NYSGA 1997.kmz

NYSGA 1997
Field Trip Guide for the
69th Annual Meeting of the
New York State
Geological Association
Edited by:
Todd W. Rayne
David G. Bailey
Barbara J. Tewksbury
Hamilton College
Clinton, NY
This guidebook was published by the New York State Geological Association. Additional copies may be obtained from the Executive Secretary of the NYSGA:
Dr. William Kelly
Executive Secretary, NYSGA
Rm. 3140, Cultural Education Center
New York State Geological Survey
Albany, NY 12230

A1: MADISON COUNTY DEPARTMENT OF SOLID WASTE LANDFILL FACILITIES
Paul W. Miller
Assistant Director
Madison County Department of Solid Waste
P.O. Box 27, Wampsville, NY 13163
General
Madison County runs a comprehensive solid waste and recycling program at its facilities located astride Buyea Road in the Town of Lincoln, Madison County, New York. Madison County is home to about 72,000 souls with numerous small villages and only one city, Oneida, with about 10,000 residents. Madison County contains the geographic center of New York State and is genuinely proud of its rural nature.
Two large abandoned marginal farmland properties have been used for landfilling activities since early 1973. Over time the changing regulatory requirements for solid waste management have necessitated numerous changes in construction standards for landfill facilities and environmental monitoring. Ideal geologic conditions have allowed operations to keep pace with changing requirements with reasonable economic investment. All expenses at the Solid Waste Department facilities are paid by user fees. No general tax revenue is used to support Department operations. In Madison County waste-pays-for-waste now but the changing world of solid waste management is creating economic pressures and challenges necessitating creative management.
One Field Trip Stop, but no Road Log

Madison County Landfill
The methods and procedures described provide a general overview of environmental and geotechnical drilling and sampling. These methods and procedures are used to provide critical subsurface data on many projects. The references that follow are just a partial list of the many publications currently available about Environmental and Geotechnical drilling.

A2: DRUMLIN-BLUFF AND BAYMOUTH-BARRIER EROSION ALONG THE SOUTHEASTERN SHORE OF LAKE ONTARIO, NEW YORK
Paul R. Pinet
Charles E. McClennen
Department of Geology
Colgate University
Hamilton, NY 13346
email: ppinet@center.colgate.edu
email: cmcclennen@center.colgate.edu
Introduction
The coastline of Lake Ontario has a long, complicated history of rapid morphologic development. The purpose of the trip is to examine the principal landform elements of this lakeshore sector -- drumlin bluffs and baymouth barriers (Fig. 1) -- in order to understand the shoreline's primary evolutionary history. Our focus will be on the coastal geomorphic changes that have occurred since the beginning of this century. The specific goals are 1) to assess our proposed model of coastal bluff evolution, 2) to examine the sedimentological interconnections between bluffs and baymouth barriers, and 3) to appraise the long-term impact (50-100 years) of engineered shore-stabilization structures.

A2: Drumlin-Bluff And Baymouth-Barrier Erosion Along the Southeastern Shore of Lake Ontario, New York

STOP 1. McIntyres Bluff and Environ
We will walk out to the edge of McIntyres Bluff in order to get an overview of this coastal compartment. BE VERY CAREFUL, BECAUSE THE EDGES OF THE BLUFF TOP ARE SUSCEPTIBLE TO COLLAPSE. Also, because the western end of McIntyres Bluff is private property with a home owner who values his privacy, we will confine our field inspection to the eastern half of the drumlin. Please comply with this request.
McIntyres Bluff is a north-south-trending drumlin. Note (Fig. 11) that the drumlin's axis lies oblique to the shoreline which east of Little Sodus Bay runs northeast-southwest. Gravel and sand eroded from the drumlin till are transported by longshore currents to the northeast and nourish a low-lying baymouth barrier. Notice how narrow the gravel and cobble beach is at the base of the bluff. This is due to two factors: 1) the lake level this past spring has been unusually high which has resulted in more beach erosion than usual, and 2) the gullies are not supplying much debris to the beach, as indicated by the dearth of mud flows that typically fill the V-shaped gully floors as well as by the absence of colluvial fans at the gully mouths. Nine years ago when we began our field studies of the southeast Lake Ontario shore, a well-developed colluvial terrace occupied the upper beach which at that time was much broader than at present due to the availability of sediment.
Features that are important to note as we inspect the gully network include:
i. the bowl-shaped head region of the gully system and its control by jointing;
ii. wet spots on the headwall that identify zones of groundwater seepage;
iii. the ubiquity of fresh slump masses at the base of the headwall;
iv. the steep-flanked ridges and narrow throat of the gullies;
v. the presence of mud flow deposits in the stretch of gully below the headwall, and of cobble and boulder lag deposits in narrow channels that are cut into the gully fill;
vi. the angularity of the large clasts in the till and mud flow deposits and the subrounded to rounded shapes of cobbles and gravel on the beach proper;
vii. the presence of a wave notch at the bluff's base and the development of a slump scarp running the full extent of the bluff.
Given the above morphological characteristics, what stage of bluff development does McIntyres Bluff represent?
Now we will hike along the beach to the east, in order to examine the morphology and sediment character of the downdrift baymouth barrier (Fig. 11). Note that the surface of the beach that runs the length of McIntyres Bluff is dominated by boulders, cobbles, and gravel, despite the large quantities of sand in the glacial till of the drumlin. Clearly, these coarse sediments are lag and storm deposits, the sand having been eroded and transported offshore by high-energy waves. Unlike ocean beach systems, there is no long-period swell in Lake Ontario, and fair-weather waves are incapable of transporting the nearshore storm-deposited sands back onto the beach (McClennen and Pinet, 1993). As such, the beach deposits are out of equilibrium with the fair-weather conditions that prevail along the lakeshore.
Note the numerous trees that have been uprooted and fallen as a consequence of the rapidly transgressing shoreline. During the past four years, this baymouth barrier has been relatively stable as indicated by our profiling surveys (Fig. 8C). This past spring, the barrier has been overtopped by storm waves forming overwash deposits. The onset of overwash along this barrier seems related to the unusually high lake levels during the past spring, in conjunction with a significant reduction in the amount of longshore drift due to the recent inactive nature of the gullies that are incised into McIntyres Bluff. Although McIntyres Bluff is in a mature morphological state, it is temporarily behaving as a bluff in old "age," because of the inactive nature of its gully system as indicated by 1) the lack of mud flow deposits and colluvial fans, and 2) the narrowness of the beach at the bluff's base. Presumably, the temporary reduction of sediment influx eroded from the bluff that supplies the coastal compartment is related to diminished groundwater seepage and surface runoff.
About 0.5 km from the eastern edge of Mcintyres Bluff (Fig. 11), we encounter a small bluff with a steep face and a wave notch. This small drumlin has just begun to be eroded by storm waves, and is a classic example of a bluff in a young stage of morphological development. Note the numerous fresh slumps at the bluff's base, the steep face (>45 degrees), and the presence of rills and a few channels etched into the bluff's surface. Given that the crest of this small drumlin does not rise more than 20 m above the lake level, which is slightly below the geomorphic threshold that we believe separates young from mature bluffs (Montesi and Pinet, 1997), we expect that it will not undergo extensive gully development as the drumlin is cut back by the advancing shoreline; rather it will remain a young bluff and eventually pass directly (i.e., bypass the mature and old stages) into the final terminal stage of morphological development. However, if the gullies of Mcintyres Bluff are reactivated and supply copious sediment to this downdrift barrier such that the beach widens substantially, then its possible that this small bluff will be protected from frequent wave attack, allowing the onset of gullying despite its small topographic stature.

