McLelland in the highlands (McLelland and Isachsen, 1986) and Carl, Foose, deLorraine, and others in the lowlands (see Carl et al. 1990 for references). In this stage, metavolcanics played a burgeoning role in the interpretation of Adirondack layered sequences (Whitney and Olmsted 1989). In addition, structural investigations documented the existence of large fold-nappe structures within most of the Adirondacks (deWaard 1964, McLelland 1984). A characteristic of this stage of Phase II was to downgrade the importance of widespread igneous intrusion and to substitute for it metamorphic processes involving recrystallization of stratified volcanics and sediments into layered gneisses modified by local anatectic effects (Carl et al. 1990; Whitney and Olmsted 1989).
Phase III of Adirondack geology began in the mid-1970's when Eric Essene, together with his students John Valley and Steve Bohlen (see Bohlen et al. 1985 and Valley et al. 1990) established a quantitative framework for Adirondack pressure, temperature, and fluid conditions during metamorphism. This approach has been significantly augmented by the oxygen isotope studies of John Valley and Jean Morrison (see Morrison and Valley 1988 for complete reference) and the U-Pb studies of Klaus Mezger (1990), all of which have provided critical data that constrain models of Adirondack evolution. Simultaneously, McLelland and Chiarenzelli (McLelland et al 1988, 1991a,b; McLelland and Chiarenzelli 1990, 1991) have conducted a U-Pb zircon study of the Adirondacks in order to follow up on Silver's (1969) pioneering, landmark investigations. The quantitative results of these research programs have provided unequivocal boundary conditions with which any interpretations of Adirondack geology must be consistent. These results and associated boundary conditions are presented in the text that follows. Significantly, and interestingly, the numbers demonstrate that Phase I interpretations were much closer to the truth than the elaborate stratigraphic models of Phase II. It has become increasingly clear that, in the Adirondacks, intrusive igneous rocks greatly dominate over metavolcanics, or even possible candidates for metavolcanics. Accordingly, it has become evident that layering in orthogneisses is not of primary origin but represents examples of tectonic layering of the sort described by Davidson (1984) in tectonites referred to as straight gneisses. Highly strained rocks of this sort have been described for the Piseco anticline by McLelland (1984) and are common throughout the region.
In conclusion, modem quantitative petrologic and isotopic data strongly indicates that the early, and classic, interpretations of the Adirondacks were, in the main, very nearly correct and herein lies a lesson worth pondering. These results offer additional support for the well documented thesis that granites are plutonic, intrusive rocks and that attempts to form them by circuitous, non-intrusive mechanisms are both outdated and destined to failure. This assessment applies equally well to long discredited examples of granitization and to more modern attempts to account for granites by metamorphosing acidic volcanics. Among the critical observations related to these conclusions are quantitative results from geothermometry, geobarometry, geochronology, and petrochemistry. Combined with a modem understanding of tectonic layering, these considerations can greatly constrain the interpretation of complex geologic terranes, as described below for the southern Adirondack region.

A6: STRUCTURE AND ROCK FABRIC WITHIN THE CENTRAL AND SOUTHERN ADIRONDACKS

STOP 1. Peck Lake Fm.
This exposure along Rt. 29A just north of Peck Lake is the type locality of the sillimanite-garnet-biotite-quartz-feldspar gneisses (kinzigites) of the Peck Lake Fm. in addition, there are exposed excellent minor folds of several generations. Note that the F1 folds rotate an earlier foliation. The white quartzo-feldspathic layers in the kinzigites consist of quartz, two feldspars,and garnet and are believed to be anatectic and have been folded by F1 indicating pre-F1 metamorphic events. Spinel has been found enclosed in garnets at this outcrop. The similarity of the Peck Lake Fm. to the Major Paragneiss of the Lowlands suggests that the Adirondacks were contiguous at the time of deposition of the rocks.

STOP 2. Irving Pond Fm.
0.5 mile north of Nick Stoner's Inn, Canada Lake. Very near hinge line of F1Canada Lake isocline.
The outer portion of the Irving Pond Fm. is exposed in low cuts along the east side of Rt. 29A just prior to the crest in the road heading north.
At the southern end of the cut typical, massive quartzites of the Irving Pond are seen. Proceeding north the quartzites become "dirtier" until they develop sillimanite-garnet-biotite-feldspar (kinzigites) layers along with quartzite.
At the northern end of the cut, and approximately on the Irving Pond/Canada Lake Fm. contact there occurs an excellent set of F1minor folds. Polished slabs and thin sections demonstrate that these fold an earlier foliation defined by biotite flakes and flattened quartz grains.
At the southern end of the outcrop dark, fine grained metadiabase sheets crosscut the quartzite. Near the telephone pole erosional remnants of diabase appear to truncate approximately horizontal foliation in the quartzite suggesting that the diabase was emplaced after an early metamorphism. At the north end of the cut a diabase sheet of variable thickness is folded in the F1fold. The folding is interpreted as Ottawan, the diabase as AMCG in origin, and the early foliation as Elzevirian. This is consistent with the presence of quartzite xenoliths in the ca. 1300 Ma tonalites.
The Irving Pond Fm. is the uppermost unit in the lithotectonic sequence of the southern Adirondacks. Its present thickness is close to 1000 meters, and it is exposed across strike for approximately 4000 meters. Throughout this section massive quartzites dominate.

STOP 3. Canada Lake Charnockite
(>1233 Ma, table 1, sample AM-87-13. Now fixed at 1251±43)
Large roadcuts expose the type section of the Canada Lake charnockite. Lithologically the charnockite consists of 20-30% quartz, 40-50% mesoperthite, 20-30% oligoclase, and 5-10% mafics. The occurrence of orthopyroxene is sporadic. These exposures exhibit the olive-drab coloration that is typical of charnockites. Note the strong foliation in the rock. Farther north along the highway there are exposed pink leucogranitic variants of this unit. The chemical composition of these is given in table 3 (ab-6). The whole rock chemistry of the charnockitic phase is similar to AM-86-17 in table 3. The lateral continuity of the Canada Lake is striking (fig. 2) but the presence of xenoliths reveals an intrusive origin.

STOP 4. Royal Mt. Tonalite
(> 1301 Ma, table 1, sample AM-86-12, now fixed at 1307±2 Ma).
Steep roadcuts, exposed across from the Canada Lake Store, expose typical examples of the early tonalitic rocks that occur within the southern and eastern Adirondacks and that manifest the presence, throughout the region, of collisional magmatic arcs of calcalkaline chemistry that existed along the eastern margin of Laurentia from ca. 1400-1200 Ma. Amalgamation of these arcs culminated in the Elzevirian Orogeny at ca. 1250-1220 Ma.
he whole rock chemistry of the tonalitic rocks is given in table 3 and important chemical trends are portrayed in figures 8, 9, and 10. Figure 7 shows the εNd characteristics of these rocks and emphasizes their petrologically juvenile character, i.e., they are not derived from any crustal rocks with long-term crustal residence but are essentially mantle derived (including derivation from melting of basaltic rock derived from the mantle at ca. 1300-1400 Ma). The εNd characteristics are compared with those from lowland tonalites and granitoids of similar age, and the similarity suggests that they are essentially the same, strongly suggesting contiguity across the entire Adirondacks at that time ( ˜1300 Ma).
A disrupted layer of amphibolitic material runs down the outcrop to road level at the east end of the outcrop. This, and other mafic sheets in the outcrop, are interpreted as dikes and sheets coeval with the tonalite. In the eastern Adirondacks it has been possible to document mutually crosscutting relationships between these rock types. Also documented there are xenoliths of kinzigitic rock in the tonalites. Within the southern Adirondacks xenoliths of quartzite similar to the Irving Pond Fm. have been found in the tonalite.

STOP 5. Rooster Hill megacrystic gneiss at the north end of Stoner Lake
This distinctive unit belongs to the AMCG suite and is widespread in the southern Adirondacks. Here the unit consists of a monotonous series of unlayered to poorly layered gneisses characterized by large (1-4") megacrysts of perthite and microcline perthite. For the most part these megacrysts have been flattened in the plane of foliation, however, a few megacrysts are situated at high angles to the foliation and show tails. The groundmass consists of quartz, oligoclase, biotite, hornblende, garnet, and occasional orthopyroxene. An igneous rock analogue would be monzonite to quartz-monzonite (see table 3 for chemical composition) and the presence of orthopyroxene makes the rock mangeritic to charnockitic.
The contacts of the Rooster Hill megacrystic gneiss are everywhere conformable, but the presence of xenoliths of kinzigite indicate its intrusive nature. Rocks such as the Rooster Hill are interpreted as derived from melting of ca. 1300 Ma tonalitic and lower crustal granitoid rocks with heat derived from large AMCG gabbroic intrusions that would ultimately differentiate to anorthosite. This suggestion is consistent with the εNd trends of AMCG and tonalitic rocks in figure 7a and with the REE distributions shown in figure 19, where it appears that melting of tonalite so as to leave a plagioclase-rich residue can give the AMCG REE-trends.

Low roadcut in kinzigites

Long roadcuts of pink quartzofeldspathic gneisses and metasediments
Long roadcuts of pink quartzofeldspathic gneisses and metasediments of intruded metagabbro and anorthosite metagabbro. The igneous rocks are believed to belong to the AMCG suite.

Roadcut of anorthositic metagabbro and metanorite of AMCG suite.

Excellent examples of anorthositic gabbros
Roadcut on west side of highway shows excellent examples of anorthositic gabbros intrusive into layered pink and light green quartzofeldspathic gneisses.

Roadcuts of quartzites and other metasediments of the Sacandaga Fm.
Mezger (1990) obtained a U-Pb garnet age of ca. 1154 from these rocks.

Red-stained AMCG quartzofeldspathic gneisses that have been faulted along NNE fractures.

STOP 6. Core rocks of the Piseco anticline.
This stop lies along the hinge line of the F2 Piseco anticline near its domical culmination at Piseco Lake. The rocks here are typical of the granitic facies of quartzofeldspathic gneisses such as occur in the Piseco anticline and in other large anticlinal structures, for example Snowy Mt. dome, Oregon dome.
The pink "granitic" gneisses of the Piseco anticline do not exhibit marked lithologic variation. Locally grain size is variable and in places megacrysts seem to have been largely grain size reduced and only a few small remnants of cores are seen. The open folds at this locality are minor folds of the F2 event. Their axes trend N70W and plunge 10-15° SE parallel to the axis of the Piseco anticline.
The most striking aspect of the gneisses in the Piseco anticline is their well-developed lineation. This is expressed by rod, or pencil-like, structures which are clearly the result of ductile extension of quartz and feldspar grains in a granitic protolith. The high temperature, grain size reduction that has occurred results in a mylonite. Where recognizable, early F1isoclinal fold axes parallel the lineation.
These rocks are similar in age and chemistry to other AMCG granites and are considered to be part of that suite.
Smooth outcrops of Piseco Core rocks showing exceptionally strong mylonitic ribbon lineations.

STOP 7. 3.3 miles east of Speculator, New York.
Typical Adirondack marble is exposed in roadcuts on both sides of the highway. These exposures show examples of the extreme ductility of the carbonate-rich units. The south wall of the roadcut is particularly striking, for here relatively brittle layers of garnetiferous amphibolite have been intensely boudinaged and broken. The marbles, on the other hand, have yielded plastically and flowed extensively during the deformation. As a result, the marble-amphibolite and marble-charnockite relationships are similar to those that would be expected between magma and country rock. Numerous rotated, angular blocks of amphibolite and charnockite are scattered throughout the marble in the fashion of xenoliths in igneous intrusions. At the eastern end of the outcrop tight isoclinal folds of amphibolite and metapelitic gneisses have been broken apart and rotated. The isolated fold noses that remain "floating" in the marble have been aptly termed "tectonic fish". The early, isoclinal folds rotate on earlier foliation. The garnetiferous amphibolites have typical igneous compositions and are interpreted as flows or sills.
Near the west end of the outcrop a boudin of charnockite is well exposed. McLelland and others (1987) have presented evidence that boudin represents a local example of charnockitization by carbonic metamorphism. However, granites of similar composition outside the marble do not develop orthopyroxene, demonstrating the local nature of the process and the limited permeation of the fluid phase.
Exposed at several places in the roadcut are crosscutting veins of tourmaline and quartz displaying a symplectic type of intergrowth. Other veins include hornblende-and sphene-bearing pegmatites.
Almost certainly these marbles are of inorganic origin. No calcium carbonate secreting organisms appear to have existed during the time in which these carbonates were deposited(> 1200 Ma ago). Presumably the graphite represents remains of stromatolite-like binding algae that operated in shallow water, intertidal zones. This is consistent with the presence of evaporitic minerals, such as gypsum, in Lowland marbles.
At the eastern end of the outcrop coarse diopside and tremolite are developed in almost monomineralic layers. Valley et al. (1983) showed that the breakdown of almost Mg-pure tremolite to enstatite, diopside, and quartz in these rocks requires low water activity at the regional P,T conditions. Similarly, the local presence of wollastonite requires lowering of CO2 activity, presumably by H2O. These contrasts demonstrate the highly variable composition of the fluid phase and are consistent with a channelized fluid phase within a largely fluid-absent region.

Extensive roadcuts in lower part of Blue Mt. Lake Fm.
Quartzites, kinzigites, and leucogneisses dominate. Minor marble and calcsilicate rock is present.

Large roadcuts in well-layered, pink quartzofeldspathic gneisses
Large roadcuts in well-layered, pink quartzofeldspathic gneisses with subordinate amphibolite and calcsilicate rock. The layering here is believed to be tectonic in origin, and the granitic layers represent an intensely deformed granite. The calcsilicate layers may be deformed xenoliths.

STOP 8. One half mile south of jct of old Rt. 30 and with new Rt. 30.
Anorthositic rocks on the SW margin of the Oregon Dome.
On the west side of the highway a small roadcut exposes typical Adirondack anorthosite and related phases.
The glacially smoothed upper surface of the roadcut reveals the presence of three major igneous phases: 1) a dark, pyroxene-rich dike that crosscuts the anorthosite, contains anorthosite xenoliths, and contains a large irregular, disrupted mass of sulfidic calcsilicate; 2) a coarse grained, Marcy-type anorthosite facies with andesine crystals 6-8" across; and 3) a fine grained anorthositic phase. Some of the coarse grained facies has been crushed and these portions bear some resemblance to the finer grained phase (note, for example, those places where fractures cross large andesine grains and produce finer grained material). However, close inspection of the finer grained material reveals the presence of ophitic texture with pyroxenes of approximately the same size as the plagioclase, and this texture and association are much better explained as igneous in origin. Therefore, the texture of the fine grained phase is interpreted as igneous in origin and may be due to chilling near the contact of the Oregon Dome massif. By contrast, large (3-4 em) rafts of coarse grained, ophitic gabbroic anorthosite seem to be "flat" in the fine grained phase. Analyses of typical anorthositic rocks are shown in Table 5.
The pyroxene-rich ferrogabbro dike shows "soft" contacts with the anorthosite and is interpreted as essentially coeval. Zircons from it give a minimum age of 1087 Ma and, by comparison with other Adirondack anorthosites, its emplacement age is set at ca. 1135 Ma. The composition of the ferrogabbro is shown in table 5 where it is seen to be rich in TiO2 and P2O5. Similar rocks occur together with other Adirondack anorthosites and are interpreted as late, Fe-enriched differentiates of a Fenner-type fractionation trend (see fig. 14). It is suggested that further differentiation within these rocks can result in liquid immiscibility and the production of magnetite-ilmenite liquids.
The upper, weathered surface of the outcrop affords the best vantage point for studying the textures and mineralogy of the anorthositic rocks. In several places there can be seen excellent examples of garnet coronas of the type that are common throughout Adirondack anorthosites. These coronas are characterized by garnet rims developed around iron-titanium oxides and pyroxenes. Recently McLelland and Whitney (1977) have succeeded in describing the development of these coronas according to the following generalized reaction:
Orthopyroxene + Plagioclase + Fe-bearing oxide + Quartz = Garnet+ Clinopyroxene.
This reaction is similar to one proposed by de Waard (1965) but includes Fe-oxide and quartz as necessary reactant phases. The products are typomorphic of the garnet-clinopyroxene subfacies of the granulite facies (de Waard 1965). The application of various geothermometers to the phases present suggests that the P,T conditions of metamorphism were approximately 8 kb and 700±50°C respectively.

Minor marble, amphibolite, and calcsilicate rock.
Predominantly very light colored sillimanite-garnet-quartz-feldspar leucogneisses interpreted as minimum-melt granitic due to anatexis of kinzigite near Oregon dome anorthosite. Enclaves of spinel-and sillimanite-bearing metapelite are present.

Roadcuts in garnetiferous metasediment.
To the west of the intersection are roadcuts in garnetiferous metasediment. A large NNE normal fault passes through here and fault breccias may be found in the roadcut and the woods beyond.

Entering granitic-charnockitic gneiss
Entering granitic-charnockitic gneiss on northern limb of the Glens Falls syncline. Note that dips of foliation are to the south.

Downdropped block of lower Paleozoic sediments.
Entering town of Wells which is situated on a downdropped block of lower Paleozoic sediments. The minimum displacement along the NNE border faults has been determined to be at least 1,000 meters.

Silver Bells ski area to the east.
The slopes of the ski hill are underlain by coarse anorthositic gabbro that continues to the west and forms the large sheet just south of Speculator.

Entrance to Sacandaga public campsite.
On the north side of NY Rt. 30 are quartzo-feldspathic gneisses and calcsilicates. An F1recumbent fold trends sub-parallel to the outcrop and along its hinge line dips become vertical.

Gabbro and anorthositic gabbro.

Stop 9. Pumpkin Hollow.
Large roadcuts on the east side of Rt. 30 expose excellent samples of the Sacandaga Fm. At the northern end of the outcrop typical two pyroxene-plagioclase granulites can be seen. The central part of the outcrop contains good light-colored garnet-microcline-quartz gneisses (leucogneisses). Although the weathered surfaces of these rocks are often dark due to staining, fresh samples display the typical white vs. grey color of the Sacandaga Fm. The characteristic and excellent layering of the Sacandaga Fm. is clearly developed. Note the strong flattening parallel to layering. Towards the southern end of the outcrop calc-silicates and marbles make their entrance into the section. At one fresh surface a thin layer of diopsidic marble is exposed.
At the far southern end of the roadcut there exists an exposure of the contact between the quartzo-feldspathic gneisses of the Piseco anticline and the overlying Sacandaga Fm. The hills to the south are composed of homogenous quartzo-feldspathic gneisses coring the Piseco anticline (note how ruggedly this massive unit weathers). The Sacandaga Fm. here has a northerly dip off the northern flank of the Piseco anticline and begins its descent into the southern limb of the Glens Falls syncline.
The pronounced flaggy layering in the Sacandaga Fm. is not of primary sedimentary or volcanic origin. Instead it is tectonic layering within a "straight" gneiss. Hand specimen and microscopic inspection of the light layers, particularly, reveals the existence of extreme grain size reduction and ductile flow. Long quartz rods consist of rectangular compartments of recovered quart and annealed feldspar grains occur throughout. The rock is clearly a mylonite with its mylonitic fabric parallel to compositional layering.
The chemistry of the light colored layers in the Sacandaga Fm. indicates that they are minimum melt granites. As one proceeds away from the core of the Piseco anticline, these granitic layers can be traced into less deformed sheets and veins of coarse granite and pegmatite. In the most illustrative cases the granitic material forms anatomizing sheets that get grain size reduced and drawn into parallelism as high strain zones are approached. The Sacandaga Fm. is interpreted as an end result of this process and represents a mylonitized migmatite envelope developed in metapelites where they were intruded by AMCG granites at ca. 1150 Ma and then intensely strained during the Ottawan Orogeny at ca. 1050 Ma. This interpretation is consistent with field relationships, the presence of spinel and sillimanite restites in the leucosomes, and with the fact that similar metapelitic rocks are crosscut by ca. 1300 Ma tonalites. The latter observation makes the Sacandaga Fm. protoliths older than the ca. 1150 Ma granitic rocks in the Piseco anticline and makes an intrusive relationship obligatory despite the conformable contact at the south end of the roadcut.

Core of the Piseco anticline.
All exposures are within the basal quartzo-feldspathic gneisses at the core of the Piseco anticline.

Re-enter the Sacandaga Fm. Dips are now southerly.

Southerly dipping quartzo-feldspathic gneisses
In long roadcuts of southerly dipping pink, quartzo-feldspathic gneisses with tectonic layering. The coarse grain size of the gneissic precursors can be seen in many layers.

A3: PATTERNS OF PHYLETIC EVOLUTION IN THE TRENTON GROUP
ROBERT TITUS
Department of Geology
Hartwick College
Oneonta, New York 13820
INTRODUCTION
The Trenton Group has been the subject of 60 years of continuous biostratigraphic study since Marshall Kay began his work during the early 1930's. The unit displays a remarkably rich and diverse fossil assemblage which, along with its facies patterns, is now quite well known. The Trenton Group is thus very well suited for research directed at understanding patterns of evolution. This ongoing research program is currently focused on recognizing patterns of intraspecific clinal variation and relating these to facies change in order to develop comprehensive records of evolutionary events. Two case studies are now complete. Evolution within the crinoid genus Ectenocrinus has been described (Titus, 1989). A similar paper on the brachiopod genus Sowerbyella is pending (Titus, in preparation). This field guide will serve as an appendix to both of those papers.

A3: Patterns of Phyletic Evolution in the Trenton Group

STOP 1. The Napanee Limestone
The strata here are primarily thick bedded, black micrites which have been interpreted as a nearshore lagoon facies. Sowerbyella is abundant. This form has traditionally been identified as S. punctostriatus. It is herein reinterpreted as the quiet water, mud bottom clinal variant of S. curdsvillensis. Examine the brachial interiors and see the preponderance of simple forms. Medial septa can be seen but muscle scars, vascular markings and sub-peripheral thickenings are uncommon. The pedicle exterior usually displays alate corners and a well developed medial fold giving this valve a peaked appearance.

STOP 2. The Kings Falls Limestone
The lower strata here are primarily thick bedded, coarse grained biosparites which have been interpreted as a barrier shoal facies, lying offshore of the Napanee lagoon facies. Toward the top of the outcrop the facies grades into a shallow shelf facies. While, at this location, these strata overlie those of the lagoon facies, presumably there were contemporaneous lagoonal facies elsewhere. Sowerbyella is very abundant at the base of this outcrop. This form is larger than the one in the lagoon facies. Most of the brachial interiors are quite ornate. They have medial septa, muscle scars, vascular markings and sub-peripheral thickenings. The pedicle exteriors have blunt corners. They don't often show medial folds. Instead they tend to be flattened or broadly rounded. This form has traditionally been identified as S. curdsvillensis. It is herein reinterpreted as the agitated facies clinal variant of a more broadly defined S. curdsvillensis. By collecting enough brachial interiors at these two outcrops, an assemblage of forms can soon be put together that show the gradation between the lagoon and the shoal clinal variants.

STOP 3. The Sugar River Limestone
There is a prominent re-entrant below the bridge which can be traced downstream. This is a thick bentonite which lies 10 m (32 ft.) below the top of the Sugar River Limestone. Sowerbyella is rare in the deeper water facies above this bentonite, but common in the shallower facies below. Examine the strata several meters below the bentonite. The common form of Sowerbyella here is the mud bottom, quiet water form. Although some of the brachial interiors are ornate, most are plain. Medial septa are seen but other features are uncommon. The form is relatively small. Pedicle exteriors often display a medial fold and alate corners. This location records a return of both quiet water conditions and the quiet water clinal variant of Sowerbyella. Upstream into the highest strata of the Sugar River Limestone, and for a considerable distance into the overlying Denley Limestone, Sowerbyella is quite scarce. This zone is part of the first Sowerbyella bottleneck.
The next location is designed to display an example of model 3 clinal variation. The middle Trentonian crinoid genus Ectenocrinus occupied a shallow to deep shelf range of environments. Because of bank margin steepening, the deep end of its range was deteriorating while shallowing facies offered opportunities for Ectenocrinus in shallow water environments. The result is a kind of clinal orthoselection.

