NYSGA 1998

NYSGA 1998
Field Trip Guidebook for the 67th Annual Meeting of the
New York State Geological Association
Edited by
John I. Garver
Jacqueline A. Smith
For the combined meeting of the
67th Annual meeting of the New York State Geological Association &
26th Annual meeting of Eastern Section of the American Association of Petroleum Geologists
Sponsored by
New York State Geological Survey/State Museum and Union College
Hosted by
Geology Department, Union College, Schenectady, NY
This guidebook was published by the New York State Geological Association. Additional copies may be obtained from the Executive Secretary of the NYSGA:
Bill Kelly, Executive Secretary
New York Geological Survey
Room 3140 Cultural Education Center
Albany, New York, 12230
Introduction
The Geology Department at Union College is proud to co-host the sixty-seventh annual meeting of the NYSGA and AAPG Eastern Section. Together this combined meeting has brought together academics, students, teachers, petroleum geologists, and environmental geologists. The Geology Department at Union College is happy to be a co-sponsor and co-host of the NYSGA meeting this year because it has been quite some time since we have been able to host such an event. The last time the annual NYSGA meeting was held at Union was 30 years ago when the principal movers in the Department were Philip Hewitt and Leo Hall (see Shaw, this volume). Shortly after this meeting the Geology Department was effectively disbanded and only recently came to life in 1985 after the College received a generous donation from a Geology alumnus. Now ,with four permanent faculty, one visitor, and several adjuncts, it is safe to say that we're back for good!
Since the last time NYSGA was held at Union, the Geological Sciences have experienced a major transformation. The first twenty of those years were clearly dominated by the then new plate tectonic paradigm. During the last ten years, and especially the last five years, we have seen quite a dramatic change in the focus of the geological sciences. Due in part to a number of external factors, the geologic community has placed a greater emphasis on issues associated with global change, hydrology, and environmental geology. This shift, has been away from the more traditional areas in geology such as field based studies in stratigraphy, sedimentology, paleontology, and structural geology. There is little doubt that NYSGA guidebooks reflect the focus and ideas of the times and if one compares the focus of the trips in this meeting to those in the guidebook for the previous NYSGA meeting at Union in 1965, you will see just how much our science has changed.
In assembling the trips for the 67th annual meeting of the NYSGA we tried to balance three important factors: our unique and exciting geologic setting; the breadth of interests of our members; and, importantly, the availability of field trip leaders. Certainly in terms of the first two factors, I feel that we have succeeded and we are looking forward to an exciting meeting!
The geologic setting of the lower Mohawk Valley is ideal for teaching and research. The lower Paleozoic stratigraphy has a regional tilt to the south so one can examine the Grenville-aged rocks of the Adirondacks to the north (Whitney and Olmstead, this volume; McLelland, this volume; Kelly and Hill, this volume), Cambro-Ordovician carbonates and clastic rocks in the Mohawk Valley (Garver, this volume; Kidd and others, this volume), and Silurian and Devonian rocks in the Catskill Mountains to the south (Friedman, this volume; Wolosz and Paquette, this volume; and Ver Straeten and Brett, this volume). The lower Paleozoic stratigraphy in eastern New York is punctuated by thick synorogenic sequences deposited during the Ordovician Taconic Orogeny and the Devonian Acadian Orogeny. To the east of the tilted lower Paleozoic stratigraphy, deformation in the section increases towards the front of the Taconic Allochthon where it and underlying rocks were deformed during the Taconic Orogeny (Hollocher, this volume; Kidd et al., this volume; Friedman, this volume). In the Taconic Allochthon, Cambro-Ordovician deep-water sedimentary rocks have been placed over shallow-water sediments of the Cambro-Ordovician margin.
The surficial and glacial geology in this area is as impressive and diverse as the bedrock geology. Ice dynamics related to the obstructing effects of the Adirondack Mountains resulted in non-uniform migration of ice sheets that coalesced south of the Adirondack massif only to encounter the Catskill Mountains to the south (Dineen and Hanson, this volume). During retreat, glacial Lake Albany (and its relatives) covered much of the Capital District and left an impressive sequence of glacial lacustrine deposits (Wall and LaFleur, this volume). During and after the most recent glaciation, karst and caves formed in the local carbonate strata, resulting in important landform development and interesting problems in hydrogeology (Rodbell and Hays, this volume; Rubin et al., this volume; Rubin, this volume). The long history of industry in the area and demand for space and water have resulted in some unique environment geology (Smith and Eslinger, this volume; Maswick and Snow, this volume; Hewitt, this volume). Finally, the neotectonic setting of eastern New York is examined in a unique trip aimed at understanding the seismic hazards in Columbia County (see Nottis and Cadwell, this volume).
It has become clear that the education of educators is and will continue to be an important priority for both Academe, government, and industry. With this in mind we have included several trips aimed at earth science teachers (Garver, this volume; Hollocher and Hollocher, this volume; and Kelly and Hill, this volume, and others) and we hope that many of the teachers who go on the field trips participate in the special one day workshop which will be held in Monday following the NYSGA meeting.
Finally, I would like to personally thank all of the people that have made this guidebook happen. First and foremost are the trip leaders who have spent a considerable amount of time putting together these trips and to my co-editor, Jacquie Smith, who helped with the endless task of reigning in papers and keeping track of things. Thanks are also due to the Organizing Committee and Committee Chairs for the combined NYSGA & AAPG Eastern Section meeting for their support, suggestions, and recommendations. Among them, Ken Johnson (Skidmore College), the General Chair, deserves special praise for keeping us together and focused for the committee meetings which periodically met for more than one year prior to the actual meeting. He also deserves special recognition for finally getting on EMail during that time (you can personally thank him at kjohnson@skidmore.edu). I would also like to thank Bill Kelly (N.Y. State Geological Survey), who is the N.Y.S.G.A. Executive Secretary, for his important contributions which include, but are not limited to advice and planning the logistics of putting this guide together. Finally, Gretchen Turner did a superb job handling the problems that editors face in terms of communication with authors and the nuts and bolts of putting the guide together.
To all who are just about to embark on the 67th annual NYSGA meeting -- have a great time and enjoy the fall foliage!
President, NYSGA, 1995
Geology Department
Union College
October, 1995
PLEASE READ THIS!
Landowners and outcrop defacement
In preparing this guidebook we asked the authors to do their best to ascertain landowner status at each stop. We have done this because increasing number of problems have arisen between landowners and geologists - in some cases geologists are no longer permitted to visit some localities in the Mohawk Valley. Additionally, a number of classic and important outcrops have been defaced, marked, or heaped with garbage by overzealous visitors who must leave their mark on every outcrop visited. Please pay careful attention to landowner status and cultivate a relationship with the landowner if you plan to use an exposure from year to year. Also remind your trip participants to use have respect for the future geologists who will visit these geological sites. Thank you.

A1: DUTCHESS COUNTY BARROVIAN METAMORPHISM FIELD TRIP
Thomas W. Donnelly
Dept. of Geological Sciences
State University of New York
Binghamton NY 13902-6000
INTRODUCTION
In the late 1800's, the notion of metamorphism was beginning to occupy a distinct place in geological thinking. The first studies linking metamorphic phenomena to distinctive mineralogical assemblages were of contact metamorphic aureoles around shallow igneous bodies. In the English speaking world, the notion of regional metamorphism, with "depth zones" marked by distinctive mineral assemblages is at attributed to George Barrow, who described such zones in 1893 in a Proterozoic meta-sedimentary sequence in the Scottish Highlands. We call these zones "Barrovian Zones" in his honor.
Dutchess County, New York, was mapped by Robert Balk of Mt. Holyoke College during a nine year period in the 1920's and 30's. Balk's paper (1936) is a classic study of the structural geology and stratigraphy of a highly deformed area. Balk's attractive sketch map has been included as Figure 1. His paper was accompanied by a mineralogical- petrological paper by his friend and colleague Tom Barth (1936) of the Mineralogical Institute, Oslo. Although this area has been recognized as perhaps America's pre-eminent Barrovian terrane, these authors interpret the metamorphism as due to igneous activity, including "emanations".
This area remains one of the clearest Barrovian sequences in the world. Unfortunately, possibly because of the eminence of these two papers, later geologists seem to have been somewhat intimidated, and relatively few studies have been undertaken here in subsequent years. Garlick and Epstein (1967) attempted to determine maximum temperatures of metamorphism using oxygen isotopes of metamorphic mineral assemblages. Vidale (1973, 1974) investigated the vein mineralogy and showed that vein-mineral assemblages were a good indication of metamorphic grade. Hames et al (1991) used 40/39 dating of micas to determine an age of about 445 m.y. for the Taconian metamorphism.; they also found an eastern zone overprinted with younger (390-400 m.y.; interpreted as Acadian) metamorphic modification. Two papers by Whitney et al. 1996a and 1996b) have investigated detailed mineralogy, concentrating on the garnets.

A1:Dutchess County Barrovian Metamorphism Field Trip

STOP 1. Parking lot on NE corner
The rocks examined at STOP 1 are the outcrops on Raymond St. adjacent to the parking lot. The rocks are highly deformed but not metamorphosed (no quartz veins, no indisputably metamorphic minerals). There are some traces of stratification and there are coarse blocks of more resistant sandstone.
Grand Ave and Main Street, Poughkeepsie. We will stop in the small parking lot on this corner. (The large rectangular block is a fine example of a New England two-mica granite). The rocks exposed along Grand Avenue (beware of cars!) are highly tectonized versions of the wacke seen just on the other side of the river. They have not developed a metamorphic mineralogy, and we refer to these as highly tectonized, unmetamorphosed sedimentary rocks. The fine-scale graded bedding is still visible in crumpled fragments. Original sedimentary sulfide has recrystallized as larger pyrite crystals, which weathering has created brown streaks on the outcrops. There is a slight wash of gypsum also present. Some original resistant sandy beds are visible as detached blocks within this melange. These blocks are dark brown because of the weathering of iron-rich dolomite crystals.

STOP 2. Biotite Grade
Note: Be extremely careful with traffic on this road! The outcrop is very overgrown (Also watch out for abundant poison ivy.). Here the rock is shiny and coarser grained than at STOP 1. There are no traces of stratification. There are many quartz veins.
Wingdale Road, just east of the junction with Highway 55. This stop is our first glimpse of metamorphism. The fine-grained biotite visible in thin section (and in hand specimen if you have good eyes and imagination!) gives the rock a shiny luster. The evidence for metamorphism is mainly the quartz veins, which are themselves deformed by physical deformation as they developed. The ubiquity of these veins in metamorphosed terranes results from the fact that most metamorphic reactions of sedimentary protoliths commonly produce water and SiO2. Vidale (1974) has studied these vein assemblages in Dutchess County and found that, at increasing higher metamorphic grades to the east, the vein mineralogy includes mica and feldspar, in addition to quartz. Stratigraphic layering is not visible here.

STOP 3. Garnet Grade
Outcrop (park on right side) is just before guard rail on right. Rocks examined are a few hundred feet up the road on the left side. Again be careful with traffic. There is visible garnet (mainly fairly small) and some chloritoid (black, shiny diamond-shaped crystals).
Wingdale Road, just beyond intersection to Clove and Clove Valley. Here the rock represents metamorphism of a fine sandy sediment. Both chloritoid and garnet are visible in rocks seen along the highway in outcrop. Chloritoid is an aluminous mineral which is stable apparently within a fairly narrow range of T and P, passing to staurolite at slightly higher grades. There is abundant tourmaline in thin section. Evidently this is detrital, because xenotime, apatite, and thorianite have also been found.

