NYSGA 2003
NEW YORK STATE GEOLOGICAL ASSOCIATION FIELD TRIP GUIDEBOOK
75th Annual Meeting
Eric Lee Johnson, Editor
HOSTED BY:
Geology and Environmental Science Department
Hartwick College
Oneonta, New York
13820
and
Department of Earth Sciences
State University of New York College at Oneonta
Oneonta, New York 13820
PUBLISHED BY THE NEW YORK STATE GEOLOGICAL ASSOCIATION.
Additional copies may be obtained from the Association by contacting Dr. William Kelly
c/o New York State Geological Survey
Rm. 3140 CEC Albany, NY 12230
or by visiting http://www.nysgaonline.org/
DEDICATION
This guidebook is dedicated to all of those who helped make this meeting possible. Special thanks are given to Dr. Jim Ebert at SUNY Oneonta for his tireless efforts in making the arrangements for the meeting, and to Dr. Harrison Schmitt for coming and helping us celebrate NYSGA's 75th anniversary. Finally, this meeting would not have been possible without the dedication and hard work of all of the field trip leaders who have spent countless hours both in the field and at the keyboard constructing the fabulous array of trips presented on the following pages.
The tops of mountains are among the unfinished parts of the globe, whither it is a slight insult to the gods to climb and pry into their secrets, and try their effect on our humanity. Only daring and insolent ones, perchance, go there.
Henry David Thoreau

A1: LATE MIDDLE DEVONIAN BIOTIC AND SEDIMENTOLOGIC EVENTS IN EAST CENTRAL NEW YORK-TULLY FORMATION CLASTIC CORRELATIVE SUCCESSION IN THE SHERBURNE-ONEONTA AREA
Gordon C. Baird
Dept. of Geosciences, S.U.N.Y. Fredonia
Fredonia, NY 14063
Carlton E. Brett
Alexander J. Bartholomew
Dept. of Geology, University of Cincinnati
H.N. Fisk Laboratory of Sedimentology
500 Geology/Physics Bldg.
Cincinnati, OH 45221-0013.
ABSTRACT
Terrigenous clastic deposits equivalent to the upper most Moscow Formation and Tully Formation record a number of key sedimentological-paleontological events that occurred during the latest part of the Middle Devonian Givetian Stage. These deposits, last mapped in detail by G. Arthur Cooper and H.S. Williams in the 1930s, have been the subject of extensive stratigraphic mapping work by the present authors over the past five years. Goals of this mapping project are: 1) to identify divisions and boundaries correlative to units identified in the Tully Limestone further west through the classic work of Phillip Heckel in the sixties and seventies; 2) to identify key facies that can be interpreted in the context of onshore-offshore paleoenvironmental models; 3) to reconstruct the faunal succession in this thick interval of strata so as to elucidate in greater detail two key levels of faunal change associated with the global Taghanic Bioevent; and 4) to use stratigraphy and sedimentology to identify structural features that were active or present during deposition of these units.
Preliminary results of our work show that most of the lower and medial parts of the Tully Formation clastic correlative succession (here in referred to as "TFCCS") between Sherburne, NY and New Lisbon, NY consist of relatively unfossiliferous strata that record the development of a structural basin (New Berlin Trough) probably centered in what is now the Unadilla Valley. Condensed units overlying regional discontinuities developed at several stratigraphic levels both in the basin and on its western flank; sediment-starvation associated with condensation led to the formation of "oolitic" chamosite under conditions not fully understood. Basin center deposits recorded dysoxic conditions and development of low diversity faunas dominated by auloporid corals and small brachiopods; eastward passage of sparsely fossiliferous deposits near New Berlin to richly fossiliferous equivalent units near Laurens reflects a shoreward basin-to-shelf transition. This contradicts the standard view that the lower and medial Tully record primarily shallow and even lagoonal conditions. Our model does vigorously support (and significantly enhances) Heckel's (1973) view of a "Tully clastic trap" controlled by structure in east-central New York.
Herein, we show that typical oxic Hamilton facies ends with onset of dysoxic, highstand conditions associated with deposition of the New Lisbon Member. We exclude the upper, Tully-equivalent part of Cooper and Williams' (1935) "New Lisbon Member", which is separated from our restricted New Lisbon Member by a basal Tully erosion surface. Currently, we believe that the restricted New Lisbon succession is equivalent to the Gage Gully submember of the Windom Member, although this is still uncertain. The TFCCS contains six internal mappable discontinuities; these are, in upward-succession: a, a base-Tully (DeRuyter Bed-equivalent?) contact; b, a sub-"Laurens" contact, possibly equating to the base of the Fabius Bed to the west; c, a sub-Symrna Bed disconformity traceable westward into central New York; d, a top-Smyrna flooding surface contact; e, an erosion surface flooring the West Brook Bed; and f, a diastem (base-Moravia Bed? discontinuity) above the West Brook Bed. The upper Tully, base-West Brook contact marks the greatest lowstand event with development of a prolific Hamilton faunal association characterized by diverse brachiopods, and corals, above the contact. East of Otselic, the Tully Formation contact with the overlying Geneseo Black Shale is conformable.
Initial Tully Fauna incursion occurs at the base-Tully (DeRuyter Bed-equivalent?) contact at New Lisbon and Laurens. The Tully Fauna persists above the Smyrna Bed into the middle part of a sparsely fossiliferous post-Smyrna Bed highstand-succession probably correlative with the Taughannock Falls Bed within the Tully Limestone. The uppermost part of the highstand succession is characterized by a return of the pre-Tully Hamilton fauna. The West Brook Bed records a major, anomalous acme ("last hurrah") of the Hamilton Fauna prior to onset of dysoxic and, finally, anoxic conditions associated with uppermost Tully and Geneseo deposition. Preliminary work on eastern sections of the TFCCS suggests that nearshore, Tully equivalent "Chemung-type" facies may yield several taxa of the Ithaca fauna absent from outer shelf, Tully platform carbonates.

A1: LATE MIDDLE DEVONIAN BIOTIC AND SEDIMENTOLOGIC EVENTS IN EAST CENTRAL NEW YORK-TULLY FORMATION CLASTIC CORRELATIVE SUCCESSION IN THE SHERBURNE-ONEONTA AREA

STOP 1. Tully Shale Borrow Pit at Tully Center (20 minutes).
Windom Fauna and Basal Tully Contact.
This well known stop for professionals and amateurs alike is best known for diverse taxa of the Hamilton Fauna that occur in strata of the Windom Member. Windom divisions exposed here include; in ascending order: The Taunton submember (25 feet); the Lansing Coral Bed (1.5 feet); the Spezzano submember (5 feet); the Gage Gully submember (8 feet); and the Sheds submember (6 feet). The Taunton submember is a shoaling (RST) interval. The Lansing probably marks the base of a subsequence interval. All of it (or part of it) commences a transgressive (TST) succession that continues through the Spezzano submember and into dysoxic early highstand facies of the lower Gage Gully submember. The Sheds marks the return of shoaling condition followed by another transgression recorded by the newly defined Highland Forest submember (here absent due to erosion) that is represented by dysoxic Emanuella-bearing strata (Baird and Brett, 2003). Diverse Hamilton taxa (too numerous to discuss here) are characteristic of the shallower facies. Dysoxic taxa, represented by Eumetabolatoechia ("Leiorhynchus") multicostatum, Emanuella praeumbona, and diminutive Allanella tullia are characteristic of the dark gray Gage Gully submember. Not discussed in most earlier reports is the occurrence of the basal layer (DeRuyter Bed) of the Tully Formation at the top of this section. This bed is represented by a 0.3 foot thick layer of impure limestone abounding in the characteristic lower Tully strophomenid Rhyssochonetes aurora and scattered phosphatic pebbles. This layer rests disconformably on the Sheds submember; at Junes' Ravine, 1.3 miles east-northeast of this section, the topmost division of the Windom Member (here eroded) is represented by the Highland Forest submember. At Junes' Ravine it is represented by 0.9 feet of silty, calcareous shale rich in Emanuella, small Tropidoleptus, small Allanella, and Mucrospirifer.

View of Tully Lakes (Kettle Lakes) to left (south).
Song Mountain Ski complex in distance. For the next 1.8 miles we will cross the Tully Valley on the Valley Heads Moraine. Notice the elevated drift-floored plain to the south and the approaching bedrock valley wall to the west.

STOP 2. Tully Limestone Divisions in Driveway Section Next to Carrs' Creek- Quarry Locality.
The Carrs' Creek-Quarry section serves to illustrate the typical, prominent carbonate succession of the Tully Formation in west-central New York. Fortuitous excavation of the new driveway and the kindness of the owners who built it, preclude our need to climb the slippery creek section. Moreover, the driveway cut reveals weathered joint surfaces that show excellent textural details and fossils. The driveway section exposes in upward succession: the Fabius Bed (3 feet visible), the Meeker Hill Bed (3.75 feet), the Tully Valley Bed (1.5 feet), the Vesper Bed (1.5 Feet), the Carpenters Falls Bed lower part (3.4 feet); the upper part of the Carpenters Falls Bed separated from the lower part by a discontinuity (0.5 feet); the Taughannock Falls Bed (2. 7 feet); the Bellona Coral Bed (0.3 feet); and the Moravia Bed (2.75 feet visible). Basal Tully units visible in the creek but not along the driveway include: the DeRuyter Bed, the Cuyler Bed, and the lower part of the Fabius Bed.
The present authors believe that the lower part of the Tully Formation succession (DeRuyter Bed-to-the-intra Carpenters Falls discontinuity) records an interval of overall shoaling with increased Tully Fauna diversity upward (see text; Baird and Brett, 2003). The topmost Carpenters Falls Bed and the Taughannock Falls Bed (medial Tully succession) records a time of relative deepening following the intra-Carpenters Falls lowstand event. This is roughly timed with the demise of the Tully fauna and redevelopment of the Hamilton Fauna (see text; Baird and Brett, 2003). A second regressive episode is recorded by the Bellona Bed. The Bellona "coral plantation" (sensu Heckel, 1973), yielding large rugosans and tabulates such as Favosites and Alveolites at this locality and others further west, records maximum return of the diverse Hamilton Fauna Biota. Actually, the basal 0.5 feet of the overlying Moravia Bed is composed of dark limestone yielding corals and Spinatrypa; we believe that is interval correlates eastward to a limestone bed that caps the West Brook Shale (Figures 5, 6).
The five-six inch thick topmost part of the Carpenters Falls Bed is characterized by "oolitic" grains somewhat resembling chamositic "ooids" in the Smyrna Bed to the east. Moreover, the top of this interval displays fracture systems resembling the "mudcracks" reported by Heckel (1973). Most significant is the presence of laminated carbonate texture resembling algal (microbialite) stromatolites (see text discussion). We believe that these structures underlie Heckel's (I973) regional "intra-Tully unconformity" which we interpret as a flooding surface. We also argue that the topmost part of the Carpenters Falls Bed, as presently defined, overlies a lowstand disconformity and may be the western equivalent of the Smyrna Bed (see discussion).

Type Tully Formation section (Junes' Ravine) to the north (left) in the uphill woods.

Excellent view of glacial U-shaped profile of Labrador Valley to right (south).

STOP 3. Tully Fm sections adjacent to Highland Forest Metropark HQ.
The fortunate location of this section allows for the use of lavatory facilties and time for box lunch consumption. Two Tully sections can be studied, time permitting. The first, below a trail bridge near the parking lot and bordered by paths, will be the main object of attention. This exposes all of the Tully Formation except a portion of the Fillmore Glen Bed-interval near the top of the Tully. In this region, a prominent regional disconformity, first characterized in detail by Heckel (1973) is observed to divide the Tully succession. The lower Tully succession (DeRuyter Bed-into-Carpenters Falls Bed interval) is observed to be overstepped rapidly eastward from the meridian of Highland Forest Park such that no lower Tully is observed in sections east of the meridian of Sheds. This erosion was believed to commence eastward from the level of the post-Carpenters Falls "intraformational unconformity" southeast of Fabius (Heckel, 1973; Figures 2, 3). From DeRuyter Reservoir eastward, this disconformity surface was believed to coincide with the base of the Smyrna Bed (Heckel, 1973) that is traceable to the Sherburne area. The present authors connect the regional, sub-Smyrna disconformity westward, not to the post-Carpenters Falls contact, as did Heckel (1973), but to the base of what Heckel termed "Carpenters Falls Bed" in the Fabius-Labrador Valley area (Baird and Brett, 2003). The thin "Carpenters Falls Bed" interval of the Fabius-Labrador Valley region is a sandy, abraded grain encrinite characterized by a sharp irregular base marked by giant hypichnial burrow prods and minor phosphorite. Careful examination of sections between Fabius and DeRuyter shows that this unit is present in most sections across the region and that it connects laterally to the Smyrna Bed. At this locality, this bed is present, but is of variable thickness; under the walk bridge it is only a few inches thick, but, at the adjacent south fork section, it is almost a foot-thick. This may reflect local Smyrna Bed cutout beneath the sub-Taughannock Falls Bed flooding surface contact. The intra-Carpenters Falls discontinuity at Stop 2 may connect to the sub-Smyrna disconformity.
Another major aspect of the lower Tully succession in this region is that the Fabius Limestone Bed at Stop 2 has transformed into a calcareous siltstone-silty limestone unit and the underlying Cuyler Bed has thickened considerably. In contrast, much of the upper Tully succession consists of clean limestone and clay shale. This is a major facies change described in detail by Heckel (1973). He related it, in part, to terrigenous sediment-sourcing from the region of lower Tully erosional cutout between DeRuyter and Sherburne (Heckel, 1973).

Cross Limestone Creek. Morainal deposits exposed in creek banks.

East valley rim of Limestone Creek.
Valley head moraine plain for next 0.4 miles.

STOP 4. Tully Formation Roadcut and Road Ditch Section
Showing Complete Overstep of Lower Tully Member Division by Sub-Smyrna Bed Disconformity.
This section displays the two uppermost submember divisions of the Windom Member overlain by units of the upper Tully Formation succession. The upper part of the Sheds submember is visible in the lowest 2-3 feet of the bank and yields numerous Hamilton Fauna brachiopods and clams. This is abruptly overlain by 2.5 feet of Highland Forest submember that is characterized by siltstone beds and a low diversity association of Emanuella praeumbona, small Tropidoleptus and Mucrospirifer. This section exposes the highest strata of the Windom below the Tully in central New York, although considerably thicker Highland Forest sections are observed in central Pennsylvania (Baird and Brett, 2003).
The visible Tully succession rests disconformably on the Windom. At this locality the basal 16 inches of the Tully is represented by the Smyrna Bed, a hard overhanging ledge composed of abraded calcarenite and minor "oolitic" chamosite. This bed displays a sharp base with development of large hypichnial burrow prods filled with encrinite, minor sand, minor phosphate, and chamosite. Lucky persons will find clusters of 3-D Tullypothyridina in some of the basal protruding burrows. As noted by Heckel, 1973, the lower Tully succession is absent due to regional overstep by the sub-Smyrna disconformity. This overstep is characteristic of Tully sections eastward to the Sherburne area. Post-Smyrna strata are represented by an expanded shale-limestone phase of the Taughannock Falls Bed interval. In the road ditch, strata of the West Brook Shale and its capping limestone unit can be seen. These units yield many Hamilton Fauna taxa. Excellent trilobite fossils, camerate crinoids and blastoids are found in this interval.

Pass Spirit House on left.
Special ornamentation and rounded corners were devices designed by owner to fool the Devil.

Notice major exposure of medial Windom Member/Cooperstown Member
Notice major exposure of medial Windom Member/Cooperstown Member strata on floor of Pleasant Brook to the right.

More visible Moscow Formation strata on floor of Pleasant Brook.
Major Tully Formation exposure (Upperville section: Locality 1) is across brook on along a tributary.

Pass Rogers Nature Center.
Geese, giant carp, and snapping turtles can be viewed from a bridge near the parking lot.

Cross Chenango Creek.

STOP 5. Thin, condensed Tully Formation Section
On "Sherburne High" Structural uplift.
This is an important section (Locality 9) described in Cooper and Williams (193S) and Heckel (1973) publications. This is the type section of the West Brook Shale, a nearly three foot-thick unit of calcareous shale abounding in diverse Hamilton Fauna taxa recorded in extensive faunal lists (Cooper and Williams, 1935, Heckel, 1973). It is the "last hurrah" for the Hamilton Fauna prior to the gradual onset of transgressive anoxia associated with Geneseo black shale deposition.
Eastward erosive downcutting of Hamilton Group strata by the sub-Smyrna Bed disconformity places the Cooperstown Member-Tully Formation contact on the Spezzano submember interval of the Cooperstown (Figures 5, 6). The Smyrna Bed at this locality is a 4-5 inch-thick unit of"oolitic" chamosite and subsidiary sandy, bioclastic abraded calcarenite that rests disconformably on underlying Hamilton Group deposits (Figures 5, 6). Discoidal dark gray to black, sand-sized chamosite grains dominate over subsidiary phosphorite pebbles and rare shells within the bed. Though not encountered by the present authors, Heckel (1973) reports a cluster of Tullypothyridina within the Smyrna Bed. Above the Smyrna Bed is a 2.2 foot interval of hard, slabby siltstone or fine sandstone facies yielding few fossils. Heckel (1973) referred to this unit as the "unnamed sandstone." Above this unit is a complex interval of "oolitic" chamosite, tabular siderite and limestone comprising the interval between the "unnamed sandstone" and the basal contact of the West Brook Shale (Figures 5, 6). Heckel (1973) referred to this interval of the "Taughannock Falls Oolite Bed" and indicated that it was a condensed lateral equivalent of the Taughannock Falls Bed (Heckel, 1973).
Careful work by Baird and Brett (this paper) shows the probability that Heckel's (1973) correlation is correct. However, the disjunct character of section match between localities 3 and 4 (Figure 5A) and localities 12 and 13a, b (Figure 6) offer the possibility that the Smyrna Bed near Smyrna (Localities 1, 2) and Sherburne Four Corners (Locality 3) may actually connect to the "Taughannock Falls Oolite Bed" rather than the lower chamositic layer (Figure 5A). Thus, from Cush Hill Road (Locality 4) to Harrisville (Locality 12), it is possible that a part of the lower Tully Member (Fabius Bed?) may be represented by the lower chamositic layer and the "unnamed sandstone" (Figure 5, 6). Both chamositic beds are well developed at this locality. Chamosite, an iron-rich silicate, of somewhat problematic origin is characteristic of several Tully (and even pre-Tully) levels in east-central New York. The absence of chamosite in western New York and east-central New York sections shows the occurrence of this mineral to be restricted to a particular depositional belt (see text discussion). The absence of common robust fossils in chamositic beds, the mud-rich character of some chamosite-rich layers, and the occurrence of chamosite in dysoxic facies yielding "leiorhynchid" -type rhynchonellid brachiopods suggest that the origin of this unusual deposit may be some physical (or diagenetic) process occurring across a range of outer shelf and dysoxic, sediment-starved environments (see discussion).

Cross Mad Brook. Upper part of Cooperstown Member exposed below bridge.

Major "Perrytown" section (Locality 17)
Fairkit Farm to left. Major "Perrytown" section (Locality 17) of Tully clastic correlative succession (see Cooper and Williams, 1935; Heckel, 1973) high up on hill to the south.

STOP 6. Thick Tully Clastic Correlative Succession Along Unnamed Creek
(Locality 18) parallel to, and east of Balcom Hill Road.
Cooper and Williams (1935) and Heckel (1973) devoted considerable discussion to the first thick clastic Tully outcrop ("Perrytown" section) that was encountered proceeding east from the area of thin Tully occurrence on the "Sherburne High". That outcrop (Locality 17) is too far from any road to be visited on this trip. However, another adjacent section (Locality 18; Figure 7) that is almost as complete is herein included as a stop. We will intersect this creek at the topmost exposed Cooperstown Member above a small falls. The upper coral layer of the Lansing Bed is 15 feet below the falls lip and the strata comprising the falls is probably equivalent to the Spezzano submember in the Windom (Figures 7, 11). A six foot covered interval follows the Cooperstown succession that is, in turn, succeeded by 5 feet of lower Tully strata and about 45 feet of continuous upper Tully section. At the adjacent "Perrytown" section (Locality 17) part of the New Lisbon Member is exposed, as is the base of the Tully clastic succession above the New Lisbon (see text; Figure 7). The basal Tully contact on that creek is characterized by chamosite, nodular diagenetic siderite, and Tully Fauna taxa (Hypothyridina, Schizophoria, Spinatrypa), probably corresponds to what we call the "base-Laurens" disconformity (see text discussion). Heckel, 1973, (Figure 3) tentatively correlated this bed to (or near) the position of the Smyrna Bed in the Sherburne area.
However, on that same creek, and on this section as well, there is a second, more complex and prominent bed characterized by calcareous, auloporid-rich sandstone in the lower part and nodular chamositic limestone and siderite in the upper part (Figures 7, 10). The present authors correlate the upper part (and tentatively the lower part) of this unit to the Smyrna Bed instead (Figures 7, 11 ). The coarse, hashy mix of fossil debris and the nodularity of the carbonate component are texturally similar to the Smyrna Bed though the unit is sandier overall.
Although the "base-Laurens" disconformity is covered here and the underlying New Lisbon is concealed (or absent due to overstep: see text discussion and Figures 7, 11), we will see higher divisions of the lower Tully clastic succession. A massive two foot-thick siltstone bed exposed above the covered interval may correspond to part of the Fabius Bed. Above this, participants can view the 8-inch-thick nodular, calcareous auloporid-pelmatozoan-rich sandstone ledge of the lower Smyrna Bed. Calcareous chamosite and diagenetic siderite are characteristic of the upper Smyrna (Figures 7, 10). This is a condensed, presumably transgressive, phase probably corresponding to a major mid-Tully deepening event. In this interpretation the lowstand (sub-Smyrna) erosional contact is understood to be at the base of the chamositic bed or at the base of the 8-inch auloporid-rich sandstone (Figures 7, 10).
Above the Smyrna Bed is a 37 foot-thick interval of flaggy to slabby thin siltstone beds with shale interbeds. With exception of a thin medial layer, the Walt Phillips Bed, this interval yields very few macrofossils. Auloporid corals, sparse pelmatozoan debris, plant fragments and the occasional mollusk comprise the observed biota. We will view this interval at greater advantage at Stop 7. The present authors correlate this thick "barren" unit to the Taughannock Falls Bed in the Tully Limestone, although the upper part of it may be missing due to regional overstep to the west of this area (Figure 11).
One may properly question correlations based mainly on lithologic and textural parameters and minimally on distinctive guide fossils. We acknowledge this problem with numerous caveats, but have to point out that lower and medial Tully deposits in this area are conspicuously poor in fossil content. Although Tully Fauna taxa occur both in the lower and medial Tully, they are uncommon to rare. Even in the Sherburne area, the chamositic beds contain few shells and these are mostly small. The only unifying taxon that characterizes this facies is the auloporid coral that abounds in the condensed units. In fact much of the Tully clastic correlative succession in this area resembles portions of the Penn Yan and Sherburne members of the Genesee Formation. It is, thus, not surprising that Thayer (1974) described the biota of the dysoxic prodelta slope facies of the Genesee Group as the "Cladochonus" (auloporid) biofacies. As such, we interpret this region to be a structural subsiding basin relative to a platform setting for the Tully Limestone further west.
Another line of evidence for subsidence of this area is the dramatic thickening of Tully units and the reappearance of lower Tully beds east of the "Sherburne High" (Heckel, 1973). Heckel (1973) correctly envisioned a "down-to-the-east" fault and a clastic trap within the basin that prevented clastics from overspreading the Tully platform. We confirm this model (Figures 5 - 8) but show that lower Tully beds are, again, overstepped to the southeast of this section before re-emerging to the east of New Berlin (Figures 6, 8). At the upper end of this transect, we will see the West Brook Bed and a capping discontinuity probably corresponding to the base of the Moravia Bed. The West Brook is richly fossiliferous and astonishingly thin relative to surrounding units. Corals, diverse brachiopods, bryozoans and other stenotopic Hamilton Fauna taxa can be found here. The West Brook has an erosional base marked by burrow prods and encrinite; this contact marks a significant lowstand event and it is responsive for conspicuous overstep of the medial Tully section west of this locality (Figures 6, 7).