STOP 4. The Sugar River And Denley Limestones
At the base of the falls a re-entrant ca be seen. This is the same bentonite which was observed at Mill Creek in Lowville. Here it is also 10 m (32 ft.) below the top of the Sugar River Limestone. The bentonite is a marker bed which is found at the same level at all outcrops from Mill Creek, Lowville to Sugar River. It has not been found at Deer River or at any of the Mohawk Valley locations.
Examine the several meters of strata above the falls. Ectenocrinus is represented by moderately abundant columnals. Look carefully as these columnals are small. It is very helpful to bring a water container and pour water on columnal rich beds. This brings out the contrast between the columnals and their micritic groundmass. Ectenocrinus is easily recognized by its trimeric morphology. Each columnal is composed of three elements fused together along faintly visible sutures. The most common form is a triangle with very rounded corners. At first glance the lumina appear to be triangular, and many are. But you will soon notice that many also have very poorly developed 4th and 5th points, along with three long and attenuated points (Fig. 5i-5m). All specimens at this level belong to the species E. triangulus.
Climb to the level above where the trail intersects the stream. Here and for quite a distance upstream Ectenocrinus columnals are different. Now they are generally rounder and the lumina are composed of more equal sized points (Fig. 5s-Sv). Forms at this level are intermediate between E. triangulus and its descendent E. simplex. Continue upstream and observe the increasingly abundant columnals of Ectenocrinus. The transition from E. triangulus to E. simplex is not just one of morphology. E. triangulus was a relatively stenotopic form which never became especially abundant. The descendent, E. simplex, was altogether different. It was eurytopic and one of the dominant forms in the upper Trenton Group, a "weed" crinoid. This evolutionary event records an ecological transition from the K-strategist ancestor to the r-strategist descendent.
The last locations are designed to illustrate model 2 examples of clinal variation. Sowerbyella passed through two facies bottlenecks, one at the base of the middle Trentonian and one at the base of the upper Trentonian (Fig. 3, arrows). The best place to study the Sowerbyella of the middle Trentonian is at Mill Creek, Lowville, but as that outcrop extends for such a long distance along the stream, it is not practical to include this location in a one-day trip. We will examine middle Trentonian forms at other locations, stops 5 and 6.

STOP 5. The Denley Limestone
At this level the transition to Ectenocrinus simplex is complete. Virtually all Ectenocrinus columnals, upstream from the bridge, are round with pentagonal lumina (Fig. 5z-5aa). Sowerbyella is not common here but the specimens which can be seen are of interest. They are nearly identical to the quiet water lower Trentonian forms. They have plain brachial interiors; the pedicle exteriors have medial ridges and alate corners.

STOP 6. The Denley Limestone
The exposure here is of the middle Denley Limestone. The facies are the shallow shelf possibly grading into the barrier (Fig. 3). Several types of Sowerbyella can be recognized from the brachial interior structure. Many have plain interiors. While medial septa are present, none of the other internal structures are seen. The other type has an ornate brachial interior. Medial septa, muscle scars and sub-peripheral thickenings can be seen, some or all, on the same interiors. The plain forms found at this outcrop are typical of the Denley's deeper shelf facies while the ornate forms, although never very common, are typical of the more shallow and agitated facies. Significantly, vascular markings are quite unusual and these ornate forms are never as large as their lower Trentonian equivalents. Pedicle exteriors are more inflated and rounded. This is Sowerbyella n. sp.

STOP 7. The Upper Denley Limestone
The best location to see Sowerbyella next is at the outcrop of the lower Steuben Limestone on Rt. 177 just west of Rodman, but this is too far away for our trip today. The Rodman exposure shows Sowerbyella within the second bottleneck (Fig. 3, arrow). Specimens there are small and have a well developed, broadly rounded medial fold. At Talcottsville we can see Sowerbyella in the barrier facies just below this second bottleneck. The form is abundant up to the base of the falls and then quite a bit less common in the bottlenecking beds above. The specimens here are, for the most part, characterized by well inflated pedicle exteriors which sometimes display the broadly rounded medial folds typical of the second bottleneck. Brachial interiors are generally plain. Notice the magnificent display of large symmetrical ripple marks (pararipples) here.

STOP 8. The Hillier Limestone
The Hillier Limestone is exposed all along Wells Creek and Big Brook. The environment of deposition was quiet deep shelf and the deposits are mostly biomicrites. This location is above the second facies bottleneck and a third species of Sowerbyella is exposed. The form, S. subovalis, is characterized by a greatly inflated pedicle exterior with a broadly rounded medial fold. Brachial interiors are mostly plain. Medial septa are seen, and sometimes poorly developed sub-peripheral folds can be found, but muscle scars and vascular markings are absent. These traits are apparently inherited from the clinal variants of S.n. sp. which made it through the second bottleneck. The ones on Big Brook are larger than their middle Trentonian ancestors.

A4: PAC STRATIGRAPHY OF THE HELDERBERG GROUP
PAC STRATIGRAPHY OF THE HELDERBERG GROUP: CYCLE DEFINITION, ALLOGENIC SURFACES, HIERARCHY, CORRELATION AND RELATIONSHIP TO "VAIL" SEQUENCES
E.J. ANDERSON AND P.W. GOODWIN
GEOLOGY DEPARTMENT, TEMPLE UNIVERSITY
PHILADELPHIA, PA 19122
ABSTRACT
Sections at Thacher Park and Schoharie demonstrate our method of defining 6th order rock cycles (PACs) on the basis of facies patterns (as opposed to key lithologies). Cycle boundaries are placed at surfaces where deeper facies abruptly overlie shallower facies in a disjunct relationship. Several genetically distinct kinds of stratigraphic surfaces are well illustrated in these sections, including PAC boundaries, intracycle surfaces ( sea-level fall surfaces) and cryptic unconformities. Criteria for correlation of cycles (PACs) in the Manlius Formation between the two localities include: degree of facies change at PAC boundaries, tracing distinctive litho and bio facies, matching patterns of facies change through several cycles and matching the hierarchic pattern of cycles. In the Manlius Formation alone 14 or more 6th order cycles, as many as six 5th order cycles, parts of two 4th order cycles and an incomplete 3rd order B cycle are represented. The Manlius.Coeymans formational boundary is a cryptic unconformity that represents a 3rd order B "Vail" sequence boundary. The stratigraphic section at Schoharie from this unconformity to the Oriskany unconformity represents a nearly complete 3rd order B sequence and is tentatively divisible into a hierarchy of three 4th order and numerous 5th and 6th order sequences (rock cycles) . In terms of systems tract analysis the section equivalent to the Coeymans and Kalkberg Formations is identified as shelf margin wedge, the lowermost New Scotland (equivalent) rocks are the transgressive systems tract and the remaining New Scotland and Becraft Formations are the highstand systems tract.
Our purpose on this trip is to introduce participants to field application of allogenic, hierarchic cyclic stratigraphy. Participants are asked to begin with the assumption that the internal structure of the stratigraphic record is a product of a hierarchic set of cyclic processes that caused sea-level fluctuations. If this assumption is correct then the significant components of the stratigraphic record may include rock cycles arranged in a hierarchy, a variety of allogenic surfaces both within cycles and at their boundaries and a systematically arranged set of unconformities. We will attempt to provide a basis for discriminating the various types of cycles and surfaces and in doing so make the case that field analysis is dictated by assumed models. We would like to emphasize that our focus on these stratigraphic components of the rock record is different from the traditional focus on grains, laminations, beds and formations. New stratigraphic models therefore require new field methods and a different approach to recording data (see Anderson and Goodwin, 1990; Goodwin and Anderson, 1988 and Goodwin et al., 1986).
In this field exercise we will be applying a hierarchic stratigraphic model consisting of 5 orders of allogenic cycles:
3RD ORDER A (Supersequence) 8-10 million years
3RD ORDER B (Sequence) 1-2 million years
4TH ORDER (eccentricity) 400 thousand years
5TH ORDER (eccentricity) 100 thousand years
6TH ORDER (precession) 20 thousand years
Choice of this particular hierarchy is justified by the large number of sequence stratigraphic studies documenting the existence of the two "third order" natural groupings and by the growing number of studies confirming the presence of Milankovitch cyclicity throughout the stratigraphic record.

A4: PAC Stratigraphy of the Helderberg Group

STOP 1. Thacher Park
This locality has been selected to illustrate the fundamental, 6th order, meter-scale, allogenic cycle. This cycle is thought to be produced by the 20 ky, precessional Milankovitch signal. At this locality these rock cycles range from less than one meter to almost three meters in thickness. Some of them are totally subtidal and others (e.g. PACs 9 and 12) have tidal flat facies at their tops. Cycle boundaries are placed at surfaces where deeper water facies abruptly overlie shallower water facies independent of the occurrence of specific facies. Cycles are pattern-defined and do not require the presence of tidal flat facies or any other recurrent lithology.
Five partially complete fifth order cycles are recognized at Thacher Park in the interval represented by the Rondout and Manlius Formations (PAC 3 by itself; PACs 5-6; PACs 7-9; PACs 10-12; PACs 13-14). In that fifth order cycles are thought to be the product of the 100ky eccentricity cycle this section at Thacher Park is incomplete in terms of the number of 6th order cycles within each 5th order cycle. Reasons that could explain this incompleteness include hiatus, vacuity, failure to recognize 6th order cycles and incorrectly designated 5th order boundaries. PACs 5-14 are potentially the record of a 4th order cycle, the product of the 400ky eccentricity cycle.
The final purpose of using this section is to illustrate our approach to correlation at the level of 6th order cycles. For example, distinctive widespread lithologies are introduced at the bases of two PACs in the sequence of PACs at Thacher Park. A subtidal well-sorted calcarenite appears in PAC 7 at all localities and at all localities PAC 11 is thick bedded and contains well developed stromatoporoids. These two PACs in which relatively large facies changes occur represent the response to larger deepening events in the basin. PAC 7 is the first 6th order cycle in a 5th order sequence of cycles and PAC 11 contains the deepest facies in its 5th order sequence. In summary distinctive lithologies in certain cycles, the magnitude of facies change at cycle boundaries and the sequential pattern of cycles are all keys to correlation.

STOP 2. Rondout to Coeymans
This locality illustrates a record of 6th order cycles in a different, slightly more offshore set of Manlius facies than those seen at Thacher Park. For example, PACs 9 and 12, in which high tidal-flat facies were developed at Thacher Park, are here represented by nearly all subtidal facies. Second the record of 6th order cycles associated with unconformities is more complete. PAC 2 (actually a 5th order sequence), PAC 6b and PAC 15, recognized here at Schoharie, are missing in the more onshore Thacher area. PAC 2 and 6b are missing owing to hiatus (non-deposition) at Thacher; PAC 15 is missing due to erosion (i.e. vacuity).
In addition to PAC boundaries and unconformable surfaces, the section at Schoharie contains good examples of intracycle allogenic surfaces. Whereas PAC boundary surfaces are thought to be produced by the eustatic sea-level rise forced by the precessional (20ky) cycle, intracycle allogenic surfaces are produced by the complementary sea-level fall on the precessional eustasy curve. Thus each 6th order cycle potentially may be composed of a set of sea-level highstand facies overlain abruptly by a set of lowstand facies. The surface between the two facies sets is thus an intracycle allogenic stratigraphic surface. Such surfaces are well illustrated in PACs 9, 11 and 12, in which massive ostracod bearing calcarenites sharply overlie thin bedded ribbon calcarenites.
Finally the Manlius-Coeymans boundary is an erosional unconformity. At Syracuse, 100km to the west, an additional two 6th order cycles are preserved below this erosion surface (PACs 16 and 17) while at Kingston, 100km to the southeast, PACs 14 and 15, present at Schoharie, have been removed by erosion (fig. 1). This major erosion surface is the basis for defining this surface as a third order "Vail" sequence boundary. The section from the Brayman-Cobleskill unconformity to this unconformity represents the second 3rd order B sequence in the Helderberg supersequence. The section from this surface up to the top of the Becraft Formation (seen in the next two stops) is the third 3rd order B sequence. Erosion associated with the Oriskany unconformity (last stop} removes most of a fourth 3rd order B sequence in the Schoharie area.

STOP 3A. Kalkberg to Becraft
This final stop is in three parts. The first and stratigraphically lowest part is on the north side of I-88 two miles west of STOP 2. The lower 25 feet of this exposure are equivalent to the upper half of the Hannacroix Member of the Kalkberg Formation in the Kingston-Catskill area and the upper five feet are equivalent to the lower part of the Broncks Lake Member (fig. 2) . The base of this section is 2-3 feet stratigraphically above the top of rocks exposed in the section seen at STOP 2. The section continues in the large road cut on the south side of I-88. There is a seven foot covered interval between the two exposures. The thickness of these two covered intervals can be exactly determined because the entire stratigraphic interval is continuously exposed three miles to the east in the Schoharie Quarry and all exposed cycles and bounding surfaces can be correlated between the two localities. The Becraft and Oriskany Formations are exposed at the top of the large roadcut; however a more accessible exposure of the upper Becraft and Oriskany rocks occurs in I-88 roadcuts one mile farther west.
Our purpose in looking at the stratigraphic interval between the Manlius-Coeymans unconformity and the top of the Becraft is to lay out a tentative interpretation of the hierarchic cyclic structure of a fairly complete 3rd order B sequence representing 1-2 million years of accumulation. The entire section is in subtidal shallow to deep ramp facies. It is divisible into three 4th order sequences associated with the 400ky eccentricity cycle. These sequences in turn are each divided into four 5th order cycles related to 100ky eccentricity. Examples of 6th order cycles (PACs) in these facies will be pointed out, but analysis at this level is incomplete. Analysis of 4th and 5th order cycles is based on pattern and magnitude of facies changes and depends on correlation of those patterns to other localities. In this case larger scale cycle boundaries have been correlated to the Cherry Valley area to the west and to localities at Broncks Lake, Catskill and Kingston in the Hudson Valley. In short 3rd, 4th and 5th order cycle boundaries in the Schoharie sections have been correlated with prominent surfaces in sections over 100km away.
4th Order Cycle I
This cycle begins at the Manlius-Coeymans unconformity, also a 3rd order B boundary, and ends at the 161 foot mark on the log (fig. 2). The upper boundary is a surface between massive calcarenite and chert bearing calcisiltite (near the base of the large cut on the south side of I-88. A prominent shale occurs on the surface. The same surface is found on top of the shallowest facies in the Broncks Lake Member in the Hudson Valley (e.g. 16 feet below the New Scotland at Kingston). In systems tract analysis this 4th order cycle represents the shelf margin wedge of the third order B sequence. Fourth order cycle I can be divided into four 5th order cycles:
5th Order Cycle 1 ends at 108. 5 feet on the log at the shallowest point in Coeymans facies. This surface is the Coeymans-Kalkberg Boundary at Kingston and Catskill.
5th Order Cycle 2 ends at 135 feet on the log at a surface separating coarser favositid rich calcarenites from gypidulid bearing, more argillaceous calcarenites.
5th Order Cycle 3 ends at 140 feet on the log at a surface just below a one foot thick shale interval with lime turbidite beds. This surface is the base of the Broncks Lake Member in the Hudson Valley.
5th Order Cycle 4 ends at the 4th order boundary.
4th Order Cycle II
This cycle contains the deepest facies in the third order sequence and at Schoharie its thickness is highly condensed (32 feet versus 66 feet at Kingston). The lower boundary (described above) is at 161 feet; the upper boundary, at 193 feet on the log, is the surface at the top of a six-foot-thick massive calcisiltite unit. In systems tract analysis of the third order B sequence the lower two 5th order cycles in this 4th order cycle represent the transgressive systems tract and the 5th order boundary at 181 feet would be the maximum flooding surface. Fourth order cycle II is divisible into four 5th order cycles:
5th Order Cycle 1 is 10.5 feet thick and in the lower part is characterized by several amalgamated lime turbidites with cherts. These amalgamated units are thought to represent 6th order cycles. The upper 6th order unit is represented by 2.5 feet of calcisiltites in discontinuous lenses. The top surface of this first 5th order cycle is correlated to the Kalkberg.New Scotland formational boundary in the Hudson Valley.
5th Order Cycle 2 ends at 181 feet on the log; a concentration of chert and a thick shale occur at this boundary.
5th Order Cycle 3 ends at 187 feet on the log; beds representing condensed 6th order cycles thicken up through this 5th order cycle.
5th Order Cycle 4 is the thick calcisiltite unit.
4th Order Cycle III
This is the thickest of the three 4th order sequences at Schoharie. It is 66.5 feet thick and in general a regressive sequence of 5th order cycles, beginning in deep ramp and ending in shoreface facies. This unit is equivalent to the highstand systems tract of sequence stratigraphy. The top of the cycle is placed at a surface above the coarsest calcarenite facies in the Becraft, at 271 feet on the log. This surface is also a third order B sequence boundary. The Oriskany unconformity, 8 feet above this surface, has eliminated most of the last 3rd order sequence of the Helderberg supersequence (third order A). At Kingston the upper 3 feet of the Becraft and the Alsen-Port Ewen Formations comprise this 3rd order sequence. Fourth order cycle III consists of four 5th order cycles:
5th Order cycle 1 ends at 204.5 feet on the log and is an 11.5 foot thick set of amalgamated lime turbidites.
5th Order cycle 2 ends at 218.5 feet on the log just below a 4-6 inch shale and like the cycle just below, is a 14 foot thick set of amalgamated turbidites.
5th Order Cycle 3 ends at 237 feet on the log just above a massive calcisiltite bed. Rocks above this surface are coarser and would generally be described as Becraft.
5th Order cycle 4 extends to the top of the 3rd order sequence and consists of five coarsening upward, calcarenite shallow ramp cycles and is the thickest of the 5th order cycles in the set.

STOP 3B Oriskany/Esopus
The last part of STOP 3 is one mile farther west on I-88. Here the surface between Becraft/Alsen calcarenite and fossiliferous quartz sandstone of the Oriskany represents superimposed 2nd (i.e. Sloss sequences), 3rd, 4th, 5th and 6th order boundaries. There are two 6th order cycles in the Oriskany followed by a major deepening into basinal Esopus facies.

A5: MIDDLE DEVONIAN NEAR-SHORE MARINE, COASTAL AND ALLUVIAL DEPOSITS, SCHOHARIE VALLEY, CENTRAL NY STATE
J.S. BRIDGE AND B.J. WILLIS
Department of Geological Sciences
State University of New York
Binghamton, New York, 13902-6000
The purpose of this field trip is to visit five outcrops which display the diversity of rock types exposed in Schoharie Valley. These deposits record a wave-and tide-influenced deltaic shoreline where the tidal range and degree of wave influence varied in space and time. During this trip we will focus on the relative influence of wave, tidal and fluvial currents in shaping this ancient coastline and nearshore sea bed.

A5: Middle Devonian Near-shore Deposits

STOP 1: Hardenburg Falls
Description The lowest 6 meters of the exposure at Hardenburg Falls (Fig. 5) contains interbedded gray sandstones, dark-gray siltstones and shales. Lenticular-wavy bedded parts contain very-fine grained sandstone beds which fine up to mudstone. These sandstone beds are wave-ripple cross-laminated and have wave-ripple forms on their tops. Thicker sandstone beds (typically 0.1 to 0.5 meters-thick) have sharp, erosional bases overlain by intraformational shale fragments and disarticulated shelly fossils. They are fine to very-fine grained and fine upwards. Internal structures change upwards from hummocky cross-strata or planar-strata to wave-ripple cross-lamination. Tops of beds are covered with wave-ripple marks (3-D, 2-D, and interfering types). Two-dimensional wave ripples indicate SW-NE paleocurrent oscillation. Some low-angle, hummocky cross-strata with superimposed wave-ripples and abundant shell fragments show a preferred paleocurrent direction to the NE. Rarely, such relatively coarse grained, shell-rich strata contain angle of repose cross-strata with SE paleocurrents. At the base of the section, amalgamated hummocky cross-stratified beds fill N-S oriented channels up to decimeters deep and meters wide. These beds contain rounded extraformational chert and quartzite pebbles, and intraformational shale clasts within a medium sandstone matrix.
Shell accumulations include the brachiopods Spinocyrtia, Mucrospirifer, Mediospirifer, Orthospirifer, Cupularostrum, the bivalves Goniophora, Actinopteria, Palaeoneilo, and crinoid ossicles. Most of the disarticulated shell fragments lie concave down along stratification surfaces and show little evidence of breakage or abrasion of fine detail. Some shell concentrations are markedly monospecific (e.g. Cupularostrum). Small, centimeter diameter vertical burrows occur in the tops of sandstone beds.
The 5 to 6 meter thick sandstone body in the middle of this section has a sharp base which is overlain by 2 meters of planar-stratified and swaley cross-stratified, fine to medium grained sandstone. These beds contain horizons of symmetrical and asymmetrical wave ripples with shale drapes. Parting lineation indicates NNE-SSW paleocurrents and asymmetrical wave ripples indicate a NE-SW oscillation with a preferred migration to the SW. These beds are overlain by about 2 meters of trough cross-stratified sandstone. Paleocurrents arc towards the NE mainly, with the exception of a rare, oppositely directed cross-set. Wave and current ripples are superimposed on the large-scale cross-strata. One set of small-scale cross-strata indicates an easterly paleocurrent. Large-scale planar cross-strata occur above the trough cross-strata and climb up low-angle lateral-accretion surfaces; that is, inclined bedding surfaces dip to the NW (normal to the underlying trough cross-strata) and the planar cross-strata within beds dip to the SE. Vertical, meniscate burrows of varying diameter occur in this part of the sandstone body. The upper part of the sandstone body is wave-ripple laminated, very-fine grained sandstone, which contains sparse shell fragments. This is overlain by several beds of hummocky cross-stratified and wave-rippled, fine-grained sandstone with abundant shell fragments.
Interpretation Interbedded shales and sandstones containing hummocky cross-strata, planar-strata, wave-ripple cross-laminae, wave-ripple marks, and abundant shelly-fossil concentrations are interpreted as nearshore marine deposits formed below fair-weather wave base (cf. Dott and Bourgeois, l 982; Harms et al., 1982; Hunter and Clifton, 1982; Swift et al., 1983). The sandstones are generally interpreted to be formed during storms by combined wave and unidirectional currents, whereas the shales generally represent fair-weather deposits or those formed below storm wave base. Oscillatory flows are generally directed alongshore, whereas the stronger unidirectional currents represented by low-angle to angle-of-repose cross-strata are directed onshore (easterly direction). Although the shelly fossils in sandstone beds are clearly transported, their preserved detail precludes long periods or distances of transport. McGhee and Sutton (1985) believe that such shelly-fossil concentrations reflect the diversity of the indigenous infauna and epifauna (in this case typically nearshore marine), although their relative abundances do not reflect original population densities. Monospecific shell concentrations may well reflect only slightly disturbed original communities, as it is unlikely that hydraulic sorting could so perfectly separate shells of one species from other shells with similar hydrodynamic properties.
The thick sandstone body in the middle of this exposure is interpreted as deposits of a channel mouth bar associated with a tidal channel which prograded into a storm-wave dominated sea. Swaley-hummocky cross-strata, planar-strata and wave-ripples in lower beds but angle-of-repose cross-strata in upper beds, are evidence for variation in wave and unidirectional currents upwards. Lower down within the sandstone, amalgamated hummocky-swaley cross-stratified beds and wave-ripples overlain by shale drapes indicate periodic variation in combined-flow strength. Oscillatory currents associated with wave-ripples are generally to the NE and SW (alongshore). Upper beds with trough cross-strata directed alongshore (NE) to landward (E) (but rarely in the opposite direction) are most likely associated with sinuous crested dunes formed by strongly asymmetrical tidal flood currents. The SE-directed planar cross-strata may represent wave-formed swash bars or straight-crested tide formed dunes (e.g.. Hayes and Kana, 1976; Boothroyd, 1978; Fitzgerald, 1984). Superimposed wave- and current-ripples indicate periodic variation in current strength, and the presence of weak wave currents. Inclined bedding surfaces exposed near the top of the sandstone testify to lateral migration of the upper surface. The sparse fauna and low degree of bioturbation is typical of sandy deposits in coastal areas (Gould, 1970; Howard and Reineck, 1981).
Fossiliferous sandstones and mudstones above the channel mouth bar deposit indicate rapid return to marine conditions. Depositional environments at the shoreline probably depended on the position of tide-influenced channels which distributed sediment to the shoreline and near-shore areas. Avulsion of the channel from this area abruptly reduced local deposition rates, allowing the sea to transgress locally over this location.