STOP 4. Staurolite grade
(park on left. Careful!). We will walk back up the hill to examine a good outcrop on the north side of the road. Here there is abundant garnet, common staurolite prisms, and coarse muscovite. There are several quartz veins, some of which are deformed.
Wingdale road, vicinity of Bald Mountain, beyond Pheasant Ridge. Here the protolith was more quartz rich and less aluminous. It is evidently one of the sandy Cambrian units. Garnet and Staurolite are both visible in this quartzose, micaceous schist. There are abundant quartz veins.

There is a good marble here, but we will not examine it on this trip.

Pass vivid (but altered) marble outcrops along road.

STOP 5. Kyanite grade
Continue up hill to where outcrops of schist appear on the right side of the road. The exact spot is somewhat difficult to locate, but there is a concrete drain next to the highway at the best place to see kyanite. Here there is good kyanite (some with faint blue color) and abundant, small, very pale garnets
The protolith is very aluminous. Here kyanite crystals and pale pink garnets are visible. Beware of black spots on this outcrop which are bits of tar spattered during road paving!

STOP 6. Sillimanite grade, pegmatite
Quaker Hill Road, east of Pawling. The protolith of this rock is evidently very similar to that seen at the Fifth stop. The rock is very garnet rich, with staurolite visible in thin section. There are many patches of pegmatite here, suggesting that the grade of metamorphism is approaching the beginnings of development of a hydrous silicate melt. Larger pegmatite bodies occur within a few miles of this locality. Tourmaline is sometimes visible in outcrop; it is conspicuous in thin section. Large crystals of staurolite are also visible in thin section. The sillimanite is seen only as fibrolite in mica-rich parts of the thin section.
The rocks here are coarse grained schist with abundant, coarse garnet. Staurolite is unidentifiable in outcrop but present in many sections. The sillimanite is very fine grained fibrolite. There are several pods of pegmatite, and tourmaline can be found in places.

Observe the magnificent woodwork on the barn here.

Center of the tiny hamlet of Mizzentop
Center of the tiny hamlet of Mizzentop, with its beautiful church and
handsome stone public library

Cross into Connecticut
Cross into Connecticut (road quality deteriorates noticeably! Note that you have just passed through a narrow strip of real estate that New York got from Connecticut in 1685 when Connecticut said it wanted to keep the bit around Stamford. You do remember your history, don't you?

STOP 7. Sillimanite Orthoclase grade; gneiss
Our next stop is a pull-off on the left (again!) at the telephone pole. There is no good outcrop here but there are several large fragments provided by the telephone company. The rock here is a gneiss, with good migmatitic veins.
Western Connecticut, west of Mizzen top. Here the meta-sediment has become a gneiss, which means extensive recrystallization of a highly weakened rock close to its melting point. There is abundant evidence of venitic separation of a granitic melt. In thin section sillimanite is visible as fibrolitic patches in mica-rich parts of the rock. Both plagioclase and K-feldspar reflect the high grade and partial melting of this rock. Although the rock resembles a granite, the chemical composition is clearly sedimentary.

STOP 8. Marble
Park on 292 on the left just beyond the junction. This is a fresh marble outcrop, with very good diopside prisms visible in outcrop. There is good tremolite and phlogopite also, but these cannot be seen in outcrop.
West Patterson. We have passed, without stopping, many seemingly excellent marble outcrops already. However, most of these marbles have been extensively hydrothermally altered, and the calc-silicate minerals are in a very poor state of preservation. This recently exposed outcrop at West Patterson has largely escaped this overprint, and there is excellent diopside, tremolite, and phlogopite in these rocks.

Magnificent exposures of pre-Cambrian gneisses
Just at the entrance of the rest stop there are magnificent exposures of pre-Cambrian gneisses that are startlingly similar the gneisses seen at STOP 7.

A2: CARTHAGE-COLTON ZONE TRANSECT
W.D. MacDonald, S.U.N.Y., Binghamton
This is a trip emphasizing contacts, rock fabrics, and regional structure. It examines rocks of the lowlands/highlands boundary zone spaced along Rt 812 over a distance of about 6.5 miles starting 2.5 mi. NE of Harrisville and ending 4 mi. SW of Harrisville. This trip examines structures, fabrics, and lithologies of Adirondack lowlands rocks and of the Diana complex adjacent to their mutual boundary near Harrisville, NY. The trip starts at the Route 3 bridge over the Oswegatchie River in Harrisville, NY. It proceeds E to where Rt 812 branches N from Rt 3, about 1.1 mi E of Harrisville and continues on Rt 812 to where Stone Road intersects Rt 812, about 1.1 mi N of the Rt 812/ Rt 3 junction. It follows Stone Rd to the north 0.3 mi to stop 1.

A2: Carthage-Colton Zone Transect

STOP 1. Gabbro/Marble Contact
This locality is characterized by multiple sills of gabbro, intruding marble, subsequently metamorphosed and multiply folded.
Meta-gabbro lies along the east side of the road, both north and south of this outcrop. Here at Stop 1 we find gabbro in the central upper region of the outcrop. The gabbro-marble contact rises both north and south from near road-level in the middle of the exposure. Marble foliations trend approximately N70E, 30 N at both ends of the outcrop. Near the south end of the outcrop erosional relief shows excellent internal details of the marble structure. Here there are isoclinal folds of both large (meter) and small (centimeter) scale, with mainly easterly plunges in the plane of the foliation. Marble foliations also envelope boudins of more competent marble and of calc-silicates. A more open north-plunging fold can be found in the marble near the north end of the outcrop. The gabbro-marble contact is obscured somewhat by surface obstructions and by shearing but can generally be located within a meter or so, defining a gentle basin in the N-S profile along the outcrop. The meta-gabbro has a moderately well-developed foliation. A few hundred meters south near the pipeline clearing it contains large (1 cm.) sub-equant hornblende porphyroclasts.
According to Wiener (1983) this meta-gabbro lies in a structural basin east of Geers Corners at the intersection of the NE-trending Baldface Hill syncline (third-order fold) and the NW-trending Meadow Brook syncline (fourth-order fold); he (ibid.) interprets the gabbro as having been intruded after the first regional folding phase and mainly before the second phase.

STOP 2 - Marble/Quartz-syenite Contact
This exposure clearly displays the contact between a sill of quartz-syenite and the graphitic marble. A sketch (Fig. 2) shows the view east at this locality. The marble is discordantly and isoclinally folded adjacent to the more gently folded quartz-syenitic gneiss. The intrusive contact shows several prominences and re-entrants; foliation in the marble generally follows the irregular syenite contact. Glacial grooves and striae of approx. N-S trend can be found all along the top of the outcrop, but are better preserved in the marble.
This outcrop is at the west end of a large sill of quartz syenite gneiss, one of several which crop out west of Baldface Hill (Wiener, 1983). The marble/syenite contact is quite irregular and is consistent with an intrusive boundary. The fabric of the syenitic gneiss is quite subtle relative to that of the marble. Some graphite crystallization has occurred along fractures in the gneiss. Disseminated sulfides are also found in the marble, contributing to rusty stains. Marble foliation here is defined by color banding, silicate rich bands, and erosional relief on weathered surfaces. The foliation dips about 45 northwesterly.
These sills of quartz syenite are believed to be genetically related to the Diana syenite complex (Buddington and Leonard, 1962). They were intruded during the second phase of regional folding. These sills appear to be less intensely deformed than the Diana Complex syenite, possibly because they are enveloped by less competent marbles which flowed around them during regional deformations.

STOP 3 - Marble/Diana Syenite Contact
This site reveals isoclinal meta-carbonates and related rocks now down- folded into the Diana syenite. The location is approximately in the common limb between the Harrisville syncline (to the NW) and the Harrisville anticline (to the SE) (Wiener, 1983). The structures here dip NW and are overturned to the SE.
The exposure has three main parts of equal width, from N to S: 1) a zone of rusty calcsilicate gneisses grading into white marbles, 2) diopsidic marble, and 3) Diana syenite. Faults and shear zones of WNW trend and steep NE dip separate these three units. The foliation within the diopsidic marble and Diana syenitic gneiss on either side of the 2-3 shear zone trend NE and dip steeply NW.
Zone 1) The foliation of the rusty calc-silicates is approximately conformable with that of adjacent marbles. Transition across this zone from rusty calc-silicates at the north end of the outcrop to white marble near the contact with diopsidic marble is gradual. The zone 1 - zone 2 contact between light-colored marbles and grey diopsidic marble is quite abrupt and probably is a fault or shear zone.
Zone 2) Adjacent to its north limit, the diopsidic marble shows strongly isoclinal folding (streaks of light grey bands in dark grey marble). Isolated diopside clusters and masses, some of which resemble metamorphosed thin-bedded limestone fragments, 'float' in a more massive dark grey matrix of calcite with abundantly scattered dark green granular diopside grains. Some of the light-green diopsidic masses have a darker green 'halo'. Numerous small faults of variable trends cut through the diopsidic marble; some have calcite slicken-fibers and others are plated with serpentine (antigorite?).
Zone 3) A fault or shear zone at the S limit of the diopsidic marble separates it from zone 3 which is dominantly Diana syenite. This shear zone (approx. 105, 70N) is strongly discordant to the foliation of the adjacent diopsidic marble which trends approx. 035 and dips 65NW. However, on south side of that shear zone are brown to rusty brown well foliated gneiss with foliation approx parallel to the shear zone. The contact of these light brown weathering gneisses with the darker greenish grey massive Diana syenite is transitional over a meter or so. This contact appears to be the intrusive contact between the paragneisses of the lowlands and the orthogneisses of the Diana complex. A very weak foliation in the Diana syenite (deformed by sub-parallel lenticular grains of K feldspar) here is about 035, 70 NW, consistent with that in the adjacent diopsidic marble

STOP 4 - Mylonitic Diana Syenite
Diana syenite is the dominating lithologic unit in the roadcuts on both sides of the road. The outcrop on the north side of the road is emphasized here. It is cut by numerous thin mylonite veins, the 'late ductile shear zones' of Wiener (1983). There are also a few small faults and larger shear zones, and some thin meta-dikes(?) of 10 to 20 cm thickness. In the eastern quarter of the outcrop is a tightly folded meta-dike(?). Abundant glacial grooves and striae at the top of the outcrop trend approximately along azimuth 175. The principal features to examine here are the Diana syenite, the mylonites, and the metadikes(?).
Diana Syenite
The Diana syenite here is weakly foliated, dark greenish-grey and relatively uniform in appearance. The weak foliation in the Diana syenitic gneiss here is defined by parallel lenticular grains of K feldspar up to about 1 cm long and up to a few mm thick. This foliation is most clearly seen in the weathered rinds along the drill-holes (drill-holes used for blasting the roadcut). The Diana syenitic gneiss foliations trend approximately N50E and dip somewhat variably but approx N50 to 60 NW. A little further south (e.g. stop 5) the foliation becomes more strongly developed and augeniferous. The Diana syenite was intruded early in the second phase of regional deformation and folding, and its protomylonitic foliation developed during the second phase of regional deformation (Wiener, 1983).
Mylonites
This is an excellent outcrop to examine the development of mylonitic zones in the Diana syenite. Mylonitic bands are of several different types:
a) thin, planar, of significant extent (to 5 m or more) mainly 1 to 3 cm thick
b) curved and branching, of irregular thickness and relatively limited extent
c) broad banded zones (to 20 cm) less abundant and typically of limited extent
The mylonites terminate by pinch-out, by truncating at high-angle junctions with other mylonites, or by branching from other mylonites at more acute angles. Although the mylonites ate clearly later than the foliation in the syenite, there is relatively little offset of the syenite foliation by the mylonites. Internally, the mylonites have a well-developed foliation, with abundant porphyroclasts of rounded subequant sphene and occasional elongate pyroxene. Sulfides occur along the margins of some mylonites. In the syenite adjacent to many mylonites are marginal alteration zones of variable width in which pinkish K feldspar dominates, and mafics are less abundant. Some of the alteration zones are several times as wide as the associated mylonite vein. In some alteration zones, prominent alignment of pyroxene grains or clusters can be seen, plunging moderately to steeply N to NNW.
The mylonites do not appear to have been folded. However, their distribution is symmetric with respect to the regional fold axis of the second phase of regional folding (see Fig. 4 and discussion). A small z-fold is evident about 200 m east of the west end of the outcrop. This small isoclinal fold pair (anticline/syncline) occurs in the fabric of the mylonite; the mylonite zone here is about 10 to 15 cm wide and its boundaries have not been folded. The sense of shear indicates W side up. Such folds in the mylonite are scarce.
Meta-dikes(?)
Several meta-dikes(?) now largely altered to pyroxene cut the syenite. They are mostly relatively thin, 10 to 30 cm thick. A tight fold has developed in one of these, in the eastern third of the outcrop. This kind of fold is not common. The theoretical axial surface plane of this fold trends about N10W and dips 75 W, and lies at a high angle to the foliation of the Diana syenite. The theoretical hinge (found by interpolating the intersection of the limbs) plunges steeply N. As a mylonite adjacent to this fold is not itself folded, it is thought that the mylonites post-date deformation of these 'meta-dikes'.
Discussion
Some structural relationships at this outcrop are summarized in Fig. 4. Here we see that the foliation ('protomylonitic' foliation) of the Diana syenite dips moderately NW. The mylonite veins appear to have a common axis of orientation which plunges 62 toward 347. According to Wiener (1983), this is the common direction of lineations for the second phase of regional deformation. That pole and the associated 95% confidence oval (Bingham statistics) represent the direction of the eigenvector corresponding to the minimum eigenvalue for the distribution of about 30 axes or normals to the mylonite veins. These are for measurements in the western third of this outcrop only. Both pyroxene lineations and the theoretical axis of the tightly folded 'meta-dike' also lie close to this direction. This suggests that although the mylonites are of greatly variable orientations, they are mainly coaxial about an axis near the foliation plane of the Diana syenite and plunging NNW.The pyroxene lineations adjacent to mylonites are therefore near the intersections of the mylonite foliations with the Diana syenite foliation, as might be expected.