STOP 2. Beach at Base of Drumlin to the W of Sitts Bluff
The lakeshore from this point westward to Little Sodus Bay (Fig. 9) has undergone extensive engineering in order to stabilize the shore of North Fair Haven and Fair Haven Beach State Park. The 1.5-km-long strand has been armored with jetties, seawalls, revetments, riprap, and groins. All of these structures are accessible by walking along the beach to the west. If we look to the northeast, we can see Sitts Bluff with its towering, unvegetated cliffs (Fig. 9). To the southwest, we can see a large drumlin that lies within the State Park. Note that it is vegetated and is stabilized by a revetment along the shore.
After examining a groin (pier) that has been reattached to the shore with riprap, we will proceed to the northeast in order to examine a baymouth barrier that is retreating at a rapid rate, at almost 2 m/yr (Figs. 9 and 10). This is not surprising, given that the net longshore drift is to the east and the barrier lies downdrift of jetties and groins, and that the base of the two bluffs between The Pond and the baymouth barrier are protected by seawalls, riprap, and revetments. This section of the lakeshore is sediment starved. Therefore, the baymouth barrier is retreating landward rapidly as storm waves overtop the beach ridge causing extensive overwash. Note the well-rounded character of the beach material.
One reason why we believe that at least segments of the southeastern Lake Ontario shoreline consist of discrete coastal compartments is that sediment starvation here is confined to the drumlin bluff-baymouth barrier couplet. Downdrift of us lies Sitts Bluff, an unprotected drumlin in a late mature stage of morphological development that is supplying coarse sediment to its down drift baymouth barrier. Our beach-profiling studies (Fig. 8) indicate clearly that the barrier to the east of Sitts Bluff has been in a quasi-equilibrium state since 1950, and has not retreated appreciably in contrast to the barrier that we are currently examining. In other words, the past 50 years of severe sediment starvation that characterizes the segment of the lakeshore that we are now examining has not affected the adjoining downdrift compartment, suggesting that bluff-barrier couplets are self-contained entities to a large degree.

STOP 3. Chimney Bluffs
Chimney Bluffs (Fig. 12), an undeveloped State Park, is located about 22 km to the west of Stop 2 (Fig. 1), and about 3 km to the east of Sodus Bay. This excursion adds a considerable amount of time (1-2 hrs) to the trip. Chimney Bluffs is a high, broad drumlin with enormous, deeply-incised gullies that are unusually long for the Ontario lakeshore. According to our classification, it looks to be in a very late mature stage of morphological development, as judged from its appearance on topo maps (almost half of the drumlin has been eroded) and clear signs of old "age" in the field. As we walk along the beach to the west, note how the steep, planar, low sides of the bluff which are controlled by gravity slides are replaced by small and then large gullies as the bluff height increases toward the drumlin's center. Also, note that the gully floors are devoid of thick accumulations of mud flow deposits. However, there are remnants of colluvial fans at the gully mouths that have been eroded back by waves to the bluff face; observe how difficult it is to distinguish glacial till from mud flow deposits which grade into one another at the gully mouths. Wave notching and a slump scarp at the bluff's base are clear evidence of a lack of sediment supply to the adjoining beach. If our interpretation that Chimney Bluffs is in the old stage of morphological evolution is correct, we expect that the rate of retreat of the top surface of the bluff will diminish with time, whereas recession of the base of the bluff will likely accelerate as the sediment supply to the beach is reduced, allowing storm waves to undermine the bluff and cause increasingly more slumping with time.
As we approach the west flank of the drumlin, we will encounter varved lake sediments deposited in glacial Lake Iroquois. Typically the varves are millimeters thick, and alternate between clay-rich and silty sand-rich laminations, the former representing low energy winter conditions when the lake surface was frozen, the latter higher-energy conditions associated with spring melting. Ice-rafted dropstones are common. Some of the varve sequences are distorted, crumpled into tight folds, that represent either soft-sediment deformation, iceberg keel groundings, or minor glacial ice readvancements during the retreat of the Ontario Lobe of the Laurentide Ice Sheet.
We can return to our vehicles by climbing up the bluff along a crestal trail which provides a superb panorama of the gullies, embayments, fans, and beach and nearshore features. BEWARE OF TREACHEROUS FOOTING AND UNSTABLE CLIFF EDGES. If you prefer, you may return to our starting point by retracing your steps along the beach. For those of us hiking the ridge line, note the appearance of the gullies from above and the nature of the shoulders of adjoining gullies. Also, the map view of Chimney Bluffs shows a flat top to the drumlin, a feature that is attributed to wave planation when the drumlin was immersed beneath Lake Iroquois. Support for this interpretation is provided by irregular lag deposits of gravel that are visible at the very top of some of the gully headwalls. Presumably, the glacial gravel was sorted and the pebbles were rounded by shallow-water wave activity on this former shoal which at that time was located several kilometers offshore of the southern Lake Iroquois shore.