STOP 2: Stevens Mountain Quarry
This quarry exposure is important because it provides continuous exposures from wave-dominated marine deposits to tide-dominated channel deposits. The extensive exposures also provide a better record of lateral variations than were exhibited at the last stop.
Description Much of the quarry exposure is composed of 2 to 6 meter thick, gray sandstone bodies separated vertically by mudstone dominated intervals of dark-gray shales interbedded with centimeter-to decimeter thick gray sandstone beds (Fig. 6). The lowest 4 to 5 meters exposed above the quarry floor contains decimeter-thick beds of fine-grained sandstone with some centimeter-thick shale horizons. The sandstone beds have sharp bases overlain Ln places by drifted plant remains and shale chips. Internal structures are difficult to discern because of intense bioturbation by roots and burrows. However, large-scale cross-stratification (westerly paleocurrents) occur, and tops of beds have small-scale cross-strata and wave-ripple marks. Burrows are mainly vertical, sand-filled tubes which range in diameter from 2 to 10 centimeters: some have meniscate fills.
Within the overlying 11 to 12 meters of section (Fig. 6a, meters 5 to 17) occur several sharp-based, 2 to 3 meter-thick, fine-grained sandstone bodies separated by shale-dominated intervals with thinner sandstone beds. Bases of sandstone bodies show tool marks, and are commonly lined with plant remains, shell fragments and shale chips. The sandstone bodies are composed of decimeter-to meter-thick beds containing hummocky-swaley cross-stratification and planar-stratification. Individual beds in such amalgamated sequences are recognizable by their erosional bases overlain by shell concentrations of transported brachiopods, bivalves and Tentaculites. The upper beds within these sandstones are commonly shell-rich and contain intraformational shale chips and extraformational quartz pebbles. Angle of repose cross-stratification also is common in these upper beds, normally indicating an E to NE paleocurrent orientation. Wave ripples cap most sandstone bodies. The sandstone bodies vary in thickness and sediment type laterally. This change is most dramatic in the lowermost sandstone body in this succession. To the north, the sandstone body is about 3 meters thick and contains mainly large-scale cross-stratification and planar-stratification, indicating a range of paleocurrent directions (Fig. 6a, meter 5-8). To the south, this body thins to about 2 meters and contains mainly swaley cross-strata (Fig. 6b, meter 6 to 8). Shale-rich zones in the sequence contain decimeter-thick hummocky cross-stratified sandstone beds with wave-rippled tops. Load casts occur in distinct, laterally extensive horizons, particularly immediately beneath sandstone bodies.
The next beds contain 4 to 5 meters of interbedded shale and sandstone (Fig. 6a, meters 17.22). The lowermost decimeter-thick sandstone beds in this sequence (immediately above a sandstone body) are markedly lenticular, and rich in shells, intraformational breccia and extraformational quartzite pebbles. They contain hummocky cross-strata and, in places, angle of repose cross-strata with easterly paleocurrents. Wave ripple marks cap these sandstone beds. Most of the remainder of this interval is wavy and lenticular bedded, with very-fine grained sandstone occurring in centimeter-thick, wave-rippled beds. In the north end of the quarry, these lenticular-wavy beds coarsen and thicken upwards to decimeter-thick, hummocky cross-stratified, very-fine grained sandstone beds with wave-rippled tops, and to amalgamated hummocky cross-stratified, fine grained sandstone beds with abundant plant remains and intraformational breccia.
These beds are sharply overlain by a 5 to 6 meter-thick fine to medium grained sandstone body (Fig. 6a, meters 23 to 28). It is dominated by large-scale trough cross-stratification (NW paleocurrent), with the exception of planar stratified beds above the erosional base and some small-scale cross-strata in the finer beds near the top. The top most decimeter-thick beds have abundant drifted plant remains, shells, intraformational shale and extraformational pebbles in their bases. Paleocurrents in these upper beds are to the SE. Large vertical meniscate burrows (5-10 centimeter diameter) occur in the lower parts of the sandstone body, but smaller (1-2 centimeter diameter) vertical burrows occur in the upper most beds. In the south end of the quarry this sandstone body is thinner (Fig. 6b, meters 19 to 23). Lower parts of the sandstone body are replaced by amalgamated swaley cross-stratified, fine grained sandstone beds, and the upper two meters are replaced by hummocky cross-stratified sandstone beds with wave-rippled lops and shale interbeds. Underlying beds are finer grained than to the north. The basal erosion surface contains flutes (E paleocurrent) and grooves, and is overlain by bivalve and brachiopod shells and plant remains. Paleocurrents in the central, coarsest part of the sandstone body are to the east.
Approximately 4 meters of interbedded sandstone and shale overlie this sandstone body (Fig. 6a, meters 29 to 33). The sequence is mainly wavy-lenticular bedded with centimeter-thick wave-rippled, very-fine grained sandstone beds. Within this succession to the south is a 0.9 meter-thick, hummocky cross-stratified sandstone bed with an erosional base and wave-rippled top (fig. 6b, meter 18). It occurs at the top of a coarsening/bed-thickening upward sequence. To the north, this bed changes laterally to a 1.3 meter-thick, fine to very-fine sandstone bed with large-scale cross-strata at the base (westerly paleocurrent) and small-scale cross-strata at the top (Fig. 6a, meter 31 to 32). This bed is in turn cut out laterally by a fine-grained channel fill. At the top of this sequence, immediately below the top most sandstone body, is an extensive zone of load casts. The top most sandstone body in the quarry is approximately 6 meters thick (Fig, 6a, meters 33 to 39). The erosional base is overlain by intraformational breccia (shale blocks up to decimeters across) with brachiopod and bivalve shells in a medium sandstone matrix. The sandstone body is medium to fine grained and fines upwards in places. It is mainly large-scale trough cross-stratified, with some lenses of planar-stratification. Reactivation surfaces in the cross-strata of the lower part of the sandstone body are mainly erosional. Between laterally adjacent reactivation surfaces, cross-strata vary in dip. In places oppositely directed, asymmetrical ripple marks are preserved on foresets and reactivation surfaces, and shale drapes may also occur. Large vertical meniscate burrows also occur in these lower beds. Paleocurrents associated with large-scale cross-strata are towards the WNW in the lower half of the sandstone body, but towards the ESE in the upper half. Low angle, lateral-accretion surfaces extend through the sandstone body, and a sandstone filled channel cuts through the upper one third.
Interpretation The sandstone body at the top of this exposure is interpreted as the deposits of a laterally-migrating tidal channel (see also Johnson and Friedman, 1969). River channels preserved in coeval Catskill facies to the east have generally north-westerly paleocurrents, but range from northerly to southerly (Willis and Bridge, 1988). Therefore paleocurrents with a westerly component are in the tidal ebb and river flow direction, whereas those with an easterly component are in the tidal flood direction. Variation in ebb paleocurrent orientation in various examples of this type of sandstone body suggest that channels had a range of orientations due to local channel curvature and/or a distributive channel pattern. Dominance of ebb-directed paleocurrents lower in the sandstone bodies, with flood currents and channel fills in upper parts is typical of the strongly asymmetrical tidal currents in tidal inlets, estuaries, and tide-influenced deltaic distributaries.
Variable dip angles of large-scale cross-strata, erosional reactivation surfaces and superimposed current ripples are further evidence of strong tidal-current asymmetry. They indicate growth of sinuous-crested dunes to 'full vortex' stage during the dominant tidal current, slackening of the current and modification of dune geometry, then erosion of the dune by the subordinate tide. Such features are commonly reported from mesotidal and macrotidal estuaries (e.g. Terwindt, 1981; Boersma and Terwindt, 1981; Dalrymple, 1984; DeMowbray and Visser, 1984). General absence of tidal bundle sequences and rarity of current ripples on reactivation surfaces suggests an erosional or nondepositional subordinate flood-tidal current and a dominant ebb-tidal current that may be reinforced by a fluvial current. Bioturbation is characteristically rare in such coastal sandstone bodies (e.g. Howard and others, 1975). Large, vertical meniscate burrows are similar to those reported in fluvial and coastal channel deposits elsewhere (Thorns and Berg, 1985; Miller, 1979; Bridge, Gordon and Titus, 1986) and may be due to upward escape of bivalves. Brachiopod-bivalve fauna sparsely distributed through the sandstone bodies is also typical of nearshore communities (McGhee and Sutton, 1985; Sutton and McGhee 1985).
Sandstone bodies just below this top most sandstone body, with swaley-hummocky cross-strata, planar-strata and wave-ripples in lower parts, but angle-of-repose cross-strata in upper parts, are comparable to the sandstone body exposed in Hardenburg Falls (Stop 1). They are also interpreted to record evidence for combined wave and unidirectional currents of varying importance and intensity. Detailed interpretation of these deposits hinges critically on their lateral transition to thicker sandstone bodies dominated by angle-of-repose cross-strata and thinner sandstone bodies dominated by amalgamated hummocky-swaley cross-strata. It is also important that their bases are sharp in many places, but in others they occur at the top of coarsening-upward sequences with interbedded hummocky cross-stratified sandstones and shales immediately beneath. Furthermore, any evidence of channel fills or lateral-accretion bedding is restricted to their upper parts. These features suggest deposition on channel mouth bars which were prograding rapidly into a marine, storm-wave-dominated area. Such bars may be on the seaward side of tidal inlets (i.e. ebb-tidal deltas) associated with estuaries or barrier-beach shorelines, or associated with tide-influenced deltaic distributaries. Landward-directed (E, NE) trough cross-strata are most likely associated with sinuous crested dunes formed by strongly-asymmetrical tidal flood currents. Sets of planar cross-strata migrating to the NE, E or SE may represent wave formed swash bars or straight-crested tide-formed dunes (e.g. Hayes and Kana, 1976; Boothroyd, 1978; Fitzgerald, 1984). Wave and current ripples on dune/bar slipfaces and troughs indicate periods of tidal slack water and decreased tidal current strength, respectively. Sparse fauna and low degree of bioturbation (but dominance of the Skolithos ichnofacies) is typical of sandy deposits in coastal areas (Gould, 1970; Howard and Reineck, 1981).
Similar sandstone bodies in somewhat younger Frasnian rocks in the Binghamton (NY) region were also interpreted as channel mouth bar deposits by Halperin and Bridge (1988). However, similar sequences reported from the Cretaceous Western Interior Seaway have been interpreted as due to deposition on prograding strandplains cut by estuaries (McCrory and Walker, 1986). McCrory and Walker (1986) and Hamblin and Walker (1979) interpreted amalgamated swaley-hummocky cross.stratified beds as beach and shoreface deposits, even though they rest sharply on interbedded hummocky cross-stratified sandstones and shale interpreted as offshore storm-dominated deposits. They interpreted this sharp contact as due to rapid progradation of the shoreface. Similar sandstone bodies occurring at the tops of gradational coarsening upward sequences have been interpreted as due to progradation of a wave-dominated deltaic shoreline (Chan and Dott, 1986; Swift et al, 1987). Swift et al.(1987) interpreted the amalgamated hummocky cross-strata as middle shoreface deposits, and angle-of-repose cross-strata as deposits on the upper shoreface by combined currents with an alongshore unidirectional component. A beach face origin for the sandstone bodies studied here is considered unlikely because of a lack of characteristic seaward-dipping planar-laminae with heavy mineral layers; alongshore directed cross-strata overlain by landward-directed cross-strata and planar-strata typical of ridge and runnel systems; and eolian cross-strata. The sharp base to many (but not all) of the sandstone bodies is hard to explain by rapid progradation of a beach face, but may be readily explained by rapid progradation of a storm-wave modified channel mouth bar which formed rapidly as a result of a channel diversion. Coleman and Prior (1982) show examples of distributary mouth bar sands that lie abruptly on offshore silts, and state that bases of sandstone bodies become sharper closer to the distributary channel.
Relatively thin sandstone bodies dominated by amalgamated swaley-hummocky cross-stratified beds are interpreted as more distal parts of channel mouth bars, which are completely dominated by storm waves (combined currents). The lack of mud suggests deposition above fair weather wave base. The common occurrence of load casts near their bases suggests rapid deposition of sand on mud. Such soft-sediment deformation features are common offshore from Mississippi delta distributaries (e.g. Coleman and Prior, 1982). Paleocurrent indicators in the base of the sandstone bodies suggest a dominantly offshore directed unidirectional current, as is common in similar Frasnian rocks in the Catskill region (Craft and Bridge, 1987; Halperin and Bridge, 1988). Lenticular sandstone beds on top of these sandstone bodies, with abundant shelly fossils, intraformational breccia, extraformational pebbles, and local channelling, are evidence of relatively strong unidirectional currents directed to the east and southeast (i.e. shoreward) superimposed upon wave currents. As there is no evidence of offshore directed currents, the unidirectional currents are not considered to be tidal flood currents but rather storm-wave associated currents. As these sandstone bodies are sharply overlain by interbedded sandstones and shales interpreted as deeper water nearshore deposits, these uppermost shelly sandstone beds are apparently associated with rapid abandonment of the channel mouth bar and reduction in sand supply. Johnson and Friedman (1969) interpreted these sandstone bodies as nearshore subtidal bar deposits, and Miller and Woodrow (in press) assigned an estuary mouth shoal origin.
Similar sandstone bodies elsewhere have been interpreted as wave-formed offshore bars or shoals (e.g. De Raaf et al, 1977; Cotter, 1985). Wright and Walker (1981) have interpreted thin pebble layers at the tops of such sequences as transported by density currents. This is considered unlikely here. If such sandstone bodies do represent distal parts of storm-wave modified channel mouth bars, then their occurrence within sequences of deeper water sandstone and shales indicates channel switching (see also Swift et al, 1987). The overall sequence at Stevens Mountain Quarry indicates sandstone body deposition in progressively more landward parts of the channel-mouth bar complexes. Thus, this is an overall regressive sequence with superimposed variations in deposition related to local channel avulsion.
As implied above, interbedded shales and sandstones with hummocky cross-strata, planar-strata, wave-ripple cross-laminae wave-ripple marks, and abundant shelly-fossil concentrations are interpreted as nearshore storm-dominated marine deposits. These are interpreted in a similar way to those exposed within Hardenberg Falls (Stop 1). Sandstone beds near the base of the quarry were deposited mainly by N-NW directed unidirectional currents moving sand as dunes. Wave ripples in the upper parts of sandstone beds indicate deposition in standing water with weak wave currents. These characteristics, plus the abundance of burrows and roots, and the absence of marine fossils, suggest deposition by fluvial or ebb-tidal currents on a near coastal plain.

STOP 3: Manorkill Falls, Schoharie Reservoir
Description Most of this section is composed of gray sandstone interbedded with red, green and gray siltstone and mudstone (Fig. 7). The lower part of this section contains two 10 meter-scale fining upward sequences (Fig. 7, meters 1 to 24). The lower 4 meter-thick interval within the first sequence has a high proportion of sandstone which sharply overlies intensely disrupted mudstones. These sandstone beds generally thin and fine upwards. Individual sandstone beds are sharp-based, sheetlike to lenticular, and centimeters to a meter thick. Thinner sandstones are very-fine grained, and small-scale cross-stratified with associated current ripples or (mainly) wave ripples. Thicker sandstone beds are mainly fine grained, and contain large-scale trough cross-strata, planar-strata, or (unusually) hummocky cross-strata passing upwards to small-scale cross-strata and associated current or wave ripples. Some thicker sandstone beds lower in the sequence show low-angle (<10°) inclined bedding surfaces extending throughout most of the thickness of the bed which dip in the direction the bed thins (i.e. to the east). Upper parts of these inclined surfaces are covered with straight crested and interfering wave-ripple marks, indicating oscillatory paleocurrents NE-SW and E-W. Oscillatory paleocurrent directions are quite variable throughout the sequence; however, a NE-SW orientation is common. Large-and small-scale cross-strata and current ripples consistently indicate an easterly paleocurrent.
Sandstone beds contain vertical burrows (including Skolithos), which range in diameter from a few millimeters to 7 centimeters; larger ones commonly showing meniscus structure. Horizontal meniscate burrows occur in places in the tops of sandstone beds. Arenicolites burrows also occur. Drifted plant remains are common in sandstones, and root casts occur in the top of some sandstone beds. There are two important occurrences of tree-trunk casts. Shelly fossils are rare. One shell-rich layer near the base of the section occurs near the top of a hummocky cross-stratified sandstone bed. This layer contains bivalves, brachiopods (e.g. Spinocyrtia, Allanella tullius) juvenile crinoid fragments and bryozoa. Fish fragments are found in places.
Mudstones are either relatively unbioturbated and fissile, or bioturbated with abundant rootcasts, small vertical and horizontal burrows, and desiccation cracks. A nodular carbonate bed occurs near the top of the first sequence where calcareous rhizoconcretions are particularly abundant. This bed is similar to the one occurring at the New York City Water intake building.
The base of the second fining upward sequence is another sandstone-dominant interval (about 3 meters-thick), in this case dominated by large-and small-scale cross-strata (Fig. 7, meter 16 to 19). This interval is in turn overlain by an interval of overlapping channel fills. These channels are oriented E-W and are filled with alternating mudstone and sandstone beds. The sandstone beds are dominated by wave-ripple lamination, with rare planar and hummocky strata. Mudstones occurring adjacent to and above these channel fills are highly disrupted by mudcracks, burrows, rootcasts and in one location a tree trunk cast.
The sandstone body at the top of this section is similar to those occurring near the top of Stevens Mountain Quarry and the lower Grand Gorge route 30 section, and shows lateral-accretion surfaces. Paleocurrents near the base of the sandstone body are NW-SE.
Interpretation In the Manorkill Falls section, desiccated, bioturbated siltstones with abundant root casts, rare tree trunk casts, and horizons of calcareous concretions, clearly represent flood basins with a well developed flora which were periodically exposed and subjected to soil-forming processes. Sandstone beds overlying such deposits near the base of the section are distinctive and show evidence of easterly progradation of wedges of sand into at least temporarily standing water in which wave currents were active. The lowest bed buried and preserved the trunk casts. Two origins for this bed are considered possible: (1) a fluvial overbank sand splay, or (2) a back barrier washover deposit. Paleocurrent evidence supports the storm washover origin and deposits show many features in common with modern examples (e.g. Schwartz, 1982). However, there is little evidence in these rocks of the presence of beaches, so barrier islands seem unlikely. Although lateral migration of tidal inlets or other types of coastal channels may have removed the evidence of beaches, an overbank sand splay from a major channel is probably the most viable explanation.
Associated with the sandstone-dominated interval near the base of the section are sandstone beds showing evidence of strong wave currents, containing marine shelly fauna, and with a diverse assemblage of burrowing organisms. These deposits testify to the close proximity of the sea, and the onshore transport of marine fossils by storm waves. Furthermore, the overlying fissile siltstones with isolated hummocky cross stratified and wave-rippled sandstone beds indicate deposition of mud in quiet, standing water which was periodically agitated by waves. All of this evidence together suggest an upward transition from fluvial floodbasin/lacustrine conditions to a brackish interdistributary bay with access to fully marine waters.
A return to subaerial floodbasin conditions is signaled by desiccated mudstones with root casts, rhizoconcretions, fish fragments and a nodular carbonate-rich horizon (beds directly under the bridge; Fig. 7, meters 15-16). The overlying sandstone beds at the base of the second 10 meter-scale sequence record periodic deposition by temporally-waning unidirectional currents, probably associated with river floods. Desiccation cracks in the siltstones capping these sandstone beds and the lack of wave ripples, suggests subaerial exposure immediately following deposition. This sequence of interbedded sandstones and siltstones is interpreted as a crevasse-splay complex, but it is hard to dismiss flood tidal currents (storm enhanced?) as a depositional agent in view of the overall coastal setting.
The overlying sequence of interbedded sandstones and shales also has features indicative of mud deposition in an area of shallow water (but with periodic subaerial exposure) in which weak wave currents periodically deposited sand, and transported marine fossils into the area. The east-west oriented channel in this part of the section were apparently filled mainly by weak wave currents. The lack of evidence of reversing unidirectional sedimentary structures precludes a tidal-channel origin, and suggests instead a fluvial channel cutting across a floodbasin/interdistributary bay. Indeed, deposits typical of intertidal flats in mesotidal or macrotidal settings are notably absent throughout this outcrop. The sandstone body at the top of this section is similar to others viewed at previous stops and are interpreted as a channel-mouth bar deposits. Johnson and Friedman (1969) interpreted these deposits as interfingering alluvial marsh, tidal flat and channel deposits. They are also somewhat similar to the "muddy shoreline" deposits of Walker and Harms (1971).
The 10 meter-scale sequences exposed at Manorkill Falls reflect relative sea level changes. Mudstone beds immediately below sandstone-rich intervals reflect floodplain emergence. Abrupt onset of sand deposition following fining indicates an abrupt change in depositional environment and the onset of a deepening trend. Such sequences may be related to channel switching which locally varied deposition rate at the shoreline.

STOP 4: Schoharie Creek outcrop
Description The lowermost 11 meters of interbedded sandstone and mudstone in this stream and hillside section is, with the exception of the 2 to 3 meter thick sandstone body (Fig. 8, meters 4.5 to 7), similar to the lower parts of Manorkill Falls section (Fig. 8). Centimeter-thick, very-fine grained sandstone beds are mainly wave-rippled, whereas the decimeter-thick, fine-grained sandstone beds may also contain large-scale cross-strata (paleocurrents to N), planar-strata, and hummocky cross-strata. Tops of sandstone beds are normally wave-rippled and bioturbated by root casts and/or vertical, horizontal and U-shaped burrows of various sizes (millimeters to centimeters in diameter). Gray siltstones are bioturbated by roots and small burrows, and desiccation cracks are common. Calcareous concretions occur locally, and buff calcareous shale commonly fills desiccation cracks. There are several carbonate-rich mudstone beds near the top of this interval, in one case associated with pseudo-anticlines. Some siltstones are relatively undisturbed and fissile.
The 2 to 3 meter thick, fine-grained sandstone body near the base of the section is somewhat unusual. It is dominantly large-scale trough cross-stratified, with some large-scale planar cross-strata and planar-strata. Large-scale cross-strata indicate paleocurrents to NNW and in a generally easterly direction. Wave- and current-ripples occur on reactivation surfaces, the current-ripples oriented in the opposite direction to the adjacent large-scale cross-strata. Ripple-marked surfaces throughout the sandstone body show vertical and horizontal meniscate burrows of various sizes centimeters in diameter.
Above poorly exposed sandstone bodies between meters 11 and 19 (Fig. 8) there occurs a 3 meter thick interval of sandstone and mudstone beds that are similar to those lower down. Above this interval is approximately 12 meters of section dominated by decimeter-thick fine to very-fine grained sandstone beds, with only small amounts of gray mudstone. Sandstone beds have erosional bases and are mainly large--scale cross--stratified, with minor amounts of planar and low-angle cross-strata. Upper parts of the beds are commonly wave-ripple laminated, with wave rippled tops. Paleocurrents from large-scale cross-strata are mainly to the NNW, but some are to the SE. Beds are intensely bioturbated with vertical and U-shaped burrows. A tree trunk cast occurs near the top of the sequence.
Interpretation The lower part of the Schoharie Creek outcrop (Fig. 8, meters 1 to 11) is interpreted in a similar way to deposits exposed in the Manorkill Falls outcrop; that is, fine grained deposition in semi-permanent, near-coastal lakes or bays, with periodic emergence, plant growth and soil formation. Sand was introduced by unidirectional and/or wave currents and reworked by waves. Although trace fossils are abundant and diverse, marine shelly fossils are notably absent.
Evidence for strong tidal currents is generally lacking, with the exception of the 2 to 3 meter thick sandstone body near the base of the section. In this sandstone body bimodal paleocurrent directions in large-scale cross-strata, and wave-and current-ripples on reactivation surfaces all indicate deposition by tidal currents. However, the topmost part of the sandstone body shows more wave influence. Its stratigraphic position between coastal bay or floodbasin deposits suggests deposition in a flood-tidal delta and/or tidal channel (c.f. Boothroyd, 1978). Wave-rippled and extensively bioturbated sandstones lower down in the section along Schoharie Creek are probably sandy tidal-flat deposits.
Sandstone beds in the upper 12 meters of section (Fig. 8) were deposited mainly by N-NW directed unidirectional currents moving sand as dunes. Wave-ripples in upper parts of many sandstone beds indicate deposition in standing water with weak wave currents. Stronger wave currents are indicated by swaley cross-strata near the base of some sandstone beds. These characteristics, plus the abundance of burrowing, lack of roots (except at the top of the section), and absence of marine fauna, suggests deposition by fluvial sheet-floods in a coastal body of standing water. Despite rare southeasterly paleocurrents, there is little evidence to suggest tidal currents. Johnson and Friedman (1969) assigned these rocks to the boundary between tidal and alluvial facies.