STOP 5 - Augeniferous Diana Syenite
Augeniferous well-foliated syenite gneiss is exposed in low outcrops on both sides of the road. A few thin mylonites are visible here. The foliation, wrapping around centimeter-sized feldspar augen, trends approximately N45E and dips 45NW. Dark-cored feldspar augen (greenish) with lighter tan rims and curved cleavage are abundant. Foliation in the Diana syenite is best seen in the drill-holes used for blasting, where the syenite weathers to a light brown color. The rock is quite uniform here, and there is no indication of interlayering with other lithologies.
Southward from this outcrop the foliation becomes better developed approaching flaser gneiss, and layers of other rocks, such as shonkinites, develop. Still further south, hornblende granite appears. This apparent regional zonation led Buddington and Leonard (1962) to conclude that the Diana complex is essentially a large isoclinally folded and metamorphosed differentiated sill complex, overturned to the SE such that higher levels of the sill lie to the SE.

A3: FINGER LAKES GORGES REVISITED
Peter L.K. Knuepfer
Tim K. Lowenstein
Department of Geological Sciences and Environmental Studies
Binghamton University
Binghamton, NY 13902-6000
email: knuepfr@binghamton.edu
email: lowenst@binghamton.edu
Introduction
The Finger Lakes region of central New York is justifiably world-renowned for its beauty, especially the spectacular deep gorges -the Finger Lakes glens- that characterize tributary streams draining into the central troughs. The origin of the troughs themselves has been the source of debate for as long as geologists have ventured into upstate New York (as summarized by Mullins and others, 1989). The recent geophysical studies by Mullins and his colleagues (Mullins and Eyles, 1996) have resolved much of this controversy, demonstrating that the modern troughs, at least, must have been formed by a combination of glacial scouring and the action of high-pressure sub-glacial meltwater. Seismic stratigraphic relationships suggest that the stratigraphic and morphologic record left behind by this event-including the Valley Heads Moraines (VHM) as well as much of the infill of the Finger Lakes troughs themselves--must have occurred between (and including) VHM time, around 14.9-14.4 ka 14C yr. B.P. (Muller and Calkin, 1993). Late-stage trough-fill sediments were deposited until 13.6 ka 14Cyr. B.P. (Wellner and others, 1996), by which time ice had retreated north of the modern Finger Lakes. Our goals on this field trip (Figure 1) are to consider the implications of this record for development of the modern Finger Lakes glens/gorges.

A3: Finger Lakes Gorges Revisited

View area on left overlooking Cayuga Lake trough.
Recent construction has impaired the view of geologic features.

STOP 1. Overview.
The Ithaca College campus is built on hanging deltas of Six Mile Creek and commands a spectacular view of the Cayuga Lake basin. Features that can be seen on a clear day include the oversteepened margins of Cayuga Lake trough to the north and hanging deltas of Coy Glen and Enfield Glen to the west and southwest. Similar hanging deltas characterize most of the streams that drain into the Finger Lakes troughs and the filled troughs to the south of the lakes. This stop affords us the opportunity to set the stage for features to be observed and discussed during the remainder of the trip.

Note increasing, incision
Note increasing, incision of Taughannock Creek as we drive downstream. Note also our position on a relatively flat surface, one of the highest hanging deltas of Taughannock Creek.

STOP 2A. Upper Taughannock Falls.
We walk onto the old rail bridge above (upstream of) Taughannock Falls.
Although incision of Taughannock Creek has been increasing as we drove downstream from Hwy. 96, here is the first significant knickpoint in the channel- a drop of some 15 m.
The view downstream is instructive: the lake-ward terminus of the high delta on which we have parked is clearly visible about 0.8 km (0.5 mi) downstream. Some of the incised meanders of the gorge upstream of the main Taughannock Falls also are visible.

STOP 2B. Taughannock Falls Overlook.
The view area affords an excellent perspective of the main Taughannock Falls. Here the creek drops some 65 m through late Devonian Genesee Group sediments. The top of the falls is formed in resistant siltstone of the Sherburne member (Grasso and others, 1986); the notch at the lip of the falls has remained little changed since a rockfall in the late 1880s or early 1890s. One question to consider is how much the waterfall has retreated upstream. A possible clue is afforded by the retreat of the amphitheater in which the modern waterfall is located. Another possible clue is the position of the delta front, downstream of this position, that corresponds to the elevation of the crest of the waterfall. Bloom (1972) has suggested that the broad valley downstream of the waterfall is indicative of re-excavation of an interglacial valley. We are unaware of any exposed pre-latest Pleistocene unconsolidated sediments anywhere within the lower gorge. A related issue is how much of the incision we see here has occurred in post-glacial time, and how much may have occurred while the ice was retreating northward through the Cayuga Lake trough (and other Finger Lakes troughs) and providing the dams for progressively lowering lakes (and accompanying deltas).

STOP 2C. Gravel Quarry.
Cross Taughannock Park Rd. from the Falls Overlook parking area and walk onto the unpaved road on the left side of the small parking area. Walk about 100 m into an old quarry. This quarry is cut into the front edge of the most prominent high hanging delta of Taughannock Creek. Recent working of the northeast wall of the quarry has resulted in a spectacular exposure of the delta-front foreset beds. Sediments include well stratified sands and gravels with occasional clayey and silty interbeds.

Note the small delta surfaces across which we drive.
The youngest delta, of course, is the active delta complex on which much of the developed area of Taughannock Falls State Park is located. This is one of the largest deltas on Cayuga Lake. If we use the analog from Canandaigua Lake and the evidence that Mullins and others have provided, we can assert that this delta probably began to form in latest glacial time, after base level for Taughannock Creek had dropped to approximately its modern position. Again, we can consider how much of the incision of bedrock occurred during this drop in lake levels, and how much is post-glacial.

STOP 3. (LUNCH) Buttermilk Falls and Buttermilk Glen.
Time permitting, we'll walk up the lower park of Buttermilk Creek. Much of the incision is represented by the two main waterfall in the lower creek; this contrasts sharply with the situation at Taughannock Creek. Delta surfaces are not as well preserved here, but again hanging deltas do occur. Why is there such a difference in incision pattern? Was there no pre-last glacial channel to re-excavate here? Joint systems playa minor role in erosion by Buttermilk Creek, but spectacular potholes are excavated into the rocks at the gorge bottom. We may meet vans at the parking lot above the gorge, or we may continue from the lower entrance to the State Park. The road log assumes the latter.

Note an entrance to an abandoned gravel quarry
Note an entrance to an abandoned gravel quarry on the right 0.2 miles before the entrance to the park. This quarry, formerly operated by the Town of Enfield, exposed foreset beds of the uppermost Enfield Glen hanging delta. These beds were overlain by a poorly consolidated, thin (less than 2 m exposed) till. We interpret the deltaic sediments as marking the input of Enfield Creek sediments into the highest level of the pro-glacial Cayuga Trough lake, with a nearby ice margin overriding the sediments during a brief interval of readvance.

STOP 4. Enfield Glen.
No tour of the Finger Lakes gorges around Ithaca would be complete without a stop to tour the upper part of Enfield Glen. The parking area and picnic ground are in a relatively broad valley containing an alluvial floodplain over bedrock. Note that bedrock is exposed in the valley walls and in tributary channels such as Fish Kill just upstream from the mill. Poorly exposed colluvium mantles other slopes of this valley wall. As we walk into Enfield Glen, note the abrupt change to a narrow gorge deeply incised into well jointed bedrock. Here jointing has played a key role in facilitating excavation of bedrock by the stream. We walk down to Lucifer Falls, below which the valley again opens into a much wider, bedrock-walled gorge. The traditional interpretation is that the wide gorge sections indicate reoccupation of a buried interglacial gorge, whereas the bedrock gorge represents post-glacial incision into the Devonian bedrock (e.g. von Engeln, 1961). The reason for post-glacial diversion through a bedrock gorge is assumed to be a mass of Wisconsin glacial sediment choking the interglacial valley, such that Enfield Creek was diverted and flowed across (and incised into) bedrock. Again, to what extent has the incision occurred in post-glacial times (which is what the "conventional" interpretation would imply) and how much occurred during ice retreat up Cayuga Lake trough? If the stream achieved a base level approximately at the modern Cayuga Inlet altitude, then virtually all of the base level fall that has driven incision occurred during late glacial time. How much knickpoint retreat has occurred? How can we assess this?

STOP 5. Landstrom Gravel Pit.
Get permission at the office before entering.
This pit is excavated into the lowermost delta of Enfield Creek and represents the last depositional event before Enfield Creek had stabilized at more or less its current base level. Note the similarity of overall stratigraphy to what we viewed at Stop 2C. Here we have a higher clay content and the development of clay balls and climbing ripples (which may or may not be visible on the day of the field trip).

We soon climb into the Valley Heads Moraine
We soon climb into the Valley Heads Moraine complex of West Danby, NY. This is one of the longest and southernmost valley-filling reaches of Valley Heads Moraines in central New York.

STOP 6. Gravel Pit.
Walk past the gate and down the hill, noticing poorly exposed well rounded gravels on right. Note hummocky morainal topography visible ahead on to the left. Turn right at road at bottom of hill and enter main gravel workings. Here the Valley Heads Moraine is comprised of well rounded, sorted, and weakly stratified sands and gravels. This is typical of VHM deposits in those exposures we have visited in central New York, as well as in exposures discussed by Yang (1992). Clearly water played an active role in abrasion and deposition in this glacial-marginal environment! We'll discuss the implications of this for development of the VHM and its relationship to the Cayuga Lake trough, the hanging deltas, and the exposures we've already seen.