A3: TIMING OF INTRUSION, ANATEXIS, AND METAMORPHISM IN THE PORT LEYDEN AREA OF THE WESTERN ADIRONDACKS
Frank P. Florence
Science Division, Jefferson Community College
Watertown, NY 13601
frank florence@ccmgate.sunyjefferson.edu
Robert S. Darling
Dept. of Geology, SUNY College at Cortland
P.O. Box 2000, Cortland, NY 13045
darlingr@snycorva.cortland.edu
Introduction
The purpose of this field trip is to examine evidence of the igneous, metamorphic, and tectonic history of the western Adirondack Highlands preserved in rocks in the vicinity of Port Leyden, NY. We will visit and discuss rocks and structures at six sites that include interlayered metasedimentary and metaigneous lithologies and a distinctive nelsonite (apatite-magnetite-ilmenite) dike. Rocks at these sites display structural fabrics and crosscutting relationships formed through a sequence of metamorphic, hydrothermal, and melting events. Recent geochronological studies confirm that this area has undergone a complex tectono-thermal history.
In particular, this trip will focus on uncertainties concerning the anatectic or intrusive origin of leucogranitic gneisses and melt veins in the gneiss surrounding the nelsonite dike. By clarifying structural and field relationships, we hope to discriminate between interpretations of igneous activity and metamorphism in these gneisses. Potentially, this may establish useful constraints on regional interpretations of orogenic sequence.
Trip Objectives
This trip highlights some of the controversial structural and chronological relationships among rock units. When visiting the stops on this trip, you are encouraged to consider these questions:
1. What was the timing of penetrative deformation and to what extent, if any, can mineral assemblages and textures formed during the intrusion of the AMCG suite be recognized in the Port Leyden area rocks?
2. Unstrained phases that crystallized from melts occur in veins, reaction zones, and in equigranular granitic rocks. Is granitic intrusion implied by these features? Which features might be anatectic? Can anatexis explain the 1032 Ma zircon age date from granitic gneiss?
3. Are the quartz-sillimanite nodules and stringers (or their protoliths) relatively early features that were present in the surrounding gneiss prior to penetrative deformation or did they form relatively late due to hydrothermal activity in a crystallizing magma?
4. Did pegmatites originate from late-magmatic fluids at the time of granite intrusion or are they a result of fluids evolved from in situ rocks at the time of the peak of metamorphism?
5. What tectonic significance can be placed on the moderate (vs. high) pressure estimates obtained from this region of the Adirondack Highlands?

A3: Timing of Intrusion, Anatexis, and Metamorphism in the Port Leyden Area of the Western Adirondacks

Low roadcuts of Ordovician Black River Group limestone.

Barrett quarry.

View of historic Black River canal quadruple locks.

STOP 1. Rt. 12. Calc-silicate and Aluminous Metasedimentary Rocks
These rocks comprise one of two narrow but mappable bands of calc-silicate gneiss that extend northeastward, under the Pleistocene deltaic sands, and reappear along the Moose River (Fig. 1). Foliation in both calc-silicate bands strikes northeast and dips moderately northwest.
A variety of bulk compositions can be observed in this exposure, but the dominant minerals are clinopyroxene, biotite, K-feldspar, and quartz. Other minerals include grandite garnet, titanite, wollastonite, calcite, and scapolite. Compositional data are limited, but EDXF spectra from the garnet show mostly grossular with a small andradite component. The scapolite is mostly meionitic. In thin section, the margins of wollastonite grains contain thin rims of retrograde garnet identical to that described by Valley et al. (1983) from drill cores just 3.0 kilometers NE of this location. This garnet is unusual because it contains stoichiometric fluorine (up to 0.76 wt.%; Valley et al., 1983). Wollastonite occurs in apparent textural equilibrium with calcite and quartz.
As noted in some other Adirondack marbles, the occurrence of regional metamorphic wollastonite suggests an H2O-rich fluid composition (Valley and Essene, 1977; 1980b). We interpret these rocks similarly. Specifically, at the PT conditions in this region (700-750 °C and 4-6 kbar), a fluid composition of XH2O ≈ 0.80-0.90 is necessary to stabilize these minerals (Jacobs and Kerrick, 1981). The margins of some quartz ± pink K-feldspar veins are characterized by an assemblage of epidote + specular hematite, suggesting a higher O2 fugacity compared to that in the country rocks.
At the northernmost part of the roadcut there is a small exposure of leucocratic K-feldspar + quartz + sillimanite gneiss. Minor amounts of biotite, magnetite and almandine garnet occur locally. The presence of matrix sillimanite in these exposures, coupled with the bulk composition of the calc-silicate gneisses, strongly argue in favor of a sedimentary origin for both rock types.

STOP 2.Leucogranitic Gneiss with Quartz-Sillimanite Stringers
Outcrops in the woods on both sides of the road consist of relatively homogeneous, granitic gneiss. Abundant, cm wide stringers of quartz and sillimanite, with minor oxides, biotite, and garnet, are distributed throughout these rocks. Orientation of these stringers parallels the gneissic foliation defined by biotite and oxides, which locally varies between N20E and N50E. These exposures lie within a 1.5 km wide band of leucogranitic gneiss that can be traced N40E from the Black River to exposures along the Moose River at Lyonsdale (Fig. 1).
Immediately west of the road, and about 5 m uphill, are a few exposed knobs that also contain quartz-rich layers (leucosomes?) parallel to foliation. Biotite and magnetite occur in mm thick margins on these layers. Note that tight to isoclinal folds deform the quartz layers and are parallel to foliation and the quartz-sillimanite stringers.
Walk across to the east side of the road, enter the woods, and continue walking to the northeast for about 100 m to where a large dome of pink, granitic gneiss is exposed. On the southwest side of the outcrop there is a boudinaged quartz pegmatite layer. Sheared layers of leucogneiss parallel the margins of the boudin and flow inward around the necks. Coarse K-feldspar, possibly precipitated as a melt or hydrothermal phase, extends from margins through the pegmatite. Boudin margins contain bladed sillimanite, biotite, and magnetite. Additional folded and disrupted quartz veins are displayed around the right side of the outcrop.
Quartz-sillimanite stringers are distributed throughout the rock. Their orientation relative to penetrative deformation is well displayed across the top of the dome, where these features can be seen to form parallel to the plane of foliation. NE trending isoclinal folds deform some quartz-sillimanite segregations.

Outcrop of leucogranitic gneiss.