STOP 5: Grand Gorge Route 30
Description The lowermost 28 meters of this section (Fig. 9) is closely comparable to the upper 30 meters or so of Stevens Mountain Quarry. In particular, sandstone bodies increase in thickness upwards and contain more angle-of-repose cross-strata relative to swaley cross-strata and planar-strata. There arc also lateral transitions within the lower parts of the upper two sandstone bodies between angle-of-repose cross-strata, swaley cross-strata and planar-strata. Large-scale cross-strata in the upper parts of sandstone bodies indicate paleocurrents to the NNE and NE, whereas those lower down indicate S to SSW paleocurrents. However, solitary sets of planar cross-strata within the planar/swaley parts of these sandstone bodies indicated paleocurrents ranging from NNE to SE. Oscillatory current directions are commonly NW-SE. Large (5-10 centimeter-diameter) vertical meniscate burrows occur commonly in sandstone bodies, along with millimeter-wide Skolithos, Arenicolites and Chondrites. Load casts occur beneath sandstone bodies low in the section. Beds at the base of the section contain marine fossils.
Further up section, above a 25-26 meter covered interval, a 17 meter-thick sequence of red mudstones and red-gray very-fine to fine-grained sandstones are exposed (Fig. 9, meters 53 to 70). Mudstones are either fissile, or intensely disrupted and blocky-weathering with slickensided clay-lined surfaces between blocks. Desiccation cracks are common. Horizons of calcareous concretions (rhizoconcretions) occur rarely, as do pseudo-anticlines. Bioturbation is pervasive and includes root casts and vertical burrows. One burrow type is 1-2 centimeters in diameter and has a meniscate fill. Others are 3 to 5 millimeters in diameter. The sandstone beds are sheetlike to lenticular, and range in thickness from centimeters to decimeters. Thinner beds are small-scale cross-stratified and have wave- or current-rippled tops. Thicker beds may be dominantly small-scale cross-stratified but can have large-scale cross-strata or planar-strata immediately above their erosional base. Sparse paleocurrent data indicate flow to the N or SSW. Large vertical, meniscate burrows (centimeters in diameter) occur in these sandstones but bioturbation is dominated by smaller burrows (1 centimeter diameter). Upper parts of sandstone beds may have root casts and rare calcareous concretions. Meter-thick upward fining and coarsening sequences occur in places (e.g. Fig. 9, meters 58 to 60).
The upper part of the Grand Gorge road cut contains two erosional-based, fine-to medium-grained sandstone bodies. The lower one is dominated by low-angle cross-strata and planar-strata, but also contains angle of repose cross-strata. Along erosion surfaces occur drifted plant remains and a rare occurrence of the nonmarine bivalve Archanodon (internal mold, transported). The upper sandstone body shows lateral-accretion bedding passing laterally into a channel fill. Paleocurrents lower in the sandstone body are to the south. Vertical burrows of varying diameter occur in these lower parts.
Interpretation The lower succession exposed along route 30 is comparable to the top of Stevens Mountain quarry, and is interpreted similarly. Red beds in the middle of the Grand Gorge Route 30 roadcut indicate a fluvial overbank sequence (sec also Miller and Woodrow, in press). Desiccated, bioturbated mudstones represent floodbasin deposits. Blocky weathering mudstones with local calcareous concretions and pseudo-anticlines indicate repeated welting and drying, illuviation, and formation of ped structure typical of calcareous vertisols. Fissile mudstones indicate relatively minor influence of bioturbation and repeated exposure, and may represent relatively high deposition rates in perennially ponded areas. Wave-rippled sandstone lenses indicate periodic wave activity on such ponded areas. Sandstone beds with unidirectional sedimentary structures indicate periodic sand deposition during floods, either as individual sheet flood deposits within the floodbasin, or associated with progradation of levees or crevasse splays. The coarsening-upward sequence at meters 58 to 60 (Fig. 9) suggests progradation of a levee or crevasse splay into a floodbasin. Calcareous concretions and pseudo-anticlines in overlying beds indicated subsequent rapid abandonment, reduction in deposition rate, and soil formation. The overall vertical sequence at Grand Gorge, Route 30 therefore is generally regressive, representing an upward transition from nearshore marine through coastal and fluvial deposits.

A6: INTERACTION BETWEEN KARST AND GLACIATION IN THE HELDERBERG PLATEAU, SCHOHARIE AND ALBANY COUNTIES, NY
ARTHUR N. PALMER
State University College
PAUL A. RUBIN
Environmental Sciences Division
Oak Ridge National Laboratory
MARGARET V. PALMER
R. D. 4, Box 82
Oneonta, NY 13820
Karst topography is well developed in a narrow band along the Helderberg Escarpment in Schoharie and Albany Counties, New York, where highly soluble Silurian and Devonian carbonate rocks are exposed. This is one of the finest examples of glaciated karst in the country. The purpose of this field trip is to examine some of the major karst features of the region and to show how they have interacted with Wisconsinan glaciation.
Karst features, particularly solutional caves, preserve the record of past groundwater flow patterns and geologic events very clearly -- much more so than surface topography, whose relict features tend to be easily destroyed or buried. Deposits in caves can also be dated radiometrically or paleomagnetically. Paleontological studies are beginning to show the colonization history of certain species, some of which are now extinct. Karst features may therefore aid in unraveling the complex glacial history of the area. Conversely, a knowledge of the glacial history helps to explain certain unusual aspects of the Helderberg karst. Several open-ended questions will be posed on the field trip, and participants are encouraged to contribute to the discussion.
Four sites will be visited on this field trip: from west to east these are (1) the Cobleskill Plateau, (2) Barton Hill, (3) the Knox area, and (4) the Clarksville area. Although in general they share a common geomorphic setting, each has unique characteristics. Further details about the Helderberg karst are provided by Kastning (1975), Baker (1976), Palmer (1976), Mylroie (1977), Mylroie and Palmer (1977), Palmer et al. (1991), and Rubin (199Ia, 1991b).

A6: INTERACTION BETWEEN KARST AND GLACIATION IN THE HELDERBERG PLATEAU, SCHOHARIE AND ALBANY COUNTIES, NY

STOP 1. View of the Cobleskill Plateau
The Cobleskill Plateau contains the largest karst drainage systems in the Northeast, of which the tourist caves Howe Caverns and Secret Caverns are part. A general view of the plateau is seen at Stop 1. The low, broad plateau consists mainly of limestones of the Helderberg Group, overlain in places by younger strata and by glacial till. The beds dip an average of 1.5 degrees to the SSW toward the entrenched valley of Cobleskill Creek, into which most of the karst systems drain. The bedrock floor of the valley is buried beneath as much as 30 m of late Quaternary glacial/alluvial sediment. The creek now follows a route slightly different from its deep-stage pattern, and in places it has migrated laterally onto exposed bedrock that once flanked the valley. Where the creek crosses the Coeymans Limestone (Helderberg Group), much or all of its flow is lost to underground solution conduits and emerges nearly a kilometer downstream. The present course of the creek is postglacial, so the diversion conduits have apparently developed within the past 14,000-15,000 years. The conduits are well adjusted to the present flow pattern, so it is doubtful that they are reactivated older features.

Cross bridge over partially dry bed of Cobleskill Creek.
During low flow the entire creek flows underground for
about a kilometer

STOP 2. Doc Shaul's Spring
Doc Shaul's Spring, at junction with road to right. The spring is on the left, on private property, and is visible from the road.
Many karst springs are located in the exposed limestones at the eastern end of the plateau around Howe Caverns. Farther west, however, the erosionally truncated limestones have been covered by valley sediment. Doc Shaul's Spring, the main outlet for water in the western part of the plateau, rises upward from the limestone subcrop through a conical pit in the overlying sediment.
The original spring appears to have issued directly from the exposed down-dip edge of the limestone prior to aggradation. The depth to the original spring level has not yet been determined. Divers have found the opening to be nearly choked with logs and sediment. Did the valley sediment accumulate slowly enough that it was continually swept away by the up-welling water, so that the spring remained open? Or was the spring inactive during aggradation and reactivated only by water forcing its way upward through the sediment? In McFail's Cave, which feeds the spring, calcite deposits show evidence for re-solution as much as 25 m above the present spring level, indicating prolonged or frequent flooding to that height. Saturated sediment has an effective specific gravity of about 1.1 (accounting for buoyancy and porosity), so this pressure head could have balanced the weight of about 23 m of sediment.

Drumlins and glacially deranged drainage.
In this area the latest movement of glacial ice was west-southwest,
owing to deflection by the Helderberg Escarpment.

Road makes abrupt detour around the nose of a drumlin.
The drumlins in this area are superposed over a north-south preglacial valley in the Helderberg limestones, now almost entirely filled with glacial till.

"Drumlin Sheep Farm" on right.
Continue straight past road to left. (Road to left goes up crest of drumlin, providing a good view of the surrounding terrain.)

STOP 3. Brown's Depression
This huge closed depression, which contains a second-order sinking stream, looks like a large sinkhole. In reality it is an unfilled or exhumed remnant of the north-south bedrock valley filled elsewhere by glacial till. The depression is on private property.
Surface drainage was deranged by glaciation in many parts of the Cobleskill Plateau. Prior to glaciation a prominent north-south stream valley extended through the western part of the plateau, reaching downward through almost the entire Helderberg Group in places. Deflection of glacial ice by the Helderberg Escarpment imposed a local west-southwest movement of glacial ice. As this direction was nearly perpendicular to the valley, the valley was almost completely filled with glacial till. As a final flourish, the ice camouflaged the valley with transverse drumlins that stand well above the surrounding terrain. Surface streams now follow a circuitous route around the drumlins, losing themselves here and there in swamps.
The buried valley, shown in Figure 4, was detected with gravity surveys, refraction seismology, and well logs by Palmer (1976) and with reflection seismology by Mylroie (1977). Its average depth of fill is 60-70 m.
Brown's Depression appears to be an enormous sinkhole but is in fact an unfilled part of the preglacial valley. A second-order stream sinks into the limestone at the western edge of the depression. Either this part of the valley was never filled by glacial sediment or it was later exhumed by sapping through underground conduits. Laminated clays in the bottom of the depression at an altitude of 1100 ft (335 m) are probably deposits from glacial Lake Schoharie, in which case the depression was present during the last glacial retreat (Mylroie, 1977).

Another view of Brown's Depression on the right.

Cross the valley draining into Brown's Depression.

Barrack Zourie
Barrack Zourie, an outlier of Middle Devonian strata, is
visible on the right

Shale pit in the Esopus Formation on right.

STOP 4. Sinkholes and Shafts above McFail's Cave
Pull off road into grassy parking lot on left. Various sinkholes, shafts, and sinking streams can be viewed along the path that leads from the far end of the parking lot. Permission to visit the property must be obtained from the National Speleological Society, which owns the woods in which the karst features are located. The local address is c/o Emily Mobley, Box 10, Schoharie, NY 12157. Please do not cross the field, as it is private cultivated land.
Sinkholes are clustered in areas of the Cobleskill Plateau where groundwater recharge is most abundant. Stop 4 is located in a relatively low part of the plateau where the lower Kalkberg and Coeymans Limestones are exposed and is surrounded by higher areas of relatively impermeable overlying strata. Deep shafts, sinkholes, and sinking streams are well developed here. Their origin is mainly solutional, although some enlargement has taken place by erosion and subsidence of overburden and bedrock blocks. The effect of prominent vertical joints (with strikes of N 22° E and S 70° E) is clearly shown by the linearity of the shafts. These recharge points feed underlying McFail's Cave, which contains about 10 km of accessible passages. The property is owned by the National Speleological Society (see address in Road Log for access permission). The cave is closed to visitors, but similar features can be seen in Howe Caverns, which is open to the public.
The main part of McFail's cave consists of a long sinuous canyon passage that descends from the recharge area in the direction of the stratal dip. It intersects a low-gradient tributary tube that follows roughly along the strike of a low-angle thrust fault of northeasterly dip. It appears that the tube was once the upstream part of the main passage of Howe Caverns to the southeast, but late Pleistocene collapse and sediment fill have blocked the connection. The McFail's Cave water now follows an irregular diversion passage southwestward to Doc Shaul's Spring.
Many of the inactive passages in the cave are blocked by gravel, sand, and clay. Even the active stream passages show evidence of having been nearly filled with sediment at one time. The position and character of the sediments are well out of adjustment with the present flow regime, which suggests that they are related to glaciation. Laminated clays occupy the very lowest levels of stream entrenchment in the main passage. They are found only in caves in the drainage basin of Schoharie Creek (including tributaries Cobleskill Creek and Fox Creek), so they were probably deposited during a phase of glacial Lake Schoharie, when groundwater was rather stagnant. Laminated clays in McFail's Cave range from 970 to 1130 ft (296.344 m) in altitude.
The geomorphic sequence thus includes (1) fluvial entrenchment of the region, exposing the limestones; (2) development of cave passages during a preglacial (or interglacial) time, when the water table stood about 20-25 m higher than today; (3) further fluvial entrenchment, causing diversion of groundwater away from Howe Caverns to Doc Shaul's Spring; (4) filling of large parts of the cave with sediment, probably during glacial advance or retreat, and probably contemporaneous with aggradation of stream valleys; (5) excavation of much of the sediment by cave streams, possibly with some enlargement of passages by glacial meltwater; (6) deposition of lake clays and thin overlying gravels and sands; and (7) postglacial erosion of clays and other sediments. Some of these stages may actually overlap in time or represent sequential phases of the same event. Does the intervention of stage 5 between stages 4 and 6 indicate multiple glaciation, or simply different phases of a single glaciation? The close relationship of the sinkholes and cave to the present topography shows that the basic morphology of the plateau is fairly ancient. The cave is adjusted to the bedrock configuration, including the deep stage of fluvial erosion, whereas the glacial features are simply superimposed.
Despite the great concentration of groundwater recharge in this area, during most of the year the water entering the cave is slightly supersaturated with dissolved carbonate from the soil and the upper part of the bedrock. Degassing of carbon dioxide from the water in the cave reduces the ability of the water to hold dissolved carbonates. It is therefore unable to enlarge the cave by solution at these times, although small amounts of abrasional enlargement may take place. Only during periods of heavy overland flow does solutionally aggressive water enter the sinkholes. Most of the present cave enlargement takes place at these times. During the early stages of cave development, before the cave became aerated, groundwater was able to penetrate through the entire plateau before becoming saturated, since it retained all the carbon dioxide it acquired from the soil. Total hardness of the cave water during low flow is now about 250 mg/liter calcium carbonate equivalent. High-flow measurements show an irregular scattering of lower solute concentrations.

STOP 5. Howe Cave Quarry
Park along the right side of the road and climb the right bank to the quarry overlook. Please be careful! There are good exposures of Manlius and Coeymans Limestones, with a low-angle thrust fault in the far wall. The downstream end of Howe Caverns is located in the residual rock peninsula projecting from the far wall of the quarry.
The Manlius and Coeymans Limestones are exposed in the wall of the Howe Cave Quarry, which was once a source for crushed rock. The Rondout Dolomite beneath the quarry floor was also mined for cement. Please watch your step so you don't slip over the edge --we have not arranged for permission to enter the quarry. The Manlius is very thin bedded, with much differential resistance, whereas the overlying Coeymans is massive and nearly uniform in character; but these differences are almost totally masked in this artificial exposure. Compare this view with that at Thacher Park (Stop 10).
The peninsula of rock projecting into the quarry contains the downstream end of Howe Caverns. This part of the cave is not open to the public. Note the trace of a low-angle thrust fault in the opposite wall of the quarry. Howe Caverns intersects the fault, and the fault's effect on the origin of the cave is described in the following section.

STOP 6. Howe Caverns
Howe Caverns, which is open to the public, contains many of the features described for McFail's Cave. A map of the cave is shown in Figure 9. A 45-meter elevator ride leads to a large tubular conduit that originally formed at or near the water table in the upper part of the Manlius Limestone, nearly parallel to the local strike. A drop in the water table caused the cave-forming water to entrench the floor as a free-surface stream, which is seen today along the tour route. Vadose tributaries enter on the up-dip side. The largest of these, now dry, is the Winding Way, a narrow canyon-like passage formed mainly by the downcutting of a free surface cave stream. Its sinuosity was initiated by small structural irregularities in the limestone and has been smoothed by stream action.
Banks of laminated clay in the main passage are apparently remnants of material deposited at the time of Glacial Lake Schoharie. Their elevation is 867-908 ft (264-277 m) above sea level. The clay is horizontally laminated, rather than parallel to the passage contours, showing that the clay beds are not simply the result of periodic flooding by the cave stream. The clay extends below the deepest level of the cave stream. In places it is interbedded with dark gray organic-rich beds 1-2 mm thick at intervals of 0.5-1.0 cm. It is uncertain whether they represent seasonal deposits, but if so, the entire sequence could represent about 1000-3000 years. Carbonate minerals comprise up to 25% of the sediment (by weight). Some of the clay bodies show remnants of a thin overlying carapace of calcitic flowstone.
The fault exposed in the Howe Cave Quarry cuts through the cave at the "Old Witch," just below the Bridal Altar. The significance of the fault to the cave origin has been debated. No passages actually follow the fault, and there has been no apparent solutional enlargement along it. However, the main passage splits into two levels where the fault intersects it. Apparently the fault was used as the guiding fracture when the cave stream began to entrench its channel and abandoned the original passage level, probably triggered by entrenchment of the Cobleskill Creek valley.

Cross Schoharie Creek.
Deep entrenchment into the Helderberg limestones by this river and its tributaries, Cobleskill and Fox Creeks, has been responsible for most of the karst development in Schoharie County.

Terrace Mountain, on the right
Terrace Mountain, on the right, shows fine exposures of
the gently dipping Manlius and Coeymans Limestones.

Barton Hill rises to the left.
Barton Hill stands in bolder relief than the Cobleskill Plateau. Fox Creek has entrenched entirely through the Helderberg limestones, leaving their down-dip edges exposed far above the valley floor. Virtually all its karst drainage is perched above base level. Many springs issue from the southern (down-dip) side of the hill, as well as along the eastern edge, where King Creek has entrenched partly through the Helderberg Group.
Most of the caves are simple down-dip conduits, which implies that the limestones were entirely exposed before significant cave development took place. As shown in Figure 10, the caves are concentrated in the upper Manlius Limestone. Prominent bedding-plane partings at that horizon were probably the initial pathways for the groundwater that formed them. The underlying limestone is relatively shaly and dolomitic, and so is less favorable for cave development. A few narrow diversion routes for water have formed in the underlying Rondout and Cobleskill at the southern edge of the plateau, but they are merely tight crawlways. They may be post-glacial, following fractures and partings opened by glacial loading and unloading, but since the limestones were exposed long before the Wisconsinsan glaciation, it is possible that they are much older.
Gage Caverns, shown in profile in Figure 10, reaches almost to the northern edge of Barton Hill. Its former catchment area was apparently much larger than at present. It is uncertain whether this is an indication of glacial modification of the drainage basin, or meltwater recharge to the cave, or simply the result of truncation of the headwaters by fluvial erosion.

STOP 7. Schoharie Caverns Spring
Entrance to Schoharie Caverns. Glacial till has partly blocked a perched karst spring. Postglacial solution channels several centimeters deep are visible in the rock face. The cave is not open to the public.
Schoharie Caverns consists mainly of a single high, narrow canyon passage about 600 m long and containing a substantial stream. Divers have recently penetrated through water-filled sections, nearly doubling this length. This is private property whose ownership has recently changed. Contact the authors for information on access policy. The cave is located mainly in the upper part of the Manlius Limestone, and its water emerges high above the Fox Creek valley. Bedrock is exposed in the floor in the upstream part of the cave but is obscured by sediment in the downstream part. Carbonate rocks extend much farther downward, and a nearby smaller spring issues from the base of the Cobleskill Limestone a full 30 m lower.
This site provides a good example, though a rather simple one, of how caves can be used to interpret the sequence of geomorphic events. Schoharie Caverns contains remnants of calcite travertine that has been deeply corroded by solution. U/Th dates for the travertine exceed 350,000 years (the limit of the technique at the time), showing that the deposits are considerably older than Wisconsinan (Stein-Erik Lauritzen, Bergen, Norway, personal communication, 1983). The cave, of course, must be older still. Because the cave is mainly a down-dip vadose canyon, it must have originated after Fox Creek had deepened its valley below the spring elevation. Corrosion of the travertine was probably caused by partial blockage of the spring by late Pleistocene glacial till, which flooded much of the cave with water. Erosion of the till, as well as artificial excavation in the late 1950s, allowed the cave to drain. A culvert was laid in the excavated trench to prevent the cave stream from being ponded by later subsidence of the till.
The limestone face on the western side of the cave entrance contains solutional rills up to a centimeter deep formed by water draining out of the overlying soil. They are clearly postglacial, but their exact age depends on when the face was exposed from beneath its till cover.

STOP 8. Calcite Cemented Glacial Deposits
Along the southeastern edge of Barton Hill is an exposure of glacial outwash cemented with calcite. The calcite was contributed by groundwater seepage from the adjacent limestone plateau. Degassing of carbon dioxide from the emerging groundwater caused supersaturation with respect to calcite.

Karst spring on left.

STOP 9. Limestone Rise (Karren Topography)
Climb onto the Helderberg limestones.
Limestone Rise, a small property owned by the Nature Conservancy, displays good examples of a bare karst "pavement" with solutionally enlarged joints in the Coeymans and Manlius Limestones. The soil is very thin and entirely absent in places. Such fissures constitute what is known as karren, an exposed form of epikarst. It is typical only of glaciated regions, which suggests either that the soil has been partly stripped off by glacial action, or that fissure enlargement has been enhanced by glaciation. Loading and unloading by glacial ice could have widened the joints and made them more susceptible to solution, perhaps by subglacial meltwater, and allowing soil to subside into the enlarging fissures.

STOP 10 Sinkhole Entrance of Knox Cave
Park on right of road, follow path about 1/4 mile to sinkhole entrance of cave
Knox Cave is owned by the Northeastern Cave Conservancy (see information in Road Log). The cave and property can be visited only by permission, which includes signing a liability waiver. Its entrance is a sinkhole caused mainly by the collapse of several high fissures and rooms. Several other sinkholes are located nearby, one of which formed rather recently, as shown by its steep, fresh walls of overburden. The cave is a nearly inactive relic consisting of a network of fissure passages superimposed on a crudely branching system of tubes and canyons. Although the entrance is clearly a collapse sinkhole, into underlying passages, it has undergone much of its enlargement by solution, subsidence, and erosion.
Knox Cave is located only about half a kilometer from the up-dip edge of the Helderberg Plateau. It appears that the present recharge area is not large enough to account for the great size of the cave and its sinkhole, in contrast with the broad area feeding the rather narrow shafts at McFail's Cave. The lack of active streams in Knox Cave (except during heavy runoff) suggests that it may be a survivor of a time when the plateau extended much farther to the north, providing a larger recharge area. On the other hand, runoff from the shaly New Scotland Formation feeds sinkholes farther west in the plateau, and perhaps the cave is now dry simply because of shifting of sink points. An intriguing possibility is that cave enlargement could have been enhanced by glacial meltwater. The network pattern of many of its passages indicates rather severe flooding, which may have required much more water than could have been provided by the existing catchment area.