Note asymmetric valley walls.
As is typical of valley walls in east-west valleys in the glaciated Appalachian Plateau, the south-facing walls have relatively gentle slopes and the north-facing walls are steep and truncated. Coates (1966) interprets these as till shadows: south-facing walls have thick lodgement till deposits resulting from loss of basal stress as the glacier descended into the valley, whereas the south wall has thinner, compressed till due to high pressure as the south-traveling basal ice "piled up" against the valley.

STOP 7. Robinson's Gravel Pit.
Get permission at office before entering pit.
As a final stop, we look at deltaic, deposits not associated with the pro-glacial lakes of the Cayuga trough. Here we are downstream of the Valley Heads Moraine of Willseyville Creek. The landform is a kame surface, and we interpret the sediments as primarily deltaic foreset beds. However, they dip back toward the valley wall rather than toward the valley center, suggesting a complex mode of origin. Local slump features and contorted bedding attest to the kame origin of this deposit.

A4: DEVONIAN FLUVIAL TO SHALLOW MARINE STRATA, SCHOHARIE VALLEY, NEW YORK
John Bridge and Scott Jarvis
Department of Geological Sciences, Binghamton University,
Binghamton, New York 13902-6000
Introduction
The Middle to Upper Devonian Catskill clastic wedge in New York State is part of a foreland-basin fill that has relatively thick non-marine deposits to the east and thinner marine deposits to the west. The overall succession of strata is indicative of marine regression: however, there is ample evidence in coastal and fluvial strata of periodic marine transgressions. There are several superimposed scales of cyclicity in these strata, at least three of which can be readily distinguished:
(1) 10- to 100-meter thick, asymmetrical coarsening-upward to fining-upward sequences, representing on the order of hundreds of thousands to millions of years, that have been interpreted as recording relative sea-level changes associated with eustasy, tectonically induced changes in sediment supply and subsidence rate, and possibly climate change (e.g., Dennison and Head, 1975; Dennison, 1985; House, 1983, 1985; Johnson et aI., 1985; Ettensohn, 1985; Brett and Baird, 1986; Van Tassell, 1987; Willis and Bridge, 1988; Dennison and Ettensohn, 1994).
(2) Meter- to 10-meter thick fining-upward or coarsening-upward sequences of sandstone and shale (representing thousands to hundreds of thousands of years) that have been interpreted to have been formed by lateral migration, abandonment and :filling of channels; progradation and abandonment of channel-mouth bars, crevasse splays and tidal flats; and :filling of coastal bays (e.g., Johnson and Friedman, 1969; McCave, 1968, 1969, 1973; Miller and Woodrow, 1991; Bridge and Willis, 1994). Although these sequences could have been caused by autocyclic processes such as channel switching (avulsion), allocyclic causes such as eustatic sea-level changes must also be considered (Van Tassell, 1987; Bridge and Willis, 1994).
(3) Centimeter- to decimeter-thick, sharp-based sandstone stratasets capped by shale that are commonly interpreted as the deposits of individual floods in rivers and floodplains, or storm events at sea (e.g., Gordon and Bridge, 1987; Craft and Bridge, 1987; Halperin and Bridge, 1988; Willis and Bridge, 1988; Bridge and Willis, 1994; and many others).
The validity of these interpretations of the different scales of facies sequence hinges on describing them in detail, and being able to correlate them between different outcrops. Physical tracing of strata between outcrops is made difficult by the wide spacing of outcrops (and cores), variable dip of strata, lateral changes of facies, lack of distinct marker horizons, and the absence of seismic profiles. Biostratigraphic correlation is very crude because biozones extend for on the order of a million years (100 m of strata) and only recently has it been possible to correlate marine and non-marine rocks using miospores. Stratigraphic correlation of 10 meter thick to 100 meter thick sequences in marine strata is accomplished using laterally extensive black shales and limestones (Sutton et al., 1962; Sutton, 1963; Woodrow and Nugent, 1963; McCave, 1969, 1973; Rickard, 1975, 1989; Brett and Baird, 1985, 1986, 1990). However, their correlation with coastal and non-marine rocks remains uncertain (Halperin and Bridge, 1988; Bridge and Willis, 1994). In coastal and fluvial strata, it is difficult to correlate 10 meter thick and thinner stratasets for more than on the order of kilometers. This inability to correlate any but the thickest sequences across the basin makes it difficult to interpret these deposits in a sequence stratigraphic context, and difficult to assess whether the controls on their formation are local, regional or global.
Examples of the different scales of sequence in fluvial, coastal and shallow-marine deposits will be examined in the Schoharie Valley (in the vicinity of Gilboa, Grand Gorge and Prattsville). Bridge and Willis (1994) studied selected outcrops and cores through the interleaved fluvial and marine deposits of the Schoharie Valley (Figure 1), in order to expand upon previous descriptions and interpretations of these strata (e.g., Johnson and Friedman, 1969; McCave, 1968, 1969, 1973; Miller and Woodrow, 1991). Attention was focussed on the influence of wave, tidal and fluvial currents in shaping the coastal region, and on how these currents and coastal morphology changed during marine transgressions and regressions. Basin-scale controls on deposition were evaluated by considering coeval fluvial deposits to the east (Willis and Bridge, 1988) and offshore marine deposits to the west (Brett and Baird, 1985, 1986, 1990).
Ten-meter thick sequences were correlated between cores and outcrops over distances of up to 10 km (Figure 2; Bridge and Willis, 1994). Stratigraphic subdivision of these rocks (Figure 3; modified from Rickard, 1975, 1989) is due mainly to the work of Cooper (1930, 1933, 1934) and Cooper and Williams (1935). Some formations were defined based on fossils in fully marine deposits, but key zone fossils are absent in coastal and fluvial deposits. Nevertheless, the formations defined by Cooper can generally be recognized in Schoharie Valley. Outcrops to be visited span the upper Moscow Formation (Manorkill Falls), the Gilboa Formation (Route 30 near Grand Gorge, Hardenburgh Falls; Stevens Mountain Quarry), and the lower Oneonta Formation (Route 30 near Grand Gorge). These Formations are mid to late Givetian in age. This stratigraphic interval includes a major transgression and regression, upon which are superimposed smaller scale changes in relative sea level.

A5: Middle Devonian Near-shore Deposits

STOP 1. Hardedburgh Falls
Description
The Hardenburgh Falls outcrop (Figure 4) is assigned to the middle-upper part of the Gilboa Formation, and is equivalent to the middle part of the Stevens Mountain Quarry section (Figure 2). The lowest 6 meters of this exposure (Figure 4) comprise gray sandstones, dark gray siltstones and shales. The mudstones are sparsely fossiliferous and typically bioturbated. The fine to very-fine grained sandstones are sharp-based, cm- to dm-thick sheets and lenses. Thinner ones contain wave-ripple cross strata and are capped by wave ripple marks (2-D, 3-D, and interfering types). Thicker, coarser grained sandstones have hummocky-swaley cross strata and/or planar strata in their lower parts, and abundant intraformational shale fragments and disarticulated shelly fossils. Near the base of the section (e.g. Figure 4, meter 2) relatively thick sandstone stratasets have unidirectionally dipping low-angle to high-angle, medium-scale cross strata with easterly paleocurrents, and extraformational pebbles. At the base of the section, amalgamated sandstone stratasets fill N-S oriented channels that are up to decimeters deep and meters wide. These relatively coarse grained strata contain rounded extraformational chert and quartzite pebbles.
Shell accumulations include the brachiopods Spinocyrtia, Mucrospirifer, Mediospirifer, Orthospirifer, Cupularostrum the bivalves Goniophora, Actinopteria, Palaeoneilo and crinoid ossicles. Most of the disarticulated shell fragments lie concave down along stratal surfaces and show little evidence of breakage or abrasion of fine detail. Some shell concentrations are markedly mono specific (e.g Cupularostrum). Small, centimeter-diameter vertical burrows occur in the tops of sandstone strata.
The 5.5 meter thick, medium- to fine-grained sandstone body in the middle of this section has a sharp base that is overlain by planar strata and swaley cross strata with horizons of wave ripples draped by shale. Asymmetrical wave ripples indicate a NE-SW oscillation with a preferred migration to the SW. The overlying medium-scale trough cross strata dip to the NE mainly (with a rare, oppositely directed set) and contain superimposed wave- and current-ripple marks. Medium-scale planar cross strata with a SE dip, occurring above the trough cross strata, climb up low-angle surfaces that dip to the NW. Vertical, meniscate burrows of varying diameter occur throughout the sandstone body. Several stratasets of shell-rich sandstone with hummocky cross strata and wave ripples occur immediately above the sandstone body.
Interpretation
Interbedded mudstones and sandstones with hummocky cross strata, planar strata, wave-ripple marks and associated cross strata, and abundant shelly-fossil concentrations are interpreted as nearshore marine deposits formed below fair-weather wave base (e.g. Dott and Bourgeois, 1982; Hunter and Clifton 1982; Swift et al., 1983). Sandstones were formed during storms by either dominantly wave currents or by combined wave and unidirectional currents (Craft and Bridge, 1987; Southard et al., 1990, Duke et al., 1991), whereas mudstones represent fair-weather deposits or those accumulated below storm-wave base. Amalgamated sandstone stratasets indicate deposition mainly above fair-weather wave base. The preferred easterly paleocurrent orientation of some low- and high-angle cross strata, and the association of increasing grain size with increasing inclination of these cross strata, suggest onshore directed currents (asymmetrical wave, and/or flood tidal) that in places were strong enough to change hummocky bedforms to dunes (e.g. Nottvedt and Kreisa, 1987; Amott and Southard, 1990; Myrow and Southard, 1991; Cheel and Leckie, 1992). The shallow channels may be cut by either storm-induced or tidal currents (e.g. Cacchione et al., 1984; Craft and Bridge, 1987).
The sandstone body in the middle of this exposure is interpreted as the deposits of a tidal channel mouth bar that prograded into a storm-wave dominated sea. The vertical variation in sedimentary structures indicates wave currents or combined wave and unidirectional currents lower down, but dominantly unidirectional currents higher up. Paleocurrents associated with wave ripples are generally alongshore (NE-SW). The alongshore to landward-directed trough cross strata (NE-E) are probably associated with sinuous-crested dunes formed by tidal flood currents, and ebb tidal currents were relatively unimportant. Planar cross strata migrating to the SE may represent wave-fanned swash bars or tide-formed straight-crested dunes (e.g. Hayes and Kana, 1976; Boothroyd, 1978; Fitzgerald, 1984; Sha and De Boer, 1991). Superimposed wave and current-ripples indicate periods of reduced wave- and tidal-current strength. The sparse fauna and low degree of bioturbation, but dominance of the Skolithos ichnofacies, is typical of sandy deposits in coastal areas (Howard and Reineck, 1981).
The relatively sharp base and top of this sandstone body could be interpreted as due to diversion of a channel into this area followed by abandonment, reduction in sand supply, and marine transgression. However, an alternative explanation could be one of progradation of a sandy shoreface during a relative sea-level fall (a forced regression), followed by modest erosion of the top of the sandstone body during marine transgression. These alternatives are discussed below.