STOP 3. Lyonsdale.
Metapelitic Gneiss, Leucogranitic Gneiss, Sillimanite Nodules, and Undeformed Pegmatites
Walk down to Moose River. Follow trail upstream approximately 250 m to northward bend in river. Continue around bend, then walk down to the river's edge to metapelitic gneiss.
The river's orientation through this scenic section is parallel to lithologic layering. A low angle wall of homogenous quartzofeldspathic gneiss is exposed across the river. Garnet-rich metapelitic rock (kinzigite) is exposed for over 200 m along strike on this side, close to the water's edge. There is a sharp compositional boundary between the metapelitic gneiss and adjacent leucocratic gneiss about 5-15 m away from the river. From this contact back to the bridge, the gneiss contains less biotite and garnet.
A meter scale isoclinal fold is exposed in the metapelite layer on the southern end of the outcrop, with fold layers made up of gray, biotite-rich and pinker, perthite-rich layers. Garnet is ubiquitous; some layers contain coarse garnets > 1 cm. Elsewhere, pods of abundant small garnets form in quartz veins. Sillimanite is distributed throughout some layers and absent in others. Florence et al. (1995) found undeformed melt veins along grain boundaries locally in the metapelite. Small (0.1 mm) inclusion-free, euhedral garnet neoblasts occur with plagioclase and K-feldspar-quartz segregations surrounding embayed biotite. These textures indicate partial melting according to the biotite-dehydration reaction: Bt + Sill = Grt + Kspar + melt.
Walk back towards the parking area along the trail. Notice the pink leucogranitic rock at the canoe launch by the bridge. This gneiss clearly displays nodules of quartz and sillimanite. Two orientations can be observed, and a deformed pegmatite is seen to truncate and cross-cut the quartz-sillimanite veins.
Return to the road, cross it to a dirt road, and walk along the road in the downstream direction. After about 100 m, walk left through the woods and out onto a broad exposure of orange-pink, perthitic, leucogranitic rock with many quartz-sillimanite stringers, veins, and lenses. At least two orientations of quartz-sillimanite veins are displayed here: N50E veins are cross-cut by N20E veins (Cunningham and Willis, 1996). En echelon gash veins indicate extension and rotation within a shear zone and some small-scale granite melt veins truncate quartz-sillimanite veins. Deformed pegmatite and quartz layers are common. Large (0.5 m width), undeformed pegmatites dated at 1027 Ma cross-cut all other features.
Scattered calc-silicate inclusions are present in the gneiss throughout the outcrop area. 250 m below the bridge, biotite-rich and garnetiferous horizons become common. These layers contain sodic plagioclase, as well as microcline. Quartz-sillimanite nodules are absent in these layers, although sillimanite is present in the rock. Notice that quartz-sillimanite nodules reappear in pink, leucogranitic layers.
Postel (1952) and Whitney and Olmsted (1988) have described rocks in the northeastern Highlands that are similar to the leucogranitic gneiss. Whitney and Olmsted (1988) referred to these rocks as the Lyon Mountain Gneiss and interpreted them to be metamorphosed products of altered volcanic rocks and interlayered basin sediments. Chiarenzelli and McLelland (1990) suggest the alternative hypothesis that the interlayered lithologies may have been the result of "emplacement of sheet-like intrusions combined with high strain to yield tectonic layering." If the leucogranitic gneiss at Lyonsdale is properly correlated with Lyon Mountain Gneiss, then correctly establishing its origin is has regional consequence.
Undeformed pegmatites cross-cut all lithologies in this area. These demonstrate fluid flow at and immediately after the peak of granulite facies metamorphism and anatexis, or during late crystallization of granite intrusives. If the PT estimates of Florence et al. (1995) represent conditions of granulite facies metamorphism, then these dikes likely formed under thermal and stress conditions that permitted fractures at between 4 and 6 kbar pressure.

STOP 4. Agers Falls. Leucogranitic Gneiss
Extensive outcrops of leucogranitic gneiss are exposed along the Moose River in this vicinity. This is the same leucogranitic gneiss seen upriver at Lyonsdale, although here it appears more homogeneous. The rock is composed largely of quartz and perthitic feldspar, with minor amounts of iron oxide, biotite, sillimanite, and garnet. Through much of the outcrop the rock has an equigranular texture and is largely lacking a gneissic fabric. Numerous pegmatites and coarse, irregular granite veins transect the gneiss.
Scattered grains or cm long veins of sodic plagioclase are recognizable in thin sections from this exposure. Plagioclase is sometimes myrmekitic and commonly exhibits erose grain margins, suggestive of resorption. Locally, oxide grains surround plagioclase margins. Garnets are almandine-rich and show irregular grain margins. Some garnets are surrounded by small biotite grains. This texture suggests the retrograde reaction: Kfs + Grt + H2O = Bt + Sill.
Elongate and nodular quartz-sillimanite segregations are abundant throughout the outcrop. Two sets of nodules are recognizable: one is aligned with the NE trending regional structure and a second, cross-cutting set is oriented at about N20E. Discontinuous layering defined by these features sometimes exhibits swirl patterns. Hercynite spinel is commonly present in the quartz-sillimanite segregations.
The single zircon grain dated at 1031± 8 by Orrell and McLelland (1996) was collected at this outcrop. They describe the sample grain as clear and prismatic and the determined age is concordant. They consider the leucogranitic gneiss to be a late- to post-tectonic intrusion emplaced into relatively shallow crust. Essentially all vein features are viewed as having formed during a single hydrothermal event that accompanied and ultimately outlasted magmatic intrusion. In this interpretation, the quartz-sillimanite features must have formed at about the time of emplacement or closely afterwards, when the intrusion was still partly molten or at a temperature sufficiently high to permit ductile deformation.
Our alternative interpretation suggests that anatexis of in situ lithologies has produced the many magmatic features observable in the Port Leyden region and that the zircon age reported by Orrell and McLelland (1996) dates the formation of granite veins in this rock. The quartz-sillimanite±spinel veins and nodules are considered to be relict features that have been rotated into their current configuration. The zircon age obtained at this outcrop and mineral ages from nearby sites define the time of a widespread thermal event that coincided with the development of the pervasive fabric and compositional layering seen at the regional scale. It needs to be emphasized that other zircon grains from here and at Lyonsdale demonstrate that cores and overgrowths do not record the same ages. Abraded cores yielded ages up to 1119 Ma that Orrell and McLelland (1995) view as inherited from older granitic rock present in the magma source region. We consider it more likely that these older zircon ages reflect crystallization within this crustal horizon in response to an earlier episode of igneous or metamorphic activity.

Low outcrop of foliated, pink and gray felsic gneiss.

STOP 5. Goulds Mill, Quartz-rich Gneiss
Walk along a dirt road parallel to the river for approximately 50 m, then cut down through trees to broad exposures along the river banks. A quartz-rich unit is exposed beginning under the bridge immediately below the dam and continuing downstream for over 50 m. Locally, the rock appears to be conglomeratic. In other places, it gives the appearance of being composed of many parallel quartz veins. The rock displays a red to yellowish coloration indicating presence of hematite and iron hydroxides. Portions of the rock are tinted green, probably due to formation of chlorite. Weathering of sillimanite and/or feldspar has produced a white wispy streaking throughout the rock.
Downstream from the bridge sillimanite needles occur in a quartz-rich matrix. Individual grains are often aligned with the dominant fabric in the rock. Some coarse sillimanite also forms fan-shaped sprays. Some pegmatites and potassium feldspar-rich veins are aligned with the rock fabric. Also present are relatively undeformed pegmatite veins that cross-cut lithologic layering. These veins contain commonly potassium feldspar and sillimanite associated with quartz.
The protolith for this quartz-rich gneiss is uncertain. Possibilities include arkosic sandstone, hydrothermally altered felsic volcanics, or a hydrothermally altered felsic intrusion.
Possible brittle deformation features crop out on the right (northern) side downstream of the bridge. A continuous 1 to 5 m thick, apparently horizontal rock layer displays a crackled texture, in contrast to the massive character of the rock on either side. This layer cross-cuts the steeply NW dipping foliation. Cm scale patches of dark colored, cryptocrystalline material is present on outcrop surfaces and along small scale joints or shear planes. Currently, studies are underway to determine the composition of this material.