STOP 7. Walt Phillips Bed and Associated Medial Tully Strata
On Roadcut Along Walt Phillips Road (Locality 16).
This short stop serves to show the sparsely fossiliferous medial Tully clastic interval in good light. In particular, this outcrop is the type section of the Walt Phillips Bed a 16 inch-thick double-ledge interval of bioturbated siltstone that can be regionally correlated across the Columbus area (Figures 6, 7). At this section, this unit is at least 10 feet below the sub-West Brook disconformity. At locality 15, 0.5 miles northwest of this section, the Walt Phillips Bed is 18 inches below that contact and at locality 17, 0.9 miles east of this section, it is 17 feet below the West Brook. This pattern shows that there is a major trend of westward erosional overstep of the medial (Taughannock Falls Bed equivalent?) Tully succession below the sub-West Brook disconformity (Figures 6, 7, 11).
The Walt Phillips Bed is notable for the highest occurrence of Rhyssochonetes aurora in this region. It is a rare component along with Emanuella subumbona and scattered auloporids within the bed. This distribution shows that a diminished Tully Fauna component did survive within the Tully well past deposition of the Smyrna Bed. In the creek below the road is an incomplete in-place succession of lower Tully beds probably corresponding to the Fabius Bed succession to the west. In addition to this, however, are large loose Tully blocks derived from the till upstream from the actual outcrop. One of these blocks, resembling the lower Smyrna auloporid-rich sandstone unit at Stop 6, is capped by a thin, chamositic limestone bed yielding abraded calcarenite with auloporids, and rare Rhyssochonetes. Moreover, this bed displays "mudcrack" -like fractures and well-developed microbialite stromatolites similar to those originally described by Heckel (1973) along the intra-Tully discontinuity at Bellona and those observed by the present authors in the Tully Valley (see Stop 2). We believe the loose block to be a Smyrna Bed erratic and that the stromatolitic bed is associated with the top Smyrna transgressive flooding surface (see text for discussion).

Strata of lower part of Cooperstown Member
Strata of lower part of Cooperstown Member in stepped ramp falls in floor of Center Brook to the right.

Cross Otselic Creek

Cross Unadilla Creek, leave New Berlin and enter Otsego County.

Creek to left displaying Cooperstown Member
Creek to left displaying Cooperstown Member strata including two separated ledges of the Lansing Coral Bed.

Borrow pit on right displaying strata of Sherburne Member.

Upper Tully Fm clastic correlative succession
Upper Tully Fm clastic correlative succession in creek (Locality 26) below road on left.

Borrow pit on right displaying Sherburne Member strata.

Butternut Valley to right.

Cross Butternut Creek.

STOP 8. New Lisbon Member Type Section
And Higher Tully Clastic Correlative Units in the Vicinity of Otto Stahl Road.
This section (Locality 34 of present paper) was the focus of detailed study by Cooper and Williams (1935) particularly for definition of the New Lisbon Member. However, a long and nearly complete section of the Tully clastic correlative succession is also present along and near this creek. Cooper and Williams (1935) established the New Lisbon Member as a 60-foot-thick succession of flaggy siltstone beds and shale interbeds rich in the brachiopod Camarotoechia ("Leiorhynchus") mesacostale. The lower-middle part of the type section of this unit is exposed on adjacent Stony Creek (Locality 32) and the upper 38 feet of Cooper and Williams (1935) original "New Lisbon Member" is present at Stop 8. Cooper and Williams (1935) and Heckel (1973) both recognized that this unit was pre-Tully in its lower half and Tully equivalent (yielding Rhyssochonetes and Tullypothyridina) in its upper half. Thus, this unit filled the interval between the underlying Cooperstown Member, yielding Pustulatia ("Vitulina"), and fossil-rich strata of the overlying lower-middle Tully succession ("Laurens Member" of Cooper and Williams, 1935; Figures 1, 3).
As noted in the text, the present authors have found a chamosite-bearing bed on this creek yielding Rhyssochonetes, rare Hypothyridina, conularids, and abundant Emanuela 28.8 feet below the base of the "Laurens Member" and 10 feet above the lower end of this outcrop (Figure 7). This unit, located nearly at road level in the creek, is herein taken to be the base-Tully contact (Figures 7, 11). Moreover, a chamosite-rich contact bed is now observed to floor the New Lisbon interval at nearby locality 31 allowing us to formally bracket the unit with physical contacts (Figures 7, 11). With this new information, we herein restrict the New Lisbon Member to include only pre-Tully highstand deposits between the two chamositic beds (Figures 7, 11). The "upper New Lisbon" of Cooper and Williams (1935) is herein included in the basal Tully clastic succession. We agree with Heckel (1973) that the 28.8 foot-thick succession of C. mesacostale-rich slabby siltstones and sandstones in the abandoned quarry and creek walls near the car pull off probably corresponds respectively to the DeRuyter and Cuyler beds in the Tully Limestone. Also significant is the fact that this pre-"Laurens Member" Tully interval appears to be absent in all localities between this section and Sheds (Figures 7, 11).
A few hundred yards east (uphill) from the car pull off are a few small road ditch cuts in the "Laurens Member" succession. The first of these exposes a 20 inch-thick silty mudrock unit abounding in Emanuella, Echinocoelia?, Hypothyridina, Schizophoria, Spinatrypa, and Mucrospirifer. This corresponds to Cooper and Williams, 1935, "second shell bed" of section parlance. 14 feet further up the road is another bed yielding abundant, well preserved Tullypothyridina and Pseudoatrypa. This is Cooper and Williams' (1935) "third shell bed" of the section. Finally, 12 feet above that unit is a thin, shell hash bed rich in Emanuella, Echinocoelia?, large burrow prods and phosphatic pebbles. We tentatively link the first shell bed in the road ditch to part of the Fabius Bed and one (or both) of the two higher shell beds to the Carpenters Falls-Smyrna bed interval in the Tully Limestone (Figures 7, 11).
Higher strata of the clastic Tully succession can be viewed only from private land in the creek. Hence, we will not see this part of the section. It is worth noting that the West Brook Bed is spectacularly developed on the south fork of this creek and is floored by an eleven-inch-thick clean, crystalline calcarenitic limestone bed yielding rugose and tabulate corals. It appears that there is eastward overstep of medial Tully strata by the sub-West Brook disconformity beneath this limestone layer.

Strata of medial Tully Fm clastic correlative succession visible in creek falls to the right.

Cross north fork of Stony Creek tributary.

Entrance to Gilbert Lake State Park on the left.

"Houghtelings Glen" section of Cooper and Williams, 1935, to right.
This outcrop (Locality 38) exposes an excellent lower Tully Formation clastic correlative succession (see text) and was intended as a stop. However, a major windstorm has tumbled trees blocking the stream course for the immediate future.

Region of smashed and toppled trees.
This was the result of the same tornado (or downburst) that flattened trees at "Houghtelings Glen."

Hamlet of Milford Center. Susquehanna Valley to left.

View of Goodyear Lake to left.

STOP 9. Tully Clastic Correlative Succession
(or Tully equivalent Gilboa Formation) on South Side of Route 28 by Goodyear Lake.
This is a "floating section" probable corresponding to part of the middle-upper part of the "Laurens Member" (Cooper and Williams, 1935). About 35 feet of section is exposed; spectacular flow rolls (seismites?) are visible in the lower third of the section; about two thirds-up is a 20-inch bed rich in Tullypothyridina, Echinocoelia, Mucrospirifer, Spinatrypa and bivalves. In the topmost 5 feet, pelmatozoan debris is prominent at several levels. This section, and, to a lesser extent, Stop 8, stand in stark contrast to the sparsely fossiliferous sections (Stops 5, 6, 7) of lower-middle Tully facies in the Sherburne-New Berlin area (Figures 7, 8, 11). We hypothesize that a structural, bathymetric trough centered at New Berlin separated the Tully carbonate platform in west-central New York from a sandy shelf area in eastern New York (Figure 12). Hence, the coarse, fossil-rich, Tully clastic deposits of the Otego and Susquehanna Valleys corresponds to the "Chemung" facies belt for the Taghanic Stage (Baird and Brett, 2003). Current explorational work is directed to localities in this valley and further east to characterized both the biofacies and stratochronology of this poorly known interval. We hope to determine whether inner shelf biotas at the Taghanic level have a unique taxonomic composition or correspond to the Ithaca Fauna of overlying units. A great final mystery is centered on the relationship of various marine units (Cooperstown Member, clastic Tully succession, and Geneseo Member) to the Gilboa Formation. This problem is beyond the scope of the present report, but is the subject of ongoing investigation (see Bartholomew and Brett ???).

Sunday extra stop

STOP 10. Tully Fm Section on Taughannock Creek.
Type Section for the Taghanic Stage.
This outcrop provides an excellent look at the Tully Formation section and adjacent units. Several feet below the base of the Tully is the westernmost occurrence of the Lansing Coral Bed; Favosites, Heliophyllum and other corals can be found as well as large fistuliporoid bryozoans and diverse brachiopods. The Spezzano submember is recessive below the disconformable base of the Tully. Large hypichnial burrow prods characterize the basal Tully surface where the Hamilton shale has eroded away below it. The remaining Tully succession to the base of the 5-foot-thick Fillmore Glen succession is massive, dove-white weathering fine limestone that is clean enough for quarrying as is done at Portland Point across the lake. Key interval units and contacts will be pointed out at this stop.
This is the key reference section for the Taghanic provincial stage, the Taghanic Unconformity and the Taghanic bioevent (see Johnson, 1970; Heckel, 1973; Aboussalam and Becker, 2001; Baird and Brett, 2003). However, it is significant that this Tully section is very incomplete due to westward erosive beveling of most lower Tully strata between Skaneateles Lake and this locality (Heckel, 1973). As such, the Carpenters Falls Bed, yielding Tullypothyridina and other lower Tully taxa, rests disconformably on the Hamilton. This is Johnson's (1970) "Taghanic Unconformity" which was believed to mark the start of a major craton-ward stratigraphic onlap event. Careful examination of higher Tully in western New York suggest, however, that the lowstand base-line for the transgressive Taghanic onlap event is at a higher level within the upper Tully Member or near the top of the Tully (Baird and Brett, 2003). The lowstand base-line for the upper Tully TST interval is the regressive Bellona Coral Bed and the actual transgressive onlap is understood to commence at one or more horizons bounding (or within) the Fillmore Glen Bed (Baird and Brett, 2003). The top-Moravia Bed omission horizon (knobbly surface) and five calcilutite-shale repetitions comprising the Fillmore Glen succession are well displayed at this section. Note the black, fissile strata of the Geneseo Member in the bank walls. This unit records major transgressive deepening in the foreland basin owing to combined effects of eustatic transgression and thrust loading to the east (Johnson et al., 1985; Ettensohn, 1998). Higher outcrops of this unit and overlying deposits, exposed in the great fall on this creek, are worth seeing, but are too far upstream to be accessed on this trip.

STOP 11. Tully Limestone section on Kashong Creek
Stromatolites Above Mid-Tully Unconformity; Bellona Bed; Undulatory Top-Tully Erosion Surface.
We will descend into Kashong Creek upstream from the bridge and proceed to a breached up-fold of Tully in the creek floor. Along the contact between the underlying Carpenters Falls Bed and the overlying Taughannock Falls Bed is a layer of fracture networks, domal limestone knobs and stromatolitic carbonate deposits filling low areas peripheral to the knobs. This is Heckels', 1973, stromatolitic horizon associated with his regional intra-Tully unconformity. Close examination of the stromatolite texture shows that the lamination is arrayed concentrically around the structureless limestone knobs and parallel to fracture networks. Very close examination of the laminite shows that it has a perpendicular "pillared" microstructure that is normal to the visible lamination, a characteristic that Heckel, 1973, also noted. The stromatolitic layer is usually only millimeters-thick in swales and it is erosionally breached over knobs. The base of the stromatolitic layer contains skeletal debris, pyrite, phosphate grains and some fish bones; this marks the horizon of the discontinuity surface. It is important to compare the stromatolitic surface at this locality with material from Stops 2 and 7 as these latter occurrences are believed to be essentially the same level.
Several feet above the stromatolitic horizon is the Bellona Coral Bed which is thin, muddier than surrounding Tully facies, and profusely fossiliferous. Where it is swept clean on bedding surfaces, hundreds of corals can be seen. Easily spotted taxa within the bed include: Heliophyllum, Cystiphyllum, Favosites, Alveolites, and the brachiopod Spinatrypa. This is the "coral plantation" of Heckel, 1973, that serves as a key datum in mapping over huge areas. Above the Bellona Bed are clean carbonate deposits of the Moravia Bed. However, Heckel, 1973, correctly noted that the top-Tully contact at this locality is undulatory. Up to a foot of Moravia is present above the bridge, but, below the bridge, the Geneseo black shale rest directly on coral-rich Bellona deposits.
The Geneseo Member rests on the Tully with knife-sharp definition. Tully carbonate is locally separated from the well-jointed Geneseo by a thin lag bed of pelmatozoan calcarenite, but not detrital pyrite or bone debris. However, at Gorham, the next section to the west, detrital pyrite and bone debris characteristic of the Leicester pyrite rests on the Bellona Bed. Still further to the west, the base-Geneseo corrosional discontinuity is observed to cut out the Tully Formation and truncate the topmost beds of the Windom Member (Baird and Brett, 1986). The occurrence of channel-like grainstone lentils within the Geneseo Shale 16 inches above the Tully contact at Bellona, suggest the possibility that the real Leicester horizon is actually within the basal Geneseo on this creek and that the observed Tully-Geneseo contact at Bellona is a pre-Leicester surface that may project eastward to some level within (or at the base of-) the Fillmore Glen Member. This issue is still being investigated.

A2: COASTAL MARGIN INTERFLUVE PALEOSOLS AND THEIR STRATIGRAPHIC RELATIONSHIPS WITH TIDALLY-INFLUENCED DELTAIC DEPOSITS IN THE SONYEA GROUP (FRASNIAN) OF NORTHWESTERN DELAWARE COUNTY, NEW YORK
Daniel Bishuk Jr.
EA Science and Technology
6731 Collamer Road
East Syracuse, New York 13057
danbishuk@eaest.com
Jon Hairabedian
New York City Department ofEnvironmental Protection
West-of-Hudson Office ofEngineering
P.O. Box370
Shokan, New York 12481
James R. Ebert
Earth Sciences Department
SUNY College at Oneonta
Oneonta, New York 13820-4015
Introduction
The Catskill Clastic Wedge has been broadly subdivided into three magnafacies: Chemung, Cattaraugus and Catskill. These magnafacies have been interpreted as representing offshore, coastal/nearshore and terrestrial (fluvial) paleoenvironments, respectively. Of these magnafacies, the Cattaraugus has been the most poorly characterized and the nature of the Catskill shoreline has long been the subject of debate. Tectonic, eustatic and climatic processes interacted in complex ways to produce nearshore paleoenvironments in the Catskill Wedge which had aspects of both marine and terrestrial settings.
In this study, we describe tidally-influenced sediments of the Cattaraugus Magnafacies in parts of the Sonyea Group (Late Devonian, Frasnian) in Delaware County, New York (Fig. 1). However, our main focus is on paleosols in these strata as key features in understanding the complex nearshore environments which bordered the Catskill Sea. Several types of paleosols are documented, which yield insights into depositional and climatic processes at the Sonyea shoreline. More significantly, we show that paleosols, a typical feature of the Catskill Magnafacies, developed directly on sediments deposited under brackish, deltaic/estuarine conditions and on definitively marine sediments. In addition, many of the paleosols that we describe are immediately overlain by sediments deposited under marine or brackish conditions. Such abrupt paleoenvironmental changes provide insights into processes of tectonic subsidence, eustatic changes in sea level and autocyclic processes, such as avulsion, which controlled the nature and position of the Catskill shoreline.

A2: COASTAL MARGIN INTERFLUVE PALEOSOLS AND THEIR STRATIGRAPHIC RELATIONSHIPS WITH TIDALLY-INFLUENCED DELTAIC DEPOSITS IN THE SONYEA GROUP (FRASNIAN) OF NORTHWESTERN DELAWARE COUNTY, NEW YORK

STOP 1. Skytop Lane Quarry (outcrop 43)

Maywood Historical Depot -- Lunch

STOP 2. Sheetz Quarry (outcrop 51)

STOP 3. Sidney Mountain Quarry (outcrop6)

A3: GLACIAL REGIME AND DEPOSITIONAL ENVIRONMENTS ALONG THE RETREATING LAURENTIDE ICE SHEET, NORTHEASTERN APPALACHIAN PLATEAU, NEW YORK
P.J.Fleisher
Earth Sciences Department
SUNY-Oneonta
ABSTRACT
Deglacial environments along the retreating Laurentide Ice Sheet throughout of the upper Susquehanna region were primarily controlled by an ice front configuration draped across the high relief, Appalachian Plateau terrain. Valley ice tongues extended 20 km beyond upland positions commonly terminated in ice-contact lakes. Active ice flow nourished by the ice sheet persisted within through valleys, where as ice tongues in non-through valleys became detached and stagnated. Temperate conditions and hydrologic connection with the ice sheet supported highly turbid englacial and subglacial tunnel discharge directly into through valley ice-contact lakes. Subaerial streams favoring positions lateral to ice tongues transported sand and gravel outwash to prograding deltas. Thus, bottomset and toeset sediments interfinger with lacustrine silt and sand. Post-glacial dam incision followed by lake emptying led to the modern landscape. A finer facies of thick silt and sand lies beneath the modern floodplain and is typically bound by deltaic terraces (commonly called kame terraces) consisting of the coarser facies sand and gravel.
In contrast, downwasting over non-through valley headward divides led to loss of nourishment and near separation of the ice tongue from the ice sheet, thus causing the combined effect of hydrologic disconnection and stagnation. Consequently, non-through valley environments of deposition were less uniform, characterized by a) local pending behind ephemeral ice-cored dams, b) kilometer-size, detached ice masses partially covered by outwash and lake sediments, and c) meltwater sources limited to the volume of the detached ice tongue. These collectively describe an environment significantly different from through valleys. Non-through valley sediments are far more diverse, less well sorted, and lack the laterally uniform subsurface conditions characteristic of through valleys. In addition, the landform assemble also differs through the lack of a continuous lacustrine plain, segmented terraces graded to local base levels, isolated small esker-forms, kame and kettle topography unrelated to systematic moraines (called kame fields), and mega-kettles that span the valley floor (known as dead ice sinks).
The Bering Glacier, Alaska, depicting sedimentary conditions in ice-contact lakes and along peripheral drainage systems is cited as an analog for the retreating Laurentide Ice Sheet in central New York.

A3: GLACIAL REGIME AND DEPOSITIONAL ENVIRONMENTS ALONG THE RETREATING LAURENTIDE ICE SHEET, NORTHEASTERN APPALACHIAN PLATEAU, NEW YORK

I-88 parallels the Susquehanna River for the next 2.4 miles.
Active floodplain aggradation mantles lake sediments of Glacial Lake Otego that was dammed during retreat by the Wells Bridge Moraine, our first stop.

Good view of the modern flood plain and change in valley trend
The highway rises above the valley floor and provides a good view of the modern flood plain and the abrupt change in valley trend that is a remnant of a preglacial ingrown meander. A bedrock promontory on the horizon to the right (north) protrudes into the valley along the inside of the meander bend.

Stop 3 is on the left (south), but we won't stop now.
Access to this area is possible from County Rd. 48 (locally referred to as the Otego-Wells Bridge Rd.), which we will take on our return to the Oneonta area from Wells Bridge.

Continue west on I-88 past Rt. 7 & Otego exit.

View to the west of valley plug formed by the Wells Bridge Moraine.

Back of the Wells Bridge Moraine and an associated kame terrace.
View to the right (north) across the valley includes the back of the Wells Bridge Moraine and an associated kame terrace. The next 1.2 miles provides excellent overviews of the moraine and the gap cut by the Susquehanna River.

STOP 1. Wells Bridge Moraine (Franklin and Unadilla Quadrangles).
The hummocky relief of this moraine is common for other moraines in the upper Susquehanna drainage basin. This moraine completely blocked the valley following glacier retreat thus forming the dam for Glacial Lake Otego that extended about 25 km upvalley to Oneonta. Logs of water wells from the floor of the valley penetrate more than 400 feet of silt without encountering bedrock. Rates of sedimentation from similar ice-contact lakes at Bering Glacier, Alaska, related to vertical settling of suspended sediment between 0.6 and 1.2 m/yr. Therefore, Glacial Lake Otego may have been relatively short-lived.
The configuration of the Laurentide ice front during retreat is commonly thought to have consisted of valley ice tongues. Early work portrayed relatively short ice tongues, where as more recent evidence suggests that they may have reached 20 km in length. This raises implications of sediment source areas and transport mechanisms, while opening the topic of mechanisms of moraine formation. Considering that most moraines in the region consist of stratified drift and lack topographic expression other than in valleys, alternatives to conventional moraine formation may be entertained.
The Wells Bridge Moraine is assumed to have been emplaced about 15,000 years BP and breached about 14,000 years BP. We will consider the field evidence for an 1140 feet lake level at stop 3. Field work in the Unadilla and Sidney areas indicates that the Upper Susquehanna Lake Chain has greater downvalley extent than will be covered in this road log.

Cross Ouleout Creek.

Highway parallels the river for one mile.

Highway climbs onto outwash terrace near mouth of Sand Hill Creek.

Highway drops into Sand Hill Creek incision of outwash
Highway drops into Sand Hill Creek incision of outwash and immediately climbs to follow the contact of the Wells Bridge Moraine and outwash.

Crest of moraine on the left, breach on the right.
It was within the breach of this moraine that archeological excavations uncovered charcoal in silt covering a buried river point bar. This was dated at 13,000 to 14,500 years BP, thus providing limiting age of Lake Otego.

Road parallels river for 0.6 miles before rising onto an outwash terrace.
A correlative terrace can be seen across the valley to the north at an elevation of about 1140 feet.

Suggest a second level for Lake Otego below 1140 feet.
White house across the valley to the north is situated on a terrace at about 1120 feet. Other planar landforms can be found along the valley that suggest a second level for Lake Otego below 1140 feet.

STOP 2.
Massive red sandstone outcrop on south side of parking lot is covered with glacial abrasional features on joint faces that parallel the valley wall. Note that striae and grooves are inclined in the downvalley direction, thus indicating topographically controlled ice flow consistent with valley ice tongue model. Basal flow of overriding ice would not have produced such features. Several joint block surfaces that do not parallel the valley trend are also polished, but not by ice flow. These very smooth and polished surfaces are the product of abrasion by highly turbid water flow and support the notion of a highly charged subglacial hydrologic system.

Gravel excavation on right contained deltaic foreset and topset beds
Gravel excavation on right contained deltaic foreset and topset beds indicating a downvalley current direction.

Gravel excavation on left contained deltaic foreset and topset beds
Gravel excavation on left contained deltaic foreset and topset beds indicating a downvalley current direction.

Road drops to modern flood plain
Road drops to modern flood plain, which is superimposed on Lake Otego lacustrine plain. Lacustrine plain continues to the left and right.

Kame on the right.
Prior to construction of I-88 a bit smaller feature could also be seen to the northeast.

STOP 3.
Park on the right beyond the overpass and walk south between I-88 and the forested slope. About 300 meters south of overpass and upslope from the fence, at an elevation of approximately 1120-1140 feet (2/3 the way up the forest-free slope), pebbly coarse sand, fine sand, silt, and a few clay seams were exposed during construction. Fluvial channel structures with small scale foreset beds inclined into the slope and down valley rested upon finely laminated, rippled and cross bedded silt and fine sand. These sands are interpreted to have formed along the strandline of Lake Otego by wave generated currents that moved into this valley wall alcove. The sand about 1140 feet lacks pebbles, is considered to be of eolian origin, blown up slope from the beach.