STOP 2. Stevens Mountain Quarry
Description
Stevens Mountain Quarry exposes a largee proportion of the Gilboa Formation (Figure 2). Attention will be directed to the 2 to 6 meter thick sandstone bodies that occur in the upper 30 meters of the quarry (Figure 5). These fine-to medium-grained sandstone bodies have sharp, erosional bases with tool and flute marks, overlain by intraformational breccia comprising shale clasts, plant fragments and shelly fossils (bivalves, brachiopods, Tentaculites and crinoid fragments). Relatively thin (2 to 3 meters) sandstone bodies are composed of amalgamated, dm- to m-thick stratasets containing swaley cross strata and planar strata. The erosional bases of individual stratasets are commonly overlain by transported shell fragments. These sandstone bodies vary little in grain size throughout their thickness: however, they tend to be capped with concentrations of shells, intraformational shale fragments and rounded extraformational pebbles (e.g. Figure SA, 15-17m, Figure 5B, 7-9m and 12-15m). As the thickness of these sandstone bodies increases, stratasets of medium-scale cross-stratified sandstone begin to dominate their upper parts (e.g. Figure 5B, 19-23m). This kind of sandstone body may fine upwards, coarsen upwards or show little vertical variation in grain size. Like the thinner examples, they are capped by concentrations of shells, intraformational breccia and extraformational pebbles. The thickest sandstone bodies are medium grained and composed almost entirely of medium-scale cross strata, with lenses of planar strata. Erosional reactivation surfaces in cross sets are common. In places, asymmetrical ripple marks on reactivation surfaces and cross stratal surfaces indicate a paleocurrent direction opposite to the medium-scale cross strata (e.g. Figure 5A, 33-39m, Figure 5B, 29-34m). Stratasets in these sandstone bodies may be inclined at several degrees relative to the basal erosion surface (i.e. large-scale inclined strata) and sandstone-filled channels may occur in their upper parts. Immediately above some sandstone bodies, sandstone stratasets may have unidirectionally dipping low-angle and angle-of-repose cross strata, abundant shell accumulations and intraformational and extraformational pebbles. Such stratasets may also be markedly lenticular and fill channels up to dm deep and meters wide (as at the base of Harden burgh Falls). Laterally extensive horizons of sandstone load casts occur immediately beneath some sandstone bodies.
Paleocurrent directions from medium-scale cross strata in and above the upper parts of all sandstone bodies are generally in an easterly direction (range from NNE to SE). Medium-scale cross strata in the lower parts of the thickest sandstone bodies are generally oriented approximately 180 degrees to those higher up. Tool marks at the base of sandstone bodies give paleocurrents to the Wand NW or to the E and SE.
Skolithos burrows (mm to cm diameter) are common in tops of sandstone stratasets. Arenicolites and Chondrites also occur. Large (up to O.1m diameter) vertical, sand-filled meniscate burrows are characteristic of thicker sandstone bodies. Some sandstones are so bioturbated that primary sedimentary structures are difficult to discern (e.g. base of quarry). Fossils in the shell accumulations include brachiopods (Spinocyrtia, Mucrospirifer, Tropidoleptus carinatus, Cupularostrum (Camarotoechia), Mediospirifer, Orthospirifer mesastrialis), bivalves (Goniophora, Grammysia, Actinopteria, Cypricardella, Palaeoneilo), Tentaculites and crinoid ossicles (Cooper and Williams, 1935).
The thickness and sedimentary characteristics of sandstone bodies and mudstone intervals change laterally over hundreds of meters (compare sandstone bodies e and b in Figures 20A and B). Also, the thickest sandstone bodies dominated by angle-of-repose cross strata occur at the top of the quarry, whereas relatively thinner ones with mainly swaley cross strata and abundant shelly fossil accumulations occur lower down.
Interpretation
The thickest cross-stratified sandstone bodies with erosional bases overlain by intraformational breccia with shelly fossils, large-scale inclined strata and channel-fills, and opposing paleocurrent directions, are interpreted as deposits of laterally migrating tidal channels (see also Johnson and Friedman, 1969). Dominance of ebb-directed paleocurrents lower in sandstone bodies and flood dominance in the upper parts (in places associated with channel fills) is typical of strongly asymmetrical tidal currents in tidal inlets, estuaries, and tide-influenced deltaic distributaries. Variable dip angles of medium-scale cross-strata, erosional reactivation surfaces and superimposed current ripples are further evidence of strong tidal-current asymmetry. They indicate growth of sinuous-crested dunes to 'full vortex' stage during the dominant tidal current, slackening of the current and modification of dune geometry, then erosion of the dune by the subordinate tide. Such features are commonly reported from mesotidal and macrotidal settings (e.g. Terwindt, 1981; Boersma and Terwindt, 1981; Dalrymple, 1984; DeMowbray and Visser, 1984). General absence of tidal-bundle sequences and rarity of current ripples on reactivation surfaces are due to an erosional or non-depositional subordinate flood-tidal current and a dominant ebb-tidal current that may have been reinforced by a fluvial current.
Sandstone bodies in which lower parts have swaley cross strata, planar strata and wave ripples but upper parts have angle-of-repose cross strata, suggest wave currents or combined wave and unidirectional currents lower down but dominantly unidirectional currents higher up. Interpretation of the depositional environment of these sandstone bodies hinges critically on their lateral transition to thicker sandstone bodies dominated by angle-of-repose cross strata and thinner sandstone bodies dominated by amalgamated swaley cross strata. In some locations, sandstone bodies have sharp erosional bases, but in others they occur at the top of coarsening-upward sequences with interbedded hummocky cross-stratified sandstones and shales immediately beneath. These features suggest deposition on channel mouth bars that were prograding into a marine, storm-wave dominated area. Such bars may have been on the seaward side of tidal inlets (i.e. ebb-tidal deltas) associated with estuaries or barrier-beach shorelines, or associated with tide-influenced deltaic distributaries. Landward-directed (NE to SE) trough cross-strata are most likely associated with sinuous-crested dunes formed by tidal flood currents, but also perhaps with asymmetrical shoaling wave currents.
The relatively thin sandstone bodies dominated by amalgamated swaley cross- stratified strata are interpreted as more distal parts of channel mouth bars, that were completely dominated by storm waves. The lack of mud suggests deposition above fair weather wave base. The common occurrence of load casts near the base of these sandstone bodies suggests rapid deposition of sand on mud. Such soft-sediment deformation features are common offshore from Mississippi delta distributaries (e.g. Coleman and Prior, 1982). Paleocurrent indicators in the bases of sandstone bodies suggest a dominantly offshore directed unidirectional current, as is common in similar Frasnian rocks in the Catskill Region (Craft and Bridge, 1987; Halperin and Bridge, 1988; Bishuk et al., 1991) and in many other ancient storm-dominated shelf deposits (Leckie and Krystinik, 1989). Offshore-directed currents may be associated with coastal set-up during storms (e.g. Craft and Bridge, 1987; Halperin and Bridge, 1988; Leckie and Krystinik, 1989; Duke, 1990; Duke et al., 1991). However, fluvial and tidal currents may also have had an influence. Lenticular, pebbly and shelly sandstone strata, locally in channels, at the top of these sandstone bodies are evidence for relatively strong shoreward directed currents, possibly due to tidal flood, "storm-tides", or asymmetrical shoaling-wave currents (e.g. Cheel and Leckie, 1992). Thus, these coarse, shelly strata on top of sandstone bodies are associated with rapid abandonment of the channel-mouth bar, reduction in sand supply, and marine transgression (see, for example, Boyd and Penland, 1988; Boyd et al., 1989; Sha and DeBoer, 1991; Penland et al., 1988).
Bioturbation is characteristically rare in coastal sand bodies such as these (Howard and others, 1975; Howard and Reineck, 1981). Large, vertical meniscate burrows are similar to those reported in fluvial and coastal channel deposits elsewhere (Thoms and Berg, 1985; Miller, 1979; Bridge, Gordon and Titus, 1986), and may be due to upward escape of bivalves. The brachiopod-bivalve fauna is typical of Devonian nearshore communities (McGhee and Sutton, 1985; Sutton and McGhee, 1985).
The vertical variations in sandstone body characteristics within the upper 30 meters of Stevens Mountain quarry suggest a decreasing marine influence, with progressively more proximal channel bar deposits higher up. This trend is continued upwards into the non-marine Oneonta Formation.

STOP 3. Manorkill Falls
Description
The Manorkill Falls section (Figure 6) exposes the upper part of the Moscow Formation where a pronounced grain-size fining occurs (Figure 2). When the water level in Schoharie Reservoir is unusually low, up to 10 meters of sandstone is exposed (below the log in Figure 6). Many sandstone casts of tree trunks (part of the famous Gilboa fossil forest) are seated in the gray-green disrupted mudstones immediately above these sandstones. The casts stand upright within the lowest, decimeter-thick sandstone strataset of a 2-m thick sandstone body (Figure 6, 3.5-5.5 m). Within this upward-fining sandstone body, sheet-like stratasets contain mainly medium-scale cross strata with easterly paleocurrents. Reactivation surfaces with ripple marks are superimposed on these cross strata. Sandstone stratasets between meters 6 and 7 have low-angle (swaley?) cross strata, wave-rippled upper surfaces, and rare shelly fossils (bivalves, brachiopods, crinoid ossicles, bryozoa). The distinctive wedge-shaped sandstone body between meters 7 and 8 has low-angle bedding surfaces that dip in the direction the sandstone body thins. Upper surfaces of these large-scale inclined strata are covered with symmetrical and interfering wave-ripple marks. Ichnofossils are abundant and diverse throughout these sandstone bodies, and include Skolithos, Meunsteria, Arenicolites and Diplocraterion.
The overlying mudstone-dominated sequence (meters 8-24, Figure 6) contains relatively undisrupted and fissile mudstones that grade up into intensely disrupted mudstones with desiccation cracks, root casts, burrows, and calcareous rhizoconcretions. Interbedded sandstone stratasets are mainly wave-ripple cross stratified, but may have planar or hummocky-swaley strata at the base. Between meters 18 and 21 (Figure 6) such stratasets occur in E-W oriented, fine-grained channel fills. Some upward-fining sandstone stratasets have medium-scale and small-scale stratasets with easterly paleocurrents (e.g., 16-18 m; Figure 6) Rare transported shelly fossils and tree trunk casts occur again near the top of the section. At the top of the measured section, a 6 meter thick sandstone body has stratasets inclined at up to 5 degrees relative to its basal erosion surface (i.e. large-scale inclined strata). Stratasets lower down in the sandstone body have planar and swaley stratification with superimposed ripple marks. Higher up, stratasets of medium-scale cross strata fill channel forms.
Interpretation
The mudstones at the base of the measured section record flood basins with soils that supported a forest of trees. The sandstones that overlie these deposits, including those that buried and preserved the bases of tree trunks, show evidence for landward progradation of wedges of sand into standing water. This and the fossil content indicate a back-shoal washover origin. The overlying fissile mudstones with wave-rippled sandstone indicate deposition in quiet standing water that was periodically agitated by waves, possibly an interdistributary bay or back-barrier lagoon. A return to subaerial conditions is signaled by the occurrence of desiccation cracks, root casts, and calcareous paleosol nodules. Overlying mudstones and sandstones with evidence of landward-directed unidirectional paleocurrents, wave currents, periodic exposure and fine-grained channel fills may have been deposited on sandy to muddy tidal flats that changed upwards to subaerial plains. The landward-directed paleocurrents may be associated with tidal floods and/or storm overwash. The sandstone body at the top of this section is interpreted as a tidal-channel mouth-bar deposit.

STOP 4. Gilboa
At the side of the road at the northern tip of Schoharie Reservoir, immediately to the west of Schoharie Creek, there is an exhibit of sandstone casts of tree trunk casts that were recovered during excavations associated with the construction of the dam. There are a number of places in this area where such tree trunk casts can be seen in situ. One of these is Manorkill Falls.