STOP 6. Port Leyden. Nelsonite Dike and Metapelitic Gneiss
This is the only described occurrence of titanium-rich nelsonite that does not occur within or in close proximity to large volumes of anorthositic rocks. Because of the strong field and experimental evidence supporting the formation of nelsonite by magmatic immiscibility during evolution of anorthosite suite rocks, we infer that anorthosite must have once been present in the western Adirondacks.
From the parking spot, cross to the south side of the road, walk over a rounded outcrop with quartz-sillimanite nodules and stringers, and down to the left. Continue walking for approximately 40 m to the southeast. Notice the low outcrops of foliated , light pink pelitic gneiss along the way. Abrupt transition to dark rocks delineates the contact with the nelsonite. Look at your compass to see the effect of standing over a body of rock that is about 40% magnetite! Good specimens of nelsonite material may be obtained from the waste pile on the south side of the mine shaft.
The nelsonite occurs as an elongate, nearly parallel sided body. Width is between 3 and 4 m and the dike is 30 m along strike at the surface. Contacts with the surrounding metapelitic gneiss are generally sharp, and locally parallel to foliation. At the northern end of the dike, the contacts truncate regional foliation defined by quartzofeldspathic layering. Detailed mapping and geophysical studies indicate that the nelsonite strikes north-south and has a near-vertical dip (Laforce et al., 1994). This attitude cross-cuts the N30E to N50E foliation in surrounding outcrops of gneiss.
The rock is equigranular and homogenous except towards the margins, where the relative abundance of pyrite increases shows a weak foliation. Garnet crystallizes interstitially to magnetite and apatite at the expense of ilmenite in this zone. A silicate-rich, 5 cm wide outer portion of the contact zone contains abundant plagioclase and is essentially sillimanite-free. A variety of textures indicate that plagioclase crystallized from a melt, including myrmekitic and symplectic plagioclase, and intergranular veins of plagioclase with associated magnetite. Anorthite content in plagioclase decreases with distance away from the dike margin. This suggests plagioclase grew in response to apatite breakdown and Ca flux outward from the dike. A model whole rock reaction for the contact zone is Mag + Ap + Hm + perthitic Kfs + Qtz + Sill = Grt + Mnz + titaniferous fluoro-Bt + PI-rich melt.
Florence (1997) proposed that halogen release associated with apatite breakdown lowered the solidus temperature, leading to partial melting. Did this melting event occur at the time of nelsonite intrusion or during later granulite facies metamorphism?

A4: STRATIGRAPHY, SEDIMENTOLOGY, AND GEOCHEMISTRY OF SENECA LAKE, NY
John D. Halfman
Department of Geoscience
Hobart and William Smith Colleges
Geneva, NY 14456, halfman@hws.edu
Donald L. Woodrow
Geneva, NY 14456, woodrow@hws.edu
Introduction
The purpose of this field trip is to look at the stratigraphy, sedimentology and Chloride geochemistry of Seneca Lake, and how they relate to deglaciation of the area. Our field trip (cruise), aboard the H-WS Explorer, will focus on representative sediments that have accumulated in the northern part of the lake since deglaciation which are typical of what is found in the other Finger Lakes. We will also highlight our ongoing efforts to collect and interpret high-resolution (2 - 12 kHz) seismic reflection profiles that image down to the glacial drift, and highlight Mike Wing's and Bill Ahrnbrak's recent understanding of the high chloride concentrations in Seneca Lake in comparison to the other Finger Lakes.
The Finger Lakes of central New York State consist of 11 elongated, north-south trending basins just south of Lake Ontario (Fig. 1). The basins are glacially scoured into the northern edge of the Appalachian Plateau (Coates, 1968. 1974). At Seneca Lake, the bedrock is primarily Devonian shales (Hamilton Group in the northern end of the lake), and lesser amounts of sandstones and carbonates that gently dip to the south-southwest. Silurian carbonates, shales and most importantly evaporites (mostly halite) are found below the Devonian section.
Seneca Lake is the largest (by volume) and deepest of the Finger Lakes (Fig. 1, Table 1). Only Cayuga Lake immediately to the east of Seneca is longer (61 km) and almost as deep (132 m). The other basins are smaller. ranging in length from 5 to 32 km and maximum water depth from 9 to 84 m. The present day lake is drained to the north-northeast through the Seneca River (New York State Cayuga-Seneca Barge Canal).
Cruise Log
This field trip starts and stops aboard the H-WS Explorer, and investigates the sediment character and water chemistry at selected locations (Fig. 1). We will look at the X-Star seismic stratigraphy of the sediments while underway between stations. Surface grab samples will be collected at the first two stops, and a short piston core at the third stop. We will also deploy the CTD and collect water samples to analyze at the third station as well. We will not concern ourselves with the lacustrine marls because they compose a small fraction of the sediments in the lake.

STATION 1. Shallow-Water Sandy Silts
Northern end of Seneca Lake
The lake floor is covered by sandy silts, with the coarsest sediments in the northwest margin of the lake. Shell and plant debris and the occasional ice-raft pebble make up the other minor components. We believe that these sediments must be derived from the reworking of and erosion of glacial drift. Sub-bottom images suggests that the sand forms a thin wedge of sediment above the pink clays and/or glacial drift. In many places the seismic images are attenuated by gas (biogenic methane?).
Suitable substrate (coarse materials) is covered by zebra mussels. Zebra mussels are an exotic species that has been recently (1991?) introduced in the lake. They are prolific filter feeders, filtering, on average, a few liters of water each day, extracting the plankton from the water column. One possible measure of their impact is historical Secchi disk data. Secchi disk are used to measure water clarity. To a first approximation, deeper Secchi disk depths correspond to less turbid water (e.g.. smaller plankton concentrations). Historical Secchi disk data have shown increasing Secchi depths over the past decade from a few meters to over 5 meters deep during the productive early summer months. This change is consistent with a decrease in chlorophyll concentrations. The exact reasons for the increase in water clarity (and decrease in chlorophyll) are not completely understood but the Secchi disk data suggests two hypotheses: that the burgeoning population of zebra mussels has become large enough to significantly reduce the standing crop of plankton, and/or an increase in the quality of sewage treatment by lake-shore residents has reduced the anthropogenic "nutrient loading" to the lake.