Road drops back down to the lacustrine plain at an elevation of 1,060 feet.

The highway traverses an outwash/alluvial bench
The highway traverses an outwash/alluvial bench between 1100 and 1120 feet at the confluence of Otego Creek and the Susquehanna River.

Dead-ice sink on right

Unadilla River

STOP 4.
Through valley landforms (New Berlin South Quadrangle)- high terrace (1180') consisting of deltaic valley train marking the ice margin (no moraine) at the head of which is a dead-ice sink. Although excavation in this quarry has been inactive for several years, large-scale, deltaic foresets are still expressed through the colluvium. The nearest base level to which these deposits may be graded is 22 km downvalley at Rock Wells Mills (Guilford Quadrangle), where kame and kettle, morainic topography may have served as a drainage plug in a constricted segment of the valley. A partially preserved hummocky, kame terrace at White Store (16 km downvalley) also suggests ice burial.
Earlier exposures here included interbedded, very coarse, clast supported boulder gravel and pebbly, coarse sand. A carbonate cement in some units is related to limestone clast weathering, leaching, and reprecipitation at depth, thus indicating the "bright" nature of this outwash.
The Unadilla River is incised within a valley train and flows from a dead-ice sink upvalley to a lacustrine plain downvalley. These landforms and their internal features are indicative of active ice retreat with glaciofluvial sediment transport into an ice-contact lake. Partial burial of a detached, large remnant ice mass caused retarded melting until after meltwater sedimentation ceased, thus allowing collapse and topographic preservation of a dead-ice sink. Water well logs within the sink verify the lack of subsurface, stratigraphic continuity.

Road rises on gravel terrace above lacustrine plain

A high-level terrace at 1280' to 1300' on west side of valley
A high-level terrace at 1280' to 1300' on west side of valley continues for 1.5 miles west of South Edmeston. It appears to be an older kame terrace with no apparent downvalley base level, thus suggesting ice served that purpose.

Road rises on pitted planar gravel and kame field

Inwash sediment source and topographic expression of kame field, Columbus Quarter (New Berlin North Quadrangle). The kame field and associated pitted plain may have served as a local, temporary dam (1220') for a lake upvalley. Pebble counts from upland "drab" drift now resting on the "bright" drift of the kame field suggest tributary inwash from Tallette Creek as a source of sediment (see plot of pebble count data, after Yuchniewicz, 1996).

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 Oaksville Moraine.

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.

Quarry operation in sand and gravel valley train
Quarry operation in sand and gravel valley train with classic assemblage of glaciofluvial sedimentary structures on the left (east). The road continues along the floor of a dead-ice sink

STOP 6. Cassville-Cooperstown moraine, at Index
(5 km south of Cooperstown) (Cooperstown Quadrangle).
This classic through valley landform assemblage (a moraine linked with an extensive valley train) represents active ice deposition. Old excavations in a nearby borrow pit exposed stratified sand and gravel. Test borings indicate that the gravel of the moraine interfingers with silt upvalley, toward Cooperstown, before yielding entirely to silt (120+ feet thick, Fleisher, 1992). The classic kame and kettle expression of this moraine loses all topographic expression as it rises to the uplands, as is characteristic of moraines in central New York State. This has implications for a moraine forming mechanism that would involve sediment transport dominated by glaciofluvial processes.
Based on topographic criteria, Krall (1977) correlates the Cassville-Cooperstown Moraine with moraines in the Hudson Valley, thereby suggesting a readvance circa 14,000 years BP. Ridge, et al., (1991) makes a tentative correlation with the West Canada Readvance of late Wisconsinan, pre-Valley Heads age (circa 15.5 ka) based on elevation projections of ice-marginal deposits in the Mohawk Valley and the limit of readvance inferred from subsurface stratigraphy (Fleisher, 1986).
Although high enough, this moraine did not dam Glacial Lake Cooperstown, although landforms and stratigraphy suggest that a proto-Lake Cooperstown may have extended to the moraine by "swamping" a dead-ice sink.

Dead-ice sink
Dead-ice sink between Cassville-Cooperstown Moraine and Cooperstown occupies the valley floor for the next mile.

STOP 7. Lake Front Park on Doubleday Ice Margin
Well data from the crest of the moraine (Bassett Hospital) indicate that the morainic dam for Otsego Lake consists of 180 feet of bouldery silt. Subtle hummocky terrain in the village of Cooperstown may be traced northwestward, where it is well expressed in a golf course at the lake shore. Hanging deltas along the western lake shore and lake clay on the northwestern shore indicate that while the moraine dammed Glacial Lake Cooperstown at an elevation of 1250'. The view northward includes a conspicuous cross-sectional valley asymmetry, which played an important role in determining the primary sediment source for Glacial Lake Cooperstown.
A seismic reflection survey shows that Otsego Lake occupies a rock basin that has been eroded as much as 132 m (433') lower than present lake level and infilled with up to 88 m (289') of sediment. Sediment thickness is greatest in the southern half of the lake basin and thins by -50% to the north where maximum water depths occur. As an alternative to the traditional view of deglaciation, which depicts continuous backwasting of an active ice margin with short valley lobes, subsurface and land-based evidence suggests backwasting retreat of 20 km long ice-tongues that remained in the valleys, subject to collapse while adjacent upland regions were ice-free. Proglacial sediments were largely transported to the south end of the lake basin as alluvium from ice-free western tributary streams (Fleisher, et al., 1992).

Highway traverses Cassville-Cooperstown moraine.

From here south for several miles, planar gravels
from here south for several miles, planar gravels appear related to upland tributary inwash sources.

Valley floor consists of lacustrine clays
Valley floor consists of lacustrine clays from here south to a moraine at Portlandville at 68.9.

STOP 8. 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, sand and gravel outwash over 300+ ft. of silt interpreted to be of lacustrine origin. Does the stratigraphy indicate two stades (readvance) 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.

STOP 9. Moraine at margin of Dead Ice Sink
Approval for entrance to quarry must be obtained in Office.
Many of the landforms that characterize the through valley assemblage may be found in the vicinity of this location. They include a kame moraine and associated outwash features, dead ice sink, and lacustrine plain. During the many years of operation, a variety of ice-contact features have been exposed. Upper-most units are flow tills (debris flows) that randomly interfinger with interbedded sand and medium gravel, frequently including silt and sand of local significance. Blocks of lodgment till have been observed in piles with sub-meter boulders that come off screening equipment. Massive foreset beds of gravel that range in size from bouldery gravel, through cobble and pebble gravel, to pebbly, course sand are commonly found beneath the overlying flow till at an intermediate, vertical position through the operation. Well sorted, laminated lake sand is very common on the downvalley, lower margin of the quarry.
The hummocky, kame and kettle topography common to the quarry area fades downvalley to join with the surface of extensively developed outwash terraces. Over the years, quarry operations in these terraces have uniformly exposed massive gravel foreset beds indicative of meltwater sedimentation in an ice-contact lake that expanded upvalley to follow the retreating ice front. MacNish and Randell (1982) describe such a setting to depict slow, active-ice retreat.
These extensive paired terraces formed as "lateral deltas" into the expanding ice-contact lake, thus growing in the upvalley direction as space became available during retreat. Episodes of slow retreat would accommodate growth until delta terraces from either side of the valley coalesced to fill the valley, thus forming a valley train with internal deltaic structure.

Entrance to Brae Pit, Cobbelskill Products quarry.
Time permitting, we will drive in to view massive foresets of delta terrace.

Road parallels Charlotte Creek through the moraine on the right.

STOP 10. 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 this transverse valley was clogged by remnant ice. These landforms are typical of valleys in which the stagnated ice was subject to inwash burial and many local ponds and lakes existed. Here, the Davenport Moraine dammed Charlotte Creek valley at 1280 ft. Many shoreline landforms along the valley are graded to this lake.

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

Town of Davenport Center quarry on right
Town of Davenport Center quarry on right (no longer active). Past exposures contained well-developed, large-scale deltaic foreset beds graded to 1280 ft. water level. Forty foot deep kettles indicate progradation that incorporated grounded ice.

STOP 11. Hanging delta at the confluence with Middle Brook
Clark Company Stone Products. Although quarrying activity continues, much of this original hanging delta into Glacial Lake Devenport (1280 to 1300 ft.) has been excavated. Remnants of evidence supporting this interpretation may still be seen. Exposed at the upper quarry level are several meters of well stratified, strongly imbricated topset gravel. Discontinuous exposure to lower positions eventually lead to multiple packages of lacustrine sand and silt interspersed with small-scale deltaic infill and cut and fill structures.

Stop 12. Very localized hummocky topography reaching elevations of 1360 ft.
Turn right on Olive Branch Road (Dead End) at Brandow's Trailer Sales. Access by permission of Brandow Family.
Recent excavations exposed several, 2-3 m thick, alternating units of cross bedded coarse sand, tan laminated lake sand, and sand and fine gravel, each with unconformable contacts. The entire sequence was overlain by 2-3 m of moderately well sorted pebble gravel. Isolated from potential tributary sources, this ice-contact landform appears to have originated as meltwater infill of localized ponding on ice-cored terrain, much as would be anticipated in contact with a remnant ice masses in a transverse valley.

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.

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

OPTIONAL STOP 1. Village of Laurens Quarry
This is a good example of inwash sourced aggradation associated with a stagnant, downwasting ice tongue that occupied this non-through valley and other like it. At various times, excavation has exposed tilted and truncated sequences of interstratified gravel and pebbly sand with fluvial sedimentary structures and local silt layers of lacustrine origin. The basic feature is an ice-contact, hanging delta (Mt. Vision Quadrangle). Landform expression and large-scale foreset and topset beds indicate deposition into an ice-contact lake. Initial aggradation here was into a lake graded to a moraine dam at West Oneonta, 5 km downvalley. A local inwash source provided adequate sediment for additional topset, alluvial aggradation above lake level.
Otego Creek passes through a narrow breach that interrupts the continuation of this landform at the same elevation across the valley floor. The localized concentration of accumulated sediment here is attributed to inwash from Lake Brook. Similarly, other tributaries fed sediment from ice-free slope to local lakes and ponds distributed along the valley and gravel plains are graded to these water bodies.

Gravel plain remnant for next half mile.

Road rises onto pitted planar gravel remnant.

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

OPTIONAL STOP 2. Esker in dead-ice terrain
Once again we see concentrated aggradation at a stream confluence, here where West Branch joins Otego Creek. Landforms typical of a non-through valley assemblage include a kame field (differs from a kame moraine only in that the kame and kettle topography is not linked with planar outwash, as would be anticipated at an active ice margin), gravel plain, dead-ice sink, and an esker. Eskers are rarely preserved, partly because conditions favoring formation were not common, but also because they are excellent sources of sand and gravel, thus they are typically removed by gravel mining. Here, the esker is indicative of stagnant ice conditions. Lack of excavation precludes investigation of internal features.

Turn right onto Angel Road that traverses kame field.

OPTIONAL STOP 3. Overview of meltwater channel
Overview of meltwater channel through discontinuous gravel plain 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 an esker.

OPTIONAL STOP 4.
Overview of discontinuous, planar gravel surface and terrain of ice-cored origin bordering a local, upvalley lacustrine plain.

OPTIONAL STOP 5. 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 due to thinning during glacial retreat, thus resulting in downvalley ice-tongue starvation and ultimate stagnation.

A4: KARST OF THE SCHOHARIE VALLEY, NEW YORK
Arthur N. Palmer
Department of Earth Sciences
State University of New York
Oneonta, NY 13820
Paul A. Rubin
HydroQuest
P.O. Box 387
Stone Ridge, NY 12484
Margaret V. Palmer 619 Winney Hill Road
Thomas D. Engel 16 Equinox Court
Delmar, NY 12054
Brian Morgan HydroQuest
INTRODUCTION
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 (Figure 1). This is one of the finest examples of glaciated karst in the country. This field trip focuses on the valley of Schoharie Creek, which rises in the Catskills and drains northward to the Mohawk River. Where it intersects the carbonate belt, it is joined by Cobleskill Creek from the west and by Fox Creek from the east. Deep entrenchment by these streams has allowed a great deal of subsurface drainage to take place through the carbonate rocks, accompanied by karst topography and caves. During the latest phases of glacial retreat, about 14,000 years ago, a dam of glacial ice between the neighboring plateaus blocked the drainage and formed glacial Lake Schoharie, which inundated most of the caves and karst to depths as great as 150 m. The primary purpose of this field trip is to examine some of the major karst features of the area and how they have interacted with the effects of Wisconsinan glaciation. A secondary purpose of the field trip is to address cave resource protection considerations via land-use analysis and through use of the karst principles discussed in this paper. Further details about the Helderberg karst are provided by Kastning (1975}, Baker (1976), Palmer (1976), Mylroie (1977), Cullen et al. (1979), Palmer et al. (1991), Rubin (1991), Rubin (1995), Rubin et al. (1995), and Dumont (1995).
This field trip illustrates some of the major surface karst features of New York, including both surface karst and caves. Three different plateaus are visited, each with a different geomorphic setting and subsurface drainage pattern. No special equipment is needed for this field trip. However, many of the stops are on private property, and visitors at other times need written perm1ss1on. Caves that are not open to the public are rather dangerous and require special equipment and prior experience under supervised instruction. Appropriate contacts are listed in this guide.

A4: KARST OF THE SCHOHARIE VALLEY, NEW YORK

STOP 1. WELL FIELD ON CAMPUS OF SUNY COBLESKILL
This stop gives a chance to stretch and to review the geology of the region. The zero point in the field-trip log is at the nearby traffic light on NY Route 7. Note Cobleskill Creek (a western tributary of Schoharie Creek) and the broad plateaus rising to the north and south. To the northeast is the Cobleskill Plateau, which consists mainly of limestones of the Helderberg Group, and which contains the largest caves and karst drainage systems in the Northeast. Preglacial entrenchment of Cobleskill Creek allowed tributary subsurface drainage to develop in the plateau, and so the creek is directly responsible for much of the karst in the region.
In July 1991, SUNY Cobleskill began pumping from one of several wells at this stop to support an aquaculture program. The well penetrates several meters of glacial-alluvial sediments and extends into the underlying Onondaga Limestone. It was pumped intermittently at about 30 gpm. Soon afterward many domestic wells in the vicinity ran dry or lost capacity. The SUNY well was widely held responsible. It seems to have an unusually broad effect, because wells lost capacity all over the county. Some ran dry even in Oneonta.
The blame was clearly misplaced. There are several reasons why the SUNY well could not have been responsible: (1) the volume of water pumped was many orders of magnitude less than that necessary to lower the water table over such a broad area. (2) In a cone of depression the greatest drawdown is in the center, but drops in water level in many of the affected wells were several times greater than in the pumping well itself. (3) Wells closest to the pumping well were unaffected. (4) Some of the affected wells also experienced an influx of hydrogen sulfide, which is typical during periods of low recharge of meteoric water. (5) The SUNY well was such a malignant influence that some of the domestic wells went dry even before SUNY began pumping. The N.Y. Dept. of Environmental Conservation de-fused the issue by declaring that there was not enough evidence to prove that the SUNY well had affected the domestic wells.

Quarry on left is in Onondaga Formation

STOP 2. DOC SHAUL'S SPRING
Many karst springs are located in the limestones along the southeastern end of the Cobleskill Plateau, near Howe Caverns (Figure 1 ). Farther west, however, the limestones exposed by pre-glacial erosion were covered by glacial and alluvial sediment, to depths as great as 30 m. Doc Shaul's Spring, the main outlet for water in the western part of the plateau, rises upward from conduits in the limestone through a conical pit in the overlying sediment (Figure 3 and 4). It is one of the largest springs in the state. Much of the water that feeds it can be seen in caves higher in the plateau.
The original spring appears to have issued directly from the exposed down-dip edge of the limestone. The depth of the original spring has not yet been determined, but it is likely to be in the lower Manlius Limestone, about 25 m below the present spring. Divers have found the opening nearly choked with logs and sediment. Did the valley sediment accumulate slowly enough that it was continually swept away by the upwelling 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? Saturated sediment has an effective specific gravity of about 1.1-1.2 (accounting for buoyancy and porosity), so a pressure head of about 30 m would have been necessary to balance the weight of the sediment. Dissolution of travertine in the tributary caves extends to levels at least 10 m higher than this. Thus either interpretation is feasible.

Turn up the hill to the left onto the plateau,
Turn up the hill to the left onto the plateau, which consists mainly of limestones of the Helderberg Group.

The top of the plateau has been streamlined by glacial ice
The top of the plateau has been streamlined by glacial ice, which was moving nearly westward in this local area. Deranged drainage, elongate bedrock hills, and drumlins are common.

Road makes a wild curve around the nose of a prominent drumlin.
The drumlins in this area are superimposed across a north-south pre-glacial valley in the Helderberg Limestones.

STOP 3. BROWN'S DEPRESSION
Because of the difficulty of parking and land access, this "stop" actually consists of short views from the bus at several locations, accompanied by geologic discussions. Most of the depression is on private property where visitors are not welcome.
Brown's Depression (Figure 5) looks like an enormous sinkhole, but it is actually a remnant of a preglacial valley, which elsewhere has been completely obscured by glacial till. 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 (Mylroie, 1977).
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 wetlands.
The buried valley was detected with gravity surveys and well logs by Palmer (1976) and Milunich and Palmer (1997), and with reflection seismology by Mylroie (1977). Its average depth of fill is 60-70 m. The generalized gravity profiles are shown in Figure 6. The deep part of the valley lies beneath the eastern flank of Brown's Depression, rather than directly beneath it.
The pre-glacial valley, of which Brown's Depression is part, separates two of the largest caves of the Northeast, McFail's Cave to the east and Barrack Zourie Cave to the west (Figures 6 and 7). Both discharge into Doc Shaul's Spring through water-filled passages. They probably join each other before reaching the spring, but whether the passages are open phreatic tubes or are occluded by glacial sediment is uncertain. Both caves contain extensive calcite-rich, fine-grained rhythmites that are convincingly demonstrated to have been deposited during the pending of glacial Lake Schoharie (Palmer, 1976; Dumont, 1995).

Another view of Brown's Depression on the right.

Cross the valley draining into Brown's Depression.

Barrack Zourie, a prominent hill consisting of an outlier of Middle Devonian strata, can be seen to the right.

Pass shale pit in the Esopus Formation on right.

STOP 4. MCFAIL'S CAVE ENTRANCES
Pull off road into a grassy parking lot on left. The various entrances to McFail's Cave can be visited by following the path through the woods at the northeastern end of the parking lot. The cave entrances and parts of the surrounding woods are owned by the National Speleological Society. Permission to visit the property can be obtained from the chairperson of the McFail's Cave Committee. Please do not cross the field, as it is private farmland. When you leave the parking lot, turn right onto the paved road.
This property is owned and managed by the National Speleological Society, and a waiver is required to visit it. The cave entrances are all gated, and access is limited only to groups with the proper experience and equipment.
McFail's Cave has a mapped length of 10.7 km (6.6 mi.). Most of its water is contributed by two independent recharge areas. The main passage, which consists mostly of a single long down-dip canyon, is fed by numerous shafts and sinkholes in the vicinity of this stop. Its largest tributary, the Northwest Passage, is a remnant of a lengthy strike-oriented passage that formed at a former base level (Figures 5 and 7). Its strike orientation shows that water was prohibited from continuing down-dip to the Cobleskill Valley by the lack of limestone exposure in that direction. Dye tracing shows that the Northwest Passage is fed by drainage to the northeast of Brown's Depression, and that the water that sinks in the depression does not enter the Northwest Passage, but flows through Barrack Zourie Cave to Doc Shaul's Spring (Mylroie, 1977; Dumont, 1995).
The original flow was to the southeast to Howe Caverns, forming a passage graded to an outlet level at about 275 m elevation (Figure 7). In McFail's Cave the passage is at an elevation of 305-312 m. An alternate explanation is that the passage was formed or modified by glacial meltwater at a time when ice blocked the Cobleskill Valley, and does not represent a true erosional base level. The passage is partly relict today and has been segmented by breakdown and fill. The lower part of McFail's Cave is entirely water filled (Figure 3). Divers have penetrated to a depth of more than 6 m in the final sump (Sump 5 on Figure 5).
As Cobleskill Creek cut headward into the limestones south of the plateau, a new flow route developed to Doc Shaul's Spring, and the strike passage joining McFail's and Howe was abandoned (Palmer, 1976; Mylroie, 1977). Howe Caverns is simply a beheaded remnant of this passage. The accessible passage segments still contain perennial streams, but most are underfit because of the loss of drainage area. Small scallops indicate past discharges much greater than at present (Palmer, 1976). This may indicate modification by glacial meltwater (Rubin, 1991). Rubin (1999) used the Chezy-Manning formula and physical measurements to reconstruct Howe Caverns flows for 1938 and 1996 floods, using a range of Manning's n values between 0.03 and 0.04. This provided a range of peak flows between 270 and 365 cfs for a 1996 flood and between 1135 and 1520 cfs for the flood of 1938. Infrequent floods of this magnitude may still not have been sufficient to account for the conduit size present.
The cave entrances lie in a little woodlot that contains the densest cluster of vertical shafts in the Northeast (Figure 8). Each shaft feeds a tributary of McFail's Cave. Most of the tributaries are not traversable. A maintained trail snakes through the property past some of the more significant karst features. Please do not cross the open field, as this is private cultivated land.
The tops of the entrance shafts are located in the thin-bedded Kalkberg Limestone. The Coeymans Limestone accounts for the vertical-walled sections of the shafts. The deepest shafts reach about 3-6 m into the Manlius Limestone. Much of the main passage originated at or near the Coeymans/Manlius contact and has since been entrenched downward as much as 12 m as a narrow sinuous canyon. In places its ceiling rises along prominent joints to heights up to 20 m. At the junction with the strike-oriented phreatic passage (see Figure 5) the cave stream has still cut only half way downward through the Manlius. Farther downstream the main passage descends through the rest of the Manlius and about a meter into the Rondout.
The main stream passage (Figure 9) extends almost exactly down the local dip, deviating significantly from this trend only where joint control is prominent. This orientation implies gravitational flow - hence vadose cave origin. Despite the prominent jointing in the limestone, the cave stream, as well as the original cave-forming water, was deflected laterally more than 1.5 km before reaching the potentiometric surface. From the standpoint of groundwater contamination, it is important to realize that infiltrating water in limestone, or probably any other bedded rock, rarely takes a perfectly vertical path downward to the potentiometric surface.
The density of sinkholes and shafts in this small area requires some explanation. Limestones of the field-trip area are overlain in most places by relatively impermeable sedimentary rocks and by glacial till, which concentrate runoff into the few places where the limestone is exposed. Here, for example, is a narrow embayment in the low-permeability overlying rocks, where the relatively pure limestones are exposed beneath a thin soil cover (Figure 7). Runoff converges on this area from the high areas all around, so it is not surprising that so many karst features are concentrated here. The cave contains many small tributaries in its upstream end, fed by the openings seen at this stop. Where the cave extends beneath the cap of relatively insoluble rocks the only tributaries are mere trickles that deposit travertine.
Joint control of passages is strong in the area around the entrances, but farther down dip, where the cave is overlain by as much as 100 m of bedrock, joint control is much less prominent and bedding-plane control dominates. Minor bedding-plane thrusting may have aided the apparent lack of joint control. It appears that the abundance of fissures in the entrance area is due in part to erosional unloading. Glacial stresses may have played some part, although it must be kept in mind that the cave is known to predate at least the Wisconsinan glaciation.
Despite the great amount of surface runoff, the water entering the cave is not particularly aggressive. In fact, during most of the year the cave water is slightly supersaturated with calcite and is unable to enlarge the cave by solution. Only during periods of high runoff, when sinking streams are fed by large amounts of overland flow, is the cave water solutionally aggressive.
Another small cave is located in the Becraft Limestone almost directly above McFail's Cave, which lies 60 m below. Water from the higher cave drains into one of the entrances of McFail's and thus traverses two different caves in entirely different limestone formations.