Thompson's Lake is seen on the way to Stop 11.
The lake occupies a closed basin in the Onondaga with no surface outlet. Much of the water enters -- and all of it exits -- through solution conduits in the limestone. Drainage from the lake resurges at Pitcher Farm Spring about 1.5 km to the southwest. It is probable that the lake originated as much by glacial derangement of surface drainage as by karst processes.

STOP 11. Thacher Park Overlook
This is only a brief stop, mainly for the view. The salient feature of Thacher Park is a 30-m cliff of limestone and dolomite of the Rondout, Manlius, and Coeymans Formations. The Albany lowland lies more than 300 m below. Note the differential weathering that distinguishes the various beds so well, in comparison with the artificial exposure in the Howe Cave Quarry. Several cave springs occur in the Rondout Dolomite at the base of the cliff, perched on the underlying Indian Ladder Shale. Note the gentle fold in the eastern part of the cliff.

Road cuts in limestones of the Onondaga Group.

Dry bed of Onesquethaw Creek on left.
Much of the water in the basin flows underground through the Onondaga.

STOP 12. Clarksville Cave
Park behind June's Place (diner). Please do not usurp the parking places of potential customers. The cave entrance is a collapse sinkhole reached by following the trail northwest from the parking lot for about 1000 ft. The cave drains a large part of the Onondaga bench to the north, seen over the last two miles of the field trip. Enlargement of the cave was probably enhanced by subglacial meltwater, as will be shown at the next stop.
Clarksville Cave is a large cave in the Onondaga Limestone consisting mainly of a single strike-oriented passage developed along a major thrust fault and its subsidiary faults (Figure 12; Rubin, 1991b). Its cross section is larger than that of any other cave in the Onondaga Group. Minor tributaries and diversion passages complicate its pattern. The cave takes overflow from inaccessible conduits to the west that drain the bed of Onesquethaw Creek. The main cave entrance, viewed on this field trip, is a collapse sinkhole over the main passage. The entrance is on private property; anyone who wishes to visit the cave should first check with one of the authors regarding access policy, conservation, and safety.
The cave contains considerable amounts of detrital sediment ranging from silty clay to coarse gravel. The direction and velocity of past water flow are shown by scallops, which are intersecting solutional hollows in the limestone wall. Their asymmetry, like that of current ripple marks, shows the former flow direction. Their length is inversely proportional to the velocity of the water that formed them (see Curl, 1974). Scallops are abundant in nearly all caves formed by rapidly flowing water; but in places in the upstream part of the cave they indicate velocities and discharges much larger than those expected from the presently available catchment area (Rubin, 1991b). There is evidence for enlargement of the cave by glacial meltwaters. Outwash gravels a few hundred meters southeast of the cave give an additional hint of glacial meltwater activity. Multiple passage levels in the cave are not graded to each other, as might be expected if they were formed by a gradually declining water table. They appear to have formed by floodwater enlargement of the cave. Relict infeeder passages show evidence that flow was formerly available from sources that are no longer active. These ideas are clarified at the next stop.

Anticline in the Onondaga limestone.
This structure appears to be merely a shallow flexure above a large
underlying thrust fault.

STOP 13. Subglacial Meltwater Channels
The southeastward view at Stop 13 shows a broad subglacial meltwater channel that once fed Clarksville Cave. The channel today is a broad valley containing no stream and out of adjustment with nearby stream patterns. That it is a subglacial meltwater channel is verified by the presence of upstream bedrock gorges (on posted property on the other side of the road), which have total disregard for the present topography. They average about 3-5 m wide and 2-3 m deep and have a crudely braided pattern. None are active today. They cut across hill tops and lead headward to the very up-dip edge of the Onondaga escarpment. The solubility of the limestone was favorable to the origin of such channels in a short time, whereas the structural integrity of the limestone has preserved them since they were abandoned. Their pattern is compatible only with subglacial meltwater channels.
Several cave openings lie in favorable positions to have received water from these channels, but they also are relict features today. Recent exploration shows that these caves are (or once were) tributary to Clarksville Cave. Small solutional scallops give evidence for considerable discharge in the past. A large amount of meltwater must have entered Clarksville Cave, providing a greater discharge than is available under present conditions and explaining its large size and relict flow features. According to Figure 3, it is unlikely that large caves could have formed in the few thousand years that subglacial meltwater must have been active in the area; but certainly much of the cave enlargement could have taken place during this time. Meltwater is rather low in carbon dioxide content and is rarely capable of dissolving more than a few tens of mg/liter of limestone. However, the meltwater would have come in almost immediate contact with the limestone, allowing almost its entire solutional capacity to be exerted in enlarging caves. A maximum of 0.05 cm/yr of wall retreat would be expected, which translates to as much as one meter of diameter increase per 1000 years.
About 10 km to the east, in Hollyhock Hollow Sanctuary, several bedrock pedestals up to 2.5 m high and 2-2.5 m in diameter show evidence for stream scouring around their entire bases. They are located in a valley tributary to glacial Lake Albany and appear to have been formed by meltwaters in the same manner as the channels in the Clarksville area.

Marcellus-Onondaga Contact
The contact between the Marcellus Formation and the underlying Onondaga Group is clearly visible as the break in slope where the hillside of shale meets the more resistant limestones. Runoff from the hillside sinks in many places on the Onondaga bench and resurges at springs at the base of the limestone to the east.

STOP 14. Sinking Stream at Onesquethaw Cave
Unlike the other caves on the field trip, Onesquethaw Cave is fed by a large perennial sinking stream. The entrance actually receives water only during high flow, as most of the stream sinks several hundred meters upstream and emerges into the cave about midway along its length. The sinking stream is fed by a catchment area of 3.3 km , mainly on shales and sandstones of the Hamilton Group. During floods the cave can fill entirely with water. The flashiness of the runoff has a distinct effect on the character of the cave. Chert beds in the Onondaga have created constrictions in the main passage, around which floodwaters have formed anastomotic diversion routes. The cave is presently owned by Albany County and is closed to visitors. The sinking stream follows what appears to be a postglacial channel partly perched on glacial till. The cave is so well adjusted to the position and size of the stream that it seems that the entire cave could be postglacial. Furthermore, detrital sediments, so abundant in the other caves on the field trip, are very sparse. The other caves clearly pre-date the latest glacial advance. Is Onequethaw an exception?
It is possible, however, that the present sinking stream has simply reoccupied a preglacial route, reactivating a cave that is much older than it would appear. Calcite speleothems have been dated at about 20,000 years b.p. by Stein-Erik Lauritzen (personal communication, 1983). This is an odd date, because it coincides roughly with the latest Wisconsinan maximum. Perhaps conditions beneath the continental ice sheet were not as static as most researchers have assumed.

A7: HISTORY, ECONOMY, AND GEOLOGY OF THE BLUESTONE INDUSTRY IN NEW YORK STATE
James R. Albanese
Department of Earth Science
SUC Oneonta 13820
William M. Kelly
New York State Geological Survey
3136 CEC Albany, NY 12230
ABSTRACT
Upper Devonian sandstone units that occur in southeastern New York, known as bluestone, were a source of dimension stone for the construction industry of southern New York and eastern Pennsylvania long before their descriptions were reported in the geological literature in 1886. Various parts of the Sonyea and West Falls Groups have been quarried. However, the economic significance far outweighed the geological importance of the units, so precise relationships between the minable stone and its geologic unit or environment of deposition were and are largely unknown. The extraction of bluestone, like the exploitation of other low value per volume commodities, has been driven principally by market demand rather than by geologic availability. Descriptions of the properties of the stone are primarily cosmetic rather than petrographic.
The purpose of this trip is to demonstrate the association between the geologic and mining aspects of bluestone quarrying. It will explore the connection between the geology and the mines and examine historic and modern quarrying operations. Participants will learn of the technologies used to locate, identify, and extract currently recognized bluestone resources. In addition, changes in the style of mining and in parameters of the product that could be economically recovered throughout the history of the industry will be illustrated.
Bluestone has been historically quarried in the Hudson Valley, central and southern New York, and northern Pennsylvania and continues to be quarried in some of these areas. Bluestone, or flagstone, is the trade, or commercial name for this type of dimension stone. The rock can be generally classified as a well-cemented, angular, medium-to fine-grained sandstone with its most outstanding characteristic being the bedding. Even-bedded rock that can be split in parallel layers is marketable as bluestone, regardless of its grain size, texture, color, or composition. In general, this bedding characteristic and its resistance to wear and weathering, the result of being composed primarily of quartz, give the stone its economic value. The "blue" of the bluestone name came from the color of the stone quarried in Ulster County during the 19th century, and is no longer an accurate description. The color of bluestone varies from the more common blue and gray colorations to include green

A7: HISTORY, ECONOMY, AND GEOLOGY OF THE BLUESTONE INDUSTRY IN NEW YORK STATE

STOP 1. Good evenly bedded bluestone.
Thick-bedded (up to 20') sandstone layers interbedded with sandy shale containing sandstone rubble and intraformational breccia. Shale is red and weathers rusty. Good evenly bedded bluestone, some cross-bedding, plant fossils. Well-jointed; the main joint set is roughly N-S and the subsidiary set roughly E-W. Outcrop is 0.5 miles long along the south side of the road. More thinly, parallel-bedded and cross-bedded units can be seen on the east end of outcrop.

STOP 2. Rock suitable for bluestone
Rock suitable for bluestone with 1" bedding in 6' thick layers interbedded with shale layers up to a meter in thickness. Channel deposits composed of basal conglomerates, massive and cross-bedded sandstones, grading into parallel bedded layers at the top of each sequence. Some of the shale contains pyrite, weathering to produce the rusty coloration. Toward the east end of outcrop, very thick (30 feet) sandstone, some cross-bedded, overlying a shaley basal conglomerate, and underlying good bluestone layers. Toward the west end, a channel cut into existing cross-beds and overlain by shale overbank deposits.

STOP 3. Indian Country Bluestone Company
Prominent cross-bedding is displayed in the quarry. Note: such well-developed cross-beds are of interest only from the sedimentary structure standpoint, they are a disaster for a quarry. We will see a demonstration of bluestone splitting by old methods (by hand) and stone trimming with the guillotine. Please look over the yard to see the variety of products ready for shipping-pattern stone, tread, veneer, and crushed stone produced from the waste bluestone.

STOP 4. Thick sandstone channel with shale overbank deposits.
Note: This is a dangerous place. Stay off the septum of sandstone on the south side of the road, between old and new Routes 17. The drop on the other side is as great as the height of this outcrop. Be careful on the road because vehicles move past this blind curve at high speeds. Also, stay back from main outcrop face to avoid danger of rock falls.
Thick sandstone channel with shale overbank deposits. Stratigraphic packages consist of a breccia at the base, cross-bedded layers above, changing to planar layers near the top, typical of fluvial deposition. Some of these packages may represent a single flood deposit. The cross-bedding is at such a large scale that the rock may appear to be massive. Lateral accretion surfaces produced by channel migrations are visible on the upper surface of the outcrop on the south side of the road. Breccia layers contain shale pebbles, plant material, and concretions. Storm beds and features of soft sediment deformation are present.

STOP 5. Tompkins Bluestone Company
At this stop we will be the guests of Tompkins Bluestone, one of this area's bluestone distributors. Tompkins uses laser-guided computer-controlled 3' and 10' circular saws to cut stone to whatever dimensions the market demands. They are involved in specialty stone cutting and carving as well as the production of stone for the more traditional applications as treads, flooring, and veneer.

A8: STRATIGRAPHY AND DEPOSITIONAL ENVIRONMENTS OF THE LOWER PART OF THE MARCELLUS FORMATION (MIDDLE DEVONIAN) IN EASTERN NEW YORK STATE
DAVID H. GRIFFING
Department of Geological Sciences
State University of New York at Binghamton
Binghamton, New York 13902-6000
CHARLES A. VER STRAETEN*
New York State Geological Survey
The State Education Department
Albany, New York 12230-0001
(*Present Address:
University of Rochester
Rochester, New York 14627-1001)
The "layer cake" stratigraphy of the Hamilton Group in western and central New York appears to change into a more complex and confusing sequence towards the east, where strata thicken and are increasingly dominated by coarse grained siliciclastics. These trends are exemplified by the lower part of the Marcellus Formation. In western and central New York, this part of the Marcellus Formation is characterized by poorly fossiliferous, black shales with limestone interbeds. These limestones vary in type from intervals of barren or fossiliferous carbonate concretions to complex meter-scale packages of skeletal limestone beds that can be easily correlated over most of this region. Lateral facies changes and the dramatically increased thickness make correlation difficult in strata of the lower part of the Marcellus Formation in easternmost exposures. An approximate 13-fold increase in thickness in eastern New York is accompanied by a transition to more medium-to coarse-grained siliciclastics that overlie the initial black shale deposits. However, the distinctive faunas, sedimentary fabrics, and taphonomic indicators present in some of the strata persist through the facies transition. These strata form useful marker beds that provide a more detailed understanding of the evolution of the northern Appalachian Basin during late Eifelian time.
The purpose of this field trip is to examine the stratigraphic, sedimentologic, and taphonomic changes across the limestone-shale/coarse-grained siliciclastic transition in the lower part of the Marcellus Formation. The following report is a result of two separate ongoing studies. The first is a detailed examination of the limestone-rich facies of the Cherry Valley Member and the upper part of the Union Springs Member in western and central New York state (DHG). The second study involves regional correlation of strata in the lower part of the Marcellus Formation in eastern New York.

A8: STRATIGRAPHY AND DEPOSITIONAL ENVIRONMENTS OF THE LOWER PART OF THE MARCELLUS FORMATION (MIDDLE DEVONIAN) IN EASTERN NEW YORK STATE

Note exposures of Gilboa Fm. in roadcuts north of this intersection.

Note exposures of the Panther Mountain Fm. in small roadcut on the north side of the road.

Note outcrop of lower Hamilton Group shales on W side of the road, N of the junction.

Note roadcut of lower strata of the Marcellus Fm. on the right (south) side of the road.Placemark

STOP 1. Chestnut Street Roadcut
Park at the east end of the roadcut. Note the panorama view of the Mohawk Valley and the southern Adirondacks to the north.
This outcrop features the uppermost part of the Union Springs Member, an excellent exposure of the complete Cherry Valley Member, and the lowermost part of the Chittenango Member. Two thin zones of intensely deformed shale within the Union Springs Member and a thrust fault that ramps up through the Cherry Valley Member can be observed at this locality. Bosworth (1984a, b) associated these deformation features with northwest-directed overthrusting that occurred during either Acadian or Alleghenian orogenic events.
The exposed 4.0 m of Union Springs Member shales contain several intervals of carbonate concretions at this locality. One on these concretion intervals commonly features the goniatite Cabrieroceras pleibeforme and lies approximately 1.0 m below the base of the Cherry Valley Member. The Chestnut street Beds form one 20 cm-thick ledge of skeletal wackestones and packstones around 17 to 20 cm below the base of the Cherry Valley Member. Although the individual Chestnut Street Beds are amalgamated and burrow-mottled, distinction of the uppermost darker gray, crinoid-rich packstone bed can be made on outcrop. A thin, silty, calcareous, gray to black shale interval between the Chestnut Street Beds and the Cherry Valley Member features the goniatite Agoniatites nodiferous. The clean weathered faces of this roadcut make it one of the best places to view internal stratigraphy of the Cherry Valley Member. The 1.2 m-thick section of Cherry Valley contains a 15 to 20 cm-thick lower massive packstone bed, a 60 to 65 cm thick middle nodular packstone/marlstone interval, and a 37 to 40 cm-thick package of upper packstone beds. Concentrations of whole and partially preserved goniatites can be observed in the basal bed and upper packstone intervals. Orthocone nautiloids can also be observed in the base of the upper packstone interval, especially 30 cm below the upper contact of the Cherry Valley Member. Planolites traces are common on the bedding plane surfaces of fallen blocks of the Cherry Valley Member. Calcareous black shales with small brachiopods, dacryoconarids, and small orthocone nautiloids mark the transition from the Cherry Valley Member into the overlying sooty black shales of the Chittenango Member.

Note outcrops of Kalkberg Fm. along U.S. Rte. 20 west of Sharon Springs.

Note outcrops of Onondaga Limestone.

Union Springs Mbr. of the Marcellus Fm.
Note exposures of the Union Springs Mbr. of the Marcellus Fm. and the underlying Onondaga Limestone in the drainage ditch on the right (west) side of the road for the next 0.4 miles.

STOP 2. Mineral Springs Section.
This stop exposes the upper 4.5 meters of the Union Springs Member, a complete section of the Cherry Valley Member (1.4 m-thick), and the lowest part of the Chittenango/Berne Member. Calcareous dark gray shales a meter above the base of the section contain small brachiopods and bivalves; these include the brachiopod Leiorhynchus and the large bivalve Panenka, along with nautiloid and goniatite cephalopods. A 2 m-thick interval of tectonized shale occurs in the upper part of the Union Springs Member above these fossiliferous shales. The Chestnut Street Beds form a 20 cm-thick package of two skeletal wacke/packstone beds separated by gray calcareous shales. The basal Chestnut Street bed directly overlies a 5 cm-thick dark, argillaceous concretionary lime mudstone layer with pyritized burrows. The internal stratigraphy of the Cherry Valley Member at this locality is similar in most ways to that of STOP 1. However, individual packstone beds of the upper massive interval increase in thickness and are separated by shaly partings.

Chittenango Mbr. of the Marcellus Fm.
Note hillside exposure of the Chittenango Mbr. of the Marcellus Fm. on the south side of the road.

Note Carlisle Center and Esopus Fms. in roadcuts.

Note Oriskany and Becraft Fms. in roadcuts south of I-88.

Note Becraft and Kalkberg Fms. in roadcuts south of I-88.

Note Coeymans Fm. in low roadcuts on the south side of the highway.

Manlius, Rondout, and Brayman Fms.
Note exposures of the Manlius, Rondout, and Brayman Fms. in large roadcut to the south of I-88.

Note exposures of Onondaga Limestone in roadcuts and in adjacent Fox Creek for the next 1.5 miles.

Note Onondaga Limestone exposures behind Highway Dept. Building (south of Rte. 443).

Berne and Otsego Mbrs of Mount Marion Fm.
Note that large roadcuts on Cole Hill Rd. feature the type-Berne and the complete Otsego Mbrs. of the Mount Marion Fm. (= proximal marine facies of the upper part of the Marcellus Fm.).

Note small quarry in the Berne Mbr. (south of the road).

STOP 3. Section at Long Road ravine (Thompson's Lake section of Rickard, 1952)
This section illustrates the eastward appearance of coarse-grained siliciclastics in the lower part of the Marcellus Formation. The stop exposes a relatively thick section of the Bakoven Member, a 4.5 m-thick section of the stony Hollow Member, and a 1.4 m-thick section of the Cherry Valley Member.
The Bakoven Member (= Union Springs Member) is exposed downstream from and within the lowermost waterfall. The shales contain carbonate concretion horizons and thin-bedded limestones. Rickard (1952) recognized the true concretionary Cabrieroceras bed near the base of the lowermost waterfall. A 1.75 m-thick interval of tectonized shale occurs in the face of this waterfall.
A 2.0 cm-thick bed of the Lower Proetid Unit occurs at the base of the Stony Hollow Member, just below the crest of the lowermost waterfall. The overlying 1.8 meters of buff-weathering, pyrite-rich, calcareous siltstones and shales are equivalent to the massive sandstone unit near the top of the Stony Hollow in the Hudson Valley (STOPS 4 and 5b). The overlying Upper Proetid Unit consists of a package of coarsely crinoidal packstone and grainstone beds that display relict cross-stratification. Quartz silt and sand are common in these limestone beds at this locality, in contrast to previous stops.
The Cherry Valley Member is exposed within and caps the uppermost waterfall. The lower and upper Cherry Valley intervals are heterolithic at this locality and contain thin interbeds of terrigenous siltstone/sandstone with skeletal packstone and grainstone beds. The middle nodular interval consists of a few large, sharply bounded concretions in dark gray shale. Coarse crinoidal beds with finely comminuted brachiopod valves are common in the uppermost Cherry Valley.

Note sandstone at the top of the Berne Mbr. (Mount Marion Fm.) on right past the intersection.

Note Helderberg Lake on left (north) side.

Note exposures of the Berne Mbr. behind the house at the corner.

Exposures of the lower part of the Marcellus Fm
Note the large parking area on the north side of the road, which provides parking for access to exposures of the lower part of the Marcellus Fm. in and below the falls on Onesquethaw Creek, directly to the north.

Note exposure of Onondaga Limestone in roadcut on north side.

Note quarry in the Otsego Mbr.

Upper part of the Mount Marion Fm.
Note outcrops of the upper part of the Mount Marion Fm. on the south side of the intersection

Note quarry in the upper part of the Mount Marion Fm. on the right (west) side of the road.

Note outcrops of the Plattekill Fm. along the road for the next 0.4 mile.

Note scenic views of the northern Catskills for the next 0.5 mile.

Note exposures of Plattekill Fm. red beds and fluvial sandstones in roadcuts.

Note additional outcrops of the Plattekill Fm. for the next 2.4 miles.

Note outcrops of the Mount Marion Fm. over the next 0.4 mile.

Note scenic view of Potic Mtn. to the left (west).

Note outcrops of the Onondaga Limestone in roadcuts for the next 0.2 mile.

STOP 4. BUTTERMILK FALLS
This locality exposes approximately 41 m of the Stony Hollow Member and 3.7 m of the terrigenous sand-rich facies of the Cherry Valley Member. Strata at the base of the lowest falls lie approximately 12 m above the Bakoven/Stony Hollow Member contact, which is exposed in a ravine 2.1 km to the north (Locality 8).
A 7 m-thick interval in the lowest falls features a series of thin dolomitic beds (ca. 2-22 cm-thick). This interval, which is best seen in the high bank north of the lowest falls, correlates with a similar package of thin dolomitic beds in the lower part of the Stony Hollow Member at Kingston (Locality 13/Stop 5b). Fine-grained resistant sandstones forms the caprock of the lowest falls.
A second falls 4 m upsection is capped by 1.6 m of highly bioturbated, argillaceous sandstone. The lower part of the unit appears medium bedded, and the upper part is massive. Prominent subhorizontal to vertical Zoophycus traces occur in the upper 0.9 m. This unit is also traceable to Kingston (Locality 13, STOP 5b), where it is represented by 0.9 m of fossiliferous, bioturbated fine sandstone.
The main upper waterfall at Buttermilk Falls is capped by the massive sandstone (3.3 m-thick) near the top of the Stony Hollow Member. At this locality, the massive sandstone unit is highly bioturbated and displays a prominent slaty cleavage. A 0.3 m-thick recessed interval of sandy shale separates the massive sandstone from the underlying Lower Proetid Unit. This highly fossiliferous unit is composed of 2-3 separate beds that total 0.6 m in thickness at Buttermilk Falls.
The massive sandstone that caps the main falls at Buttermilk Falls is overlain by siltstones and silty shales in the upper part of the Stony Hollow Member. The Upper Proetid Unit is covered at STOP 4. A 0.6 m-thick sandstone that forms a small cascade upstream of the main falls may correlate with a similar thin unit that overlies the Upper Proetid Unit or the Chestnut Street Beds at most localities.
The Cherry Valley Member forms a 2 m-high falls at the top of the outcrop. The argillaceous sandstone of the Member is highly cleaved and has several recessed carbonate-rich intervals with numerous small brachiopods. The sandstones appear to be highly bioturbated. A small outcrop of sandy shales that is exposed approximately 5 m upstream is apparently part of the Cherry Valley Member. An abandoned quarry in the south bank of the creek exposes the lowermost black shales of the overlying Berne Member.

Note outcrops of the Lower Devonian New Scotland and Becraft Fms. for the next 1.2 miles.

Cross Catskill Creek.

Note outcrops of the Tristates Group
Note outcrops of the Tristates Group (Esopus, Carlisle Center, and Schoharie Fms.) and the Onondaga Fro. for the next 8.9 miles

Cross Kaaterskill Creek.