STATION 2. Proglacial Rythmites
Offshore of Belhurst Castle - Geneva Country Club
To the south of station 1, pink clays may be found below a thin veneer of postglacial material. The pink clays are well stratified, very cohesive and exhibit light red to pink colors. Single pebble to granule sized grains and ostracode shells are found widely scattered in the sequence. Well defined, parallel to subparallel reflectors characterize these clays in the sub-bottom profiles. The reflectors commonly outcrop onto the lake floor in shallow water. We interpret these clays as proglacial rythmites, typical of many proglacial lakes. It is unclear whether the couplet are annual events (varves). In the northern part of the lake, the pink clays are folded. The subbottom profiles suggest that the pink clays collapsed into ice-block holes in the underlying drift and/or the ice front experienced small readvances

STATION 3. Postglacial Muds & Lake Geochemistry
Offshore of Clarks Point
Moving south across exposures of glacial drift, the lake floor descends to a flatter floor typical of the deeper parts of the lake. At this location the sediments are very fine grained, black to gray, stratified and rich in organics. These sediments blanket the older materials with thicknesses up to 8 meters in the sub-bottom profiles. The sequence thins rapidly in water depths shallower than 60 meters and on the steep slopes on either side of the basin. They contain no shell material, a few coarse grains of silt and sulfide minerals. These muds are interpreted as the postglacial sediments deposited since deglaciation. The source is the suspension of fine material carried into the lake by streams and/or erosion of older, shallow-water bottom sediments by surface and internal waves.
We will also deploy the CTD and analyze the surface water for its chloride concentration to discuss the source of the "extra" Chloride to Seneca Lake.

A5: CAMBRO-ORDOVICIAN AND MODERN CARBONATE FACIES OF THE MOHAWK-HUDSON VALLEYS, NEW YORK
Gerald M. Friedman
Department of Geology, Brooklyn College and Graduate School of the City University of New York
Brooklyn, NY 11210
and Northeastern Science Foundation affiliated with Brooklyn College
Rensselaer Center of Applied Geology
15 Third Street
gmfriedman@juno.com
A shallow epeiric sea covered most of the North American continent. The shallow water limestones and dolostones that accumulated on this platform make up the Upper Cambrian, Hoyt and Galway formations and the Lower Ordovician Tribes Hill Formation that we will examine during this field trip. These carbonates were deposited in a roughly east-west belt at 20° S latitude. Quartz sand and dolomite were deposited in varying intermixed quantities (Galway Formation) with localized algal reefs (Hoyt Limestone) fringing the southern Adirondack shore. The limestones and dolostones of the Tribes Hill Formation were laid down along the eastern edge of the craton about 35 miles west of the shelf edge. As they were deposited close to the shelf edge they were affected by the tidal fluctuations of the adjacent waters of the deep ocean. To the east of the platform, deep-water shales and carbonate clast conglomerates accumulated on the continental slope and rise.
The end of Early Ordovician time is marked in the rock record by the Knox unconformity which exposed the broad continental shelf. In Middle Ordovician, the Taconic Orogeny was initiated followed by the Acadian Orogeny in the Devonian and the Alleghanian Orogeny in the Late Carboniferous with intervening periods of quiescence and carbonate deposition. The eastern edge of the craton was subjected to deformation during the Appalachian orogenies but in the region of the Mohawk Valley and Saratoga Springs of the present field trip, the Cambro-Ordovician platform strata are relatively undeformed. Figure 1 is a map showing the field-trip route marked with the location of the stops.

A5: CAMBRO-ORDOVICIAN AND MODERN CARBONATE FACIES OF THE MOHAWK-HUDSON VALLEYS, NEW YORK

STOP 1. Modern Nonmarine Limestones (Travertine)
This stop can also be reached through the Main Gate of the State Park on Route 9.
Route of Walk
Walk downhill into a small wood with picnic tables towards a stone building. Bicarbonate-charged, saline waters issue from a faucet at the side of the building and pass through a pipe below dirt road, re-emerging on the bank of Geyser Brook. Calcite precipitates on the steep slope of this bank, forming a terrace of travertine. The water is known as Orenda Spring and the terrace as Orenda Terrace; we will get a better view of this terrace later from below. A walk of a few hundred feet along dirt road leads to the Hayes Well Spring at bridge across Geyser Brook. A taste of this water is "rewarding" for its initial effect. Prolonged drinking is not recommended. From Hayes Well Spring you see the Island Spouter "Geyser" a few hundred feet upstream. The water of this fascinating "geyser" is from a well; the water spouts from a small orifice to a height of about 30 feet. This well was drilled about 80 years ago and the large cone of travertine has formed since.
Follow the path along Geyser Brook upstream. In the bank on the left is an exposure of Middle Ordovician Canajoharie Shale, an outer shelf to slope facies. On occasion, graptolites can be found in this shale. Continue to Orenda terrace to study travertine (Fig. 2). Note the rippled surface of the travertine and the brown iron-oxide coloration that occurs in streaks. Up to approximately 4 cm of calcite may precipitate annually at the foot of the terrace. Note "caves" and dripstone at far end of terrace and search for calcite-coated twigs and leaves or impressions of leaves in the travertine. Pisolites occur in abundance on the walkway between travertine cone and bank of brook. Spheroids and ooliths are likewise present (Schreiber et al. 1981).
Discussion
Although in the classical studies on limestones, including the authoritative work of Pia (1933), nonmarine limestones are accorded some measure of importance, the literature on modern carbonates neglects nonmarine carbonates. At this stop you'll see an excellent example of travertine, a rock which, like reef-rock, crystallizes in an initially stony condition.
The term travertine is derived from the Italian word, travertino, a corruption of tiburtino, "the stone of Tibur", which is a former name of the locality now called Tivoli (see Sanders and Friedman, 1967, p. 176). The type locality at Tivoli has been classically described by Lyell (1830, p. 207-210) and Cohn (1864). Some authors make a sharp distinction between the terms travertine, tufa and sinter: others use these terms synonymously (Pia, 1933; Gwinner, 1959, Sanders and Friedman, 1967, p. 176).
Travertine in Saratoga Spa Park gathers around the orifice of wells, on terraces from which water descends, or as a cone around a "geyser". Waters enriched in calcium bicarbonate issue from the subsurface, lose their carbon dioxide and insoluble calcite precipitates:
2HCO3- + Ca++ ↔ 2CaCO3 + CO2 + H2O
Twigs and leaves of beech, maple, and oak are preserved as they become coated with calcite or leaves form impressions in the travertine. The calcium bicarbonate-enriched waters originate nearly 1,000 feet below the surface in the underlying Cambrian-Ordovician limestones and dolostones, especially in the Gailor Dolomite. The waters are confined as in an artesian well beneath a thick cover of impervious Canajoharie Shale from which drilling recovers them. In the early years, the springs issued from natural crevices in the rocks, particularly from the prominent MacGregor fault. Later, pits were dug; the present wells flow through pipes set in bore holes.
The composition of the Saratoga mineral waters is unique among waters that precipitate travertine (Back et al. 1995; New York State Department of Health 1959; Young and Putnam 1979). Most waters that make travertine, especially those of the classical areas in Europe, drain areas of karst, and are of low salinity. By contrast, analyses of Saratoga waters give salinities that geologists classify as "brackish" (approximately 11%). Inspection of tables of analyses (e.g. Kemp, 1912) indicates the closeness of the composition of these waters to that of formation waters. This is especially true of the high concentration of NaCl. As the waters in the subsurface apparently dissolve limestones and dolostones, the concentration of the calcium, magnesium, and bicarbonate is higher than that of many formation waters. As in most formation waters, the sulfate content is low. Although the origin of the mineral waters is controversial (Hewitt, McClennan, and Nilsson, 1965), this controversy parallels that of the origin of formation waters. The mineral waters are probably formation waters whose salinity has been lowered as a result of mixing with meteoric water.
Newly precipitated calcite from the Orenda Springs gives a strontium isotopic value of 0.716429 (10) 87Sr/86Sr (± 2 S.D.)*, a continental crust signature (Fig. 3). If the waters were derived Upper Cambrian formation water, an Upper Cambrian seawater signature would have been obtained (see Fig. 3). According to Siegel (1996) carbon isotopic analyses likewise hint that the carbon may be expelled from the mantle
* The Sr analyses are normalized to 86Sr/88Sr = 0.11940.
Analyses of NBS 987 averaged 0.710241 (09) (n = 39) during the period of these analyses.
Errors on 87Sr/86Sr are given as 2 sigma (95%) in the last two digits.
The aquifer containing the mineralized water is under carbon-dioxide pore pressures which are estimated to be as high as five atmospheres that spew mineral water and gas several meters high (Siegel 1996). The acidity of the mineral waters depends on how much carbonic acid they contain; pH typically ranges from 5.5 to 6.5. The mineral springs discharge cold water (Sneeringer and Dunn. 1981).
The distribution of the wells, a total of about 200, is controlled by the MacGregor Fault and its subsidiary faults. The mineral waters always occur on the eastern (downthrown side) of the fault.
Puzzling is the observation that the precipitate is calcite rather than aragonite. Travertine in other places is usually calcite, but the composition of the responsible mineral waters shows depletion in magnesium. By contrast the mineral waters of Saratoga are enriched in magnesium and because the magnesium ion inhibits the formation of calcite, aragonite should result, but apparently does not. This question deserves study.