STOP 5. HOWE CAVERNS
Howe Caverns is the largest Northeastern cave open to the public (Figure 10). The original entrance lies in the southeastern corner of the Howe Cave Quarry (Stop 6), whose operation has obliterated the connection with the main part of the cave. Howe Caverns is now entered through a 45 m elevator that descends from the visitors' center at the top of the plateau. From the elevator, the cave tour follows the main passage downstream for 500 m to a short boat ride on a dammed lake. Near the elevator, the main passage is joined by a down-dip tributary canyon, the Winding Way. The tributary water now follows a lower route, leaving the Winding Way dry during periods of low and moderate flow. However, during periods of high flow, the downstream outlet is inefficient. In response, floodwaters initially backup in the up gradient portion of the Winding Way, then aggressively flow through the Winding Way to the strike-oriented portion of Howe Caverns. It has not yet been determined whether this floodwater overflow route carries only water from the eastern portion of the Sagendorf Corners sub-basin or if it also serves as a bypass shunt around breakdown for floodwaters from upstream portions of the Inferred Strike Conduit (Figure 7). Upstream from the elevator the main passage can be followed a short distance to breakdown at the northwestern terminus of the West Passage, situated southeast of a small valley tributary to Cobleskill Creek. This passage appears to have once been the downstream continuation of the strike-oriented passage in McFail's Cave, whose terminus now lies about 3 km to the northwest (Figure 7). Tracer test results confirm that groundwater efficiently flows beneath this small valley prior to rising as the River Styx in Howe Caverns. Howe Caverns, now isolated from McFail's, receives its present drainage from those parts of the plateau updip to the north.
The main passage of Howe Caverns is a large tube up to 10 m wide and about 6 m high, with a 3-5 m canyon cut in its floor (Figure 11). The solutional ceiling is approximately at the Coeymans/Manlius contact. At Titan's Temple, the largest room in the cave, the lower-level canyon diverges from the tube. The stream, of course, follows the lower level, which assumes a more tube-like configuration farther downstream. The upper level is clay-choked in the former downstream direction. At the divergence point between the two levels, the cave intersects a subtly exposed reverse fault, also seen in the Howe Cave Quarry (Stop 6). Here the fault dips 14 degrees to the south-southwest. The fault can be seen in the chin of the bedrock feature known as the "Old Witch" (Gregg, 1974). It can also be seen extending downstream of the Old Witch along the southwestern passage wall.
The significance of the reverse fault (and smaller subsidiary ones) to the local cave development has been debated. Egemeier (1969) and Mylroie (1977) attribute a great deal of the local cave-passage orientation to the fault. Gregg (1974) and Kastning (1975) downplay the influence of the fault. The general consensus is that stratigraphy and base level exerted the main control over the Northwest-Southeast Passage in McFail's and the main passage of Howe, and that the fault controls local passage segments.
The main passage contains some noteworthy deposits. Most unusual are large remnant banks of thinly laminated clay rhythmites. These are clearly ponded-water deposits, and their abundance here and in other caves of the Schoharie Valley (but rarely, if ever, in other New York State caves) suggests that they accumulated in the late Wisconsinan glacial Lake Schoharie. As in McFail's, they occupy the lowest level of vadose entrenchment. Almost the entire solutional history of the cave pre-dates the latest retreat of glacial ice. Since their deposition, the clay banks have been largely eroded away by the cave stream. Remnants of a thin flowstone covering conform to the eroded shapes of the clay banks. Most other dripstone and flowstone in the cave appear to be older than the clay fill. Many stalactites and stalagmites have been resurrected from fallen pieces, not necessarily in their original locations.
The Winding Way is the largest of several tributary passages, and the only one that can be followed for any great distance. Like most of them, it enters the main passage from the up-dip side. Although most of the tributaries are vadose canyons, their solutional ceilings descend steeply across the strata to join concordantly with the ceiling of the main passage. This is odd behavior for vadose passages. Should gravitational water have any greater tendency to cut downward across the strata above the main passage than it would anywhere else? The angle and direction of the discordant ceilings are suspiciously parallel to the trend of the reverse fault, however; initial solution of the main passage may have opened imperceptible fractures parallel to the main fault that guided the vadose water.
The tours exit the cave through an artificial tunnel between the Winding Way and the elevators. Much of this tunnel is excavated through a fault zone marked by small slickensided exposures and related calcite mineralization.
Howe Caverns Karst Hydrology and Cave Resource Protection
Howe Caverns, the most frequently visited commercial cave in New York State, is currently reviewing all factors that may impact their underground natural resource (Rubin et al., 2003). HydroQuest recently completed the first phase of this work for Howe Caverns, Inc. Phase I of this study examined surface and subsurface hydrologic and geologic factors that control groundwater influx into the cave (Figure 7). This is important because existing and potential land-uses within the drainage basins of commercial (and other) caverns may pose a threat to air quality, groundwater quality, cave-dwelling species, and surface streams. Protection of the cave resource requires knowledge of the karst principles discussed in this paper, coupled with information on tracer test results, bedrock geology, glacial geology, topography, land-use, and karst features.
Three surface drainage basins tributary to Howe Caverns were delineated by analysis of DEM and DRG data, and field verification. Land-use within these basins was photo-identified from high resolution imagery and mapped. Karst features (e.g., caves, sinkholes, sinking streams, springs) near Howe Caverns were located with a Global Positioning System receiver and plotted. Tracer test results were also plotted. All this was brought together in two unified GIS map documents (Figures 7 and 12). The tracer test flow routes plotted assume groundwater flow in a down-dip direction, often independent of surface drainage basins. Additional tracer tests may find that sub-basin boundaries extend further up-dip within Helderberg Group carbonates to the north-northeast beyond surface watershed boundaries. The southwestern sub-basin boundaries depicted truncate nearly coincident with the inferred strike conduit, reflecting a strike-oriented flow path into Howe Caverns. Actual sub-basin boundaries may also extend a short distance southwest and up slope of the inferred strike conduit. The exact subsurface path of the inferred strike conduit represents a reasonable hydrologic interpretation based on tracer test, geologic, and hydrologic factors.
The dominant land-use throughout the three sub-basins tributary to Howe Caverns is agricultural (Figure 12). Three active dairy farms operate in the watershed, each with large hay fields and row crops. Small horse farms are also present. Remaining portions of the watershed are largely comprised of forest, low intensity residential development, and wetlands. The greatest potential threats to Howe Caverns are manure/septic odors and related water quality degradation. Land-uses that pose this threat are primarily agricultural and residential based. What best management practices should be considered within each drainage basin?

STOP 6. HOWE CAVE QUARRY
This quarry is now owned by Cobleskill Stone Products, which plans to develop part of the property into an educational center for use by school groups. The 150-year-old house on the property, once a hotel for visitors to the old Howe's Cave may become the visitors' center, with exhibits on mining and geology. The house was built next to the former natural entrance to the cave. However, quarry operations obstructed, partially filled, and removed portions of the roof of the downstream section of the cave. Cobleskill Stone Products is currently examining the feasibility of reopening this cave section for educational and commercial purposes.
The full thickness of the massive Coeymans Limestone and underlying thin-bedded Manlius Limestone is visible in the quarry walls, although their stratigraphic differences are masked in these artificial surfaces (Figure 13). A few meters of the Dayville Member of the Coeymans appears here. The highly jointed quarry floor is in the lower Manlius and is underlain by the Rondout Dolomite, which is extensively riddled with mine tunnels that extend below the quarry floor (Figure 14). The limestone was once quarried for crushed rock and the dolomite for cement.
Note the diagonal fault trace in the quarry wall (Figure 13). This is the same reverse fault that is visible in Howe Caverns. Here it has about 0.45 m of displacement, and, as in the caverns, it dips 14° to the SSW. It is actually subsidiary to a much larger bedding-plane thrust exposed in the mine below the quarry floor. Strontianite and barite veins up to a meter thick occur along the lower fault zone.
The downstream end of Howe Caverns, beyond the limit of the tours, is contained within the narrow peninsula of limestone extending from the western side of the quarry. The truncated end of the natural cave is obscured by the debris cone at the base of the quarry wall. Water from the cave drains through an artificial tunnel and shaft into the cement mine below. The resurgence of this water is located in the cliff below the quarry (Figure 14). The original drainage pattern of underground water has been disrupted considerably by the mining and quarrying.
A large system of caves - the Secret-Benson-Barytes system - once joined Howe Caverns near its downstream terminus but has since been dismembered from the latter by the Howe Cave Quarry (Figures 7 and 14). Secret Caverns, which like Howe is open to the public, connects downstream with other caves in a discontinuous series that has been traced southward to a spring below the Howe Cave Quarry.
In the cliff face southwest of the quarry is a spring (Nameless Spring) at the base of the Cobleskill Formation, perched on the shaly Brayman. Guenther and Rubin have conducted tracer tests from an enlarged joint in the floor of the quarry north of the silo complex to Nameless Spring (Figure 14). Notable hydrocarbon residue and staining (perhaps diesel fuel) are visible at the tracer injection point. Travel time to Nameless Spring has been documented in as little as 23 minutes. This spring is a popular drinking source for the local community who assume that karst springs must be pure water. This conduit provides an example of stream piracy trending in a down-dip direction away from the strike-oriented Howe Caverns conduit. During high flow, water also emerges from an ephemeral spring farther west at the Rondout-Cobleskill contact (Figure 7). The latter has a larger alcove and small cave formed by extensive dissolution and collapse, suggesting that this spring was active for a longer time, and that Nameless Spring is a rather recent diversion route for this water (Mylroie, 1977).
Glacial striae are visible in the limestone surface at the edge of the quarry, indicating that the amount of denudation since the last glacial retreat has been negligible here. This surface was covered by soil and till before the quarry was excavated, and the high carbonate content of the overburden prevented solutionally aggressive water from attacking the bedrock surface. Compare the almost complete lack of post-glacial denudation here to the rates of 20-30 cm per thousand years in some karst areas (Sweeting, 1973).

Drive or walk a short distance down Howe Cave Road
Drive or walk a short distance down Howe Cave Road, which branches to the right, to springs at base of plateau. This water is the drainage from Howe Caverns and neighboring cave systems.

Cross Schoharie Creek.
Deep entrenchment into the Helderberg Limestones by this north-flowing stream and its tributaries has been responsible for most of the karst development in Schoharie County.

STOP 7. TERRACE MOUNTAIN
We descend from the limestone plateau and follow Cobleskill Creek Valley downstream (Figure 1). Our route then crosses Schoharie Creek and follows it upstream for a few kilometers. Terrace Mountain looms to the right, with the gentle dip of its exposed limestone beds clearly visible. These strata look superficially like the Manlius and Coeymans Limestones, which are displayed so clearly in many escarpments in the region. However, the Terrace Mountain cliffs are actually composed of the Cobleskill and Rondout Formations, which reach their maximum thickness in this area. The Manlius and Coeymans occupy the higher slopes.
To the north, Schoharie Creek runs through a narrow gap in the limestone plateaus. This gap was once the site of ice blockage during the waning phases of Wisconsinan glaciation, about 14,000 years BP, which flooded the valley to create glacial Lake Schoharie. Prominent faults cut through the valley here, as shown by massive pyrite and strontianite bodies in the bedrock at the base of Terrace Mountain, but the main faults are probably obscured by the great thickness of valley sediment.
Although Terrace Mountain might seem an ideal location for karst features, today it does not have a catchment area large enough to supply recharge to more than a few shafts, sinkholes, and narrow abrasive caves, all of them small. The largest cave is recently discovered Ain't No Catchment Cave, named to poke fun at those who said there was insufficient catchment area on the mountain to form a significant cave. However, this low, wet crawlway simply proves the point. Its main virtue is that it is one of the very few caves developed in the Cobleskill Formation.
We leave the Schoharie Valley and follow the valley of Fox Creek upstream to the east along the southern edge of Barton Hill.
Barton Hill
Barton Hill is large enough, compared to Terrace Mountain, to have developed extensive karst drainage and large caves (Figure 15). Some of the springs at the southern (down-dip) edge of the plateau are used as a water supply for the village of Schoharie, and with increasing development of Barton Hill many land-use management questions have arisen. Studies of the Barton Hill drainage pattern go back more than 40 years (Gurnee, 1961).
The Helderberg limestones of Barton Hill are perched high above the nearby valleys, and the known caves have no apparent control by local fluvial base levels (Figure 15). As Fox Creek cut headward toward the east, the first opportunities for groundwater drainage to its valley were in the down-dip direction. The unrestricted down-dip drainage toward Fox Creek has prevented the development of phreatic tubes at large angles to the dip, at least as far as exploration has shown. In contrast, the limestones in the Cobleskill Valley were exposed only gradually by headward stream erosion as cave development progressed, allowing strike-oriented tubes to form before the erosion breached the limestones in the down-dip direction (Mylroie, 1977).
Glacial deposits have filled many sinkholes on the top of the plateau and have blocked some former spring outlets along the southern edge, but for the most part glacial disruption of the underground drainage pattern has been relatively minor. Virtually all of the major cave development is concentrated along the Coeymans/Manlius contact or in the upper half of the Manlius. Gage Caverns retains a large cross-sectional area upstream nearly to the northern margin of Barton Hill, suggesting that its catchment area must have been much larger when the cave was forming. Preglacial or subglacial meltwaters could have helped to enlarge the cave independently.

The steep bluffs of Barton Hill rise on the left.
Barton Hill is large enough, compared to Terrace Mountain, to have developed extensive karst drainage and large caves (Figure 15). Some of the springs at the southern (down-dip) edge of the plateau are used as a water supply for the village of Schoharie, and with increasing development of Barton Hill many land-use management questions have arisen. Studies of the Barton Hill drainage pattern go back more than 40 years (Gurnee, 1961).
The Helderberg limestones of Barton Hill are perched high above the nearby valleys, and the known caves have no apparent control by local fluvial base levels (Figure 15). As Fox Creek cut headward toward the east, the first opportunities for groundwater drainage to its valley were in the down-dip direction. The unrestricted down-dip drainage toward Fox Creek has prevented the development of phreatic tubes at large angles to the dip, at least as far as exploration has shown. In contrast, the limestones in the Cobleskill Valley were exposed only gradually by headward stream erosion as cave development progressed, allowing strike-oriented tubes to form before the erosion breached the limestones in the down-dip direction (Mylroie, 1977).
Glacial deposits have filled many sinkholes on the top of the plateau and have blocked some former spring outlets along the southern edge, but for the most part glacial disruption of the underground drainage pattern has been relatively minor. Virtually all of the major cave development is concentrated along the Coeymans/Manlius contact or in the upper half of the Manlius. Gage Caverns retains a large cross-sectional area upstream nearly to the northern margin of Barton Hill, suggesting that its catchment area must have been much larger when the cave was forming. Preglacial or subglacial meltwaters could have helped to enlarge the cave independently of the normal topography.

STOP 8. SCHOHARIE CAVERNS
This is the property of the National Speleological Society, and visitors must sign a liability waiver. Schoharie Caverns (or Shutters Corners Cave) consists almost entirely of a single large canyon passage that can be followed upstream about 600 m to a sump. The upstream 240 m of this section consists of a single joint-controlled fissure. The upstream sump and 5 others beyond have been dived (Schweyen, 1989). The traversable part of the cave terminates in two small branches containing infeeder streams. Flowstone is abundant in the cave, but it has been extensively redissolved by back-flooding and further damaged by mindless collectors. The spring opening was nearly blocked by glacial till, flooding the cave. The redissolving of the flowstone probably dates from that time. Beyond a number of sumps and upstream of a waterfall, a large stalagmite is positioned on the bedrock floor in the center of the stream passage (Schweyen, pers. comm.). This is unexpected because it could not have formed when the cave stream was flowing. Its presence suggests the cessation of water responsible for conduit formation, perhaps during an interstadial period, with long-term infiltration through the epikarst.
The former owner deepened the entrance by clearing away much of the till, still not reaching bedrock; so the cave became drained. Although subsidence of the entrenched till has begun to block the entrance again, a metal conduit installed in the bottom of the trench keeps the water level low, although during severe floods the Schoharie Caverns entrance fills nearly to the ceiling.
Uranium-series dating of speleothems from Schoharie Caverns (Lauritzen and Mylroie, 1996) gives dates greater than 350,000 years, the practical limit for the technique, suggesting that the origin of the cave is at least mid-Pleistocene. The semi-perched nature of the cave makes it difficult to relate this date to karst features elsewhere in the area or to fluvial events in the valleys below.
Solutional rills in the limestone face above and to the left of the entrance are up to several centimeters deep. The time required for this grooving was apparently not very long, although it is not certain exactly when the face was first exposed. It is almost certainly post-glacial.
Most or all of the water that feeds Schoharie Caverns has also passed through the soil, but the high degree of aeration in the cave causes calcite to precipitate inside as speleothems. The water still continues to lose carbon dioxide where it exits the cave, maintaining slight supersaturation, but not enough to cause detectable precipitation of calcite.
Another much smaller spring lies a few hundred meters to the southeast in the Rondout or Cobleskill Formation. It has deposited a large sheet of tufa on the hillside below the cave entrance. Evidently the water that feeds the cave has passed through soil rich in carbon dioxide and has reached equilibrium with dissolved limestone at these high CO2 levels. The groundwater was not aerated through cave entrances, and so carbon dioxide rapidly degasses from the water only where it comes in contact with the low carbon-dioxide levels of the outside air. The water rapidly becomes highly supersaturated and deposits some of its dissolved load.
We continue eastward, leaving the Fox Creek Valley and climbing gradually onto the plateau in which Knox Cave is located.
Karst Systems of the Knox Area
The plateau in which Knox Cave is located is a broad, low-relief upland bounded on the north by the Helderberg Escarpment but with gradational boundaries in the other directions (Figure 16). The Helderberg Limestones form most of the surface, but here the New Scotland Formation in the middle part of the sequence is shalier than in the Cobleskill area. Much surface runoff collects on this shaly limestone and flows either south into tributaries of Fox Creek, or north into sinkholes in the Coeymans Limestone. The dip is about 2.5 degrees to the SSW, about twice as steep as at the other sites.
A rather limited recharge area feeds the cave-forming limestones along the northern border of the plateau, and they are not exposed in the down-dip direction. Instead, the springs are located in the Coeymans or Kalkberg Formations, which indicates that subsurface drainage must rise upward across the strata to the springs. Nevertheless, some of the largest caves in the state have formed here. Knox Cave, with 1.2 km of passage, is accessible only to persons through written permission of the Northeastern Cave Conservancy.
The springs that drain Knox Cave and other caves in the vicinity are located in a small tributary of Fox Creek (Hesler, 1984). Because they are partly blocked by thin glacial till and other sediments, and by massive collapse in places (as at the Knox Cave entrance), subsurface drainage is rather inefficient, particularly during floods, accounting for many features in these caves attributed to severe flooding, such as long fissures that intersect in a network pattern.
To the west is a large drainage system that feeds several springs on small northerly tributaries of Fox Creek. Water draining into several deep sinkholes has been dye traced to the springs over distances of several kilometers. Dye travels this distance in less than 16 hours during all but the lowest flow conditions. On the basis of measurements in 1984, Rubin (1986) reports that the largest of these springs (Bogardus Spring) had the third highest low-flow discharge in the field-trip area (3.2 liter/sec), in comparison to 6.3 liters/sec for Doc Shaul's Spring and 5.2 liters/sec for No Admittance Spring below Howe Caverns. This underground system appears to represent one of the largest undiscovered cave systems in the state.

On the left is a bank of semi-consolidated glacial till
On the left is a bank of semi-consolidated glacial till cemented by calcite, which is apparently deposited by emerging karst groundwater.

View of the Helderberg Escarpment
View of the Helderberg Escarpment immediately to our right, but here it has unimpressively low relief. There are good examples of glacially deranged drainage on the left.

STOP 9. LIMESTONE RISE
Park in the small parking lot on the left of the road, and follow the trail, which crosses the road and climbs to the top of the limestone plateau. Fine examples of limestone pavement and fissured epikarst.
Limestone Rise, owned by The Nature Conservancy, displays good examples of a "karst pavement" consisting of limestone surfaces with solutionally enlarged joints. The Coeymans and Manlius Limestones are represented here. The soil is very thin and absent in many places. Such fissures constitute one form of what is known as the epikarst - the uppermost zone of karst in which many paths of infiltration have been enlarged by dissolution (Figure 17). Exposed bedrock such as this 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 dissolution, perhaps aided by glacial meltwater, allowing soil to subside into the underlying fissures.

STOP 10. ENTRANCE SINKHOLE OF KNOX CAVE
Knox Cave is owned and managed by the Northeastern Cave Conservancy, and visitors must sign a liability waiver. It is a complex of solutional fissures surrounding joint-controlled tubes and canyons (Figure 18). In places the passages coalesce into rooms up to 20 m high. The entrance sinkhole lies approximately in the middle of the cave and nearly blocks access to the northern half.
The entrance sinkhole extends through the upper 2/3 of the massive Coeymans Limestone and is rimmed by a slope of thin-bedded Kalkberg Limestone (Figure 19). From the pattern of underlying cave passages, it appears that much of the sinkhole's origin is due to collapse of bedrock between narrow fissures. The cave is located mainly in the Manlius Limestone, and the Coeymans/Manlius contact appears half way down the entrance canyon. Deeper, very inaccessible passages in the cave extend through 2.5 m of Rondout, about 30 cm of Cobleskill, and 5 m into the Brayman Formation, which is locally a dolomitic shale.
The cave was open to the public for several decades prior to 1961, and a few remnants of the old staircase into the entrance passage are still visible (Figure 19). The cave is basically a crude branch work system of tubes and canyons strongly aligned in the NNE-SSW direction of the major joints (Figure 18). It originated as a southward-trending tube that was later blocked by collapse and sediment. At a later time, a canyon passage formed a bypass around the blocked tube. About half the cave's length consists of fissure-like passages along joints, probably formed by diversion and injection of floodwater resulting from blockage of the main passages by sinkhole collapse and sediment fill. Multi-level and discordant passage intersections support a floodwater origin for the fissures. Note the complexity and strong joint control of the cave, as shown in Figure 18, compared to the dendritic patterns of other caves in the regions (Figures 5 and 10).

STOP 11. SPRINGS SOUTH OF KNOX CAVE
South of Knox Cave is a small park maintained by the Town of Knox. This area contains many springs fed by drainage from the caves in the plateau (Figure 16). Most are hidden by brush and are not readily visible. The springs rise from the underlying Coeymans Limestone through thin glacial, alluvial, and lacustrine deposits.
The nature of the karst drainage is cryptic. The overburden appears to average only a few meters thick, hardly enough to block large subsurface drainage paths, as in the valley of Cobleskill Creek at Doc Shaul's Spring (Stop 2). Yet, during high flow, water spurts upward in jets several decimeters high from minor fissures in the bedrock and soil, as though many subsurface conduits were semi-confined. Dye traces and geophysical field work are in progress in an attempt to evaluate the subsurface plumbing.
A short distance down the road from the parking lot is a wetland walkway that is ideal for bird-watching. It is hard to imagine a more pleasant and peaceful little park.