Glenerie Fm and Helderberg Group
Note exposures of the Glenerie Fm. and the underlying upper part of the Helderberg Group (Becraft, Alsen, and Port Ewen Fms.) for the next 0.6 miles.

Note exposures of the Schoharie and Onondaga Fms. for the next 1.1 miles.

Mount Marion visible to the right of the Thruway.
Note large quarry at the north end of the ridge which exposes the Berne and Otsego Mbrs. of the Mount Marion Fm.

Note exposure of Stony Hollow Mbr. on left (east) side of highway.

Note dark to rusty shales of the Berne Mbr. exposed on right (west) side of highway.

Note outcrops of Stony Hollow Mbr. for the next 0.7 miles.

STOP 5A. City View Terrace Roadcut
This outcrop exposes the upper 10 m of the Bakoven Member and the lower 23 m of the Stony Hollow Member. The transition between the members is gradational at this locality, and it is underlain by a 1-3 m-thick zone of tectonized shale. Calcareous nodules within this interval have yielded Cabrieroceras and another unidentified goniatite (G. Kloc, 1991, personal commun.). The lower strata of the Stony Hollow Member at this locality are thinly laminated and exhibit little to no burrow-mottling (chiefly small Planolites traces). Body fossils are rare and are chiefly small tentaculids (Nowakia?) and rare small brachiopods and bivalves. The upper part of this outcrop is exposed along the adjacent road to the northwest, where thin (ca. 3-10 cm) dolomitic beds appear in the section.

STOP 5B. Large Route 28 Roadcut
This large roadcut exposes approximately 60 m of the Stony Hollow Member. The base of the outcrop lies approximately 11 m above the contact with the underlying Bakoven Member (see STOP 5a). The thinly laminated character of the unit is visible on many of the rock surfaces, and the low degree of burrow-mottling (chiefly small Chondrites burrows) is also notable. These features characterize much of the Stony Hollow Member at Kingston. A series of thin (ca. 3-17 cm-thick), buff-weathering dolomitic beds occurs 11-21 m above the base of the outcrop. Two thicker, buff-weathering units are visible farther upsection (27 and 33 m above the base of the outcrop, respectively). The lower bed (ca. 0.8 m-thick) is underlain by a layer of rusty-weathering burrows and concretions. The upper unit (ca. 0.9 m-thick) features an upper bioturbated sandstone (with Planolites and Teichichnus traces and small rugose corals, and a lower fossiliferous unit featuring hyoliths, coleolid tubes, gastropods, and cephalopods. Small bivalves may also be found in this interval. This unit is correlatable with the lower massive sandstone at Buttermilk Falls (Locality 9, Stop 4).
The next key unit is a thick, massive, light buff-weathering unit (ca. 5-6 m-thick) that is traceable northward along the outcrop belt as far as Catskill. Auloporid corals are scattered throughout the unit, and two horizons feature small streptelasmid rugose corals. The overlying interval (ca. 12 m-thick) is characterized by darker, blocky-weathering, calcareous shale-rich strata.
The top of the outcrop is marked by the resistant, thin-bedded to massive sandstone near the top of the Stony Hollow Member. The base of this interval is marked by the Lower Proetid Unit and an underlying brachiopod-rich bed. The upper surface of the sandstone marks the base of the Upper Proetid Unit at outcrops in the Kingston area.
The overlying transition into the Cherry Valley Member is not well exposed along Route 28, but a complete section through this interval may be seen 1.7 km to the east along Route 209. At the latter outcrop, the top of the thin-bedded to massive sandstone unit is overlain by a 3.5 m-thick interval of silty shales which feature a series of thin, richly fossiliferous beds with proetid trilobites. These beds may have small calcareous nodules or thin limestone lenses. This interval represents the Upper Proetid Unit in the Kingston area. An overlying thin, calcareous sandstone bed contains numerous brachiopods, which include Productella, Kayserella, and Leptaena. The brachiopod bed is capped by a prominent, 60 cm-thick, bioturbated sandstone that lies approximately 5 m below the top of the Route 209 outcrop. The two lower beds of the Cherry Valley Member are present in the upper 2.7 m of the Route 209 cut.

STOP 5C. Cherry Valley Member Along Route 28 And Adjacent Railroad Cut
This roadcut and the adjacent railway cut expose a 10 m-thick section of the terrigenous sand-rich facies of the Cherry Valley Member. At this locality, the member is composed of five highly bioturbated arenaceous units separated by sandy shale intervals. These units range from 0.7-1.4 m.thick. The sandstones display prominent cleavage, especially in the railroad cut. (Note that the upper sandstone is missing in the highway cut and the lower sandstone is exposed downsection from the railway cut.) Body fossils are relatively uncommon in the sandstones, though carbonate-rich nodules feature shelly to styliolinid-rich hashes with abundant quartz sand. Leiorhyncus, other small brachiopods, and auloporid corals may be present near the tops or bases of the sandstone units. Uncommon to rare nautiloid and goniatite cephalopods may also be found. Teichichnus is the most prominent trace in the sandstones. The intervening sandy shales feature numerous limonitic burrows, accompanied by small brachiopods and bivalves.

A9: A SURVEY OF PRECAMBRIAN TO PLEISTOCENE GENERAL GEOLOGY IN CENTRAL NEW YORK
H. S. Muskatt
Utica College of Syracuse University
and
David E. Jones
The Ecosystems Center, Marine Biology Lab
Woods Hole, Ma.
The intent of this field trip is to introduce several aspects of geology covered in the New York State earth science syllabus. We shall examine rocks that range in age from Precambrian to Pleistocene. Many rock types from a wide variety of Paleozoic depositional environments will be seen. Opportunities to collect a wide range of fossils and minerals will be provided. The field trip will develop a tectonic framework of deposition as we move higher into the stratigraphic record. Sedimentary structures and erosional features are well represented at these stops. There are several good photo opportunities of faults and potholes in addition to the rocks. We will also be seeing a wide variety of landscape features produced during the last glacial retreat, including lakes, U shaped valleys, kettle ponds, meltwater channels, sinkholes, and drumlins. An excellent general field trip of the Utica area to the west by Tewksbury and Allers (1984) is highly recommended.

A9: A SURVEY OF PRECAMBRIAN TO PLEISTOCENE GENERAL GEOLOGY IN CENTRAL NEW YORK

Good example of glacial "U" shaped valley erosion.
You will be driving up the Otego Creek Valley, a very good example of glacial "U" shaped valley erosion.

Canadarago Lake to the right.
This lake, a fingerlet, was formed at the end of the last ice age by the blockage of glacial meltwaters downstream which formed a natural dam. Fingerlet: Muskatt's term for a smaller version of a Finger Lake.

Utica Shale and meltwater channel
Outcrop of Utica Shale (Ordovician). To the east, a good meltwater channel can be viewed.

An outcrop of Trenton Limestone (Ordovician) containing trilobites is on the east side of the road.

Precambrian basement material can be seen in an outcrop on route 5 to the east and up the hill.

STOP 1. Potholes and Moss Island
A very good outcrop of Precambrian material can be seen on the NE. side of the road just before reaching the plant. This rock is a quartzose syenite gneiss of Precambrian age. Multiple veins of quartz that have filled fractures in this material attest to a history of deformation in these rocks. Please do not use rock hammers at this outcrop, as the use of them would mar the large pothole, a very unusual erosional feature that is visible here. For some periods of the Pleistocene ice blocked the flow of water through the St. Lawrence River creating Lake Iroquois. This greatly increased the flow of water through the Mohawk R., for some 1000 yrs., from about 12,000 to 11,000 yrs ago, which carved this pothole and those across the river. Across the river is Moss Island. In May, 1976, Moss Is. was declared a National Landmark. It is one of only 400 such registered natural treasures in the U.S. The shoreline of Moss Is. displays several of these potholes, some more than 30 feet deep and about 20 feet in diameter, and makes an excellent photo opportunity. Lock 17, which is the highest lift lock on the barge canal system (40.5 ft.), can also be seen.

STOP 1A. Roche Moutonee in park
In the town park ahead is a good exposure of bedrock that displays roche moutonee glacial alteration.

STOP 1B. Photo opportunity of a well scoured boulder
555 E. John St. Boulder on display in front of the Rockton Plaza was removed from a pothole during excavation.
Photo opportunity of a well scoured boulder, an abrasion tool used by the swirling waters for the creation of potholes.

Moss Island can be viewed on the left (East).

STOP 1C. Parking Lot for church on right, W. of the church.
Walk to fence in the back center of lot. Across abandoned railroad right of way observe nonconformal contact of Little Falls Dolostone (Cambrian) overlying Precambrian syenite gneiss. Dolostone is overlain by classic man made limestone wall.

STOP 2. Outcrop of Late Cambrian Little Falls Dolostone.
This unit directly underlies the Black River Group (Ordovician) in this area. It nonconformably overlies a Precambrian syenite gneiss. The Little Falls Dolostone is a sandy, medium-grained dolostone with some sandstones occurring near the base of the unit. Except for the abundant colonial algae Cryptozoan, not seen at this site, the unit seems to be barren in this area. Quartz crystals from this site are locally known as Little Falls, Herkimer, or Middleville Diamonds. Anthraxolite is a black, lustrous, carbonaceous mineral often found in association with the quartz. Calcite, dolomite, pyrite, marcasite, galena, sphalerite, chalcopyrite, hematite, and glauconite have also been reported.

STOP 3. Paradise Road bridge over Thruway.
Park at S end. Ordovician Dolgeville Facies, a mixture of calcisiltites and black shales, overlain by the Ordovician Utica Shale can be seen in rock cuts on both sides of the Thruway. It is a deeper water facies of the Trenton limestones. The contortions seen in the Dolgeville are considered to be due to a submarine slide when the unit was in a soft sediment phase. Graptolites and trilobites from the Utica Shale exposed in the ditch at the SW corner of the bridge have been found. WARNING: Do not climb down the cliff faces. It is illegal and dangerous to go rock collecting there. Looking at the NE wall a very visible fault can be seen. It the first of many major faults that cross the Thruway further to the East. (This stop is a good camera opportunity).

STOP 4. (LUNCH) North of the road is an abandoned quarry.
This is now part of a private home, so please do not do damage to these rocks. 2-1/2 feet of Gull River Limestone underlie 18 feet of Kings Falls Limestone. These rocks are very fossiliferous, shelly and nonshelly calcarenites from the Ordovician Trenton Group. See stop #2, Cameron (1972) for more detail.

STOP 5. Outcrop of Silurian Otsquago Sandstone
Outcrop of Silurian Otsquago Sandstone, a formation of the Clinton Group. Coarse red sands, cross bedding well displayed. Interbedded green shales represent the Sauquoit Formation seen further to the west. A deltaic distributary channel environment of deposition is postulated for these rocks.

STOP 6.Intersection of routes 20 and 166.
Drive under the Route 166 bridge and park along the outcrop on Route 20. Kalkberg Limestone, Oriskany Sandstone, Esopus shale, Carlisle Center and Onondaga Formations. A 1-3 cm (Tioga) bentonite (volcanic ash) layer is found in the Kalkberg. It has been dated as being about 395 million years old. Several rock types are visible at this outcrop that represent a wide variety of Devonian depositional environments. See Fisher (1979), stop #6 for more details.

Glimmerglass State Park on right.

STOP 7. Road Cut in Dugway Rd. Otsego shales (Middle Devonian).
Epifaunal filter feeders were very common in this deltaic environment. The rise of highlands in the east (due to tectonic uplift) have altered the environment of deposition to something quite different from those seen at Cherry Valley. For more information, see Grasso ( 1977).

Start of the Susquehanna River.
A small park at the junction with Lake Street marks the start of the Susquehanna River. Here the waters of Otsego Lake flow over a Quaternary moraine that dams the lake.

A10: HYDROGEOLOGY OF GLACIAL DRIFT IN THROUGH VALLEYS NEAR DRYDEN AND CORTLAND, NEW YORK
TODD S. MILLER AND ALLAN D. RANDALL
U.S. Geological Survey
The Appalachian Plateau physiographic province encompasses nearly all of southwestern New York (fig. 1, inset). About 85 percent of this region consists of till-mantled bedrock uplands; the remainder consists of broad valleys that are partly filled with stratified drift. The stratified drift includes coarse-grained sand and gravel deposits that are the only highly productive aquifers in the region; it also includes extensive fine-grained sediments that do not yield usable amounts of water. The divide between streams that drain northward to Lake Ontario and those that drain southward to the Susquehanna River crosses some of these broad valleys, which are termed "through valleys" because they are continuous across a major topo.graphic divide. An area of unusually hummocky topography near the divide in each through valley suggests the former presence of buried ice; these areas were collectively termed the Valley Heads moraine by Fairchild (1932).
This article briefly describes current concepts of (1) the stratigraphy of stratified-drift aquifers and confining units in valleys within and south of the Valley Heads moraine, (2) recharge to those aquifers, and (3) their potential for use as sources of large water supplies during dry periods. These concepts were developed over the past few years from geohydrologic studies, cited herein, that emphasized or included valleys near Dryden, Harford, and Cortland (fig. 1). The article concludes with the log of a field trip that is intended to illustrate several of these concepts and to offer an opportunity for discussion of what is and is not known about these aquifer systems.

A10: HYDROGEOLOGY OF GLACIAL DRIFT IN THROUGH VALLEYS NEAR DRYDEN AND CORTLAND, NEW YORK

STOP 1: Site Of Former USGS Gaging Station On Gridley Creek.
This site is near a preglacial drainage divide now crossed by Gridley Creek, whose flow was reversed by glaciation. It is a good site at which to measure basin yield because the valley fill is only 500 feet wide and probably not thick; therefore, ground-water underflow through permeable sand and gravel is likely to be small, and nearly all runoff leaves the watershed in the stream, where it can be measured.

Ski area on left.

Meltwater spillway ahead on left
Meltwater spillway ahead on left, incised into south end of mound of drift that Route 392 crosses just ahead. The divide between east-flowing Gridley Creek and west-flowing Virgil Creek is here, in the middle of a valley 2,000 feet wide.

Summit, view (to rear) across valley of Virgil and Gridley Creeks.

Route follows East Branch Owego Creek. Kame terraces on right side of valley.

Kame terraces on both sides at intersection.

STOP 2: Hesitation Stop: Harford Underground Gas-Storage Project.
In 1953, a cavity was excavated by solution mining in salt beds at a depth of about 3,000 feet to provide storage space for liquified petroleum gas that would arrive here by pipeline. The brine created by mining was discharged to lagoons excavated in the permeable outwash and alluvium on the valley floor. A few years later, many residents of Harford Mills, visible to the south, began to pump salty water from their shallow wells. Water sampling, resistivity surveys, and a test well disclosed that outwash extended to a depth of 56 feet and contained salty water over a large area near the hamlet and Owego Creek, except close to the water table, where water was fresh (Dunn, 1967, 1968). Gravel layers at depths of 97 and 170 to 194 feet yielded fresh water.
The raised, lined pond visible here is now used to contain brine produced during operation of this facility.

STOP 3: Cheese Factory Brook.
Except during the spring freshet and other brief periods of unusually heavy runoff, this brook ceases to flow somewhere near or downstream from this former creamery. We hope to view the point of dryness, note exposures along the brook, and discuss the significance of seepage losses from tributaries as a source of recharge.

STOP 4: Springs Along Cheese Factory Brook.
About 0.2 mile upstream from the creamery, a reservoir (now in disrepair) was built to develop a spring at the base of the bluff along Cheese Factory Brook. Exposures upstream and downstream show that the bluff consists of sandy, somewhat silty gravel. The presence of the spring suggests that impermeable sediment is not far below. On April 25, 1991, a tributary just upstream went dry 300 feet before it reached Cheese Factory Brook. Seepage from the tributary may be a major source of water to the spring.

STOP 5: Streamflow And Water-Level Measurement Sites.
This is the location at which the water levels in figure 7 were measured. Most runoff from the 3.42 square-mile watershed above this point occurs as underflow through the stratified drift.

STOP 6: Pit In Crevasse Filling.
Locked gate ahead, permission from Harford Teaching and Research Center farm manager required to enter; continue 0.1 mile to pit. Brook to left of dirt road flowed past Route 38 April 25,1991.
This infrequently used pit is on the property of the State College of Agriculture's Harford Teaching and Research Center. About 10 feet of till that contains only rounded pebbles overlies about 20 feet of stratified, sorted but very silty gravel and, in places, deformed medium to fine sand, silt, and clay. Poorly sorted, silty gravel is characteristic of early, deglacial, proximal stratified drift near the Valley Heads moraine.

Harford Teaching and Research Center main buildings on left.

STOP 7: Daisy Hollow Brook.
Daisy Hollow brook flows farther out onto the valley-floor outwash than any other tributary near Harford. One reason why may be observed here. Also exposed is the flat-stone drab gravel that is typical of alluvium along upland tributaries.

This terrace like flat is capped by till.

STOP 8: Southworth Road Exposure Near Dryden.
The large variety of sediments in this exposure illustrate the complexity of the depositional environment at the Valley Heads moraine. Figure 4 shows the appearance of the exposure in 1984. Sediment facies include till, fluvial, and lacustrine deposits. The drab till at the top of the exposure represents a readvance of ice during the late stages of Valley Heads glaciation. The surficial till overlies a coarse, cobbly gravel that in turn overlies either a discontinuous pebbly sand or till hummocks. Till hummocks have pockets of gravel, disturbed sand stringers, and some folded bedding, all of which suggest movement of the sediments such as may occur during topographic inversion. Beneath the hummocks is rhythmic silt/clay with scattered stones, presumably dropped from floating ice. A till containing abundant bright clasts underlies the rhythmite. A bright, coarse, cobbly gravel underlies the bright till.

STOP 9: Gravel Pit At South Cortland
Coarse outwash and ice-contact sediments deposited by Valley Heads ice. Note boulders 2 to 3 feet in diameter at entrance to pit. Excavation is along the boundary between outwash deposited by northeastward-flowing meltwaters and ice-contact deposits laid down on and adjacent to blocks of ice that were in the middle of the valley. A geologic log of a test well in the gravel pit indicates the following stratigraphy, in feet below land surface: 0-45 feet sand and gravel, 45-50 till, and 50-75 sand and gravel. Mining of sand and gravel will stop 10 feet above the water table so as to minimize the effects on water resources and so that land could be used after mining operations cease.

STOP 10: Hesitation Stop: Remediation Of A Spill Of Organic Solvent.
A recovery well pumps about 970 gallons per minute to remediate the source area (O'Brien & Gere, 1991). Pumped water is routed through an air stripper, which brings sufficient air into contact with the water to allow the volatile organic solvent to evaporate. The water is then discharged into infiltration lagoons, where it seeps to the water table.

STOP 11: City Of Cortland Well Field.
Otter Creek flows through the well field but is usually dry from midsummer to late fall. During low-flow periods, when well 3 is turned on and streamflow disappears in Otter Creek because of induced infiltration. When the well is turned off, streamflow reappears. Abundant vegetation in the channel indicates a relatively large amount of nutrients in the water. In contrast, the channel of nearby Dry Creek contains little or no vegetation, which suggests that runoff from that drainage basin contains relatively few nutrients.
An air-stripping tower was installed beside the main plant in anticipation that the plume of organic solvents (source about 2 miles southwest, stop 10) would reach the well field. Concentrations of trichlorethylene within the plume decrease with distance from the source, and only trace amounts have been detected at the well field, not enough to warrant use of the air stripper. The decrease in concentrations has been attributed to volatilization, discharge to Otter Creek, and biological degradation and transformation of the trichlorethylene.
Well 4 at plant 2 is one of the most productive municipal wells in New York. It is capable of pumping 4,000 gallons per minute. The well taps outwash 63 feet in saturated thickness. At present, it is used only intermittently.
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B1: FIELD ILLUSTRATIONS OF ROCK TYPES AND GEOLOGIC FEATURES IN THE UPPER SUSQUEHANNA VALLEY ADJACENT MOHAWK REGION
by
David M. Hutchison
Hartwick College
The outcrops and surficial features observed on this trip have been selected (1) to show students the change in lower Paleozoic stratigraphy through time (2) to illustrate several sedimentary and topographic features (3) to help students gain a better understanding of the geologic framework of this area.
The trip starts at the large Upper Devonian flood plain channel behind the F.W.Miller Science Building on the Hartwick College campus in Oneonta and ends in Precambrian garnet gneiss six miles east of Canajoharie. Progressively older beds are exposed to the north because of three factors: the gentle southerly dip of the beds, the erosion by the Mohawk River and the uplift, tilting and erosion of large fault blocks associated with normal faults in the Mohawk River Valley.
The clastic sediments are the result of the Middle Ordovician Taconian Orogeny (470-435 m.y. ago) and the Middle and Late Devonian Acadian Orogeny (J85-355 m.y. ago) (Fisher. 1965). Both of these times of crustal unrest and uplift east of the Hudson River and in New England provided an influx of clays, silts and sands into the Ordovician and Devonian seas which occupied the area of this field trip. The carbonate rocks were deposited by these seas during periods of quiescence between orogenies.

B1: FIELD ILLUSTRATIONS OF ROCK TYPES AND GEOLOGIC FEATURES IN THE UPPER SUSQUEHANNA VALLEY ADJACENT MOHAWK REGION

STOP 1. Upper Devonian Oneonta Formation
STOP 1. Upper Formation Devonian Oneonta
This outcrop consists of dark grey shales, thinly bedded siltstones and sandstones. A massive sandstone which fills in part of a Devonian stream channel is well exposed for about 150 feet at the left end of the outcrop. There is an erosion surface between the underlying shale and the overlying massive sandstone. This surface which marks the bottom of the channel rises stratigraphically to the right. Interference ripple marks are exposed just below the road level. Plant materials are abundant in some layers. A few galena crystals about 1 mm across have been found in the thinly bedded sandstone.
Trilobites, marine pelecypods and brachiopods of the Gilboa Formation have been found in beds stratigraphically 75 feet below this outcrop, and marine fossils are found in thinly bedded siltstones and shales stratigraphically 200 feet below stop 1. This outcrop is next to Nick's Diner on Chestnut Street (Route 23) . We probably will not stop here today.
A clean sandstone which shows extensive trough-type cross bedding crops out in the woods between Stop 1 and Stop 1A.

STOP 1A. Red mudstones and shales
Red mudstones and shales with root casts, worm burrows and ripple marks are exposed at the top of the hill just below the Ernest B. Wright Observatory which houses a 16" telescope. Walk or drive (if the sliding chain link fence gate is open) up the hill at the far end (west) of the Binder Physical Education Building parking lot. Continue up the hill through the second gate to the outcrop below the Observatory next to the practice soccer field.
The view down the valley is looking southeast to Mt. Utsayantha in Stamford (about 30 miles map distance). The broad U-shaped valley is the result of Pleistocene glaciation. The Susquehanna River flows in from the north (out of your view) and through the city of Oneonta. The terraces on either side of the valley are kame terraces. Red beds are exposed near the top of the hills across the valley.

Small kettle hole on right.
This has been partially filled in to make a parking lot for Pyramid Mall.

View to right showing kame terraces across valley.

On right across valley gravel pit in delta kame.

STOP 2. Stop directly across from dam at Goodyear Lake.
The Upper Devonian Gilboa Formation with horizontal sandstones and siltstones is exposed here. The well developed flow-rolls are of interest. It is apparent that these are primary structures (formed while the sediment was still "soft") rather than secondary structures (formed after lithification), but there is some question as to their origin. There are brachiopods, bryozoans, crinoids and pelecypods.

Outcrop of Middle Devonian Cooperstown Shale
Outcrop of Middle Devonian Cooperstown Shale at curve in road. Good view ahead of broad U-shaped valley. The Susquehanna River flows south from its headwaters in Otsego Lake (Cooperstown) in this valley.

Cross Susquehanna River.
Clays indicate a glacial lake occupied this area. If you want to study the clays you should go past junction with County Road 35 and park in small area to the right along Route 166.

Meander scar of Cherry Valley Creek is visible to left.