STOP 5. Route 30 near Grand Gorge
Description of lower part
This road cut just north of Grand Gorge (Figure 1) exposes the upper Gilboa and lower Oneonta Formations, the overall sequence recording marine regression. The lowest 28 m of the section (Figure 7) comprises the Gilboa Formation. This Formation is composed of 2 to 6 m thick sandstone bodies separated by mudstone-dominated intervals.
The sandstone bodies have sharp bases with erosional sole marks, overlain in places by intraformational breccia composed of shale clasts, plant fragments, and shelly fossils (bivalves, brachiopods, crinoid fragments, and Tentaculites). The thinner sandstone bodies (2-3 m) are composed of amalgamated, decimeter to meter thick stratasets of swaley cross strata and planar strata. Erosional bases of individual stratasets are commonly overlain by transported shell fragments, and wave-ripples are rarely preserved beneath the thin shales draping the stratasets. Grain size of these sandstone bodies varies little through their thickness, except at the top where concentrations of shells, intraformational and extraformational pebbles occur.
As the thickness of these sandstone bodies increases, medium-scale cross strata begin to dominate their upper parts (e.g. meters 14-28). The thickest sandstone bodies are almost entirely composed of medium-scale cross strata. Medium-scale cross strata show opposing paleocurrent directions, especially in the upper parts of sandstone bodies. Symmetrical and asymmetrical ripple marks (draped with shale in places) superimposed on medium-scale cross strata also indicate divergent paleocurrents. These sandstone bodies may fine upwards, coarsen upwards, or show little vertical variation in mean grain size. They may also be capped with pebbly and shelly beds.
The dark gray, sparsely fossiliferous shales are typically bioturbated and interbedded with gray, centimeter to decimeter thick sandstone sheets and lenses. The thinner sandstones have wave-ripples and associated cross strata. Thicker, coarser grained ones have hummocky-swaley cross strata and/or planar strata in their lower parts. Sandstone load casts occur immediately beneath some sandstone bodies.
Skolithos burrows (mm to cm diameter) are common in the tops of sandstone stratasets. Arenicolites and Chondrites also occur. Large (up to 0.1 m diameter) vertical, sandstone-filled meniscate burrows are characteristic of thicker sandstone bodies. Shelly fossils include brachiopods (Spinocyrtia, Mucrospirifer, Tropidoleptus carinatus, Cupularostrum, Mediospirifer, Orthospirifer mesastrialis), bivalves (Goniophora, Grammysia, Actinopteria, Cypricardella, Paleoneilo), Tentaculites, and crinoid ossicles. Most shell fragments are disarticulated, lie concave down along bedding planes, and show little evidence of breakage.
Interpretation of lower part
The thickest, cross-stratified sandstone bodies are interpreted as deposits of laterally migrating tidal channels (Johnson and Freidman, 1969; Miller and Woodrow, 1991; Bridge and Willis, 1994). Ebb-directed paleocurrents are dominant in the lower parts of sandstone bodies, whereas there is flood dominance in the upper parts, as is typical of asymmetrical, coastal tidal currents. The general absence of tidal-bundle sequences and rarity of current ripples on reactivation surfaces in medium-scale cross strata is an indication of strongly asymmetrical tidal currents where either the flood or ebb current is dominant during deposition.
Some of the sandstone bodies indicate the influence of wave currents or combined currents lower down, but predominantly unidirectional currents higher up. Interpretation of the origin of these sandstone bodies depends on their lateral transition to the other types of sandstone bodies, as observed in nearby outcrops of coeval strata (e.g., Stevens Mountain Quarry, Figure 20). Accordingly, these bodies are interpreted as channel-mouth bars that were prograding into a storm-wave dominated sea. Such bars may have been on the seaward side of tidal inlets (i.e. ebb tidal deltas) associated with estuaries or barred coastlines, or associated with tide-influenced deltaic distributaries (Johnson and Friedman, 1969; Miller and Woodrow, 1991; Bridge and Willis, 1994). Landward-directed medium-scale cross strata may be associated with flood-tidal currents or may represent wave-formed swash bars. Ripple marks on dune/bar slip faces indicate periods of tidal slack water or reduced tidal currents.
The thinnest sandstone bodies dominated by swaley-hummocky cross strata are interpreted as more distal parts of channel-mouth bars that were dominated by storm waves and associated currents. The lack of mud suggests deposition above fairweather wave base. The occurrence of load casts near the base of these sandstone bodies suggests rapid deposition of sand on mud, as is common offshore from deltaic distributaries.
The shelly and pebbly strata at the top of the sandstone bodies are evidence of relatively strong shoreward-directed currents, possibly due to flood tides, "storm tides", or asymmetrical shoaling-wave currents. They appear to be associated with abandonment of the channel-mouth bar, relative reduction in sand supply, and marine transgression.
The sparse fauna, low degree of bioturbation, and dominance of Skolithos ichnofacies in the sandstones is typical of sandy deposits in coastal areas. Large, vertical meniscate burrows are similar to those reported in fluvial and coastal deposits elsewhere (Thoms and Berg, 1985; Miller, 1979; Bridge et al., 1986) and some may be due to upward escape of bivalves. The brachiopod-bivalve fauna is typical of Devonian nearshore communities (McGhee and Sutton, 1985; Sutton and McGhee, 1985).
The shales with interbedded sandstones are interpreted as nearshore marine deposits formed below fairweather wave base (details in Bridge and Willis, 1994). Sandstones were formed during storms by either wave currents or combined wave-unidirectional currents, whereas mudstones represent fairweather deposits or those accumulated below storm wave base.
Description of upper part
The lower Oneonta Formation (Figure 6, meters 53-83) comprises red, green and gray mudstones with subordinate red to gray sandstones. The mudstones range from fissile and relatively unbioturbated to intensely disrupted with blocky texture, slickensided clay-lined surfaces, horizons of calcareous nodules, and rare pseudo-anticlines. Desiccation cracks, root casts and burrows are pervasive. The distinctive trace fossil Spirophyton is present but hard to find. Sandstone stratasets are sharp based, cm to dm thick sheets and lenses. The thinner ones have small-scale cross strata with wave- or current-rippled tops. Thicker ones tend to have medium scale cross strata or planar strata at their bases. Centimeter-diameter, vertical meniscate burrows (Beaconites) are common, and root casts and calcareous concretions may occur at the tops of the sandstone stratasets. Groups of several sandstone stratasets with intervening mudstone may be arranged into meter-thick upward-fining or upward-coarsening sequences (e.g., Figure 6, 58-60 m). Two 4 to 6 meter thick sandstone bodies occur at the top of the sequence (Figure 6, 74-83 m). These have erosional bases (lined with intraformational shale clasts, plant fragments, and rarely the non-marine bivalve Archanodon) and large-scale inclined strata, that are channel filling in places (meters 80-83). Internal structures are medium-scale cross strata (mainly low angle), planar strata, and cm-diameter, meniscate vertical burrows.
Interpretation of upper part
Disrupted mudstones represent flood basin deposits. Blocky textures, desiccation cracks, calcareous concretions and pseudoanticlines indicate repeated wetting and drying and formation of ped structure typical of calcareous vertisols. Fissile siltstones indicate relatively less bioturbation but repeated exposure, and may represent relatively high deposition rates and/or hypersalinity in ephemeral lakes. Spirophyton in fissile siltstones suggests more permanent standing water, possibly brackish (Miller, 1979; Miller and Johnson, 1981). Wave-rippled lenses indicate periodic wave activity. The thinner sandstones with unidirectional sedimentary structures indicate periodic sand deposition during floods, either as individual sheet-flood deposits within the floodbasin, or associated with levees or crevasse splays. Coarsening-upward sequences indicate progradation of levees or crevasse splays. Root casts and calcareous concretions in and above the tops of these units indicate reduction in deposition rate and soil formation. The thick sandstone bodies are interpreted as the deposits of river-channel bars and fills, with some tidal influence possible.
Discussion of meter-scale sequences in coastal deposits of the Catskill clastic wedge
Meter-scale sequences were interpreted by Bridge and Willis (l994) to record changes in relative sea level related to processes such as lateral migration, abandonment and filling of tidal channels; progradation of channel mouth bars and tidal flats; and filling of coastal bays. These sequences are parasequences using the terminology of Van Wagoner et al. (1988, 1990) and Mitchum and Van Wagoner (1991). As average Givetian deposition rates were very approximately 0.1 mm/yr, meter-scale sequences represent 104 to 105 years on average.
Sharp-based coastal sandstone bodies similar to those in the Schoharie Valley have been ascribed to deposition on prograding strandplains cut by estuaries or tidal inlets, and to wave-formed offshore bars or shoals (references in Bridge and Willis, 1994). Sandstone bodies with swaley-hummocky cross strata overlain by angle-of-repose cross strata have been interpreted as beach and shoreface deposits, their sharp bases being related to wave-current erosion during relative sea-level fall and so-called forced regression. The coarse grained strata on the tops of these sandstone bodies were taken as so-called transgressive lags associated with relative rise in sea level. In contrast, Bridge and Willis (1994) explained the sharp bases of some of the sandstone bodies as due to rapid progradation of storm-wave modified channel-mouth bars that formed rapidly as a result of channel diversion. A beach face origin of the sandstone bodies was considered unlikely due to absence of: (1) characteristic seaward dipping planar laminae; (2) alongshore directed cross strata interbedded with landward directed cross strata and planar strata typical of ridge and runnel systems; (3) eolian cross strata. Although beach deposits can be eroded during marine transgressions, and by tidal inlets, universal removal seems unlikely. The apparent absence of beach deposits in both transgressive and regressive sequences along a shoreline that clearly experienced strong storm waves is probably related to high deposition rates of sand and mud near the mouths of a complex of channels. Thus channel switching is a viable mechanism for producing the meter-scale alternations of channel-related sandstone bodies and mudstones. Upward-fining sequences representing filling of coastal bays and/or progradation of tidal flats could be formed in areas away from active channels, perhaps landward of abandoned channel mouth bars (shoals) that were being reworked by wave currents.
Estuaries and sand barriers with muddy bays on their landward sides are expected in coastal areas during relative sea level rises (references in Bridge and Willis, 1994). Although there is evidence for washovers from sandy coastal shoals, there is no clear evidence for barrier beaches. Furthermore, no evidence was found for ravinement during regressions and estuarine filling during transgressions. The lack of evidence for muddy intertidal channel deposits with large-scale inclined heterolithic bedding, and for tidal-bundles in cross-stratified sands, indicates that the North Sea tidal flats may not be good analogues for these deposits (contra McCave, 1968; Johnson and Friedman, 1969). The strongly asymmetrical, ebb-dominated currents possibly record the dominance of river flow during deposition.
Other Givetian and Frasnian coastal and marine deposits in New York and Pennsylvania are broadly similar to those exposed in the Schoharie Valley, and interpreted depositional environments have several features in common: (1) sandy, tide-influenced channels; (2) shallow bays and tidal flats where mud and sand were deposited; (3) rarity of beaches; (4) storm-wave domination of the marine shelf (references in Bridge and Willis, 1994). Much of the variability in the deposits from different areas could be explained within the context of a wave- and tide-influenced deltaic coastline. Meter-thick sequences similar to those in the Schoharie Valley but in the Frasnian part of the Catskill clastic wedge have been interpreted to reflect eustatic sea-level changes associated with Milankovich climatic cycles (Van Tassell, 1987). Such an interpretation is considered unlikely here. Such an allocyclic control could only be justified if sequences could be traced laterally for distances greater than would be expected for individual delta lobes, and indeed across the whole basin. However, correlation of sequences for more than a few hundred meters has proved difficult in these types of rocks (Miller and Woodrow, 1991; Duke and Prave, 1991).
Discussion of 100-meter scale sequences in coastal and fluvial deposits of the Catskill clastic wedge
The Gilboa and Oneonta Formations in the Schoharie Valley can be correlated biostratigraphically with fully marine strata to the west and fluvial strata to the east (Figures 3 and 8). The correlations are approximate because the resolution of biostratigraphic zones is on the order of a million years at best. McCave (1969) associated mudstone-dominated fluvial sequences with marine transgressions based on the assumption that rising sea level would lead to increased rate of floodplain aggradation and higher preservation potential of overbank mudstones. This concept has been used by others more recently (e.g., Shanley and McCabe, 1993; Wright and Marriott, 1993). However, floodplain aggradation rate is only one of several controls on proportion of channel-belt deposits in alluvial successions (Mackey and Bridge, 1995). Within the 100-meter scale sequences represented by the Manorkill and lower Oneonta Formations, the mean grain size and proportion of channel-belt sandstones, and the size of individual channels, increase upward (Willis and Bridge, 1988). These trends were interpreted by Willis and Bridge (1988) to record progradation of seaward-fining distributive fluvial systems. Dating of these strata is not accurate enough to determine if the upward increase in channel-sandstone proportion is related to decreasing accommodation space and sediment accumulation rate in the way suggested by Shanley and McCabe (1993). However, by analogy with the Miocene Siwalik strata in the Himalayan foredeep, high proportions of channel-belt sandstone bodies are associated with high subsidence and deposition rates, increasing channel sizes, and progradation of megafans (Willis, 1993; Khan et al., 1997; Zaleha, 1997).
McCave (1969, 1973) and Johnson and Friedman (1969) postulated that thin marine limestones (e.g., Tully) were deposited during marine transgressions, because sea-level rise would cause trapping of sand and mud near the coast, starving deeper marine areas of sediment. In contrast, Brett and Baird (1985, 1986, 1990) interpreted such limestones to represent shallowing and sediment starvation-winnowing events following sea level lowstands. According to Brett and Baird (1990), the upper Moscow Formation is represented in marine areas by a Type 2 sequence boundary that is overlain by the lower Tully Limestone (Figure 3). No corresponding erosion surface has been recognized in coeval coastal deposits. The upper Tully Limestone apparently rests on a transgressive erosion surface (Brett and Baird, 1990), and may therefore correlate with the transgressive lower Gilboa Formation in Schoharie Valley (Figure 3). However, relative sea level changes are not necessary congruent in different parts of foreland basins, such that cross-basin correlation of regressive or transgressive strata does not necessarily imply coeval events.
Many authors favor eustatic sea-level changes as the main control on the large-scale sequences (references in Bridge and Willis, 1994). The transgressive deposits of the Gilboa Formation are associated with a widely accepted eustatic sea-level rise that can be recognized worldwide (House, 1983; Johnson et at., 1985). However, eustatic sea-level changes must be moderated by tectonically induced changes in subsidence and uplift rates and sediment supply. To establish whether eustasy or tectonism is the dominant control on sequence development in foreland basins, it is necessary to establish the ages of sequences, and the relative thickness of transgressive and regressive deposits, in different areas of the basin (Jordan and Flemings, 1991). It is very difficult to know if rising and falling sea level are coeval or otherwise in different parts of the marine basin, because of dating limitations. Also, it is very difficult to predict the timing and rate of deposition in different parts of the basin in relation to tectonic uplift of the adjacent mountains and development of the peripheral bulge. Jordan and Fleming's (1991) model suggests that the thicker regressive strata relative to transgressive strata in this region points to a tectonic control of some of the large-scale sequences.