A5: GEOLOGY AND GEOCHRONOLOGY OF THE SOUTHERN ADIRONDACKS
James McLelland
Dept. Geosciences
Skidmore College
Saratoga Springs, NY
and
Colgate University Emeritus
Introduction
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.The location of the Adirondacks within the larger Grenville Province is shown in Fig. 1. Topographically, the Adirondacks are divided into Highland (H, Fig 2) and Lowland (L, Fig. 2) sectors. The former is underlain principally by orthogneiss, the latter by paragneiss rich in marble, and the Carthage-Colton Mylonite Zone separates the two regions (Fig 2a,b). The region has experienced multiple metamorphic and intrusive events and large-scale ductile structures are common (McLelland, 1984; McLelland et al., 1996). U-Pb zircon geochronology (Table 2, Fig. 3) indicates that the oldest igneous rocks exposed are ca 1350-1300 Ma tonalitic, arc-related plutons (Fig.2a) intrusive into older Highland paragneisses of uncertain age (McLelland et al., 1996). The oldest igneous rocks in the Lowlands consist are ca. 1200 Ma granodiorites (Fig.2a) intrusive into older paragneisses of uncertain age (Wasteneys et al., 1999). In both the Highlands and Lowlands meta pelitic migmatites 1200-1300 Ma contain anatectic material of ∼1170 Ma age. Following intrusion ca. 1207 Ma granodiorites in the Lowlands, leucogranitic and tonalitic rocks were emplaced at ca 1172 Ma (Fig.2a,b) and were accompanied by deformation and metamorphism (Wasteneys et al., 1999) assigned to the latest, culminating phase (ca. 1220-1160 Ma) of the Elzevirian Orogeny of Moore and Thompson (1980).
From ca 1160-1150 Ma the entire Adirondack-Frontenac region was intruded by anorthosite-charnockite-mangerite-granite (AMCG, 1155 Ma, Fig. 2b) magmas that are associated with four anorthosite massifs (Marcy, Oregon, Snowy, Carthage) recently dated directly at 1155 ± 10 Ma by SHRIMP II zircon techniques. Early age determinations of the anorthosite (McLelland and Chiarenzelli, 1990; Silver, 1969) were based on multigrain dating of MCG granitoids that exhibit mutually crosscutting relationships with the Marcy Anorthosite Massif (MM, Fig. 2b) and are interpreted as coeval with it. The absence of direct dating of the anorthosite was due to the sparse igneous zircon populations in rocks of this composition, a condition that posed a serious obstacle to the study of anorthosites prior to the advent of single grain TIMS and SHRIMP II methods. Current SHRIMP II direct dating of the Adirondack anorthosite massifs demonstrates that both they and their associated ferrodiorites and granitoids were emplaced at 1155 ± 10 Ma and that the entire complex represents a classic AMCG suite.
The final major events of Adirondack evolution comprise: 1) the emplacement of the Hawkeye granite suite at ca 1095 Ma (Fig 2b) followed almost immediately by 2) high-grade metamorphism (Storm and Spear, Appendix to this article; Spear and Markussen, 1997; Bohlen et al. 1985; Valley et al., 1990) resulting in vapor-absent, peak granulite facies conditions in the Highlands (T - 750°-800°C, P - 6-8 kbar) and associated with widespread recumbent, isoclinal folding and the development of intense penetrative fabrics. This granulite facies metamorphism and deformation are assigned to the collisional Ottawan Orogeny of Moore and Thompson (1980), evidence of which occurs throughout the Grenville Province (cf. Rivers, 1997; McLelland et al., 2001b). Toward the end of this major orogenic event much of the Adirondack region was intruded by late- to post-tectonic leucogranites (ca. 1055 Ma, Fig 2a) belonging to the Lyon Mt. Granite (LMG) and thought to be related to delamination and extensional collapse of the orogen (McLelland et al., 2001 b). The Elzevirian and Ottawan Orogenies, taken together, comprise the Grenville Orogenic Cycle (ca. 1350-950 Ma) of Moore and Thompson (1980).

A5: GEOLOGY AND GEOCHRONOLOGY OF THE SOUTHERN ADIRONDACKS

STOP 1. PECK LAKE MIGMATITIC METAPELITE.
This exposure along Rt 29A just north of Peck Lake is the type locality of the Peck Lake migmatitic metapelites that consist of restitic sillimanite-garnet-biotite-quartz-oligoclase melanosome and quartz and two-feldspar leucosome of approximately minimum melt composition (Mclelland and Husain 1986). Small red garnets are common in the leucosome and, in most cases, grow across foliation. The leucosomes occur as irregular, elongate bodies generally parallel to foliation but quite commonly exhibiting crosscutting relationships with respect to the restite and to one another. Based upon these compositional and crosscutting relations, the leucosomes are interpreted as anatectites and a reasonable metapelitic source rock can be prescribed by reintegrating their composition with that of the restite (Table 3) to yield a graywacke-slate precursor. The restriction of these partial melts to within the migmatite is thought to be the result of near-solidus, hydrous melting that would cause ascending melts to intersect the solidus as they began to rise. Close inspection of the leucosomes reveals that, in their most pristine configuration, they consist of coarse granite and pegmatite. In low strain zones elsewhere in the Adirondacks it is manifestly clear that the leucosomes originally formed an anastomosing arrays of veins, dikes, sheets, and pods. Subsequent high strain resulted in rotation into psuedoparallelism and commonly produced disruption that caused separate grains of white feldspar some of which show elongate tails - ie, the rock was on its way to becoming a mylonitic "straight gneiss". At the last stop (Stop 9) of this trip we shall see the end result of this process in a series of platy, stretched, and highly grain size reduced and mylonitic equivalents of the rocks seen here. The anatectic origin of these units is further suggested by the uncommon, but not rare, occurrence of plagioclase- and/or garnet-rimmed hercynitic spinel in the restite. Recently, Bickford and Mclelland have run a U/Pb zircon pilot study on the age of the anatectites and have found that they contain 1220-1250 Ma cores, 1020-1050 Ma metamorphic rims, and relatively thick, nicely zoned mantles that fall into the interval 1190-1170 Ma. Some of the rims show zoning, but this is minor compared to the mantling zircon. These results demonstrate that majority of the anatectites were produced by partial melting during the culminating Elzevirian Orogeny dated at ca 1210-1170 Ma (Wasteneys and Mclelland, 1999) and are not of Ottawan origin. Currently, we are continuing this research by utilizing both zircon and monazite geochronology.
This outcrop provides an excellent mesoscale example of Adirondack structure (Fig. 7). The overall strike of foliation is N60-70W and dips vary from north to south and pass through the vertical demonstrating that the roadcut defines a large minor isoclinal, recumbent fold. Smaller, meter-scale isoclinal folds are also well exposed and are aligned parallel to the roadcut-scale fold. All of these are similar in style, and parallel to, the very large, regional Canada Lake isocline or nappe. It is clear that the Canada Lake nappe stage of deformation rotates an earlier foliation and has a moderately dipping axial planar foliation that intersects the earlier foliation at a high angle. The time interval between these foliations remains unspecified, but it is clearly post-anatexis and probably all Ottawan (Ca 1090-1030 Ma). It appears that the Ottawan has successfully obliterated most Elzevirian fabrics. Note that some of the smaller minor folds contain apparently terminated compositional layers that could very well represent Elzevirian isoclinal noses. If so we are looking at isoclinally refolded isoclines.
Migmatitic metapelites of this sort occur throughout the southern and western Adirondacks as well as in the Adirondack Lowlands. We interpret them as flysch sequences of shale and greywacke that were being shed from the Elzevirian magmatic arcs that dominated the region from ca 1400-1300 Ma. Tonalites and granodiorites of the arcs locally crosscut the metapelites and provide a minimum age for them. In a broad sense, they are thought to be coeval. The metapelites of the Adirondack Lowlands are compositionally similar, but are thought to represent a different, younger, and uncorrelative arc environment.

STOP 2. IRVING POND QUARTZITE.
The Irving Pond quartzite unit cores the Canada Lake isocline and is folded back on itself. At map scale it is exposed across strike for over 3000 m but its "true?" thickness is on the order of 1OOO m. Notwithstanding, it represents an enormous volume of orthoquartzites with minor pelitic intercalations. Its minimum age is unknown but is being investigated by zircon geochronology. Similar quartzites along the St Lawrence River contain zircons as young as 1300 Ma, and the same may prove to be true for the Irving Pond. It is suggested that these quartzites may have been deposited along the present eastern margin of the Adirondack-Green Mt block between ca 1250 and 1200 Ma (see Fig. 4). During that interval, this margin was passive as the block moved westward towards the Elzevirian subduction zone that dipped westward beneath Laurentia and its leading margin, i.e., the Adirondack Lowlands (Fig. 4b,c). Upon collision, (ca 1200- 1170 Ma), the accumulated sandstones were deformed and metamorphosed. Although this scheme is speculative, it is consistent with the little that is known about Elzevirian events.
Within the clearing there are three small, but informative outcrops. The first consists of 30-40 cm scale layers of pure quartzite together with 2-3 cm scale metapelitic layers of. These dip gently to the southeast. The bulk of the Irving Pond quartzite consists of layer upon layer of the pure quartzite with little, if any, intervening metapelite. The second exposure is located a few tens of feet to the west where the rocks form a low ledge running uphill. The southern termination of the ledge that faces the clearing shows steeply dipping layers that are discordant with underlying layers. The discordance is due to some sort of ductile shearing in the outcrop but does not affect the following interpretation. Looking back to the first outcrop, it is clear that the gently dipping layers seen there can be projected up above the ground surface so that they must have been situated overhead at this locality; however, they must also have suddenly dipped steeply and rotated through the vertical in order to have their present configuration, i.e., they form recumbent isoclinal folds. The trend of this recumbent, isoclinal fold axis is ∼N70W and the plunge is ∼15 deg. southwest, i.e., it is a minor fold of the Canada Lake isocline family and of Ottawan age. At the base of the southernmost outcropping, and below a small ledge, there is preserved a foot-scale isoclinal nose. Inspection of the geometry makes it clear that the outcrop preserves an isoclinally refolded isocline - perhaps of Elzevirian age.
Farther down the clearing towards the highway a clean outcrop consists of pure, glassy quartzite dip slopes exposed together with dark layer-like bodies of fine-grained pyroxene-plagioclase granulite. Elsewhere, the chemistry of the granulites approximates that of diabase, is unlike calcsilicate, and locally exhibits microscopic diabasic texture (Mclelland and Husain, 1986). The igneous nature of the granulites is further manifested by the xenoliths of quartzite present in them. Elsewhere the metadiabases are isoclinally folded and the limbs of these folds clearly truncate 10-15 cm scale intrafolial isoclinal folds that are rotated by the isocline. We interpret this older set of isoclinal minor folds to be Elzevirian in age. The metadiabases have not been dated, but they are similar to some mafic rocks of the ∼1150 Ma AMCG suite and are tentatively assigned an AMCG age.

STOP 3. Canada Lake Charnockite
Large roadcuts expose the type outcrops of the Canada Lake charnockite (Figs. 8-11). This highly deformed orthogneiss has a granodioritic composition, and consists of 20-30% quartz, 40-50% mesoperthite, 20-30% oligoclase, and 5-10% mafics including small, sporadic grains of orthopyroxene. The exposures exhibit the drab olive color typical of charnockites around the world. In the woods these rocks tend to weather pink and exhibit a maple sugar brown weathering rind. The unit is ∼500 m thick and consists throughout of relatively homogeneous granitoid with pegmatites and minor amphibolitic layers. A multigrain U-Pb zircon age of 1251±33 Ma (Mclelland and Chiarenzelli, 1990) indicating that this unit belongs with other calcalkaline rocks of broadly Elzevirian age. Mapping along its contact for a total distance of ∼300 km, has not revealed any crosscutting features, but xenoliths of country rock have been recognized and substantiate an intrusive origin. The apparent conformity is attributed primarily to extreme and ductile tectonism. In addition, an original conformable, sheet-like form is quite possible. Rocks of similar age and composition are found in the Green Mts. of Vermont (Ratcliffe and Aleinikoff, 19xx).

STOP 4. Royal Mt. Tonalite
Steep roadcuts exposed across from the Canada Lake Store expose typical tonalitic rocks that are relatively common within the southern and eastern Adirondacks, the Green Mts of Vermont, and the Elzevir terrain of the Central Metasedimentary Belt of the Canadian Grenville. In all of these occurrences, the tonalites manifest the presence of magmatic arcs that existed along the southeastern margin of Laurentia during the interval ca 1400-1200 Ma diagnostic of the Elzevirian. Within the Adirondacks, multi- and single-grain TIMS U-Pb zircon geochronology indicate emplacement of the tonalitic magmas at ca 1350-1300 Ma. The present outcrop has been dated by both TIMS methods and the single grain age is constrained at 1307&plumn;2 Ma (Aieinikoff, pers comm., 1991 ). Well-documented examples of these arc terrains extend to the Llano uplift and Van Horn areas of Texas (Mosher, 1999; Patchett and Ruiz, 1990; Roback, 1996) and attest to a global-scale system that may have been an ancient analogue of the present day East Indies arc. During the Elzevirian, various arcs must have collided and amalgamated, and continental arcs may have come into existence as well. These details remain to be unraveled, but in the meantime we note that the Elzevirian came to a close at -1210-1170 Ma with the collision of the Adirondack Highlands-Green Mt block with the Andean-style arc then existent along the southeastern edge of Laurentia (Fig 4c). Mclelland et al (1991), refer to this collisional event as the "culminating Elzevirian Orogeny" that closed out the Elzevirian interval of arc magmatism and amalgamation in the area.
The whole rock chemistry of the Adirondack calcalkaline rocks are shown in Figs (9, 10,11) where their calcalkaline affinities are clearly visible. In addition, &Nd characteristics are presented in Fig. 8 and demonstrate that the tonalites represent juvenile additions to the crust from the mantle, and partial melting of the older rocks can produce younger Adirondack granitoids. Neither the Sm-Nd data, nor any other isotopic data, give any hint of pre-1350 Ma crust in the Adirondacks and suggest that the original arc must have been of ensimatic origin.
Numerous disrupted amphibolitic sheets are present within this outcrop and within Adirondack tonalites in general. The origin of these is enigmatic but they do not appear to have been derived from local country rocks. Their elongate, sheet-like character suggests that they may represent coeval mafic dikes of the sort commonly observed in tonalites. It is also possible that they represent enclaves incorporated from an amphibolitic source region. This interesting problem needs some isotopic investigations to be applied to it.
In a few places the tonalite crosscuts the migmatitic metapelites seen at Stop 1. This fixes a minimum age for the latter at ca 1300 Ma. As indicated in the discussion at Stop 1, the tonalites and metapelitic rocks may represent an original, essentially coeval, magmatic arc-flysch system.

STOP 5. ROOSTER HILL MEGACRYSTIC CHARNOCKITE.
This deformed charnockite is characterized by the presence of 20-40% megacrysts (2-4cm) of alkali feldspar set in a groundmass of quartz, oligoclase, biotite, hornblende, garnet, and sporadic orthopyroxene. In general, these have undergone dynamic recrystallization during high temperature shear strain and have developed asymmetrical tails, although flattening reduces the degree of asymmetry in most cases. In cases where tail asymmetry permits, a southeast side up and to the northwest (N70W, 10-15 SE) is clearly the dominant displacement sense. As the degree of shear strain intensifies, both feldspars and quartz become elongated in the direction of tectonic transport and ribbon or pencil gneisses result. The orientation of these fabric-forming elements is parallel to the regional isoclinal fold axes. In Fig. 8 the εNd growth curve for Rooster Hill charnockite lies on one of the AMCG suite growth lines that pass through the tonalite region indicating that this suite can be derived by partial melting of the ca 1300 Ma arc rocks.

Low roadcut in migmatitic metapelites

Roadcut through granite, gabbro, and tonalite.

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. RIBBON GNEISS IN THE CORE OF THE PISECO ANTICLINE.
A long, low roadcut on the left (northwest) side of Rt 8 exposes outstanding examples of ribbon gneiss developed in a megacrystic facies of the Piseco core rocks. The overall composition here is similar to that of the Rooster Hill megacrystic charnockite examined at Stop 5, but here the original feldspars and quartz have been extended into ribbons on the order of 60cmx.25mmx.05mm (Mclelland, 1984). Assuming interstitial quartz aggregates of ∼1 cm diameter, the current dimensions indicate an extension of 6000%, i.e., if this strain were to be equally distributed throughout a 1 km thick block of crust that block would be 60km long, .25 km wide, and .005 km thick. Of course, such high ductile strain cannot be integrated through the entire crust, but the numbers provide some flavor for the strain involved.
The ribbon gneisses exposed here are folded into a minor anticline that rotates the foliation of these L>>S tectonites. The axis of this minor fold trends N70W and plunges 10-15 SE. This orientation is parallel to the axis of the F3 Piseco anticline, to the F2 recumbent isoclines of the Adirondacks (Figs. 6, 7), and to the axis-parallel elongation lineations associated with these isoclines. The fact that all of these structural elements are parallel to one another indicates that they all shared some common kinematic experience during their tectonic evolution. In order to identify the common experience, we first note that all of the deformation in these ca 1150 Ma rocks must be of Ottawan age. Next we note that minor isoclinal folds in the outcrop lie with their axial planes in the foliation and their axes defining rod-like structures parallel to the N70W extension lineation. The low dip of the lineation suggests that the mechanism responsible for it was likely to have been thrust faulting out of the southeast and towards the northwest. This would have resulted in the stretching that elongated the quartz and feldspar. The thrusting could also have rotated earlier fold axes into the thrust plane and parallel to the direction of tectonic vergence. However, rather than rigid rotation of the axes, it is proposed that the early folds became too ductile to buckle and their axes began to "flow" as passive markers in the direction of tectonic transport, i.e., the recumbent fold axes began to undergo increasing curvature in response to the velocity gradients in the ductile flow field. The ultimate result of this process is sheath folds, and it is suggested that the isoclinal folds with axes parallel to the ribbon lineation represent parts of sheath folds. It is further suggested that the Canada Lake isocline may itself be a large sheath fold whose southern closure is obscured by Paleozoic overburden. The large F3 folds such as the Piseco anticline and the Gloversville syncline are interpreted as structures formed by NS constrictional forces that arose in response to the EW elongation associated with the large scale, ductile thrusting of the Ottawan Orogeny. Later NNE upright folds of the F4 set were superimposed on the F2 sheath folds and the F3 corrugations by continued, but waning, NW-SE contraction associated with the NNE suture situated somewhere beneath, or beyond, the present day Coastal Plain. Finally, it should be remarked that the ribbon lineations and upright lineation-parallel corrugations described above are similar to those encountered in core complexes. The problem with this alternative is that the late- to post-tectonic Lyon Mt Granite emplaced at ca 1050 Ma is unaffected by these fabrics. Accordingly, we associate the fabrics with ductile Ottawan thrusting. Orogen collapse similar to core complex kinematics is thought to have taken place at ca 1050-1030 Ma.

STOP 7. MARBLE AND CALCSILICATE.
Exposed in roadcuts on both sides of the highway are examples of typical Adirondack marbles and their associated lithologies, i.e., garnetiferous amphibolite, calcsilicates, and various disrupted blocks of both internal and external members. The disruption and boudinage attests to the extremely ductile behavior of the marbles. Also exposed are vertical, crosscutting, and undeformed veins of tourmaline-quartz symplectite. Besides calcite, the marbles contain diopside, tourmaline, sulfides, and graphite. The graphite has a biogenic carbon signature, and the marbles are thought to have formed inorganically via stromatolite accretion in an evaporitic environment. Some calcsilicate layers consist of almost monomineralic white, Mg-rich diopside crystals up to 10 cm in length. These are probably the result of metasomatism by fluid phases. The presence of the approximately Mg-pure assemblage
tremolite => enstatite+diopside+quartz+H2O
allowed Valley et al. (1983) to calculate a fluid with X(H20) = .11-.14 for the reaction. Elsewhere in the outcrop localized occurrences of wollastonite reflect the presence of H2O as an agent lowering CO2 activity. In both cases, the metamorphic conditions were T = 710°C and P = 7 kbar. These cases provide excellent examples of how the dilution of a fluid phase by a second constituent lowers activities and allows reactions to run to the right at P, T conditions below that for the pure fluid.
In the High Peaks region there exist many well-exposed examples of places where marble and calcsilicate xenoliths occur within anorthosite. This relationship documents that the marbles are older than ca 1150 Ma. It is likely that they formed during the same shelf sequence event that was posited for the Irving Pond quartzite precursor sands, ie, ca 1220-1200 Ma.

Roadcuts of steeply dipping metasediments.

Long roadcuts in pink granitic gneiss
Long roadcuts in pink granitic gneiss inter-layered with calcsilicates. The layering here is interpreted as tectonic and the granite as intrusive.

STOP 8. MASSIF ANORTHOSITE AND FERRODIORITE OF THE OREGON DOME.
This small, but instructive roadcut has outstanding examples of two important facies of anorthosite as well as a typical ferrodiorite dike associated with massif anorthosite. The best vantage point for examining the anorthosite and ferrodiorite is on top of the roadcut. Whole rock analyses of these rocks are given in Table 1.
On climbing to the top of the outcrop at its south end, one immediately sees a distinctive dark dike of ferrodiorite filled with plagioclase grains (white, ∼An43) and andesine xenocrysts (blue-gray, ∼An52); note reaction rims on the blue-gray andesine. The dike exhibits somewhat soft contacts and irregular veinlets shoot off in a fashion suggesting that the anorthosite was not yet wholly solidified when the dike intruded. Farther up the outcrop two ∼30cm-scale xenoliths occur in a ∼5-10 m-scale ferrodiorite; one of these is fine grained and the other coarse. The eastern contact of the ferrodiorite is situated just to the highway-side of the coarse xenolith. Calcsilicate xenoliths also occur in the ferrodiorite and are characterized by sulfidic staining near road level.
Farther up the outcrop the exposed rock consists entirely of anorthosite. A fine-grained, leucocratic facies forms a matrix to large (5-20 cm long) crystals of blue-gray, iridescent andesine that is typical of massif anorthosite. In places these crystals appear broken as if disrupted from a larger mass. An example of a still intact mass is found at the far end of the outcrop where a raft of very coarse grained Marcy-type anorthosite with ophitic to sub-ophitic orthopyroxene sits in a matrix of fine-grained leucoanorthosite that clearly disrupts the raft at its edges. Close inspection of the matrix reveals that its small pyroxenes exhibit subophitic texture; hence the finer grained facies must be magmatic. Workers in Adirondack anorthosite have always noted that grain size reduction produces leucoanorthosite similar to this this fine-grained facies, and examples of this may be seen at this stop. However, it is possible to distinguish between the grain size reduced and the fine-grained igneous varieties by noting that plagioclase in the former has the same composition as the coarse plagioclase, whereas the igneous variety is invariably 5-10% less anorthitic than the coarse plagioclase; in this case 52% versus 44% (Boone et al, 1969). Genetic interpretations and further details of the anorthositic suite are given in the main text.
The smooth, upper surface of the outcrop affords excellent opportunities to examine the garnet coronas or "necklaces" that are found throughout the Adirondack anorthosites. These are accounted for by the reaction
Orthopyroxene + Plagioclase + Fe, Ti-Oxide => Garnet + Clinopyroxene ± SiO2. br /Spear and Markussen (1997) have studied this, and other garnet-producing reactions, and determined that they took place during isobaric cooling (P ∼6-7 kbar) from ∼700°-∼630°C during waning stages of the Ottawan Orogeny (1090-1030 Ma). Mclelland et al (2001) have suggested that this reaction took place at ca 1050 Ma, which is the age of small, equant metamorphic zircons associated with the coronites. The zirconium was provided by Fe-Ti oxide, which accepts significant Zr into the Ti lattice site.
Recent SHRIMP II U-Pb dating of zircons from the ferrodiorite documents its age as 1155 ± 9 Ma, and this sets a minimum age for the anorthosite. Thirty kilometers to the northeast at Gore Mt, charnockite enveloping, and showing mutually crosscutting relationships with the Oregon Dome anorthosite (Lettney, 1969), yields a SHRIMP II age of 1155 + 6 Ma. These results document that the anorthosite series of the Oregon Dome was emplaced at ca 1150 Ma. This age is indistinguishable from emplacement ages of the Marcy massif and marks the time of emplacement of the Adirondack AMCG suite.

Minor marble, calcsilicate, amphibolite
Minor marble, calcsilicate, amphibolite together with meter-scale layers of white quartz-feldspar leucosomes containing sporadic sillimanite and garnet. SHRIMP II ages of zircons from these outcrops indicate that the leucosomes crystallized from melt at 1180-1170 Ma. This interval is interpreted as the time of anatexis associated with the migmatitic metapelites of the region, and falls into the culminating Elzevirian Orogeny.

Junction of Rt 8 and Rt 30

Charnockite on the north limb of the Glens Falls syncline.