Hummocky morainic deposits
After crossing the creek NOTE THE EXCELLENT hummocky morainic deposits showing numerous kettle holes and knob and kettle topography.

Cherry Valley.
The town was settled about 1740. On November 11, 1778 over 40 people were killed by Tories and Indians in the infamous Cherry Valley Massacre.

STOP 3. Highway Department gravel and salt storage area.
Topographically this spot is a break in the Helderberg escarpment. Fairchild interpreted this as a glacial spillway formed at a time when ice filled most of the Mohawk Valley. Meltwater was blocked from draining north and "spilled over" the escarpment eroding the valley (Fairchild, 1925). Fleischer (personal communication and elsewhere in this Guidebook) feels that the valley is the result of glacial scouring of a through valley.
Three formations are exposed here. The lowest formation (exposed about .1 mile down the road) is the Esopus Shale. This is overlain by the Carlisle Center calcareous siltstone which contains numerous worm burrows Toanurus cauda-galli (Rickard and Zenger, 1964). About .l mile south along the road the Carlisle Center siltstone is overlain by the Middle Devonian Onondaga Limestone which contains abundant crinoids, corals and brachiopods.
The upper part of the Onondaga contains abundant chert. Jointing is very obvious at this outcrop.

STOP 4. The Lower Devonian Manlius Fm
The Lower Devonian Manlius Formation (lower Thacher Member) crops out on the right.
This laminated micrite contains some ostracods, tentaculitids and stromatoporids. Some mud crack marks are present. This is the intertidal facies of LaPorte, 1967. The Coeymans Formation rests on top of the Manlius Formation.

STOP 5. Three Middle Devonian formations exposed here.
The lower most formation is the Union Springs Shale which is a black fissile shale containing calcareous concretions. There is a thin limestone near the top of the shale. This is overlain by the Cherry Valley Limestone which is about 7 feet thick and contains a cephalopod fauna. More than 100 feet of the jet-black Fissile Chittenango Shale rests on top of the Cherry Valley Limestone. At the east end (left) of the outcrop the Union Springs Shale has been broken up and sheared indicating some minor faulting.

STOP 6. Lower Devonian Kalkberg Limestone
The Lower Devonian Kalkberg Limestone is exposed in a fresh outcrop. This is a medium grained, thin to medium bedded limestone with abundant chert. There are numerous brachiopods, bryozoans, some corals and trilobite fragments. A 2" thick layer of bentonite is exposed near the eastern end of the outcrop. At the east end of grassy field below outcrop there is a small creek that disappears into a sink hole in the limestones.

STOP 7. (in old quarry) Lower Devonian Coeymans Fm
The Lower Devonian Coeymans Formation consists of a coarse grained thickly bedded limestone with abundant brachiopods, crinoids and corals.

STOP 8 (STOP AT PARK NEXT TO OLD BATHS)
Stop to look at springs and tufa deposits in the city park at the north end of the village of Sharon Springs. There is a strong odor of H2S from the spring water. This is probably caused by the water passing through the underlying thick black shales.

Upper Middle Ord Shales and Sandstones in cliff at left.

Two drumlins are in view to the left.

STOP 9. Canajoharie gorge
Four Lower and Middle Ordovician formations are exposed in Canajoharie gorge.
The lower most formation is the Chuctanunda Creek Dolostone. This is unfossiliferous except for the "hippopotami backs" which are dolomitized hemispherical stromatolites (algal mounds) (Park and Fisher. 1969). Large potholes have formed in the dolostone (Canajoharie is the Iroquois name for the "Pot that washes itself") (Park and Fisher, 1969).
The Middle Ordovician Kings Falls and Sugar River black limestones overlay the dolostone forming a disconformity. These limestones and the thin black shales in them contain abundant trilobite fragments, bryozoans, brachiopods and crinoids.
The limestones are overlain by more than 100 feet of Middle Ordovician Canajoharie Shale.

Cliffs of Upper Cambrian Little Falls Dolostone.
En route to the next stop note the cliffs of Upper Cambrian Little Falls Dolostone.

STOP 10. Precambrian garnet gneiss
Follow the foot path at the east end of the outcrop to the railroad cut. Beware of Poison Ivy along the path and at the Railroad cut.
Rusty weathering Precambrian garnet gneiss exposed along the south side of the road and in railroad cut.
The folded Precambrian gneiss is overlain by the Upper Cambrian Little Falls Dolostone. The dolostone is brecciated for several feet above the contact with the gneiss. Apparently there was movement along the unconformity during Ordovician time when the normal faults in the Adirondack Mountains and the Mohawk River Valley were formed. There are also good exposures of the gneiss north of this outcrop across the Mohawk River along route 5 (5.4 miles east of Palatine Bridge).

Observe the fault-line scarp of the Noses Fault
Observe the fault-line scarp of the Noses Fault from the vehicles. There is a good view of the fault-line scarp on the overpass of Route 55 over the railroad tracks. Note that west of the Noses fault-line scarp the Mohawk River Valley has steep cliffs of Little Falls Dolostone. These are caused by the down cutting of the Mohawk River on the upthrown western side of the fault. East of the fault-line scarp on the downthrown side of the fault-line scarp there are no cliffs.

B2: ENVIRONMENTAL GEOLOGY AND HYDROLOGY OF ONEONTA AREA
BRENT K. DUGOLINSKY
Department of Earth Sciences
State University College at Oneonta
Oneonta, New York 13820

STOP 1: Wilber Lake Reservoir
Wilber Lake serves as the principal water supply for the city of Oneonta and adjacent suburban areas. The lake itself covers an area of approximately 90 acres when full and receives runoff and infiltration from a drainage basin of only about 1650 acres. Relief within this drainage basin is 393 feet, with the highest elevation being 1928 feet. Nearly 80% of the drainage basin is forested mainly with mature hardwoods and pine. Some selective lumbering has occurred over the years, but clear-cutting of trees has not, and lumbering has apparently had minimal effect on the drainage characteristics within the basin. The remaining area is comprised of fields, pastures and open bushy areas. The city owns or controls most areas adjacent to the reservoir and no development is allowed. Surface deposits around the lake consist of a silty clay glacial till matrix with shale and siltstone slabs in a layer ranging from about seven to twelve feet thick, decreasing in thickness on the surrounding hillsides to a little over five feet (Walker, 1985).
The lake contains approximately 450 million gallons of accessible water, with another 70 million gallons contained below the water supply intake pipes. According to water treatment plant personnel, the present water use is highly variable, but typically averages about 1.5 million gallons per day (roughly 100 gallons/person/day), or approximately 550 million gallons per year. This requires more than one complete recharge of the lake every year, which only occurred for five years between 1958 and 1982 (Walker, 1985). These years were fortunately when water demands were less than present, but increasing demands since that period have necessitated the search for supplementary sources of water. The largest single users of the city's water supply are the colleges (Hartwick College and the State University College) and water usage varies greatly depending on whether or not one or both colleges are in session since the student population is greater than one-third of the total city population.
The lower reservoir holds an estimated 25 million gallons maximum (probably considerably less due to sedimentation). The water treatment plant will not use water from this source deeper than about two or three feet because of the amount of suspended sediment in the water column. Even with this additional supply, it would only equal a few days supply at the present demand. Supplementary sources include a well drilled west of the city in 1941 for industrial uses but was never used. It is now on property owned by the New York State Electric and Gas Company and was leased to the city in 1982. This well has been connected to the city's water main and can provide up to about one million gallons per day (1.5 cubic feet/second) for four months (Arthur Palmer, pers. comm.), but is only used at about half that capacity. This supply is only used when necessary due to the rust scale in the pipes. Water must be pumped back against the normal flow direction in the mains, which increases the rust scale and discoloration problem. Also, any treatment (such as chlorination) must be done at the well head. Another well drilled in Wilber Park can provide up to about 250 thousand gallons per day. This well is near the water treatment plant and is connected to the water supply pipe to the plant, but is presently not in service. A last resort supplementary supply is possible from the Susquehanna River. Problems with this source are discussed later, as this is one of the stops (#3) on this trip.

STOP 2: Water Treatment Plant
The city's water treatment plant, located on East Street near the high school, treats an approximate average of 1.5 million gallons of water per day from Wilber Lake with chlorine at the rate of 12-18 pounds/day, the rate dependent upon water demand. Also, the supplementary water supply from the West End well is treated daily with 4 to 5 gallons of sodium chloride solution at the well head when this well is being utilized.

STOP 3. Mill Race Pumping Station
The mill race, originally constructed to divert water from the Susquehanna River to provide water power to mills located downstream, now serves as a flood control channel and as a possible supplementary source of water for the city. The small brick building contains a pump which can divert water from the intake immediately above the small dam on the mill race to the reservoir via the pipe visible in the streambed which enters the mill race at the dam. However, this water source is highly undesirable and is viable only in an emergency. There is a high algae content which results in discoloration and an undesirable taste. The water may be polluted with other contaminants as well. Perhaps more significantly, during a drought, when supplementary water supplies would be most needed, the river level would be too low to provide water to this channel. Also, during winter months, it is likely that the water in the above ground pipe from the pumping station to the reservoir would freeze.

STOP 4. Neahwa Park-Buried Hazardous Waste Site
From 1881 until 1953, a coal gasification plant operated in the general area now occupied by the baseball field and adjacent buildings along Gas Avenue in what is now part of Neahwa Park. The plant was originally constructed in 1881 by the Oneonta Gas Light Company, which merged with the Oneonta Electric Power and Light Company in 1887. In 1918, it became part of the Ithaca Gas and Electric Company, which itself became part of New York State Electric and Gas Company (NYSEG). With the advent of cheaper petroleum, natural gas and electric power, the plant closed in 1953 after 72 years of operation. Most of the buildings were dismantled in 1956 and the remaining buildings and property were sold to the City of Oneonta in 1966. The site has since been used for the city garage and storage of city highway equipment, the dog pound, a road salt storage facility, and presently as a minor league baseball field. While in operation, this plant produced gas for heating and lighting from coal, as well as many by-products which are now known to be hazardous. Some of these by-products were buried in the immediate vicinity and were not discovered until recent excavations to repair the Gas Avenue bridge were begun.
NYSEG hired professional consultants to analyze the soils and groundwater in the area, including those from several drilled test wells and excavated pits, in order to determine any present or potential hazard. The initial site survey was initiated in April 1986 and the final technical report was published in February 1990. Analyses of soil from the test pits indicated the presence of the following material (TRC Environmental Consultants, 1988):
A. Aromatic hydrocarbons
| | benzene
| | chlorobenzene
| | 1-2-dichlorobenzene
| | 1-4-dichlorobenzene
| | ethyl benzene
| | toluene
B. Polynuclear aromatic hydrocarbons (PAHs)
| 1. Carcinogens |
| | benzo (A) anthracene
| | benzo (A) pyrene
| | benzo (B) fluoranthene
| | benzo (K) fluoranthene
| | benzo (GHI) perylene
| | chrysene
| | dibenzo (A,H) anthracene
| | indeno ( 1-2-3 CD) pyrene
| 2. Non-carcinogens |
| | acenaphthene
| | acenaphthylene
| | anthracene
| | fluorene
| | naphthalene
| | phenanthrene
| | pyrene
C. Non-chlorinated phenols
| | 2, 4-dimethylphenol
| | methyl-4, 6-dinitrophenol
| | 4-nitrophenol
| | phenol
D. Inorganic compounds
| | cyanide
| | iron
| | zinc
| | sulfate
| | organic nitrogen
It is assumed that these various materials are derived from the estimated 220,000 gallons of dry tar buried in this area over the history of the gasification plant operation. Although the above list of materials may sound devastating, it was determined that although present, these materials occur in such low concentrations as to be considered safe in their present location deep beneath the surface (TRC Environmental Consultants, 1990). In fact, it was determined that excavation and attempted removal of these materials may present a greater risk while simultaneously presenting the problem of what to do with the material once it is removed. It is doubtful that all of the materials could be retrieved in any case. Groundwater movements are being monitored and excavations in the area are prohibited to prevent this material being exposed. The site has been taken off the hazardous site list by the New York State Department of Environmental Conservation and any remediation will come under the control of DEC's Water Division (rather than the Hazardous Waste Division). However, controversy between City officials and the DEC officials continues. One problem includes the fact that State standards are less strict than Federal standards and if Federal standards were followed, the site would still be on the hazardous site list and may therefore qualify for Superfund cleanup. However, DEC officials state that Superfunds were used only when a site had been abandoned or the responsible party was bankrupt or uncooperative. This is not the case and NYSEG and DEC officials plan to work together to develop a remediation plan if it is determined to be necessary.

STOP 5: Hodges Pond
One of the main attractions within the city's Neahwa Park is Hodges Pond. Although swimming is not allowed, it is stocked with small fish (bluegills, etc.) and becomes an ice-skating rink in the winter months. However, recent years have shown dense weed growth in the pond and it was becoming less aesthetically enjoyable as a result. Annual weed removal was expensive and was a temporary solution at best. It was decided in 1990 to drain the pond and dredge much of the bottom mud. Also, to control weed growth without resorting to herbicides, one plan is to lower the water levels in the fall and allow freezing winter temperatures to kill the roots of the shallow water weeds.
Another potential problem is that groundwater movement in this area is generally from the area of the coal tar contamination (STOP 4) toward Hodges Pond. The potential effects of this on the pond are presently unknown.

STOP 6: Southside Mall - Planned Mining
The hill of glacial gravels at the east side of the shopping mall is owned by a local developer who plans to remove the hill and use the material to fill in wetlands along the Susquehanna River about one-half mile to the west (STOP #7). The advantages include allowing more room for expansion of the mall and/or other forms of commercial development as well as providing more potential development space in the area to be filled. However, local environmental groups (City Environmental Board, Audubon Society, and others) and private citizens who realize their responsibility for the stewardship of the environment are concerned about the effect of such massive changes to the local area. Questions have been raised not so much in regards to the removal of the gravel hill but more in terms of the effect of filling in the present floodplain/wetland area along the river. These concerns will be explained at our next stop, which is at the planned fill site. Other potential problems concerning the removal of the hillside gravel include the traffic congestion and the impact on the road surface itself from the hundreds of truckloads of gravel that will have to be hauled continuously over several months. As of this writing, the planned operations are "on hold" until all required permits are obtained by the developer from appropriate State offices (Department of Environmental Conservation, Department of Highways, etc.).

STOP 7: Southside Wetlands - Planned Filling
Most of the area between the highway and the river would be filled with the gravel from the hill adjacent to the shopping mall (STOP #6). As can be seen, most of the trees have already been cut and the developer is apparently only waiting for final approval from State agencies.
One of the principal environmental concerns regarding this operation would be the possible increased potential for flooding. The Federal Emergency Management Agency (FEMA) includes this area within the 100-year floodplain of the river, defined as being at the 1090 foot elevation. Without this area available to help contain floodwaters, it only means that other areas will suffer increased flooding, including city residential and commercial areas. Also, increased stream flow may result in increased erosion downstream. Of particular concern would be the possible erosion of the stream bed adjacent to bridge foundations. Besides the obvious destruction of the wetland habitat, groundwater properties, water tables and drainage characteristics may be adversely affected. A main point of contention is the fact that the developer has not mentioned what his ultimate plans are for this area once it is filled and ready for development. Without knowing what kind of commercial development is planned for this area makes it difficult for even non-environmentally concerned citizens to decide what is best for this area. Most people appreciate the financial benefits of commercial development in their area, as long as they have input on controlling what type of development is planned.

STOP 8. River Access - Erosion Control
This river access site is at the downstream end of a section of the Susquehanna River which was straightened in order to allow the construction of the adjacent section of Interstate 88 to proceed without the added costs of building bridges over the original river meanders. This resulted in the river flowing in a nearly straight line for over a mile. Since rivers naturally develop meanders, especially when flowing over land with minimal gradient, man has had to prevent meander development by lining the river banks with large boulders ("rip-rap"). This also tends to increase downcutting in the restricted channel since flow velocity will increase during increased discharge if it is not allowed to widen its channel. A quick glance downstream should point out the potential hazards of increased erosion caused by natural flooding or from removal of upstream wetland areas which can store floodwaters for gradual release. One of the bridge foundation supports is located in mid-stream and may be weakened by erosion.
Notice also how the flow velocity differs across the river at a meander. This should help explain why there are no boulders to prevent erosion on the inside bend of a meander. Rather, sediment is deposited in these areas of lower flow velocity.

STOP 9: Wastewater Treatment Plant
The Oneonta wastewater treatment plant is classified as a secondary treatment process which involves biological treatment to break down the organic materials followed by chlorination to eliminate potentially harmful bacteria. The average daily volume of water treated in 1990 was 3 million gallons but the plant has a maximum capacity of about 3.6 million gallons daily. According to the manager of the plant, one reason for the higher volume of waste water treated than the volume of reservoir water entering the city supply mains is infiltration of groundwater into some of the older permeable sewer pipes. Also, there are some homes in the town that may be connected to the sewer lines but may have their own private wells rather than being connected to the city water supply.

STOP 10: Transfer/Recycling Station
The transfer/recycling center was opened in 1988 when recycling became mandatory in the City of Oneonta, one of the first cities in this area to do so. Recyclables are separated by the generator (individuals, businesses, etc.) and either picked up by commercial haulers during the regular garbage pick-up schedule or brought to the station by individuals. Present materials that are removed from the waste stream and recycled include newspapers, certain types of plastic, glass (clear, green and brown are separated), corrugated cardboard, cans, tires, car batteries, used motor oil, and green waste (grass clippings, leaves, branches, Christmas trees, etc.). Newspapers, cardboard, cans and plastic containers are compacted into bales for shipment to available markets for reuse. Glass bottles are crushed before shipment in order to reduce the volume. The remaining garbage is compacted and transferred by truck to a landfill in Montgomery County.
The facility was taken over by the Montgomery -Otsego -Schoharie Authority (MOSA), which controls garbage collection and recycling in the three county area. A program to begin recycling glossy paper, magazines, junk mail, and similar material is planned to begin this fall. Future plans include the mixing of green waste with sewage sludge to produce a composted material which could be used by farmers or other citizens to enrich their fields and gardens. In many larger metropolitan areas, the use of sewage sludge in this way is not feasible due to the high levels of industrial heavy metals or chemicals it contains. It is also hoped that large-scale recycling of other plastics (such as styrofoam and brittle plastics) will be feasible with further advances in recycling technology and the subsequent development of markets for this material.

B3: FRESHWATER CARBONATE FROM THE UPPER DEVONIAN CATSKILL MAGNAFACIES, DAVENPORT CENTER, CENTRAL NEW YORK
Robert V. Demicco
JohnS. Bridge
Kelly C. Cloyd
Department of Geological Sciences
State University of New York at Binghamton
Binghamton, New York 13902-6000
ABSTRACT
Non-soil carbonates are rare in fluvial deposits of the Upper Devonian Catskill Magnafacies although they are common in other fluvial rocks. A 0.5 m thick dolomitic mudstone bed occurs near the base of the Oneonta Formation at Davenport Center, NY. This laminated mudstone layer contains filament molds, rare ostracode shells, calcispheres and a variety of burrows and mudcracks. Associated strata include sandstones interpreted as channel deposits and interbedded sandstones and mudstones interpreted as crevasse-splay, levee and flood basin deposits. We interpret the carbonate bed as a freshwater-marsh/shallow-lake deposit analogous to the "periphyton" lakes and marshes of the Everglades and interior Andros Island. This marsh/lake deposit developed during local reduction of siliciclastic input. The purpose of this trip is to examine the carbonate bed and associated strata, to compare the sequence with similar coeval sequences in the region, and to discuss our interpretations.

STOP 1. The outcrop
Non-soil carbonates are rare in fluvial deposits of the Upper Devonian Catskill Magnafacies although they are common in other fluvial rocks. A 0.5 m thick dolomitic mudstone bed occurs near the base of the Oneonta Formation at Davenport Center, NY. This laminated mudstone layer contains filament molds, rare ostracode shells, calcispheres and a variety of burrows and mudcracks. Associated strata include sandstones interpreted as channel deposits and interbedded sandstones and mudstones interpreted as crevasse-splay, levee and flood basin deposits. We interpret the carbonate bed as a freshwater-marsh/shallow-lake deposit analogous to the "periphyton" lakes and marshes of the Everglades and interior Andros Island. This marsh/lake deposit developed during local reduction of siliciclastic input. The purpose of this trip is to examine the carbonate bed and associated strata, to compare the sequence with similar coeval sequences in the region, and to discuss our interpretations.

B4: ACTIVE AND STAGNANT ICE RETREAT: DEGLACIATION OF CENTRAL NEW YORK
P. Jay Fleisher
Earth Sciences Department
SUNY-Oneonta
The depositional landforms of glacial origin on the Appalachian Plateau of central New York are atypical of classic continental environments. They consist of an unconventional landform assemblage formed by retreating ice-tongues at the margin of the Laurentide Ice Sheet. Landforms representative of back-wasting occupy most through-valleys, while widespread stagnation and down-wasting occurred in non-through valleys. A modified ice-tongue model includes valley tongues 20 km long and emphasizes the significance of in-wash as a primary source of sand and gravel, and subglacial sediment flow within tunnels as the major source and transport mechanism for silt.

B4: ACTIVE AND STAGNANT ICE RETREAT: DEGLACIATION OF CENTRAL NEW YORK

Dead-ice sink on right

Unadilla River

STOP 1. New Berlin: Through valley landforms
Through valley landforms - deltaic valley train and dead-ice sink.
Large-scale deltaic foreset beds exposed here and within a currently inactive quarry in the center of the valley contain clast supported, very coarse gravels and pebbly, coarse sands; leached and re-precipitated carbonate (limestone) form partial cement in some units. This valley train is breached by the Unadilla River that flows from a dead-ice sink up-valley to a lacustrine plain down-valley. Landforms indicate detachment of large ice mass during active ice retreat.

Road rises on gravel terrace above lacustrine plain

High-level terrace (1280') on west side of valley
High-level terrace (1280') on west side of valley above road (continues for 1.5 miles) is of unknown origin.

Road rises on pitted planar gravel and kame field

STOP 2. Columbus Quarters: kame field and pitted plain in through valley.
These landforms, breached by the Unadilla River at its confluence with Tellette Creek, may have served to temporarily dam a local lake up-valley. Pebble counts from upland drab drift and kame field bright drift suggest tributary in-wash from the Tellette Creek drainage basin was a partial source of sediment (see Case # 4 for discussion of pebble count data).

Proceed east across lacustrine plain

Road descends onto the pitted and discontinuous valley train of Butternut Creek.

Road descends onto the kame and kettle topography of the moraine at Oaksville.

Road crosses the crest of the moraine
Road crosses the crest of the moraine and continues on and off the moraine for next 1.5 miles through the villages of Oaksville and Fly Creek.

STOP 3. Fly Creek: Valley train and dead-ice sinks in through valley.
Sand and gravel valley train with classic assemblage of glacio-fluvial sedimentary structures. What criteria may be used to distinguish outwash from in-wash?

STOP 4. Cassville-Cooperstown moraine at Index
This classic through valley landform (moraine and associated valley train) represents active ice deposition. It consists of stratified sand and gravel that grade up-valley into silt 120+ feet thick (see Case #1, Figure 2). Is it an end moraine, as proposed by Krall, 1977, or a recessional moraine? Although high enough, it did not dam Glacial Lake Cooperstown. Landforms and stratigraphy between here and Otsego Lake suggest detachment of a marginal-ice cleat and development of a dead-ice sink (see Case #1, Figure 4).

Dead-ice sink
Dead-ice sink behind Cassville-Cooperstown moraine occupies the valley for the next mile.Dead-ice sink behind Cassville-Cooperstown moraine occupies the valley for the next mile.

STOP 5. Council Rock Park on Doubleday Ice Margin
Well data indicate 180 feet of bouldery silt beneath a moraine that plugs the valley and dams Otsego Lake. Subtle hummocky terrain in the village of Cooperstown can be traced westward to the golf course. While ice-cored, this landform dammed Glacial Lake Cooperstown at an elevation of 1250'. The view northward reveals conspicuous cross-sectional valley asymmetry, which is an important aspect of sediment source for Glacial Lake Cooperstown (see Case #1, Figures 5-9).