STOP 2. Lester Park
Domed Cyanobacterial Cabbage Heads: (Stromatolites).
This locality is the site of one of the finest domed microbial mats to be seen anywhere preserved in ancient rocks. On the east side of the road in Lester Park a glaciated surface exposes horizontal sections of the cabbage-shaped heads composed of vertically stacked, hemispherical algal layers (Fig. 4). These structures, known as Cryptozoons, have been classically described by James Hall (1847, 1883), Cushing and Ruedemann (1914), and Goldring (1938); an even earlier study drew attention to the presence of ooids as the first reported ooid occurrence in North America (Steele,1825). Interest in these rocks has been revived as they are useful environmental indicators (Logan, 1961; Fisher, 1965, Halley, 1971). The heads are composed of discrete club-shaped or columnar structures built of hemispheroidal microbial mats expanding upward from a base, although continued expansion may result in the fusion of neighboring colonies into a Collenia-type structure (Logan, Rezak and Ginsburg, 1964). The stromatolites are part of the Hoyt Limestone of Late Cambrian (Trempealeauan) age. Their intertidal origin has been inferred by (1) observations in the rocks, and (2) by analogy with similar modern microbial heads.
The evidence for deposition under tidal conditions for the Hoyt Limestone at Lester Park includes: (1) mud cracks, (2) flat-pebble conglomerate, (3) small channels, (4) cross-beds, (5) bird's eye structures, (6) syngenetic dolomite, and (7) stromatolites (for criteria on recognition of tidal limestones, see Friedman, et al., 1993). The analogy with modern environments relates to the occurrence of cabbage-shaped microbial heads in the intertidal zone of Shark Bay, western Australia, in which the height of the domes is controlled by the degree of turbulence (Logan, Rezak and Ginsburg, 1964; Hoffman, Logan and Gebelein, 1969). With increasing wave and current energy the height of the domes increases; the relief of the domes decreases landward towards quieter water conditions.
At Lester Park the heads which are circular in horizontal section range in diameter from one inch to three feet; many are compound heads. The size of the larger heads suggests that they formed in highly turbulent waters.
The line of depositional strike along which the domed microbial mats occur was probably where the waves were breaking as they came across the deeper ocean from the east and impinged on the shallow shelf.
Several petrographic observations in these rocks permit an analogy with modern microbial mats in hypersaline pools of the Red Sea Coast (Friedman and others, 1972; Friedman and others, 1985). Mat-forming cyanobacteria secrete radial ooids, oncolites, and grapestones which occur in these rocks; interlaminated calcite and dolomite which in part compose the stromatolites of the Hoyt Limestone correspond to alternating aragonite and high-magnesian calcite laminites which modern blue-green cyanobacteria secrete. In modern microbial mats the high-magnesian calcite laminites contain abundant organic matter in which magnesium has been concentrated to form a magnesium-organic complex. Between the magnesium concentration of the high-magnesian calcite and that of the organic matter sufficient magnesium exists in modern microbial laminites to form dolomite. Hence the observation in ancient stromatolites, such as observed in the Hoyt Limestone, that calcite and dolomite are interlaminated, with calcite probably forming at the expense of aragonite and dolomite forming from high-magnesian calcite.
Note ooids, skeletal fragments, and coarse quartz particles dispersed between the microbial heads.
The dominant microbial process active in the deposition of the Hoyt Formation appears to have been binding, as microbial boundstones comprise most of the beds observed in the outcrop. There is, however, an abundance of forms in the Hoyt Formation. Logan (1961) and Logan et al. (1964) have shown that morphologic variations in microbial mats are largely related to the amount of energy or turbulence of the depositional environment. The Hoyt lithologies indicate a wide range of energy conditions.
The Hoyt Formation shows only minor evidence of the destruction (or bioturbation) of microbial mats by organisms. Evidence of both microbial binding and precipitation of calcium carbonate is present in the Hoyt Formation, with binding the dominant process during deposition.
The lower part of the Lester Park section on the west side of the road provides the most complete sequence observed in the Hoyt Formation. A vertical sequence from Lithofacies 1 to Lithofacies 4 (Fig. 5) is represented. Depositional environments ranging from the lee of an oolite shoal upward to the lower intertidal zone are shown in the lower Lester Park section. The sequence seen here may have resulted from the lowering of sea level (regression) or from the depositional buildup of carbonates (progradation).
The ooids at Lester Park were the first to have been described from North America (Steele, 1825).
Figure 6 presents a cross-section of the hypothesized Hoyt depositional model. The following observations support the model: 1) the presence of well-developed ooids in the Hoyt lithologies suggests that the offshore energy barrier was an oolite shoal perhaps similar to that of the western Bahama Bank (Ball, 1967) or to that of Abu Dhabi on the Trucial coast (Kendall and Skipwith, 1969; Friedman 1995); 2) the presence in the Hoyt of reef-like development of high-relief columnar stromatolites immediately overlying cross-bedded oolite, and the dependence of stromatolite morphology on energy (Logan et al., 1964), suggest that the large microbial heads were restricted to high-energy areas surrounding tidal washover deltas; 3) the presence of coarse calcarenite infilling the heads. Storm surges may account for mixing of carbonate grains in the different lithofacies.
Students should prepare a map of the planed, glaciated surface beneath which the domed microbial mats are exposed. An ecologic zonation may be observed based on the density of microbial heads. This distribution should be shown on the map. Sketches should be prepared of the different kinds of heads: single and compound heads, large heads vs. small heads. Is it possible from this map to determine the depositional strike of the original geologic setting? Note the kinds of particles occurring between the heads. Lie flat on your belly and use a hand lens to advantage. Note the coarse quartz particles scattered among the carbonate particles.
The vertical section on the west side of the road reveals a ledge of ooids. Examine but do not destroy this ledge. Remember, as this article points out, you are treading on hallowed ground. These oolites were the first to have been described from North America, back in 1825.