>Entering the town of Wells
Entering the town of Wells that sits on a Paleozoic inlier dropped down at least 700m by a NNE trending graben structure.

STOP 9. MYLONITIC STRAIGHT GNEISS OF MIGMATITIC METAPELITE.
Large roadcuts on the east side of Rt 30 expose excellent examples of migmatitic metapelite identical to that seen at Stop 1, but now at a high grade of strain that has resulted in extreme ductile grain size reduction that has produced long tails on the feldspars and remarkable long ribbons of quartz now consisting of annealed subgrains in the process of annealing further into smaller grain size. The high strain has produced a platy mylonite that is a "par-excellence" example of straight gneiss. The alternating light and dark layers represent porcelaneous leucosome together with restitic biotite-garnet-quartz-feldspar ± sillimanite. Numerous intrafolial isoclinal minor folds can be seen, especially in the leucosome. All of these observations make it clear that the "layering" seen in the outcrop has nothing to do with any primary or "stratigraphic" features, although some observers still try to make this assertion. The layering is strictly tectonic in origin and provides a fine learning opportunity with which to make the case with students (and others). Note the strong N70W lineation on foliation surfaces. A strong component of flattening has also affected these rocks making it difficult to interpret kinematic indicators.
At the south end of the outcrop the mylonitic migmatite is in contact with homogeneous, strongly foliated granitoids of the Piseco anticline. These granitoids have been dated at ca 1155 Ma, while the structurally overlying migmatites must be at least 1300 Ma of age since they are crosscut by the tonalites. The absence of any crosscutting contact emphasizes how ductile high strain can wipe out angular discordances.
The northerly dip of the mylonitic migmatite sequence is due to the fact that we are located on the southern limb of the Glens Falls syncline or the northern limb of the Piseco anticline. These dipping layers are locally broken across by ductile normal faults with brecciated pegmatite fillings. These features have not been dated, but a ca 1050 Ma age is expected.

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.

A6: LATE OTTAWAN DUCTILE SHEARING AND GRANITOID EMPLACEMENT IN THE HUDSON HIGHLANDS, NY
Matthew L. Gorring
Department of Earth & Env. Studies
Montclair State University
Upper Montclair, NJ 07043
David W. Valentino
Department of Earth Sciences
SUNY-Oswego, Oswego, NY 13126
Gary S. Solar
SUNY College at Buffalo
Buffalo, NY 14222
Alexander E. Gate
Department of Geology
Rutgers University
Newark, NJ 07102
INTRODUCTION
The purpose of this field trip will be to examine rocks exposed in the Hudson Highlands, NY that elucidate the late Ottawan (<1030 Ma) history of high-grade ductile shearing, migmatization, and the emplacement of a variety of granitoid plutons. Most research in the Grenville Province in eastern North America has focused on the history and processes associated with Elzeverian (1350-1180 Ma) and Ottawan (1090-1030 Ma) orogenic events and intervening periods of extension and magmatism (see McLelland et al., 1996; Rivers, 1997; Hamner, 2000 for review). In contrast, comparatively little attention has been given to late- to post-Ottawan events and their importance to the history of the Grenville Province. The period between 1030 and 960 Ma is generally characterized by a diverse set of events that include both large-scale orogenic collapse along major extensional shear zones (e.g., van der Pluijm and Carlson, 1989; Carlson et al., 1990; Culshaw et al., 1994; Ketchum et al., 1998; Streepey et al., 2000) and localized high-grade metamorphism, thrusting, and magmatism (e.g., Lumbers et al., 1990; Mezger et al., 1991; Ratcliffe et al., 1991; Gower et al., 1991; Connelly and Heaman, 1993; Owens et al., 1994; Haggall et al., 1993; Jamieson et al., 1995; Corfu and Easton, 1997). Recent geologic mapping and structural analysis in the central Adirondack Highlands and the New Jersey/Hudson Highlands has also recognized the importance of late- to post-Ottawan, high-grade, ductile transpression (Gates, 1998; Allers et al., 2001; Valentino et al., 2001; Solar et al., 2003; Gates et al., 2001a; in press). This deformation is being taken up on large transcurrent shear zones with significant amounts of displacement (Gates, 1995; 1998). In the New Jersey/Hudson Highlands, the emplacement of a chemically diverse suite of granitoid plutons is intimately associated with this crustal-scale shearing event (Gorring et al., 2002). A Middle Proterozoic escape tectonic event (e.g., Tapponnier et al., 1982) in the central Appalachians resulting from accretion to the north is interpreted to have produced this deformation and magmatism.

A6: LATE OTTAWAN DUCTILE SHEARING AND GRANITOID EMPLACEMENT IN THE HHUDSON HIGHLANDS, NY

Large roadcuts off to the right (southbound side of NY Thruway)
Large roadcuts off to the right (southbound side of NY Thruway) are hornblende granite gneiss and mafic to intermediate metavolcanic gneiss.

Large outcrops of hornblende granite gneiss on right
Large outcrops of hornblende granite gneiss on right (east) side of road (see Fig. 9A).

Stop 1a: RESIDUAL MIGMATITIC METATURBIDITE, DIATEXITE AND GRANITE SHEETS
The first of two substops is a spectacular riverbed exposure of typical migmatitic gneiss found in the Hudson Highlands. No hammers please. Pavement exposure of migmatitic metaturbidite composed of steeply SE dipping 5-10 cm-thick alternating layers of metapelite and metapsammite (Fig. 9B). Metapelitic layers are 3-5 cm thick with cm-scale, mesoscopically poikolitic garnet porphyroblasts draped by matrix minerals that form sinistral strain shadow tails defined by aggregates of Qtz+Bt+Pl+Sil. The mineral composition of the metapelite layers are consistently residual suggesting significant melt production and loss from these rocks.

Stop 1b: RESIDUAL MIGMATITIC METATURBIDITE, DIATEXITE AND GRANITE SHEETS
Walk south along road to pavement surfaces (∼0.1 miles) and vertical roadcut (∼0.2 miles) exposed on the left (east) side of the road (Fig. 9A). These outcrops are the along strike continuation of the migmatitic metaturbidite of Stop 1a. These rocks are similar to those of the stream pavement, except for the presence of residual diatexite (Grt+Bt+PI+Qtz+Sil) and concordant- to weakly discordant cm-scale granite sheets (best seen in the pavement exposures). Where present, diatexite shows local disruption of the protolith and stromatic migmatite structure (seen at Stop 1a) that apparently formed during syntectonic melting and melt movement within the diatexite. Note the local pinch-and-swell of granite sheets and the fabric boudinage of diatexite gneiss (quartzite in the interboudin partition). Residual diatexite is disrupted locally in which the diatexite has formed schollen with concordant granitic composition tails and Grt+Bt-rich melanosomatic drapes at the edges of their long dimension.

Small roadcut on left (west) side of metavolcanic amphibolite.

STOP 2: STERLING FOREST GRANITE SHEETS
This roadcut along northwest side of Long Meadow Road near the Sterling Forest Park headquarters exposes a fine example of the late- to post-Ottawan suite of pink, granite sheets found in the Hudson Highlands (Fig. 1 0). These sheets are very leucocratic (∼75 wt% SiO2), composed of almost entirely quartz, K-feldspar, and plagioclase, with only small amounts (5%) of hornblende± biotite with accessory apatite, titanite, and zircon. At this locality, the sheet has a medium-grained, equigranular igneous texture and shows no evidence of deformational fabric. However, locally these granite sheets are foliated where intersected by late-stage, ductile shear zones and thus, are clearly pre- or syntectonic with the shearing event. This particular sheet is one of several, 10-100 m thick, finger-like bodies that extend both southward and northward several kilometers from the main plutonic masses at Hogback Mountain (to the north) and Bill White Mountain (to the south). Here, the sheet is about 20 m thick, strikes N20E and dips ∼65° to the southeast and has intruded parallel to foliation in the surrounding gneisses. It intrudes amphibolites and intermediate gneisses of the metavolcanic unit (contact observed at the north end of the outcrop) and metapelitic gneisses of the metasedimentary lithofacies (exposed across the road a short distance into the woods on the other side of the road). The granite sheet exposed here (sample SF-28) has a distinctly "dished" MREE pattern, a moderately strong negative Eu anomaly (Eu/Eu* = 0.35), high HREE contents (∼20x chondritic) and very low Ba and Sr concentrations (∼20 ppm) (Fig. 58 and Table 1). These chemical characteristics indicate that amphiboles and feldspars were important phases in the petrogenesis of this particular granite. The data is consistent with partial melting of garnet-free, mafic amphibolitic source rocks and/or fractional crystallization of amphibole+ plagioclase± K-feldspar at shallower crustal levels.

Large roadcuts of calc-silicate gneisses on opposite (north) side of 17 A.

Roadcut on right (north) side and in median is rocks seen at Stop 3.

STOP 3: INDIAN HILL SHEAR ZONE
Roadcuts on north side and in median of NY Rte 17 A afford an excellent view of highly sheared quartzofeldspathic gneisses within the Indian Hill Shear Zone (Fig. 11). The Indian Hill Shear Zone is one of several deformation zones in an anatomizing system of ductile shear zones that occur in the Hudson Highlands (Fig. 2). Timing of ductile deformation on the Indian Hill Shear Zone is currently constrained (albeit loosely) by recent SHRIMP U-Pb zircon ages from one sample of sheared quartzofeldspathic gneiss within the Indian Hill Shear Zone. The sample yielded zircons with rhythmically zoned igneous cores and clear, unzoned metamorphic rims (Fig. 4a). Analyses of the cores and rims produced two clusters of concordant ages (Fig. 4b). The cores ranged from 1160 to 1230 Ma. The most concordant rims ranged from 1000 to 1060 Ma, yielding an imprecise average of∼ 1010-1020 Ma. We interpret the zoned cores and associated ages to represent the original igneous history for this rock, and the rims to represent the regional metamorphic overprint, which includes high-grade ductile shearing on the Indian Hill Shear Zone. More rigorous constraints on the timing of ductile shearing in this area await additional U-Pb SHRIMP zircon geochronology on rocks within the Indian Hill Shear Zone as well as the Sterling Forest Granite sheets (upper limit), and cross-cutting, late pegmatites (lower limit).

Stop 3a:
Spectacular ductile shear zone structures in are found on the pavement surface on top part of the outcrop in the median strip. The rock is a quartzofeldspathic mylonitic gneiss with thin (10-50 cm) interlayers of amphibolite and garnet-biotite-quartz-plagioclase gneiss locally. The quartzofeldspathic mylonite is well foliated and lineated and composed of plagioclase, quartz, K-feldspar, and biotite. This mylonite exhibits well-developed kinematic indicators including S-C fabric, reverse shear cleavage (RSC), rotated porphyroclasts, and shear bands. These kinematic indicators show a consistent dextral shear sense. The width of the zone and low S-C angle indicate significant offset. Locally there are small sinistral shear zones that crosscut the main foliation and are interpreted to be conjugate. The mylonitic foliation is commonly folded into open to tight shallowly northeast-plunging asymmetric folds. There are pegmatite dikes that are parallel to mylonitic foliation and which commonly displays pinch and swell. There are also late pegmatites that form in "gaps" in the mylonite. Cutting the mylonite at this locality is a prominent is a late pegmatitic dike (∼2-3 m thick) composed of coarse-grained granite with large crystals of hornblende that form radiating and linear aggregates. It exhibits no deformation fabric and clearly postdates the ductile deformation of the quartzofeldspathic gneisses. Unfortunately, there are no radiometric age constraints for the pegmatite.

Stop 3b:
This small roadcut again displays highly-sheared quartzofeldspathic gneisses and thin (10-50 cm) layers of amphibolite and garnet-biotite-quartz-plagioclase gneiss on both pavement surfaces and vertical cuts.

STOP 4: FINGERBOARD MTN SHEAR ZONE AND LAKE TIORATI METADIORITE
The next series of substops will require a moderate walk (∼2 miles and 300 vertical feet) along hiking trails and finally though some open forest. We will examine another spectacularly exposed ductile shear zone within quartzofeldspathic gneiss along the Appalachian Trail on the crest of Fingerboard Mountain and then look at field relations and chemistry of a sheared mafic metaplutonic rock along the shore of Lake Tiorati. Bring food and water, as we will stop near the top of Fingerboard Mountain to eat lunch.
Stop 4a: Here at the shelter, and for the next half mile northward along the Appalachian Trail on the crest of Fingerboard Mountain, a series of nearly continuous pavement surfaces exposes mylonites of another 1- to 2- m wide, 10- to 20-km long ductile shear zone. The rock is quartzofeldspathic mylonitic gneiss with mineralogy and structure very similar to that observed in the Indian Hill Shear Zone at Stop 3. Foliation is either vertical or very steeply dipping to the northwest or southeast with shallow plunging (20° NE) lineation indicating dominantly strike-slip ductile shearing (Fig. 12A). Like the Indian Hill Shear Zone at Stop 3, kinematic indicators (e.g., S-C fabric, rotated porphyroclasts) show consistent dextral shear sense and the very low S-C angle indicates significant offset.
Follow the blue trail above the shelter to the junction with Appalachian Trail. Turn right (north) and walk along the crest of Fingerboard Mountain for about ∼0.5 miles (rocks of Stop 4a above are here). Descend about 80 vertical feet into small col and then make a right (east) turn into the woods following the drainage for ∼0.1 mile. Ascend slightly up the north flank of the drainage through a series of closely spaced outcrops near the contact of quartzofeldspathic mylonite and amphibolites of the metavolcanic unit (not well exposed). Traverse northeastward for ∼0.2 miles along the upper eastern flank of Fingerboard Mountain and stop and examine one of the larger exposures of sheared metadiorite.
Stop 4b: This outcrop is still within the Fingerboard Mountain Shear Zone, but we have crossed over the contact with the quartzofeldspathic gneiss and into a distinctively coarse-grained, sheared mafic metaplutonic rock, here referred to as the Lake Tiorati Metadiorite (after Gates et al., 2001c). This particular outcrop is part of a relatively large body that is ∼200-300 m thick and is ∼5-6 km in length located on the west side of Lake Tiorati (originally mapped by Dodd, 1965). The mineralogy of the rock at this locality is essentially all hornblende and plagioclase, with minor clinopyroxene and biotite. Foliation, defined by compositional layering, strikes northeast and dips steeply (∼65°) to the northwest. Interestingly, mineral lineations plunge steeply to the northwest, indicating local thrust or normal motions on this part of the shear zone. The metadiorite is also cut by an undeformed granite pegmatite on the top of the outcrop.
Walk east-southeast, directly downhill toward Lake Tiorati for about ∼0.2 miles. Reach the base of the hill and Seven Lakes Drive and then walk a few hundred meters northward to a small roadcut marked by a stick placed into a drill hole near the top of the outcrop.
Stop 4c: This outcrop is the type locality for the Lake Tiorati Metadiorite and is located near the eastern margin of the Fingerboard Mountain Shear Zone. It is part of the same large body as mentioned in Stop 4b (Fig. 12A). At the southern end of the outcrop, the rock is essentially undeformed with good igneous textures developed (Fig. 13B). The diorite is composed mostly of plagioclase, hornblende, and clinopyroxene with minor orthopyroxene, magnetite and ilmenite. The orthopyroxene occurs as brown cores surrounded by coronas of clinopyroxene and/or hornblende. There is also a large xenolith of well-foliated biotite-plagioclase-quartz (metasedimentary) gneiss in the upper part of the outcrop. In the central and northern parts of the exposure, the diorite is cut by several anastomosing mylonite bands. The mylonite strikes northeast and is near vertical. Lineations plunge shallowly to the northeast. Kinematic indicators include rotated porphyroclasts and S-C fabrics (Fig. 13A). Where it can be determined, shear sense is consistently dextral. Recent SHRIMP U-Pb dating of small, subhedral zircons with minimal zoning obtained from undeformed metadiorite from this outcrop yielded a cluster of concordant ages averaging 1008 ±4 Ma (Fig. 6). Since this body is clearly cut by right-lateral ductile shear zones, this age provides an upper limit on the ductile deformation event that produced the Fingerboard Mountain Shear Zone.
Whole-rock major and trace element chemistry of samples from this outcrop indicate a mafic (∼50% SiO2, see Table 1) calc-alkaline, arc-like affinity for these rocks (Figs. 7A and C). We interpret the arc signature in these rocks to have been inherited from the continental lithosphere during magma generation and emplacement, and thus they do not indicate that there was active subduction zone in this part of the Grenville at the time. Therefore, based on field relations, geochronology, and geochemistry, the Lake Tiorati Metadiorite is best interpreted as a syntectonic, mafic plutonic rock that was emplaced just prior to or synchronously with a major right-lateral, ductile shearing event.

Large roadcuts of metapelitic gneisses on right (southeast) side of highway.

Classic Storm King Granite gneiss
Very large roadcuts on the right (southern) side of highway of classic Storm King Granite gneiss on the northern flank of the Bear Mountain pluton. Nearly continuous outcrop for next 0.5 miles to the east.

STOP 5: STROMATIC MIGMATITE, RESIDUAL DIATEXITE
AND SHEETS OF CANADA HILL GRANITE.
Immediately tum right on Hessian Drive. Park in nearest lots for park lodges. Walk north -0.1 mile, through middle of traffic circle, to prominent, but low, roadcut on the northern side of the circle and on the east side of NY Rte 9W.
In general, this exposure is typified by stromatic migmatite with local zones of residual diatexite in metaturbidite (Fig. 14A and B), presumably similar in protolith to that of rocks at Stop 1. The layered structure of this rock is moderately NE-dipping with a penetrative, yet variably oriented, moderately NE-plunging mineral lineation. Sub-concordant leucogranite sheets are penetrative, with one ∼0.5 m-thick sub-concordant sheet at the structural bottom, and one at the structural top of the exposure. In views perpendicular to the main fabric (best at the top and south-facing parts of the exposure, the typical tripartite structure of the stromatic migmatite is seen where the rock is dominated by the Bt+Grt-rich host (melt-depleted) to cm-scale concordant lenses and sub-concordant elongate pods of Pl+Qtz leucosome bound and draped by mm-scale concordant Bt+Grt+Sil melanosomes. Locally, these rocks have distinctive cm-scale layers of Gr+Py+Grt with subordinate Bt+Qtz+Kfs+Pl+Sil lenses. The residual diatexite in this outcrop is defined by Grt+Bt+Pl+Sil (Grt is cm-scale) in apparently melt-depleted zones of the outcrop, typically associated with cm-scale granite sheets and pods. The diatexite shows ubiquitous apparent disruption of the host-rock structure, typically marked by an abundance of granitic material. Locally, the diatexite is apparently melt enhanced (relatively more Pl+Qtz in the matrix) coincident with the most disrupted zones of the exposure, consistent with melt-enhanced granular flow of this rock during deformation. Each of these zones are delimited by a aggregate of cm-scale Grt and Bt melanosomes consistent with melt flow at the grain scale. Meter-scale dikes and sills of similar leucogranite to that found in this exposure are found in other exposures in the vicinity of West Point to the north may represent melt-escape structures to the melt lost from this rock.

Bridge over deep canyon of the Popolopen Creek
Bridge over deep canyon of the Popolopen Creek ("The Hell Hole") and views of Anthony's Nose and the Hudson River.

Canada Hill Granite
Beginning of ∼1 mile long series of roadcuts on both sides of the highway in migmatite with prominent lenses and pods of leucogranite sheets (Canada Hill Granite).

STOP 6: STROMATIC MIGMATITE AND CANADA HILL GRANITE
Park on the right side where a very wide shoulder exists at the entrance to James I. O'Neill High School. Walk down entrance ramp to large roadcuts on both sides. Mind the traffic at this stop.
Large roadcuts along exit/entrance ramps on the west side of the NY Rte 9W expose excellent examples of migmatitic metapelites and several, thin (5-10 m) inter-layered sheets and pods of distinctively blue-gray to white, leucocratic Canada Hill Granite at the northeastern margin of the Crystal Lake Pluton (Figs. 15 and 16A) (Heleneck and Mose, 1976; 1984). This outcrop was part of a Hudson Highlands field trip (Trip B-1, Stop #5) run during the 48th NYSGA in 1976 (Helenek and Mose, 1976). It is also from this locality that Aleinikoff and Grauch (1990) obtained a conventional, multigrain, TIMS U-Pb zircon crystallization age of 1010±6 Ma for the Canada Hill Granite and 1010±4 Ma for thin leucosome from the migmatitic metapelite.
The Canada Hill Granite at this locality is a coarse- to very coarse-grained equigranular rock that is composed of quartz, white K-feldspar, and white to gray plagioclase in roughly equal proportions (Fig. 16B). Locally, Pl dominates this rock. Biotite is ubiquitous as the mafic phase with accessory amounts of sphene, apatite, and zircon. Garnet is locally abundant, especially near contacts with the enclosing migmatites and is interpreted to represent undigested xenocrysts from those rocks after prolonged anatectic erosion of migmatite schollen. The host rock to the granite is stromatic migmatite with a strongly residual Bt+Grt+Pl+Qtz+Sil composition. Locally, the migmatite is schollen in the granite sheets with apparent anatectic erosion at the edges. The structure is regional (N30E, moderate SE-dip). Contacts between the migmatite and granite are typified by irregular structure and aggregates of Grt+Bt (cm-scale garnet), 1-3 cm thick. This is best seen in the exposure on the outside of the ramp (Figs 16A and B). Exposures in the median of the ramp show folded stromatic migmatite. The granite is predominantly massive textured with only local development of a weak foliation associated most commonly at contacts with migmatite host rocks. The Canada Hill Granite here is in sheet and pod form, concordant with the foliation in the surrounding moderately SE-dipping migmatite (∼N30E, 60SE). Contacts are generally gradational and diatexitic, except locally where the granite truncates foliation in the migmatite. This is best seen at south end of the northern median outcrop where granite sheets cut foliation near the hinge zone of complexly folded stromatic migmatite. Similar crosscutting relations have been documented elsewhere in the Hudson Highlands. For example, Helenek and Mose (1984) and Ratcliffe (1992) observed the Canada Hill Granite cutting gneissic fabric elements within the Storm King granite gneiss, metapelitic, and metavolcanic rocks in the Popolopen Lake and Oscawana Lake quadrangles to the south and east. In the Bear Mountain area, Helenek and Mose (1984) found xenoliths of Storm King granite gneiss enclosed in Canada Hill Granite along the southern margin of the Brooks Lake Pluton. Helenek and Mose (1984) also noted a weak axial planar foliation and the location of Canada Hill Granite in the hinges of large-scale, broad, open, upright, plunging folds (e.g., Bear Mountain synform, West Point antiform, Fig. 15) and interpreted this to indicate that the Canada Hill Granite was syntectonic to and slightly deformed during a late-stage Grenville orogenic event. This event is constrained to be syn- or post-1010 ±6 Ma, based on the current U-Pb zircon ages of the Canada Hill Granite (Aleinikoff and Grauch, 1990).

A7: PALEONTOLOGY AND STRATIGRAPHY OF THE MIDDLE HELDERBERG GROUP AT RICKARD HILL ROAD, SCHOHARIE, NEW YORK
James R. Ebert
Earth Sciences Department
SUNY College at Oneonta
Oneonta, New York 13820-4015
Ebertjr@oneonta.edu
Nancy A. Elliott
Learning Support Services
Elliotna@oneonta.edu
Introduction
Although the Helderberg Group (Upper Silurian and Lower Devonian) has been the subject of study for over 150 years, there are still many questions unanswered, particularly in the areas of regional stratigraphy, sedimentology, paleontology, and paleoecology. Rickard's (1962) stratigraphic synthesis has been the foundation for all modern studies of these rocks. The lower units of the Helderberg Group (Manlius, Coeymans, Kalkberg and New Scotland formations) have served as the "textbook" example of deposition in shallow, epicontinental seas since Laporte (1969). However, even these well-studied units still hold surprises (see Ebert and Matteson, this volume and Ebert and Matteson, 2003). Middle and upper parts of the Helderberg Group have been less well studied (Ebert, 1987). In general, the lower formations may be regarded as recording stepwise, rather than continuous transgression (Ebert and Matteson, 2003). A similar interpretation probably holds for the upper units (Becraft, Alsen and Port Ewen) as well (Ebert, 1987). However, the interval between these transgressive sequences is poorly understood in comparison. It is these middle units which will be examined on this field trip.