A5: THE PALISADES SILL, NEW YORK AND NEW JERSEY
H. R. Naslund
Department of Geological Sciences
SUNY-Binghamton
Binghamton, NY 13902-6000
(Naslund@Binghamton.EDU)
General Geology of the Palisades Sill:
The Palisades sill is a Mesozoic tholeiitic intrusion approximately 300 m thick, that extends in outcrop for 80 km in a north-south direction along the western side of the lower Hudson river (Fig. 1, Walker, 1969), and may connect under cover with the Rocky Hill and Lambertville sills to the south for a total strike length of 150 kms (Husch, 1992). The east-west extent of the sill is unknown, but if it were circular in plan view it would represent over 1750 km3 of magma. Field relations, petrography, and petrology have been described by a number of authors (Lewis, 1908; F. Walker, 1940; K. Walker 1969; K. Walker, et al., 1973; Shirley, 1987; Husch, 1990, 1992; Steiner et al., 1992).
The Palisades sill has an extensive differentiation sequence from Mg-rich gabbro at its base to Fe-rich diorite at approximately the 250 m level. The roof sequence, although much thinner than the floor sequence, is increasingly differentiated from the roof downwards (Fig. 2). The main constituents of the sill are plagioclase and pyroxene (augite, orthopyroxene, and pigeonite), which together make up over 90% of the primary phases; olivine, amphibole, ilmenite, magnetite, and minor amounts of biotite are additional phases. Igneous layering within the sill is restricted to thin (cm-scale), faint layers best seen on weathered outcrops. Calculated solidification times for the sill are on the order of 1000 years (Shirley, 1987; Gorring & Naslund, 1995).
Mineral compositions vary from the base of the sill upwards (and from the roof downwards) reflecting the increasing differentiation of the Palisades magma (Fig. 3). Plagioclase compositions vary from An78 to An28; augite from Wosub38/subEnsub51/subFSsub11/sub to Wosub38/subEnsub10/subFssub52/sub; pigeonite from Wosub8/subEnsub54/subFssub38/sub to Wosub12/subEnsub34/subFssub54/sub; orthopyroxene from Wosub4/subEnsub78/subFssub18/sub to Wosub3/subEn5sub5/subFssub42/sub and from Wo3En34Fs63 to Wosub4/subEnsub26/subFssub70/sub; and olivine from Fosub77/sub to Fosub55/sub and from Fosub13/sub to Fosub5/sub (Walker, 1969; Walker et al., 1973; Gorring, 1992; Hristov, 1995). The plagioclase, pyroxenes, and olivine of the sill are all strongly compositionally zoned (Walker, 1969; Gorring, 1992; Hristov, 1995) with augite typically zoned 20 to 25% in Mg# (Mg# = Mg/[Mg+Fe]) from core to rim, orthopyroxene 15 to 20% in Mg# from core to rim, and plagioclase in excess of 25 An% within crystal cores, with rims up to 50 An% poorer than cores. Many individual grains have relatively homogeneous central cores that make up 25 to 75% of the area surrounded by strongly zoned rims.
The base of the sill is sharply defined by a very fine-grained to glassy, phenocryst free chilled margin against the country rock. Above the chilled margin, the grain size gradually increases upward in the sill from ∼2 mm to over 10 mm at the 250 m level. The lower 2.5 m of the sill is a fine grained, compositionally homogeneous, dolerite which appears to represent liquid solidified in situ with no significant addition or loss of crystals. Approximately 10 m above the base of the sill is the well known olivine-rich horizon which is thought to have formed during the emplacement of the magma as a result of initial magma inhomogeneities, or the nearly simultaneous emplacement of two magmas of contrasting compositions. Geochemical data suggest that there were at least two additional pulses of magma injected into the chamber at the 27 m and 45 m levels. These later pulses are marked by small but abrupt regressions in the generally decreasing Mg/Fe trend, abrupt increases in Ni, Cr, and modal orthopyroxene, increases in the Crsub2/subOsub3/sub contents of augite and orthopyroxene, and abrupt decreases in grain-size (Gorring & Naslund, 1995).
Shirley (1987) has modeled the effects of intercumulus melt migration and compaction in the partially solidified Palisades sill. His model suggests that the rocks originally accumulated on the floor of the magma chamber along with 50 to 60% interstitial magma. Some of this interstitial liquid was subsequently forced upward through the crystal mush by compaction, reducing the final proportion of "trapped" liquid to between 40 and 50 %. Steiner et al. (1992) present a "cumulus-transport-deposition" model for the crystallization of the Palisades sill based on a section through the sill in Upper Nyack, NY. The model proposes that "crystals collect in convection cells or like environments at deep levels within the sheet" and are then subsequently "remelted, resorted, and/or recrystallized during epizonal emplacement due to: magma decompression, shallow level flow differentiation, and magma-wall rock interactions" (Steiner et al., 1992). In this model the first 2.5% crystallization is dominated by olivine (olivine: plagioclase: augite, 100 : 0 : 0), the next 20.5% by augite (16 : 34 : 50), and the next 26% by plagioclase (10 : 60 : 30). The substitution of orthopyroxene for olivine in the model does not significantly affect the overall fit with the observed Palisades trends; the first crystallization episode (0-2.5% crystallization) has a better fit using olivine; the second and third episodes (2.5-23% and 23-49%), however, have better fits using orthopyroxene (Steiner,.et.aL., 1992).

STOP 1.
Park in the parking lot at the headquarters building and walk down the road (away from the parkway, towards the river) about 0.1 miles to the second set of outcrops on the left. This is about the 70 m stratigraphic level in the sill. The rock here is a two pyroxene diabase with weak modal layering. Plagioclase abundances vary from 45% to 60% and pyroxene varies from 35% to 50% from layer to layer. The Palisades diabase is, in general, a rather featureless rock. This horizon is one of two in the sill that has poorly developed rhythmic layering. Along the length of the central part of the sill similar layering is found at the 70 m level near the top of the cliff face. At stop 6 we will see another example of this layering.

STOP 2.
Leave flashers on and stay on the right side of the road to avoid traffic. Remind students to stay out of the roadway. Well developed rhythmic layering is present below the olivine horizon at the 10 m stratigraphic level. The main layering "signal" is in the orthopyroxene which varies from 10% to 25%. Texturally the orthopyroxene is ophitic to poikilitic at this level.

The circle above Ross landing boat dock.
Note the crude columnar jointing in the diabase above the circle. If you desire an additional stop: above the steps under the roadway at the circle, in the second buttress of outcrops to the north, flow banding is present within the olivine horizon. Watch for poison ivy !!!

STOP 3. Chilled Margin
Chilled margin (if the road is blocked you can see similar contact relations by walking south from the circle) .. Note the minor stair steps in the lower contact and the contact metamorphism of the arkose below the sill. The sill here has a uniform composition for 3 m in from the very fine grained chilled margin. Note that in places it is hard to tell the difference in hand specimen between hornfels sediments below the sill and the Palisades fine-grained chilled margin.
Proceed south under the George Washington Bridge. A phytosaur (Clepsysaurus Rutiodon) was found in the Stockton arkose below the bridge.

STOP 4. Classic olivine zone on the left
Just before the park entrance there is a classic undercut outcrop of the olivine horizon. Although it looks really rotten you can see fresh olivine in it.

STOP 5. The entrance to Fort Lee Historic Park.
Drive in or park on the street and walk in. The first outcrops inside the park along the entrance road have sub-horizontal, irregular, plagioclase-rich layers. These appear to be layers of the type theorized by Coats (1936) to have formed by settling and compaction of a crystal mush within a viscous fluid.

STOP 6. Outcrop with rhythmic layering
Outcrop with the best rhythmic layering in the Sill is located around the corner you just passed on Bridge Plaza North. This is again at the 70 m level in the sill. This layering is very regular suggesting a mechanism involving changing conditions of crystallization at the crystal/magma interface rather than layering due to compaction processes.

STOP 7. Sandwich Horizon
Park in front of the NJ dept. of transportation lot. Outcrops of coarse, iron-rich "sandwich horizon" are exposed to the left of the cemetery entrance.

STOP 8. Upper contact of sill
Just after you pass under an impressive high bridge pull over on the shoulder. Park with your flashers on and stay off of the roadway. Be careful, if the police come, be polite and they will probably ask you to leave. Cross over the guard rail and look at the upper contact of the sill. Lots of interesting vein minerals and some grain-size layering. The roof rocks below this level (back before the bridge you just passed under) contain segregations of coarse-grained gabbro injected as dikes and sills.