Highway traverses Cassville-Cooperstown moraine.

Hamlet of Hyde Park
From here south for several miles, planar gravels appear related to upland tributary in-wash sources.

Valley floor consists of lacustrine clays
Valley floor consists of lacustrine clays from here south to a moraine at Portlandville

STOP 6. Goodyear Lake overview at Milford Center
Pitted valley train and dead-ice sink. These landforms indicate local stagnation of the valley ice-tongue during active retreat. However, the valley fill consists of 60+ ft. of ice-contact stratified drift over 300+ ft. of silt interpreted to be of lacustrine origin. Does the stratigraphy indicate two stades (re-advance) or is there a single-stade environment that accounts for both stratigraphic units? Aggradation over ground ice islands is suggested.

Dead-ice sink right and left

Road parallels Charlotte Creek through the moraine on the right.

STOP 7. Davenport Center: Pine Lake Dead-ice Sink Complex.
These kames and sinks formed in an ice-cored terrain at the mouth of Charlotte Creek where the valley was clogged by remnant ice. Case # 3 considers the origin of landforms here. Also see Case # 3, Appendix A for the results of a pebble count study that tested an in-wash origin for these materials.

Cross the lacustrine plain of Glacial Lake Davenport.
Right hand fork goes downhill and across the lacustrine plain of Glacial Lake Davenport.

STOP 8. Davenport Center: pitted hanging delta.
Current and past exposures contained well-developed, large-scale deltaic foreset beds in a hanging delta marked by kettles 40+ ft. deep. This indicates progradation onto and across grounded ice in Glacial Lake Davenport. See Case # 3 for discussion.

Highway traverses moraine at West Davenport.

Terrain to the left is what may be the only lateral moraine in this area.

Road descends to Susquehanna valley lacustrine plain.

Breach of the Oneonta moraine
Highway enters the breach of the Oneonta moraine (significantly altered by urban development).

Down-valley view south across local kame field.
Pull over and walk up Winnie Hill Road 0.1 mile. for down-valley view south across local kame field.

STOP 9. Village of Laurens landfill
Ice-contact stratified drift. At various times, excavation has exposed tilted and truncated sequences of inter-stratified gravel and pebbly sand with fluvial sedimentary structures and local silt layers of lacustrine origin. This is interpreted to indicate semi-continuous aggradation of ice-cored terrain. Tributary in-wash from the west is a possible sediment source.

Planar gravel, valley train remnant for next half mile.

STOP 10. Otsego County sand and gravel quarry: planar gravel remnant.
This landform, typical of many others within the Otego Creek valley, is a discontinuous, planar gravel deposit interpreted as a remnant of a surface of aggradation that included semi-continuous ice masses as vestiges of a collapsed ice tongue. Ice-contact fluvial and lacustrine facies are common.

Road rises onto pitted planar gravel remnant.

Road traverses terrain of ice-cored origin including a small esker out of sight to the right.

STOP 11. Overview of meltwater channel
Overview of meltwater channel through discontinuous planar gravel remnant and terrain of ice-cored origin.

Road descends to a local lacustrine plain (dead-ice sink?)

Gravel excavation to the left removed a landform thought to be a proto-esker.

STOP 12. Overview of discontinuous, planar gravel surface
Overview of discontinuous, planar gravel surface and terrain of ice-cored origin bordering a local, up-valley lacustrine plain.

STOP 13. Overview of headward area in non-through valley.
Valley floor contains a few isolated kames. Unlike through valleys, the valley floor gradually rises to upland slopes across which ice flow diminished during glacial retreat resulting in downvalley ice-tongue starvation and collapse.

B5: UNDERSTANDING THE EAST CENTRAL ONONDAGA FORMATION (MIDDLE DEVONIAN)
-AN EXAMINATION OF THE FACIES AND BRACHIOPOD COMMUNITIES OF THE CHERRY VALLEY SECTION, AND MT. TOM, A SMALL PINNACLE REEF
THOMAS H. WOLOSZ
Center for Earth and Environmental Science
SUNY College at Plattsburgh
Plattsburgh, N.Y. 12901
HOWARD R. FELDMAN
Biology Department
Touro College New York, N.Y. 10035
RICHARD H. LINDEMANN
Skidmore College
Saratoga Springs, N.Y. 12866
DOUGLAS E. PAQUETTE
Safety and Environmental Protection Division
Brookhaven National Laboratory
Building 129
The Onondaga Formation of New York and Ontario, Canada has been extensively studied (see Oliver, 1976 for references), and yet is still poorly understood. It is a unit which is, in the western part of New York and in Ontario, clearly transgressive, and yet it lacks any of the classic peritidal facies associated with shallow water carbonates. It is a "reefy" unit, but the major reef building paleocommunity of the Middle Devonian (stromatoporoids and algae) are either extremely rare or absent.

B5: UNDERSTANDING THE EAST CENTRAL ONONDAGA FORMATION (MIDDLE DEVONIAN)

STOP 1. Cherry Valley Section.
Park along road at top of west end of road cut. An almost complete section of the Onondaga Limestone is exposed along this cut. (See discussion of Stratigraphy and Brachiopod Communities).

STOP 2. Mt. Tom Reef.
Mt. Tom Reef makes up the large hill to the south of the road (See discussion of Mt. Tom).

TOP 3. Mt. Tom #6
Mt. Tom #6 forms the low, wooded ridge to the southwest of the road (See discussion in test).

STOP 4. Intersection of Collins and Geywittz Roads.
Leave cars and proceed east from the intersection. Mt. Tom #2 forms the low hill to the south of the small creek, and small outcrops may be examined along the south side of the creek valley on the hill itself. A small quarry the northwest edge of the hillside exposes bedded Edgecliff facies, a small rugosan mound is located just to the southeast of the quarry among the trees. (See text for discussion.)

B6: STORM-DOMINATED SHELF AND TIDALLY-INFLUENCED FORESHORE SEDIMENTATION
-UPPER DEVONIAN SONYEA GROUP, BAINBRIDGE TO SIDNEY CENTER, NEW YORK
DANIEL BISHUK JR.
Groundwater and Environmental Services, Inc. (GES)
300 Gateway Park Drive
North Syracuse, New York 13212
ROBERT APPLEBAUM and JAMES R. EBERT
Dept. of Earth Sciences
State University of New York
College at Oneonta
Oneonta, New York 13820-4015
The Upper Devonian paleoshoreline of the Catskill clastic wedge in New York State has been interpreted for nearly a century as a complex deltaic sequence (Barrel, 1913, 1914; Chadwick, 1933; Cooper, 1930; Sutton, Bowen and McAlester, 1970; and many others). Friedman and Johnson (1966) envisioned this deltaic complex as a series of coalescing deltaic lobes that progressively filled the Catskill epeiric sea and existed as an uninterrupted deltaic plain from New York to West Virginia. In addition, some geologists believe that such epeiric seas were tideless owing to rapid tidal wave attenuation (Shaw, 1964; and Mazzullo and Friedman, 1975). Others presume that storm (wave) processes were dominant with little or no tidal influence (Dennison, 1985).
This study offers significant departures from these interpretations, by documenting non-deltaic environments with significant tidal influence along the Catskill paleo-shoreline. The purpose of this study (Fig. 1) is to delineate sedimentary environments spanning the non-marine to marine transition in the Upper Devonian Sonyea Group and to test and challenge previous deltaic models of the Sonyea Group (Sutton, et al., 1970). Recent publications have introduced evidence for non-deltaic shoreline environments within the Catskill clastic wedge (Walker and Harms, 1971; 1975; Bridge and Droser, 1985; VanTassel, 1986; also see Seven, 1985, Table 1, p. 83). Others have shown that tidal processes were significant in the Catskill sea (Slingerland, 1986; and Bridge, et al. 1985).
Figure 2 shows the paleogeography of eastern North America during Frasnian-Fammenian time. Paleogeographic reconstructions indicate that the Acadian Mountain range supplied large quantities of sediment westward, allowing a vast alluvial plain to develop. The alluvial plain was fronted by deltaic lobes which prograded westward into the Catskill epeiric sea.

B6: STORM-DOMINATED SHELF AND TIDALLY-INFLUENCED FORESHORE SEDIMENTATION

Oneonta Formation (Genesee Group)
To your left are exposures of fluvial deposits of the Oneonta Formation (Genesee Group). This formation is similar in many aspects to the Lower Walton Fm. (Sonyea Group) included in this study.

The Susquehanna river has incised through the Wells Bridge moraine
The Susquehanna river has incised through the Wells Bridge moraine (right). This moraine served as a dam, which impounded glacial Lake Otego. Up-gradient of the moraine, 300 to 400 feet of lacustrine silts and clays underlie the valley floor. Down-gradient, an outwash terrace can be seen at the rest area (13 mi.).

A sinuous, beaded esker
Immediately after exit 11, a sinuous, beaded esker can be seen at your right (extends for approx. 0.5 miles).

Outcrops of Oneonta Fm. at left.

Glaciofluvial sand & gravel quarry to the left.

Reference marker: Helderberg Bluestone Quarry at end of road
Excellent exposure of the Oneonta Fm. (Genesee Group).

Hummocky cross-stratified sandstones
Hummocky cross-stratified sandstones interbedded with shale of the Glen Aubrey Fm. (Sonyea Gr.).

Glen Aubrey Fm.(right); ball and pillow structures predominate the exposure.

Amalgamated portion of the hummocky cross-stratified facies
Small quarry to the right consists of the amalgamated portion of the hummocky cross-stratified facies of this study (Glen Aubrey Fm.).

STOP 1A: Steele Quarry And Pasture Outcrops Of The Hummocky Cross-Stratified Facies
Location: Scattered outcrops and a small, inactive quarry are located in a cow pasture behind the Steele residence. All stop 1 outcrops can be found in the south-central portion of the Unadilla, New York 7.5 minute quadrangle, located at the intersection of Dunshee Road and Delaware County Route 35. Before visiting the Steele quarry, please ask Mr. and Mrs. Steele for permission to walk on their property. Since stop 1A is on private property, the authors ask that you do not use rock hammers on the outcrops or collect any samples.
Description: The Steele quarry and pasture outcrops (outcrop V) (Fig. 4) are stratigraphically positioned in the Glen Aubrey Formation, and are designated as the amalgamated portion of the hummocky cross-stratified facies. These exposures offer an excellent opportunity to examine hummocks and swales of HCS beds in plan view. The amalgamated portion of the hummocky cross-stratified facies is interpreted as dominantly sandy, storm deposits at or above fair weather wave base (consult text for details). Fossils are sparse and are restricted to coquinite lenses consisting of Sphenotus, Cupularostrum, Platyrachella, fish plates, and plant fragments. Sedimentary structures which cap hummocky cross-stratified beds include wrinkle marks crescentic scour depressions partially filled with very fine sand. These scours commonly exhibit asymmetrical and symmetrical ripples within the depression.
An enigmatic shale unit is found in the center of the quarry. At approximately halfway up in the shale, there is a light colored and differentially weathered horizon which has rather high concentrations of dolomite. Since this shale can not be traced laterally, it is difficult to decipher its stratigraphic context and facies relationships.
At outcrop V in the cow pasture (Fig. 4), excellent hummocky bedforms are present. Within the hummocky cross-stratified beds, climbing ripple cross-lamination are inclined at a steep angle of propagation (Fig. 9). The ripples are slightly asymmetric in the direction of propagation toward the crest of the hummock (from left to right) suggesting that they migrated slightly up the flank. Steep inclination of propagation and the form-concordant style reflects vertical growth of the hummocks with minimal migration. This supports that hummocky cross-stratification formed under high velocity oscillatory flow conditions and rapid deposition rates during storm events.

STOP 1B: Sidney Center Outcrops 1B - 1F
STOP 1B: Open Pit And Field Outcrops Of The Trough Cross-Bedded Facies
Outcrops R and S represent the lowest portion of the trough cross-bedded facies, which sharply overlies the hummocky cross-stratified facies. The trough cross-bedded facies is assignable to the lowest facies unit of the Cattaraugus Magnafacies. Locally, an erosional base defines its base and near base, which is characterized by a coarse lithoclast, shelly, and plant debris lag. Contact relationships are not observable at this stop. The outcrops are positioned within a meter or two above the contact. However, a coarse lithoclast, shelly lag is present at outcrop R. The rhynchenellid, Cupularostrum, and the bivalve, Sphenotus, occur sparsely at this horizon. The sparse marine fauna is the best way to differentiate between the trough cross-bedded facies and the strikingly similar planar-laminated and trough cross-bedded facies (Lower Walton Formation) positioned stratigraphically higher. Trough cross-beds commonly climb at low angles of propagation and interpreted as dunes within tidal inlets within a barrier island/strandplain and tidal creek system. Planar laminated interbeds are interpreted as swash bars within the tidal inlets.
STOP 1C: Logging Trail Outcrop Of The Trough Cross-Bedded Facies
Outcrop D is located approximately 40 feet higher within the trough cross-bedded facies than stop 1B. Trough cross-beds are spectacularly illustrated here owing to well developed differential weathering of the outcrop surface. Trough cross-beds exhibit normally graded, lenticular laminae, which probably reflect sediment supply pulsations induced by tides (Fig. 11). In addition, asymmetrical ripples on troughs commonly oppose the dominant, northerly paleoflow direction of trough cross-beds (Fig. 11 ). These asymmetrical ripples record a subordinate current reversal induced by tides within tidal inlets. Evidence of tidal activity in the Catskill Sea is only sparsely documented in the literature (Ebert, 1987; Walker, 1975; and Bridge and Droser, 1985) and has never been documented from the Sonyea Group.
STOP 1 D: Woodland Outcrop Of The Trough Cross-Bedded Facies
Outcrop K is also assignable to the trough cross-bedded facies. Trough cross-beds are likewise spectacularly illustrated at this outcrop owing to well developed differential weathering of the outcrop surface. Trough cross-beds represent dunes migrating within tidal inlets. These alternate with planar-laminated interbeds, which represent swash bars deposited along the margins of tidal inlets. The trough cross-beds are unidirectional to the north and climb at low angles of propagation. Rare ripple cross-lamination on weathered surfaces is suggestive of current reversals, which offers additional supporting evidence for tidal influence.
STOP 1 E: Bluestone Quarries #4 & #5 Of The Planar-Bedded Facies And The Mottled Mudrock And Sandstone Facies
The planar-bedded facies is sharply overlain at these quarries by the mottled mudrock and sandstone facies. The contact is planar and lacks incision by channels. The planar-bedded facies is extensively quarried for "bluestone" slabs used for sidewalk and construction purposes. Internal lamination within planar-beds is rare, but where it is present, it is often inversely graded. These structures are common on high energy foreshore beach/barrier setting. Also present are abraded arthropod fragments, and highly abraded and highly rounded quartz grains. The arthropod fragments may be either ostracods or trilobites. Near the upper contact of the planar-bedded facies, an association Skolithos linearis, Barroisella campbelli, and delicate plant leaves and stems strengthens a beach (foreshore) interpretation for the planar-bedded facies. Barroisella was an infaunal burrower with a long, siphon-like peduncle and lived with its long axis vertical (Boucot, 1981). It probably lived in intertidal creeks and tidal flats in the Sonyea Group.
The mottled mudrock and sandstone facies stratigraphically above has an association of abundant fish debris of Bothriolepis with Barroisella and Skolithos, occurring stratigraphically above foreshore beach deposits of the planar-bedded facies, argues in favor of brackish water conditions. The mottled mudrock and sandstone facies may represent tidal creek deposits.
Stop 1F (Optional): Bluestone Quarry #1/Transition Of Trough Cross-Bedded Facies With The Overlying Planar-Bedded Facies
Bluestone quarry #1 occurs near the gradational contact between the trough cross-bedded facies and the planar-bedded facies. Planar-bedding dominates while trough cross-bedding is rare. The purpose for visiting this outcrop is to establish the contact relationship and to have an opportunity to collect plant fossils. Fossils of carbonized and pyritized plant stems, branches, logs, and bark can be readily collected at this location. Parting lineations and aligned plant fragments are abundant, which represents a strong onshore/offshore subtidal flow pattern.

Break For Lunch

Exposures of the Triangle Fm. (Sonyea Group)
Exposures of the Triangle Fm. (Sonyea Group) consist of hummocky cross-stratified sandstone interbedded with shale.

STOP 2: Hummocky Cross-Stratified Facies. Triangle Fm(?).
Three outcrops afford ample opportunity to examine storm deposits. We will concentrate most on the largest exposure directly south of the eastbound lanes of Interstate 88. Storm beds, numerous soft-sediment deformation structures of variable scale, bioclastic mudstones and sandstones, and listric truncation surfaces are most easily identifiable at the exposure.
Please be extremely careful not to go out on the highway as this spot has limited sight distance for drivers. Also, vehicles are accelerating up the onramp onto the primary lanes. Please be alert at all times while at this stop. You may cross over to examine the middle outcrop north of the eastbound lanes. Be absolutely sure there are no oncoming cars as far as you can see down the lanes and around the bend before you cross the road. Vehicles are moving faster than they appear but if you are careful then there is plenty of time to cross. Bedding planes are exposed at the very top of the outcrop displaying wave ripples and oriented fossils. They can be reached by carefully ascending the east end of the outcrop at the large culvert that passes under the road. Note also the numerous bedding surfaces with wave (?)or current(?) oriented Mucrospirifers and Tentaculites. East of the culvert are large slabs weathered out of the outcrop that contain numerous fossils and sedimentary structures. Collect all you wish from these loose slabs, but please do not collect in situ fossils or structures as ongoing research is continuing at this time. Please do not disturb any bedrock in place that has elevation or benchmark data recorded on it. These are reference markers for future study and should not be disturbed.
As we work our way up the large exposure, note the changes in lithology. Is the sequence fining or coarsening upward? What about the coarsest sandstone beds? Do they thicken upward over the entire vertical sequence or are there no discernible trends? Notice the slump scars (?) and what's directly above them. Could loading of several storm-deposited decimeter thick sandstone beds above the mudstone have contributed to the slump, or is some other mechanism warranted here such as seismic shaking (see discussion earlier by Bishuk, this paper). Are there any channel forms? We will try to answer these questions at the outcrop.
Note the variability in the morphology of the soft-sediment features. Relative size, orientation, internal stratification, and depth of penetration into underlying beds all vary throughout the section. Does there appear to be a correlation between the size of pillows and the thickness of the overlying sandstones? Are these all just balls and pillows? Consider the possibility of load casts, foundered ripples, and channel forms. Interpretation of these structures at this locality has been a source of lively discussion among the authors. Much work remains concerning the genesis of these sandstones.
Fossil coquinites and bioclastic mudstones are common at the base of storm layers, i.e. very fine to fine sublitharenites. This gives way to planar cross stratification, wave ripples, and thin mudstones with extensive bioturbation at the top of most sequences. These deposits are interpreted as tempestites by Applebaum (in prep.). Internal morphologies vary from one storm bed to another. For a recent discussion of vertical stratification sequences in storm-deposited beds see Myrow and Southard (1991).

STOP 3A: Stream Gorge Exposure Of The HCS Facies, Glen Aubrey Fm.(?)
The exposures in the stream gorge south of Route 206 contain massive bedded sandstones, amalgamated sequences of balls and pillows, and most notably several key horizons of calcareous bioclastic conglomerates that may prove useful as lithostratigraphic or biostratigraphic "key beds". These shell beds are undoubtedly thanatocoenose assemblages and contain mostly crinoid columnals, brachiopods, gastropods, and branching bryozoans.
Numerous erosional surfaces within shell beds indicate amalgamation. Numerous reoriented geopetal structures (e.g. mud-filled gastropoda) suggest reworking of bioclastic materials by storms under waning flow conditions during the closing phases of the individual storm event. Fine mud resuspended by weak oscillatory wave currents settles out into areas of shelter porosity (underneath brachiopoda valves or inside Bellerophon chambers, see Figure 23). Similar fabrics have been documented by Kreisa and Bambach (1982).

STOP 3B: Hummocky Cross-Stratified Facies. Glen Aubrey Fm(?)
This stop has some interesting features on the south side of the road. Of particular note are the high angle dips on some of the sandstone beds. Are these slump structures or channel forms? At the present time a definitive interpretation is unavailable. We welcome any suggestions. Surprisingly puzzling is the outcrop just north of the road, which contains no such distortion of bedding with all horizontal and undisturbed stratification. Can you explain the sudden change in attitude of the bedding over such a small distance, i.e. 10 meters?

STOP 4A: HCS Facies And Dark Gray Shale Facies. Glen Aubrey Fm
Location: A huge exposure of the hummocky cross-stratified facies and dark gray shale facies can be seen on both sides of Route 8 located approximately 1 mile south of I-88 off-ramp.
Description: At this outcrop, we will look primarily at the dark gray shale, facies, which is found as three interbeds within the hummocky cross-stratified facies. The dark gray, facies records three brief transgressive periods, which represent deposition on deeper portions of the shelf. The hummocky cross-stratified facies represents storm deposits within storm base, and is similar here in many aspects to stops 2 and 3. The dark gray shale, facies is characterized by dark gray shale, a complete absence of invertebrates and biogenic structures, and locally abundant, listric truncation surfaces. The listric geometry of the truncation surfaces is consistent with submarine gravity-slide structures. The truncations are sharp and regular without any obvious local channels or erosional irregularities, which is best explained by shear rather than current scour. The rotational slumps were probably triggered by storm events or overloading causing liquefaction and foundering of the sediment.

STOP 4B. Amalgamated Portion Of The HCS Facies, Thorpe Road.
The sharp, scoured and loaded contact in the middle of the outcrop is interpreted as the contact between lower shoreface deposits (amalgamated hummocky cross-stratified facies-above fair weather wave base) and shelf deposits (hummocky cross-stratified facies-below fair weather wave base and above storm wave base). The predominance of fine sand sublitharenite above this contact, and a marked upward decrease in fauna suggest that the amalgamated hummocky cross-stratified portion of this facies occupies the lower shoreface and indicates increasing environmental stress associated with the nearshore. Siltstone and shale interbeds are rarely preserved, indicating that waves or currents (e.g., longshore-, tidal-, and storm-driven) had effectively. winnowed the fine fraction. Hummocky cross-stratification is the predominant structure in all sandstone beds at the outcrop. The top surface of the outcrop displays excellent HCS beds in plan view and a variety of biogenic and sedimentary structures which cap HCS beds. Some of the observable features include wave ripples exhibiting chaotic patterns, asymmetrical ripples with transverse ribs, conglomerate sheets consisting of flat shale intraclasts, paired vertical burrows of Arenicolites, and "figure 8"-shaped burrow of unknown ichnogenera. To reach the top of the outcrop, please walk around to the extreme right-hand side of the outcrop.
Near the base of the outcrop, an unusual interbed of grain-supported conglomerate grades to fine sand sublitharenite. The conglomerate is the coarsest unit found in the study area. Conglomeratic clasts are subrounded, with an abundance of siltstone and shale clast lithologies. These clasts commonly have Paleozoic weathering rinds. Other components include fish fragments, pockets of diagenetic galena, and plant debris occasionally replaced by galena. Graded bedding is the predominant structure consisting of alternating coarse sand and fine sand sublitharenite. Interpretation of this horizon is still pending, although it is suspected that it is a storm-generated deposit.

STOP 5. Sidney Mountain Bluestone Quarry
Planar-Bedded Facies, Mottled Mudrock And Sandstone Facies, Planar-Laminated And Trough Crossbedded Redbed Facies, And The Hummocky Cross-Stratified Facies
Location: From the gate described above in the road log, proceed on foot on dirt road for 200 meters. Walk around left side of quarry storage shed. Road will lead to the bottom of the quarry at its northern-most end. Each facies noted above can readily be seen by traversing back toward the south along the east wall of the quarry. Proceed with caution while walking on the quarry grounds. The terrain is rough and there is an open pit.
Description: Sidney Mountain quarry has the most complicated stratigraphy in the study area and exhibits a multitude of sedimentary features. As a result of its complexity, we direct you to the sections on planar-bedded facies, mottled mudrock and sandstone facies, and hummocky cross-stratified facies (especially Fig. 7).