A7: PALEONTOLOGY AND STRATIGRAPHY OF THE MIDDLE HELDERBERG GROUP AT RICKARD HILL ROAD, SCHOHARIE, NEW YORK

Rickard Hill Roadcut

B1: A VISIT TO THE LAND OF GILBOA
R.Titus
Hartwick College
Department of Geological and Environmental Sciences
Hartwick College has been built upon Oyaron Hill. The slopes here offer a number of outcrops which intermittently expose almost 200 m of strata. These record a complete facies pattern spanning environments that range from offshore marine, to coastal, to well inland on the Catskill Delta complex. All sedimentary environments are well exposed and so Oyaron Hill offers a compact opportunity to visit nearly all facies of the famed Catskill Delta, herein called the "Land of Gilboa.

STOP 1. BAGNARDI'S SHOE REPAIR
The outcrop exposed here is of the Unadilla Formation. The sequence is composed of moderately well bedded sandstones and some shales. There are two levels which represent shell hashes. These are interpreted as lag deposits, dating back to times of unusual current activity. There is at least one wave ripple marked horizon. The beds have been thoroughly searched for fossils these past 30 years and not many remain. In the past a brachiopod and clam assemblage dominated. Nautiloids, snails, trilobites, rare bryozoa, trace fossils were also found. The sequence was first described by Prosser in 1899. He listed 13 species, mostly brachiopods and clams. These strata are interpreted as representing a shallow water marine setting lying some distance offshore of the prograding delta sequence.

STOP 2. HARTWICK MAINTENANCE BUILDING
When this outcrop was exposed, about ten years ago, it was hoped that the site would yield numerous fossils. Sadly, this was not the case. Most of the sparingly exposed strata between here and Miller Hall are sparsely fossiliferous at best. These facies may represent a "dead zone" lying just offshore of the delta river mouths.

STOP 3. THE MILLER HALL OUTCROP
The Miller Hall outcrop is situated immediately above a thick sequence of mostly unfossiliferous black shales. These were exposed when Miller Hall was recently expanded. These are apparently the very first non-marine deposits in the local Catskill sequence and they thus mark the moment in time when progradation of the Catskill Delta complex had reached today's Oneonta. The facies here are therefore most likely to be the most nearshore that is possible. This is of some importance as this level thus marks the lowest level of the Oneonta Formation. This part of the sequence was not exposed when workers such as Vanuxem (1840) or Prosser (1899) studied here.
This fine exposure reveals several distinctly different nearshore lithologies. First there is the massive sandstone unit that thins out to the right. Then there are the black shales that underlie the entire outcrop. Between the two lithologies is a thick bedded, rusty sandstone with bedding gently inclined toward the massive sandstones.
The massive sandstone appears to be a stream channel, possibly a sand filled abandoned channel. The base of the deposit rises to the right, apparently up the slopes to the bank. Beyond and farther to the right is what may be an overbank crevasse splay deposit. In the past the very bottom of the exposure has displayed drag marks which are east to west oriented. These are currently buried.
The black shales are sparsely fossiliferous. A few large horizontal trace fossils have been found along with the brackish water clam Paleoneilo sp. There are several horizons of deep scouring in the otherwise very thinly laminated sequence.
The third lithology here is composed of rusty colored sandstone beds that are inclined toward the stream channel. These strata are richly fossiliferous with fragments of land plants. This deposit may represent a natural levee or bar finger deposit.
All in all, the Miller Hall outcrop is related to settings on the lower Mississippi River delta. An analogous location is found on the West Delta 15" Quadrangle. There the Tom Loar Pass distributary flows past Bob Taylors Pond and Zinzin Bay. A channel deposit along with natural levee deposits can be expected to overlay lagoonal black shales.

STOP 4. UPPER OYARON HILL
The remainder of the slopes of Oyaron Hill rise from elevations of about 435 to 525 meters. Throughout this section are a number of sandstone ledges providing good exposures of the Oneonta Formation. In between the sandstones are less well-exposed sequences of red shales and siltstones. The Oneonta was named by Conrad in 1841. Vanuxem (1842) described this location.
The sequence seems to represent the distal reaches of the Catskill delta sequence. The sandstone ledges are likely to be the deposits of meandering distributary streams while the fine-grained deposits are most likely overbank deposits. The sandstones have occasionally been observed lying with erosional truncation of underlying overbank deposits.
The sequence here displays a number of fining upwards sedimentary cycles of the sorts that have been described by Bridge and his colleagues among others (see for example Bridge, 2000, Bridge and Gordon, 1985, Gordon and Bridge, 1987, Willis and Bridge, 1988 and Bridge and Willis, 1994).
Within the sandstones are abundant trough cross beds, especially near the bottoms of the ledges. Also found are planar cross beds and horizontally laminated sandstones. Near the tops of the ledges there are occasional ripple marked horizons.
Of special interest is the presence of vertical clam escape burrows (probably Archanodon catskillensis). These are found at the "Table Rock" ledge near the top of the hill. The clams seem to have been buried by a meter of sand at the time of a waning flood (see Thoms and Berg, 1985; Bridge et al., 1986 and Gordon, 1988). They produced well-defined meniscus patterns as they worked their ways up through the fresh deposits.
The best overbank deposits are found at the top of the hill, along the sides of the soccer field. Here red shales and fined grained red sandstones are seen in abundance. The deposits are rich in root casts and also display rare Isopodichnus trace fossils. There are ripple marks and possible soil horizons. At the very top is a sequence of alternating shales and sandstones which seem to record a sequence of occasional flood events. One, exposed on the surface of an enormous displaced boulder, has several fossil logs on it. These were apparently displaced during a flood event.

B2: GRABAU'S "TRANSITION BEDS" - KEY ELEMENTS IN A RADICAL REVISION OF HELDERBERG STRATIGRAPHY
James R. Ebert and Damon K. Matteson
SUNY College at Oneonta
Oneonta, New York 13820-4015
Ebertjr@oneonta.edu; Mattesdk@oneonta.edu
INTRODUCTION AND PREVIOUS WORK
The nature of the contact between the Manlius and Coeymans formations is a recurring theme in the literature of the Helderberg Group. In the early twentieth century, Grabau (1906) interpreted the contact as gradational, with a distinctive zone of "transition beds" present between true Manlius and true Coeymans. Grabau's transition beds are best developed in eastern New York in and near the Schoharie Valley and appear to be absent from other areas (Fig. 1).
In succeeding decades, the contact was regarded as a disconformity, with views varying as to the magnitude of hiatus. Goldring (1935, 1943) indicated that as early as 1927, Chadwick saw evidence for some hiatus (see also Chadwick 1944). In central New York, Smith (1929) described the contact of the Jamesville Member of the Manlius Formation with the overlying Coeymans as unconformable. Logie (1933) felt that the erosion was significant with the Elmwood, Clark Reservation and Jamesville members of central New York absent in eastern parts of the state via erosion. Chadwick (1944, p. 152) reported the contact as "irregular, undulating" and concurred with Logie that the absence of the higher members of the Manlius in eastern New York was the result of significant erosion.
Rickard (1962) reinterpreted the Manlius-Coeymans contact as sharp, but conformable in eastern New York. In central New York, he saw the contact as complex with portions of the Coeymans Formation (e.g., Dayville Member) and higher parts of the Manlius (Elmwood, Clark Reservation and Jamesville members) and other members of the Coeymans (Deansboro/Ravena members) inter-tonguing (Fig. 2). Rickard's stratigraphy became the foundation for nearly all succeeding work on the Helderberg Group. The most notable example is Laporte's (1969) reconstruction of depositional environments.
Rickard (1962, p. 105) acknowledged the possibility that a "post-Jamesville" (subCoeymans) disconformity in central New York might extend into the middle of the Coeymans in east-central New York and to either the top or base of the Coeymans in the Hudson Valley. However, he discounted this interpretation based on the absence of an obvious faunal or physical break and the presence of the transition beds at Cherry Valley. Later, Rickard and Zenger (1964, p. 51) reported that "a layer of pebbles and abundant Favocites sp. occupies the position of the Jamesville with Dayville and Deansboro crinoidal limestones with Gypidula coeymanensis below and above" in portions of the Richfield Springs 15' Quadrangle. Whereas this may be viewed as evidence for erosional truncation of the Jamesville below the Coeymans, Rickard and Zenger (1964) retained the interpretation of Rickard (1962) of gradational facies relationships, as did Laporte (1969).
In the 1980's, Goodwin and Anderson published a series of papers outlining and applying their hypothesis of Punctuated Aggradational Cycles (PACs - a type of small scale allostratigraphic cycles) to the Helderberg Group in New York and elsewhere in the Appalachian Basin. Where the Coeymans Formation overlies the Thacher Member of the Manlius Formation, Goodwin and Anderson (1988) interpreted the contact as a minor, cryptic unconformity. They documented the progressive cutting out of more of their PACs from central New York eastward. However, they continued to support coeval relationships between higher parts of the Manlius and the Coeymans formations by correlating P ACs through these units.

B2: GRABAU'S "TRANSITION BEDS" - KEY ELEMENTS IN A RADICAL REVISION OF HELDERBERG STRATIGRAPHY

Schenevus Creek. Note the well developed cut banks on large meanders.

Outcrop of Gilboa Formation
Outcrop of Gilboa Formation with well-developed ball and pillow structures (see Hutchinson, 1977). Opposite is NYSEG power station and dam on Susquehanna River to form Goodyear Lake.

Outcrops of Onondaga Formation.

Discontinuous outcrops of Tristates Group

Outcrop of Kalkberg Formation

Outcrop of Coeymans Formation
Outcrop of Coeymans Formation in contact with Kalkberg Formation. Contact is marked by a silt-rich bed and the Punch Kill Unconformity of Ebert and Matteson (2003). See also description below for STOP 1.

STOP 1. Cherry Valley- Sprout Brook Road, Otsego County Rt. 32
This road cut is the stratigraphically lowest of a series on Sprout Brook Road (Cty. Rt. 32), Rt. 166 and, U.S. Rt. 20 that expose the entire Lower Devonian section that is present in New York State (See Brett and Ver Straeten, 1997). This outcrop corresponds approximately to Rickard's (1962) section number 94, which was measured in nearby Judd's Falls.
The section begins at the north end of the outcrop with approximately 1 meter of the Rondout Formation. The Rondout is abruptly overlain by the Thacher Member of the Manlius Formation. Near the top of the Manlius Formation, the style of bedding becomes thinner (decimeter scale) and intercalations of dark gray to black shale appear. These beds (total thickness = 0.82 m) constitute an erosionally thinned portion of the Green Vedder Member ("thin-bedded upper Thacher" of Rickard (1962)). These dark shale interbeds have yielded the carbonized biota discussed above and in Matteson, Natel and Ebert (1996).
The contact with the overlying Coeymans Formation has traditionally been placed at the appearance of the first, coarse crinoidal bed. This bed overlies the Terrace Mountain Unconformity (13.27 m from base of section). The lower Coeymans (6.03 m thick) comprises interbedded coarse and fine beds which Rickard equated with the Dayville Member. At the top of the Dayville Member, there is a change in the style of bedding to thicker and more irregular beds. The contact is marked by an erosional surface with several centimeters of relief. Sporadic, in situ favocitid corals encrust the contact. Rare, fine grained intraclasts indicate that this surface marks the position where the Elmwood and Clark Reservation members of the Manlius (which overlie the Dayville to the west) have been removed. This is the Howe Cave Unconformity (19.3 m from base of section).
Overlying the Howe Cave Unconformity, thick, irregular to vaguely nodular beds of the upper Coeymans here have been referred to both the Deansboro and the Ravena members of the Coeymans Formation (see Rickard, 1962). This unit continues to the top of the outcrop (26.6 m). Using distinct horizons of chert and beds bearing holdfasts of the cystoid Lepocrinites gebhardi, we have correlated beds from the top of this outcrop with the base of the small outcrop 0.2 miles to the south on Sprout Brook Road. At the latter outcrop, 4.7 m of the Ravena Member is exposed. Thus, the maximum thickness of the Coeymans (Dayville and Ravena combined) is slightly over 19 meters. Previous accounts of the stratigraphy in this area (e.g., Rickard, 1962, 1981; Gurney and Friedman, 1986; Liebe and Grasso, 1990) reported a covered interval between the main outcrop and this smaller exposure and a total thickness of 30 meters for the Coeymans Formation (Rickard, 1981). However, it appears that these authors did not take into account the regional dip, which is approximately two degrees in this area. When relative elevations, the regional dip and the distance between these exposures are utilized, the top of the large outcrop and the base of the smaller exposure coincide. Thus, it appears that previously published measured sections for Cherry Valley have exaggerated the thickness of the Coeymans Formation by as much as 50%. Although we will not visit the smaller exposure on this trip, the Punch Kill Unconformity (Ebert and Matteson, 2003) is well exposed here, where it separates the Coeymans Formation from approximately 2m of the Kalkberg Formation.
Taken together, the Green Vedder beds and the Dayville Member appear to constitute a gradational transition between the Manlius (proper) and Coeymans formations (Rickard, 1962, 1975, 1981). However, such a gradation cannot exist owing to the presence of the regionally traceable Terrace Mountain and Howe Cave unconformities which cap each unit respectively. Furthermore, the upper Thacher (Green Vedder Member) and Dayville Member exposed on Sprout Brook Road correlate to the west with units bearing the same names. They are distinct units and not an ambiguous zone of transition, assignable to neither formation. When correlated westward, they are separated by other members of the Manlius Formation. When traced eastward, these units thin and become the zone that Grabau (1906) designated as transition beds in the Schoharie area. Grabau's (1906) transition beds (sensu strictu) will be seen at STOP 2.

Kalkberg outcrops

Kalkberg outcrops with well-developed K-bentonites
Railroad overpass. Kalkberg outcrops with well-developed K-bentonites (see Smith, Berkheiser and Way, 1986; Ebert, Applebaum and Finlayson,1992; Tucker, et.al., 1998). Kalkberg is disconformably overlain by the Oriskany Sandstone.

Outcrops of Tristates Group
Outcrops of Tristates Group which also bears K-bentonites (see Ver Straeten and Brett, 1997).

Begin outcrop of Onondaga Formation.
Edgecliff, Nedrow and Moorehouse Members are well exposed (see also Brett and Ver Straeten, 1997).

Small outcrop of Seneca Member of the Onondaga Formation.
The Tioga-B (Onondaga Indian Nation) K-bentonite is exposed at the base of the outcrop. This outcrop has provided conodonts which are the subject of direct radiometric dating studies (Research is in progress, but see Elrick, et. al. (2002) for a preliminary report.)

Outcrop of the Marcellus Formation
Outcrop of the Marcellus Formation, including Union Springs, Cherry Valley Limestone and Chittenango members, exposed on Chestnut Street, sub parallel toRt. 20 (see Griffing and Ver Straeten, 1991; also Brett and Ver Straeten 1997).

Outcrops of the Kalkberg Formation at the edge of Leesville.
Exposed section is similar to the outcrop at Cherry Valley (40.0 in this road log).

Outcrops of Kalkberg at edge of Sharon Springs.
Section is similar to Leesville and Cherry Valley.

Outcrop of Onondaga Formation.

Outcrop of Ordovician flysch of the Schenectady/Frankfort Fm.

Note view of Terrace Mountain to the right.

Dunkin' Donuts- a convenient rest stop in this area.

STOP 2: Terrace Mountain Road, Schoharie
In the vicinity of Schoharie, there are numerous outcrops of Grabau's (1906) transition beds. This exposure on Terrace Mountain Road probably postdates Grabau's investigations, however it is one of the most accessible outcrops for examining these beds. Therefore, we selected it for this trip, rather than one of Grabau's "classic" outcrops, many of which still exist.
In this exposure, the two-fold character of Grabau's transition beds is readily apparent. The lowest beds in the outcrop comprise a thinned eastward extension of the Green Vedder Member (Rickard's "thin-bedded upper Thacher"). The Terrace Mountain Unconformity is marked by the abrupt appearance of coarse, crinoidal grainstones and packstones of the Dayville Member of the Coeymans Formation. Note that this is an extension of the term Dayville. In Rickard's (1962) stratigraphy, the Dayville did not extend eastward beyond Cherry Valley. Several of the Dayville beds display well-developed structures associated with tempestites: vertical grading, undulating to hummocky cross stratification and rare symmetrical ripple caps.
The Dayville has been substantially thinned from 6 m at Cherry Valley to approximately 1 m in the Schoharie area. This thinning is a result of the eastward descent of the Howe Cave Unconformity, which separates the Dayville from the overlying Ravena Member of the Coeymans Formation. Minor local relief on the Howe Cave Unconformity is present in the Schoharie region, where the thickness of the Dayville remnant ranges from 0.7 to nearly 1.25 m at various outcrops.
Overlying the Howe Cave Unconformity, thick beds of the Ravena Member of the Coeymans Formation comprise the majority of this outcrop and they extend to the top of the exposed section. Upper portions of the Ravena display abundant club-like holdfasts of the cystoid Lepocrinites gebhardi. We are currently investigating the lateral extent of this subdivision of the Ravena Member (See also Matteson and Ebert, 2001 ).

To the left and uphill, a large outcrop of the Helderberg Group
To the left and uphill, a large outcrop of the Helderberg Group is visible on I-88. This outcrop exposes the Kalkberg, New Scotland and Becraft formations.

The original Lester Howe Hotel and future Cave House Museum.
Also, original entrance to Howe Caverns.

STOP 3: Howe Cave Quarry
The inactive Howe Cave Quarry is the type section of the Howe Cave Unconformity. Clear truncation of the upper transition beds (eastern extension of the Dayville Member) in this quarry enabled recognition of this unconformity within the Coeymans Formation as presently defined. The truncation of beds in the Dayville is best observed in the weathered joint surface on the high wall at the entrance to the quarry, just beyond the old Lester Howe Hotel. In this face, the Green Vedder Member of the Manlius Formation (a.k.a. "thin-bedded upper Thacher") is also readily apparent. The upper part of the high wall is comprised of the massive Ravena Member, which, in parts, bears abundant holdfasts of Lepocrinites gebhardi. Elsewhere in the quarry, the Punch Kill Unconformity is exposed, overlain by approximately one meter of the Kalkberg Formation. This is visible at the top of the high wall in the main part of the quarry, but is only accessible for direct observation in a few places.
Other features of interest in the quarry include a thrust fault (Marshak and Bosworth, 1991; Mylroie and Palmer, 1977), extensive fields of ripples on the floor of the quarry (Ebert, et.al., 2000), numerous glacial striations on the surface above the quarry and, various karst and cave features associated with Howe Caverns and Barytes Cave (Mylroie and Palmer, 1977 and this volume), including the original entrance to Howe Caverns, adjacent to the Lester Howe Hotel.
The Howe Cave Quarry and Lester Howe Hotel will be centerpieces of the future Cave House Museum of Mining and Geology, a scientific and industrial educational facility, which is in the initial stages of development. In addition to various geologic, hydrologic and biologic aspects, the Museum will feature a history of mining and the mining industry, as well as multiple activities in which the public will be able to observe active mining, which is scheduled to resume in parts of the quarry. For additional information on the Cave House Museum, contact the Education Coordinator and member of the Board of Directors, Benson Guenther at the following:
The Cave House Museum of Mining and Geology
P.O. Box220
Cobleskill, New York 12043
E-mail: Cavehouse@cobleskillstone.com

B3: EXAMPLES OF ALONG-STRIKE CHANGES IN FOLD-THRUST BELT ARCHITECTURE; STRUCTURAL GEOLOGY OF THE ROSENDALE NATURAL CEMENT REGION, ULSTER COUNTY, NEW YORK
Kurtis C. Burmeister and Stephen Marshak
Department of Geology
University of Illinois
Urbana, Illinois 61801
INTRODUCTION
It was necessary for the leaders of this trip to specify a set number of vehicles to obtain permission to visit several of the following stops. For this reason, personal cars will not be allowed to follow the field trip caravan. Throughout much of the 19th century, the Rosendale natural cement region (Figure 1) was recognized worldwide as a center for the production of high-quality hydraulic cement. Although the industry collapsed in the early 1900s, the geology of this region continues to attract attention. The Rosendale area contains the northernmost portion of the central Appalachian foreland fold-thrust belt. This fold-thrust belt segment lies in the southern arm of the New York recess, a convex to the foreland, map-view curvature in the Appalachian orogen. Fold-thrust belt deformation across the New York recess involves a mechanically rigid strut of Late Silurian through Middle Devonian sedimentary strata (Figure 1; Wanless, 1921; Waines and Hoar, 1967; Rodgers, 1971; Marshak, 1986; Epstein and Lyttle, 1987; Marshak and Tabor, 1989; Marshak, 1990). This rigid strut is sandwiched between thick, relatively ductile units of Ordovician shale (below) and Middle Devonian shale (above). Near Rosendale, the Siluro-Devonian strut thins markedly from greater than 1000 m thick in central Pennsylvania to little more than 100 m thick north of Kingston. These contrasts in the pre-deformational stratigraphy of the affected units give rise to dramatic, along-strike transitions in the scale, style, and trend of structures in the fold-thrust belt across the New York recess. Preliminary balanced cross sections based on recent geologic mapping in the vicinity of Rosendale, combined with recently re-discovered historical photographs, are providing new insights into the structural relationships underlying these along-strike transitions in fold-thrust belt architecture. Specifically, the south to north thinning of the Siluro-Devonian rigid strut that takes place near Rosendale appears to trigger a south to north: 1) slight westward rotation of structural trends in the fold-thrust belt; 2) eastward migration of the pin line of deformation resulting in a narrower cross-strike fold-thrust belt width; 3) tightening of fold amplitudes and wavelengths; 4) decrease in the spacing between thrust faults and the development of lateral ramps; 5) northward dying out of thrust faults into fault-propagation folds, and 6) redistribution of regional detachment horizons.
The Rosendale natural cement region is ideally suited for the study of along-strike changes in the architecture of fold-thrust belts. In particular, the unique regional stratigraphy provides a basis for directly examining the role of changing mechanical stratigraphy in the development of along-strike changes in the internal architecture of these tectonic provinces. Further, the fold-thrust belt near Rosendale contains all of the structural complexities of regions like the Valley and Ridge Province ofPennsylvania, yet with a cross-strike width of little more than 2 to 8 km, it is possible to walk transects of the fold-thrust belt in one day. This excursion will focus on the following aspects of the Rosendale region's geology. The Rosendale region is an ideal place to examine: 1) fault mechanisms, including ramp-flat thrust fault geometries, accommodation faulting, and duplex structures; 2) fault-bend, fault-propagation, and detachment folding; 3) strain distribution in thrust sheets; and 4) the controls of rock type and structural relationships on cleavage morphology and intensity. In addition to structural relationships, the strata in the Rosendale region preserve an abundance of Early to Middle Devonian marine fossils, so we can examine examples of classic shallow-marine carbonate facies successions (Rickard, 1962; LaPorte, 1969), and outcrops of the region provide exposures of the Taconic angular unconformity (Rodgers, 1971; Toots, 1976; Epstein and Lyttle, 1987).

B3: EXAMPLES OF ALONG-STRIKE CHANGES IN FOLD-THRUST BELT ARCHITECTURE; STRUCTURAL GEOLOGY OF THE ROSENDALE NATURAL CEMENT REGION, ULSTER COUNTY, NEW YORK

Shallowly west-dipping beds of the Union Springs Formation
Road cut exposures of shallowly west-dipping beds of the Union Springs Formation along the left (northwest) side of Route 209 North

Bridge crossing Esopus Creek

Broad, open anticline in Schoharie Formation and Onondaga Limestone.