B3: CORNWALL-TYPE IRON ORES OF PENNSYLVANIA
H.R.Naslund
Department of Geological Sciences
SUNY-Binghamton
Binghamton, NY 13902-6000
(Naslund@Binghamton.EDU)
Introduction:
Magnetite-apatite ore deposits are important ores of iron in many parts of the world. They can be classified into two major groups: a) high Ti, high -P magnetite-apatite-ilmenite deposits (Nelsonites) associated with anorthosites; and b) low-Ti, variable-P magnetite +/- apatite deposits (Kiruna type ores) principally associated with volcanic rocks. The origin of the first type is generally accepted to be magmatic, with liquid immiscibility between silicate magma and iron-titanium-phosphorous-oxide magma the most commonly cited process (Philpotts, 1967; Kolker, 1982; Barton & Johnson, 1996). The origin of the Kiruna-type, however, has been a matter of controversy, with both magmatic and hydrothermal origins proposed. The Cornwall-type deposits of Pennsylvania, may be considered as an apatite-poor variety of these Kiruna-type deposits.
Many of these Kiruna-type deposits, contain originally-horizontal, laterally-extensive, stratified, iron oxide-rich units. These stratified iron oxide deposits do not appear to have formed by the same mechanism in all localities, indeed there may be a variety of mechanism that would lead to units with similar appearances. Such mechanisms might include: (1) magmatic ignimbrites, airfall ash, lavas, or sills; (2) sedimentary exhalative, lateritic, or detrital deposits, and (3) hydrothermal replacement, veins, or phreatic surface venting of hydrothermal systems. These various types of deposits should present different field relationships, textural characteristics, and trace element compositions, in accord with their mechanism of formation. Stratified iron oxide ores have been described from El Laco, Chile (Henriquez and Martin,1978, Nystrom and Henriquez,1994); the Chilean Cretaceous Iron Belt (Travisany et al.,1995); Kiruna, Sweden (Nystrom and Henriquez,1994); Cerro Mercado, Mexico (Swanson, et al., 1978; Lyons, 1988); Vergenoeg, South Africa (Crocker, 1985); and the Bafq Mining District, Iran (Forster and Jafarzadeh, 1994). In all of these localities the authors have interpreted them as pyroclastic iron ore. Similar stratiform deposits in the Missouri Iron District have been interpreted as replacement of volcanic tuff (Snyder, 1969; Ridge,1972; Panno and Hood, 1983), or as hematite tuff produced by the venting of a fumarolic system into a lake (Nold,1988; Hauck,1990). Stratified magnetite ores of the Cornwall-type are associated with a series of diabase sills in the Newark and Gettysburg Triassic rift basins. Cornwall-type deposits are characterized by laterally extensive, banded to massive magnetite ores with only traces of apatite, located adjacent to the upper contact of diabase sills. They are generally hosted in limestone or marble country rock, but locally they are hosted in the diabase sills or in clastic sediments or gneissic country rock. The Cornwall-type ores have been interpreted as hydrothermal replacement deposits (Smith, 1931; Tsusue, 1964; Lapham & Gray, 1973; Eugster & Chou, 1979).
The recent discoveries of the enormous Cu-Au-U-REE hematite breccias at Olympic Dam in Australia (Roberts & Hudson, 1983; Oreskes & Einaudi, 1990, 1992; Haynes, et al., 1995), and the giant REE iron oxide deposit at Bayan Obo, China (Argall, 1980; Wang, 1981; Wei & Shangguan, 1983; Bai & Yuan, 1983) have made iron oxide-apatite deposits of the Kiruna type of considerable economic interest for commodities other than iron. Indeed, many Kiruna-type deposits are associated with near-by Cu, Au, U, or REE mineralizations (Crocker, 1985; Kisvarsanyi, 1990; Hauck, 1990; Einaudi & Oreskes, 1990; Marikos, et aI., 1990; Vivallo, et aI., 1993,1994, 1995a, 1995b; Espinoza, et al., 1994; Forster and Jafarzadeh, 1994; Rojas & Henriquez, 1994; Foose & McLelland, 1995). The Cornwall type deposits are associated with minor, but recoverable amounts of Cu, Au, Ag, and Co, but no significant REE concentrations have been reported in the district.

B3: CORNWALL-TYPE IRON ORES OF PENNSYLVANIA

STOP 1. Pennsylvania Granite Corporation Quarry.
Diabase is exposed in the parking lot and along the road to the south. The main point of this stop is to see the diabase that is associated with the ore deposits. Two of the largest ore bodies in this district are located 1 mile to the north of this spot. Note how unaltered and unfractured the diabase is at this location. In fact, its fresh, unfractured character is the reason there is a building stone quarry here. If the iron ores were formed from a hydrothermal system associated with the cooling of the sill, it does not appear to have affected the underlying diabase.

Top of the diabase sill.
This stretch of 23 west travels on top of the diabase sill. Note the diabase stone fence on right.

Turn right and park cars.
This is state game land open to hunters. Not a good spot to visit during deer hunting season. Proceed on foot through gate on "emergency road". Turn left at every road junction.

STOP 2. Hopewell Iron Mine
Note the "blue conglomerate" in the float along the road. This Ordovician sediment forms the cap rock above the Cornwall deposit. Its relationship to the ore body here is not known. After about 10 minutes you will come to an open field. Follow the road to the left under the power lines. At the far edge of the trees follow the road to the left back into the woods. This is the Hopewell iron mine. Dumps are on the right, the open pit is on the left. Exposed in the walls of the pit is a quartz feldspar pegmatite that predates the ore body, but cross cuts the host gneiss. A similar pegmatite is present in the Ore body in St. Peters village to the east. The relationship of these pegmatites to the diabase sill is not clear. In the dumps you should be able to find all the ore textures present in the deposit.

Hopewell Furnace National Historic Site is 3.3 miles north of here
3.3 miles north of this point is the Hopewell Furnace National Historic Site. Open daily 9 to 5, except national holidays. If you are with a student group and call in advance you can take a self guided tour of the old furnace for free. Otherwise it is $4. Hopewell Furnace National Historic Site, 2 Mark Bird Lane, Elverson, PA 19520. ph. 610-582-8773. If you want to show students about iron making in the 19th century, this is a better site than the Cornwall Furnace Historic Site

The house on right is "Pennsylvania Black granite".
The house on right is constructed of locally quarried diabase "Pennsylvania Black granite".

Elverson Homes and Johes Luck Iron Mine
Several of the older homes in Elverson are constructed out of "Pennsylvania Black Granite". The Jones Luck Iron Mine is located approximately 2 miles north of this point at the intersection of Route 82 and Red Hill Road. The mine is a good example of reclamation. The open pit is now a swimming lake north and east of the intersection. The only sign of the mine are some tailings piles on the north side of Red Hill Road (south side of the lake) just past the intersection.

Grace Iron Mine half mile north
About 0.5 miles north of here is the now inactive Grace Iron Mine where calcite-dolomite geothermometry in veins adjacent to the ore body, suggests temperatures in excess of 650 degree C. The ore body is located at the upper contact of a diabase sheet, and is composed of magnetite and serpentine.

Good diabase outcrops on the right.

STOP 3. Mine dump on the left.
This is from the main pit of the NO.3 mine.

Cornwall Iron Furnace Historic Site on the right.

STOP 4. After crossing a small bridge park on the right.
Return to bridge on foot and climb down into railroad cut at the south east side of the bridge (the corner closest to where you parked). Railroad cut through the diabase sill below the ore body. Outcrops of the diabase are just north of the bridge (to your right as you climb down). Outcrops of altered sediments below the diabase sill can be reached by walking south along the railroad right of way for about 200 meters.

STOP 5. Driveway on right.
"No dumping of Refuse Allowed" Park and walk back along old railroad right of way. Slag from the iron furnace has been dumped along either side of the road. Near where you park, most of it is covered by later piles of concrete and asphalt. About 30 meters back on the right are some good chunks of furnace slag.

STOP 6. Turn into driveway on left watching for traffic.
If you are in a big van it is best to back in so you will be able to see when pulling out. This is the upper part of the No.3 mine. The contact between the ore body and the diabase is preserved on the slope to the left of the talus pile. Be careful of the talus pile, it is unstable so it is best not to climb on it. Most of this talus is probably from the underground Elizabeth mine at the top of the hill. Walk south along the road to diabase outcrops across from the Historic Marker. Note how fresh it is for a rock associated with a major hydrothermal system. Walk across the street to the overlook beside the historic marker for a good view of the pit.

Turn around at the baseball diamond and head back the way you came.
The main point of this detour was to drive through the old Cornwall mining village. There are old tailings piles from the No. 4 mine behind first base if you haven't had enough of mine tailings by this point, or if you have been scared off by no trespassing signs at earlier stops.

STOP 7. Cornwall Iron Furnace, Historic Site.
Tuesday-Saturday 10-4; Sunday 12 -4; closed Monday.

B4: FIELD ILLUSTRATIONS OF GEOLOGIC FEATURES IN THE UPPER SUSQUEHANNA VALLEY AND ADJACENT MOHAWK REGION
DAVID M. HUTCHISON
Department of Geological and Environmental Sciences
Hartwick College
Oneonta, NY 13820
The outcrops and surficial features observed on this trip have been selected (1) to show students the change in lower Paleozoic stratigraphy through time (2) to illustrate several sedimentary and topographic features (3) to help students gain a better understanding of the geologic framework of this area.
The trip starts at the large Upper Devonian flood plain channel behind the F.W. Miller Science Building on the Hartwick College campus in Oneonta and ends in Precambrian garnet gneiss six miles east of Canajoharie. Progressively older beds are exposed to the north because of three factors: the gentle southerly dip of the beds, the erosion by the Mohawk river and the uplift, tilting and erosion of large fault blocks associated with normal faults in the Adirondack Mountains which extend into the Mohawk Valley.
The clastic sediments are the result of the Middle Ordovician Taconian Orogeny (470-435 m.y. ago) and the Middle and Late Devonian Acadian Orogeny (385-355 m.y ago). Both of these times of crustal unrest and uplift east of the Hudson River and in New England provided an influx of clays, silts and sands into the Ordovician and Devonian seas which occupied the area of this field trip. The carbonate rocks were deposited by these seas during periods of quiescence between orogenies.

STOP 1. Upper Devonian Oneonta Formation
This outcrop consists of dark grey shales, thinly bedded siltstones and sandstones. A massive sandstone which fills in part of a Devonian stream channel is well exposed for about 150 feet at the left end of the outcrop. There is an erosion surface between the underlying shale and the overlying massive sandstone. This surface which marks the bottom of the channel rises stratigraphically to the right. Interference ripple marks are exposed just below the road leveL Plant materials are abundant in some layers. A few galena crystals about 1 mm across have been found in the thinly bedded sandstone.
Trilobites, marine pelecypods and brachiopods of the Gilboa Formation have been found in beds stratigraphically 75 feet below this outcrop, and marine fossils are found in thinly bedded siltstones and shales stratigraphically 200 feet below stop L This outcrop is next to Nick's Diner on Chestnut Street. We probably will not stop here today.
A clean sandstone which shows extensive trough-type cross bedding overlies this outcrop.
Stop 1A is about 200 feet higher in elevation. Follow the road to the Ernest B. Wright Observatory.
Red mudstones and shales with root casts, worm burrows and ripple marks are exposed at the top of the hill just below the Ernest B. Wright Observatory which houses a 16" telescope.
The view down the valley is looking southeast to Mr. Utsayantha in Stamford (about 30 miles map distance). The broad U-shaped valley is the result of a Pleistocene glaciation. The Susquehanna River flows in from the north (out of your view) and through the city of Oneonta. The terraces on either side of the valley are kame terraces. Red beds are exposed near the top of the hills across the valley.

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

View to right showing kame terraces across valley.

On right (South) across valley gravel pit in delta kame.

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

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

Cross Susquehanna River.
Clays indicate a glacial lake occupied this area. Grass now covers the glacial clays.

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

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

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

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

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

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

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

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

Upper Middle Ord Shales and Sandstones in cliff at left.

Two drumlins are in view to the left.

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

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

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

Overpass of Route 5S over the railroad tracks.
Note that west of the Noses fault-line scarp the Mohawk River Valley has steep cliffs of Little Falls Dolostone. These are caused by the down cutting of the Mohawk River on the upthrown western side of the fault. East of the fault-line scarp on the downthrown side of the fault-line scarp there are no cliffs.