STOP 1. Folds exposed in road cuts along State Route 209 north of Kingston
During our drive east, we have passed from the foreland of the Appalachian orogen into the core of the Hudson Valley fold-thrust belt north of Kingston. Here, road cuts along State Route 209 expose large, slightly asymmetric open folds with no visible thrust faulting or mesoscopic folding. The geology of this stop was first described by McEachran (1985) and later used as a field trip stop by Marshak (1990). Five major, map-scale fold hinges cross State Route 209 between Routes 9W and 32. The most prominent of the folds at this stop is a large anticline (Figure 3) involving the upper Becraft Limestone, the Alsen Formation, and the Port Ewen Formation. Note the differential development of cleavage in these units. The Becraft Limestone contains little to no cleavage, while the Alsen and Port Ewen Formations clearly possess a strong, southeast-dipping cleavage. Marshak (1990) notes that these folds are characteristic of structural styles north of Kingston, in that they trend roughly 015° and lack structural complexity. Return to the vehicles and carefully merge back into traffic. As we proceed a short distance further to the east on State Route 209, notice the west-dipping beds of the Manlius through New Scotland Formations.

STOP 2. Laterally ramping thrust faults in road cut along Route 32 north of Kingston
Four major thrust faults duplicate the Rondout and Manlius Formations in a series of imbricate thrust sheets in the northern half of the long road cut along the west side of State Route 32 (Figure 4). The geology of this stop was first described in detail by Waines and Hoar (1967) and was further interpreted by McEachran (1985) and Marshak (1990). The major thrust faults in this outcrop strike roughly 040° and calcite slip fibers on these faults suggest a transport direction of 060°-070°. In places, these faults have a flat-on-flat geometry (i.e. the thrust faults are parallel bedding in both the hanging wall and footwall), suggesting large displacements. In places (particularly in the upper portion of fault block 6; Figure 4), thrust faults appear to cut down section. However, this is likely an artifact of the obliquity of the road cut face to the structures, where the hanging walls are thrust westward into the outcrop. These faults could not be traced west of the road cut, but may ramp laterally up-section and die out to the north. Return to vehicle and continue south on State Route 32 South. From here, we will proceed southeast towards the Helderberg

STOP 3. Complex structural relationships of Hasbrouck Park, east Kingston
Lock cars and follow the footpath at the east end of the parking lot for roughly 75 m as it descends around the nose of a small hill of Manlius, Coeymans and Kalkberg Formations south of the trail. At the foot of this slope is an abandoned adit cut into an extensively quarried face in the Helderberg Escarpment. The geology of the Helderberg Escarpment at Hasbrouck Park was fist described by Marshak and Tabor (1989; Figure 5) and later revisited as a field trip stop by Marshak (1990). This face provides a cross-sectional view through the leading edge of a duplex composed of a stack of horses involving Rondout through Kalkberg Formation strata (Figure 5B). The basal thrust fault in this duplex, the Hasbrouck thrust, cuts laterally up section along the sloping footpath we walked along (Figure 5C), and through the woods. The Hasbrouck thrust fault reappears on the north side of Delaware Avenue, where it places the Manlius Formation over the Alsen Formation, above extensive roof-and-pillar cement mines just east of Corporate Drive. A tear fault with a trace roughly coincident with Delaware Avenue may extend to the west of the lateral ramp. The complexity of these structural relationships amongst the Silurian and Devonian strata in Hasbrouck Park is characteristic of the geology along the Helderberg Escarpment at the latitude of Kingston.
Return to main footpath and follow it further to the south, skirting along the base of the Helderberg Escarpment. Stay to the right at the first fork in the footpath and proceed up narrow ridge. From this locality, it is possible to observe steeply dipping to overturned beds exposed by quarrying of portions of the Rondout Formation. Retrace your path along the ridge back to the main footpath, return to the vehicles, and exit the park. We will now proceed south and examine some of the along-strike changes in structural architecture of the eastern margin of the northern Appalachian fold-thrust belt near Rosendale.

Note entrances to the abandoned Delaware Mine
Note entrances to the abandoned Delaware Mine on left (west) side of road. This mine was most recently used by the Knaust Brothers to grow 'Fallout Proof mushrooms during the 1950s.

Note abandoned canal barge in tidal flats off to left (east).

The Mansion House located on this corner
The Mansion House located on this corner was built in 1833 by Jervis McEntee, one of the engineers that pointed to Rosendale as a local source for hydraulic cement for construction of Delaware & Hudson Canal (Werner, personal communication, 2003).

Note various ruins of cement kilns and associated buildings along next the road over the next kilometer.

This road follows the path of what was originally a plank road
This road follows the path of what was originally a plank road and later the roadbed of the horse-drawn Hickory Bush-Eddyville railroad. A consortium of cement companies constructed and maintained this throughway to avoid the shipping fees charged by the Delaware & Hudson Canal for moving their product from Hickory Bush to Kingston (Werner, personal communication, 2003).

Abandoned roadbed for Hickory Bush-Eddyville railroad
Abandoned roadbed for Hickory Bush-Eddyville railroad diverges from course of DeWittLake Road and is visible in trees off to right (west)

STOP 4. North end of Quarry Hill, Bloomington
This stop is on private property. You must obtain permission to visit from the owners of the house near the garage before proceeding. Walk the short distance to the garage on the edge of the tree line just south of the parking lot. Upon reaching the tree line, you will notice two depressions just through the trees to the south. These are the northern end of a series of extensive abandoned cement quarries spanning the entire length of Quarry Hill. Please proceed with caution.
First, walk into the western quarry pit. Here, a small quarry exposes east-dipping strata of the Rosendale and Glasco Members of the Rondout Formation. Retrace your path back to the edge of the tree line, and circle back around to the eastern quarry opening. Watch your step as you descend into this quarry. The slope is uneven and littered with broken glass and scrap metal. Here, the Rosendale Member of the Rondout Formation was extensively mined (all that remain are support pillars) and affords a fantastic view to the south along the axis and steeply dipping eastern limb of a northwest-verging syncline (Figure 6). The eastern quarry wall contains the gradational contact between the Rosendale Member and the underlying Binnewater Sandstone. The base of the Glasco Member forms the western quarry wall.
Note the tight hinge of the syncline and recall the relationship observed in the smaller quarry to the west. A northwest dipping, out-of-the-syncline thrust fault that apparently ramps up section out of the Rondout Formation separates the two quarries at this location (Figure 7). Thus, strata exposed in the western quarry are in the hanging wall, and were thrust to the southeast over the strata in the eastern quarry. Consider the contrasts in structural style between this location and the complex duplexes and the stacking of imbricate thrust sheets near Kingston. The quarry exposures at this location afford an excellent view of the structural relationships that are characteristic along the Helderberg Escarpment between Bloomington and Rosendale. Here, the escarpment is defined by a series of left-stepping, thrust faults that cut up section to the east, out of the cores of large synclines.

STOP A. Road cut exposure of folds along Route 32, north of Rosendale
The road cut along the west side of State Route 32 at this locality exposes a tight, northwest verging anticline in the Kalkberg Formations. This structure is a continuation of the same series of folds exposed in the mines along Quarry Hill (Figure 7). The dip of the strata in the western limb of this fold continues to increase westward. Exposures of Becraft through Port Ewen Formations are steeply dipping to overturned in road cuts along the New York State Thruway.

STOP B. South end of Quarry Hill, Bloomington
This stop is on private property and permission to visit must be obtained in advance. Trespassers will be prosecuted.
Walk north along the small dirt road located across from the house, keeping to the right as the road descends into the mouth of an abandoned cement quarry. The quarry at this locality is the southernmost of the Quarry Hill mines. Both the Rosendale and Whiteport Members of the Rondout Formation are extensively quarried at this location (Figure 15). Steeply dipping beds of the base of the Rosendale Member and the Binnewater sandstone form the eastern quarry wall. Moderately dipping beds of the Rondout through Manlius Formations form the western quarry face. The large, central quarry area exposes strata of the Rosendale Member that are repeated by two major thrust faults and extensively cut by a number of smaller thrusts. The easternmost of the two major thrust faults is steeply west dipping and exhibits a hanging wall ramp on footwall flat geometry. The western major thrust fault is moderately west dipping and also has a hanging wall ramp on footwall flat geometry. The western fault contains a horse of Rosendale Member. The two major thrust faults merge near the top of the quarry. The hanging wall of these faults is involved in a northwest-verging syncline, but is apparently not faulted. Some amount of bedding-parallel slip has also occurred along the eastern quarry wall. A close examination of this surface reveals a well-developed set of non-coaxial slip fibers. The faults exposed at this location are not continuous with those exposed at the north end of Quarry Hill (Stop 4).

Well-preserved cement kilns on left-hand (northwest) side of road

Entrance to the mines on right-hand side of road
Entrance to the mines on right-hand side of road were once operated by the Lawrence Cement Company (Werner, personal communication, 2003), and are presently owed by Iron Mountain Incorporated as a future archival site.

Cement kiln ruins on right-hand (southeast) side of road

STOP 5.1. Eastern limb of the Hickory Bush anticline, Rosendale Landfill
Lock vehicles and walk approximately 100 m northeast along Hickory Bush Road to the entrance of the Rosendale Landfill and Recycling Center. Proceed through gate and to the north corner of the facility. Recent excavation along the north wall created fresh exposures of an east-dipping sequence of strata (Figure 8). The north wall of the landfill contains Martinsburg Shale, the Taconic unconformity, Shawangunk Conglomerate, High Falls Shale, Binnewater Sandstone, Rondout Formation, and Manlius Limestone.
From a distance, the Martinsburg Shale appears un-deformed, but closer examination reveals that it is strongly deformed by complex brittle faulting. Above the Taconic unconformity, the Shawangunk Conglomerate is little more than a 5 to 10 cm thick lag deposit of the characteristic milky white quarts pebbles. Overlying the Shawangunk Conglomerate is a complete and apparently continuous sequence of High Falls Shale and Binnewater Sandstone. Both the Rosendale and Whiteport members of the Rondout Formation are quarried at this location. Separating the two quarries is the Glasco Member, which contains beautiful Halycities chain corals. At the eastern end of the north wall are exposures of the overlying Manlius Limestone, which are in faulted contact.
The east-dipping strata exposed in the Rosendale Landfill form the eastern limb of the Hickory Bush anticline. The involvement of Martinsburg Shale in core of the fold suggests this structure developed as the result of slip along a detachment at depth in the underlying Ordovician strata. The large scale of the Hickory Bush anticline suggests that the underlying thrust is a master fault in the fold-thrust belt at this latitude and ramps directly from the lower detachment horizon. The thrust fault exposed at the eastern end of the exposures along the north wall is the westernmost fault in a complex imbricate fan of thrusts that ramp out of the Rondout Formation and cut through Hickory Bush Hill (the large hill southeast of the landfill).

STOP 5.2. Western limb of the Hickory Bush anticline, rail cut near Fourth Lake
Walk southwest along the rail trail from the parking lot along Hickory Bush Road for roughly 350 m until you encounter a rail cut exposing moderately northwest dipping strata on both sides of the trail. The exposed sequence includes High Falls, Binnewater, Rondout, and Manlius Formations (Figure 9). Along the northwest side of the trail, the High Falls Shale is difficult to distinguish, but it is overlain by a complete thickness of the Binnewater Sandstone. Cross bedding, ripple marks, and graded bedding are clearly visible, as is a characteristic, roughly 10 cm thick shale horizon just below the contact with the overlying Rondout Formation. This shale layer is also present in an outcrop of the Binnewater sandstone at the Snyder Estate, where it is deformed. The Rosendale and Whiteport members of the Rondout Formation are quarried at this location. Notice how air circulating through the abandoned mines keeps this area noticeably cooler during the hot summer months.
Climb onto the embankment on the south side of the rail trail using the small path through the trees located just north of the northernmost quarry opening. The embankment was once a tramway that serviced the cement quarries in this area. Watch your step. Abandoned cement kilns and sunken shed foundations are scattered along this tramway and are often difficult to see. Proceed along tramway to the southeast, skirting along an exposure of the Binnewater Sandstone. When the tramway narrows and prevents further progress, drop down the slope to the next lower abandoned roadbed and continue to the southeast. After hiking approximately 150 m from the rail trail, you will reach and intersection with a dirt road. This road is part of the Perfume Trail on the Williams Lake Resort. Follow the dirt road to the right (southwest) and you will soon see a large mine entrance. Proceed about 10m into the mine.
This mine was operated by the Lawrence Cement Company (Werner, personal communication, 2003) and is cut into the Rosendale Member of the Rondout Formation. The contact with the underlying Binnewater Sandstone forms the floor of the mine, and the contact with the overlying Glasco Member forms the ceiling (Figure 10). The quarry is in a large, asymmetric, open anticline. Note how the shallow dip of strata in the western limb of the fold gently increases to the southwest. Proceed southwest into the mine along the axis of the fold. Note the relatively rapidly increasing dip of the strata in the eastern limb of the fold. Quarrying of the eastern limb was most likely halted due to the proximity of a fault with a trace that roughly follows the path of the Perfume Trail a few meters further to the east.
Climb up and out of the mine through the opening left by the partial collapse of the roof at the southeast end of the quarry. As you climb out of the mine, pause to note the other portions of the quarried fold around the periphery of the collapsed area. It is unclear whether strata of the Whiteport Member above the quarried fold was removed by erosion or was stripped by quarrying. Follow the footpath path down to the Perfume Trail and continue to the south until you reach a intersection of several trails. Return to the vehicles by walking north along the rail trail (covered in black gravel) from this junction. Along the way, note the various quarries and outcrops as you pass through the Fourth Lake region. The Williams Lake Resort property around Fourth Lake contains some of the best exposures of structural relationships in the region, and has long been used to teach the methods of field geology.

Binnewater Road follows the trace of the axis of the Binnewater anticline.

The south quarries of the F.O. Norton Company
The south quarries of the F.O. Norton Company in east-dipping units of the Rondout Formation located in the escarpment through trees on left (east) side of road.

The Beach Mine of the Lawrenceville Cement Company
The Beach Mine of the Lawrenceville Cement Company (Werner, personal communication, 2003) in west-dipping strata of the Rondout Formation along right (west) side of road. Iron Mountain Incorporated now uses these mines as a document storage facility.

The Hoffman Quarry of the Lawrence Cement Company
The Hoffman Quarry of the Lawrence Cement Company located east of small Wallkill Valley RR trestle on left (east) side of road. A battery of abandoned cement kilns built by the Lawrenceville Cement Company is visible on the right (west) side of road (Werner, personal communication, 2003).

Stone reservoir along right (west) side of road.
This pond was created in the late 1820's to ensure a constant reserve supply of water for a feeder stream that was used to replenish water levels in the Delaware & Hudson Canal (Werner, personal communication, 2003).

Joppenbergh Hill
Joppenbergh Hill is located immediately to the left (east) of this intersection. The hill is named after Jacob Rutsen, who became the first westerner to settle in the Rosendale vicinity in 1680. The New York and Rosendale Cement Company extensively mined Joppenbergh Hill during the 1800's. Large sections of the resulting quarry faces began collapsing in 1899, and have continued to do so as recently as April 2003. Rosendale Bridge, built by the Wallkill Valley Railroad, spans Rondout Creek from the flank of Joppenbergh Hill. Locally, State Route 213 follows the northern bank of Rondout Creek, and coincides with the trace of the Delaware & Hudson Canal. A thick, east-dipping sequence of Shawangunk Conglomerate cropping out in the south banks of the creek is visible through the trees.

OPTIONAL STOP C: Butler's Lock Quarry
This is private property and permission must be secured in advance to park in this lot. Many aspects of this stop are extremely dangerous and should only be attempted by very small groups. Pull off State Route 213 into the parking lot of the Butler's Lock Quarry and the Turco Brothers Water Service, and park out of the way. Beware that semi trucks hauling full-sized water tanks enter this lot at full speed and use portions of this space to turn around. Care must be taken when parking and walking about the lot.
Century thrust fault, road cut along State Route 213 east of Rosendale Portions of this stop are on private property and permission to visit must be obtained in advance. Trespassers will be prosecuted.
Park vehicles out of the way of traffic in the lot and walk westward along the north side of State Route 213, keeping the guardrail between you and the traffic. The large quarried face to the north is the Butler's lock Quarry and includes a continuous sequence of Rondout though Kalkberg Formations. As you proceed along the road, note how the strata in the Butler's lock Quarry are folded into a small syncline. Once at the western end of the guardrail, very carefully cross to the south side of State Route 213 and quickly climb over the guardrail. Continue westward along the south side of State Route 213, again keeping the guardrail between you and traffic, until you reach the center of the large road cut along the north side of the road (Figure 16).
The road cut exposes the Century thrust fault, which places Binnewater Sandstone over Manlius Formation. The hanging wall contains a low-angle ramp which transitions into a hanging wall flat. Similar stair-step geometry of shallow ramps and flats in the Binnewater Sandstone can also be observed in a faulted sequence southwest of Tillson. The fault zone contains what is likely a horse of crushed Binnewater Sandstone and a sliver of Manlius Limestone. The footwall is a flat in the Manlius Limestone. The footwall is also cut by several minor thrust faults, the westernmost of which appears to be a wedge fault. Thus, the hanging wall of the Century Thrust contains the Butler's Lock Quarry. The footwall contains the Van Tassel's Quarry, which we will visit in Stop 6.

Van Tassels Quarry visible through chain link fence on the right

STOP 6. The Snyder Estate
This land was originally owned by the Snyder family, who were central players in the local cement industry since its beginning during the construction of the Delaware & Hudson Canal (Werner, personal communication, 2003). The Century House Historical Society now maintains the Snyder property. This non-profit educational organization is dedicated to preserving the unique history of the Rosendale natural cement region. Leave cars and walk towards the battery of cement kilns at the east end of the lot. Continue eastward along the small trail leading down a small slope and through a small stand of trees. Proceed along the base of another battery of cement kilns to the driveway of the small, adjacent house. This lot is private property. Please obtain permission from the owners before continuing along this route. Use the flight of small stone steps directly behind the house to climb up to the abandoned tramway atop the embankment. Walk eastward along the abandoned tramway for several meters until you reach a set of small cuts. These cuts expose shallowly east-dipping beds of the uppermost Binnewater Sandstone. Note the characteristic shale horizon (previously seen at Stop 5.2) and examine the well-developed cleavage duplex that suggests some degree of bedding-parallel slip.
Continue eastward on the abandoned tramway to Van Tassels Quarry, a section of extensively mined Rondout Formation in the footwall of the Century thrust fault (Figure 11A). Van Tassels Quarry is an excellent example of the classic Rosendale natural cement quarry. Mines are cut into the Rosendale and Whiteport members, which were historically referred to as the upper and lower cements, respectively. The Glasco Member, historically known as the middle ledge, is left un-quarried between the cement layers. If time permits, follow the tramway north along the front of the quarries. The Rosendale Member is several feet thicker at the Snyder Estate than is to the north and south. Continue on the tramway to the northwest, further up onto the embankment behind the small house. Return to the vehicles in the parking lot of the Snyder Estate from the north, along the top of the embankment. Pause along the way to examine the screening house (Figure 11D) and other remnants of the cement industry, but be careful to stay away from the large openings atop the cement kilns.
As we leave the Snyder Estate, note the Brooklyn Bridge ornaments atop the gateposts. Over 100,000 barrels of Rosendale Natural Cement were used in the construction of the Brooklyn Bridge because of its unparalleled strength. Other famous landmarks built with locally produced cement include the pedestal of the Statue of Liberty and the Wings of the United States Capitol Building (Werner, personal communication, 2003).

Note Shawangunk Conglomerate in road cut
Note Shawangunk Conglomerate in road cut immediately adjacent to the driveway of the Snyder Estate along the right (north) side of road.

Bridge over Roundout Creek

OPTIONAL STOP D: Lawrenceville anticline
Pull off on small turnout on left (south) side of road. It is possible to park vehicles here for a short amount of time. For longer stays, the public parking lot located just across Rondout Creek from this location should be used.
Leave the vehicles and cross to the north side of State Route 213. Walk to the center of the bridge over the Rondout Creek and look to the west. Although it may be difficult to see during the spring and summer months, this position affords a good view of the Lawrenceville anticline. The anticline is the westernmost major, map-scale fold exposed in the Rosendale region. The Lawrenceville anticline appears in numerous photographs taken of the Delaware & Hudson Canal in this region (Figure 17).

Outcropping of Shawangunk Conglomerate in escarpment across cornfield
Note the outcropping of Shawangunk Conglomerate in escarpment across cornfield on left (east) side of road. These rocks are in the hanging wall of the Snyder thrust fault.

Ruins of mill foundation along the creek
Ruins of mill foundation along the creek behind the bed and breakfast inn is supposedly the site of one of the very first cement mills in this region (Werner, personal communication, 2003).

D&H Canal Museum
D&H Canal Museum is located about 0.1 mile south of State Route 213 on east side of Mohonk Road

Blacksmith Shop where local dolostone was first burnt for cement
Blacksmith Shop where local dolostone was first burnt for cement located a few meters north of this intersection

STOP 8. Mesoscopic folding in Silurian strata, along the Rondout Creek at High Falls
Southwest of High Falls, the Rondout Creek flows roughly parallel to strike of the Rondout Formation and Manlius Limestone. At High Falls, the creek bends sharply to the east and begins cutting perpendicular to strike, down-section towards Rosendale. The highest falls are over an outcrop of the Rondout Formation. Downstream, smaller subsequent falls separated by shallow pools occur in the Binnewater Sandstone and High Falls Shale downstream followed by minor rapids caused by debris of the Shawangunk Conglomerate. As you follow the footpath past the High Falls power station northeast, note the shallowly northwest dipping beds of Binnewater Sandstone in the cut on the right. Once past the fenced-in area surrounding the power station, leave the trail and head north across the grassy area towards the wooded area long the creek. Stop to examine the millstones cut from Shawangunk Conglomerate placed as decoration in the grass. Also note the ruins of the stone building behind the cyclone fencing. This was once the processing-plant of the E.O. Norton Cement Company, which operated mines in High Falls and along Binnewater Road near Williams (Fifth) Lake (Werner, personal communication, 2003).
When you reach the waters edge examine the asymmetric anticline in the Binnewater Sandstone and High Falls shale exposed along the north bank of the creek. This fold is the westernmost mesoscopic structure in the fold-thrust belt near Rosendale. What is unusual, however, is that the axial surface of the fold is dipping to the west, a sense of vergence opposite to the regional trend. The shape and scale of this fold suggest that it is a fault-propagation fold above a blind thrust fault. The fault underlying this structure is most likely a west dipping back thrust associated with the merger of two detachment horizons in the Ordovician strata at depth.

Small road cut in gently west dipping strata of the Kalkberg/New Scotland Fms.

Old abandoned quarry in Onondaga Limestone
Old abandoned quarry in Onondaga Limestone in trees along left (west) side of road

Esopus Creek visible through trees on left
Esopus Creek visible through trees on left (west) side of road. Onondaga Limestone underlies much of this portion of the valley

Intersection with road to Town of Hurley.
It was here that the government of New York was temporarily relocated when the British sacked Kingston in 1777

STOP 9. Cleavage duplexes in Bakoven Shale
Road cut along State Route 28, west of Kingston: Our excursion has taken us full circle, returning to the foreland of the fold-thrust belt. Here we can see evidence of deformation associated with the upper detachment of the fold-thrust belt as deformation dies out westwards. The steep escarpment and road cut along City View Terrace, beneath the Skytop Motel, contains an exposure of the two lower members of the Union Springs Formation, which is the basal unit of the Hamilton Group. Here, the upper 10 m of the fissile black shale of the Bakoven Member grade upwards into the lower 23 m of the Stony Hollow Member, a sequence of the buff-weathering shale and siltstone (Ver Straeten and Brett, 1995). The structural relationships of this location have been described by Bosworth (1984) and Nickelson (1986). Numerous examples of mesoscopic-scale deformation are visible in the Bakoven Member along the road cut. These cleavage duplexes and shear zones (Figure 14), often less than 20 cm thick, suggest that blind thrusts cut the underlying strata.