Quartz-cobble conglomerate
Pass basal Paleozoic quartz-cobble conglomerate (a possible talus deposit) on weathered Pre-Cambrian gneiss 1/2 mi. east of Cymbal's Comers (NY 147).

STOP 3. Fort Hunter Quarry
Alight at slight bend in road and walk to Fort Hunter Quarry which is across a former railroad track (now a bicycle pass) close to Mohawk River. (Fort Hunter Quarry cannot be seen from road; another small quarry visible from road is approximately 0.1 mile farther east, but will not be visited on this trip).
Products of Tidal Environment: Stromatolites
Stromatolites in the Fort Hunter quarry consist almost entirely of dolomite in the form of irregularly bedded, finely laminated, undulating structures. The rocks in this quarry are part of the Tribes Hill Formation of earliest Ordovician (Fisher, 1954). The lithofacies of the Tribes Hill Formation have been studied in detail by Braun and Friedman (1969) within the stratigraphic framework established by Fisher (1954). Figure 8 is a columnar section showing the relationship of ten lithofacies to four members of the Tribes Hill Formation. At Fort Hunter we will study the lowermost two lithofacies of the Fort Johnson Member (see column at right (east) end of section. in fig. 8.
Two lithofacies are observed: (1) lithofacies 1, mottled feldspathic dolomite, and (2) lithofacies 2, laminated feldspathic dolomite. Lithofacies 1 is at the bottom of the quarry, and lithofacies 2 is approximately half way up.
Lithofacies 1
This facies occurs as thin dolostone beds, 2 cm to 25 cm but locally more than 50 cm thick, with a few thin interbeds of black argillaceous dolostone which are up to 5 cm thick. In the field, the dolomite shows gray-black mottling and in places bird's eye structures. In one sample, the infilling of the bird's eyes shows a black bituminous rim which may be anthraxolite. In the field, trace fossils are abundant, but fossils were not noted. Authigenic alkali feldspar (microcline) is ubiquitous throughout this lithofacies. The insoluble residue makes up 22 to 54% by weight of the sediment in samples studied with most of the residue composed of authigenic feldspar.
Lithofacies 2
This lithofacies is mineralogically identical to the previous facies but differs from it texturally and structurally in being irregularly bedded and in containing abundant undulating stromatolitic structures ("pseudo-ripples") (fig. 9) as well as disturbed and discontinuous laminae. In places there are a few thin interbeds of black argillaceous dolostone. The thickness of the laminites of this facies ranges from 1/2 mm to 2 or 3 mm; on freshly broken surfaces the color of the thinner laminae is black and that of the thicker ones is gray. The insoluble residue, for the most part composed of authigenic feldspar, constitutes between 35% and 67% by weight in samples studied. These two lithofacies which form the basal unit of the Ordovician, were formed on a broad shallow shelf.
Stromatolites, bird's eye structures, scarcity of fossils, bituminous material, syngenetic dolomite, and mottling suggest that these rocks were deposited in a peritidal environment (Friedman, et al1992). Based on analogy with the carbonate sediments in the modem Bahamas, Braun and Friedman (1969) concluded that these two lithofacies formed under supratidal conditions. However in the Persian Gulf flat microbial mats prefer the uppermost intertidal environment, and along the Red Sea coast they flourish where entirely immersed in seawater, provided hypersaline conditions keep away burrowers and grazers (Friedman and others, 1973). Hence on this field trip we may conclude that the stromatolites indicate peritidal conditions without distinguishing between intertidal and supratidal. For more details on these lithofacies refer to Braun and Friedman (1969).
Data obtained from the analyses of fluid inclusions in calcite-healed fractures of these Lower Ordovician carbonate strata, in which microcline crystals are so prominent, indicate higher paleotemperatures and greater depth of burial than have previously been inferred for the rocks of this region (Urschel and Friedman 1984; Friedman 1987a,b). Average fluid- homogenization temperatures range from 96° C to 159° C . These high paleotemperatures are supported by oxygen-isotope and conodont-alteration data (Harris et al. 1978). A former depth of burial > 7 km is implied when a geothermal gradient of 26° C/km (Friedman and Sanders 1982; 1983) is used.
The study of the authigenic alkali feldspar from this quarry yielded an age of uplift of approximately 320 Ma (Friedman 1990), Carboniferous in age. The observation that authigenic feldspar in Cambro-Ordovician carbonates occurs along the proto-Atlantic shelf from the Appalachian Basin to Newfoundland, Scotland, and Greenland paralleling the Taconic belt suggests active crustal epeirogeny in Pennsylvanian-Permian time.
Thus, following subsidence to great depth, Pennsylvanian to Permian epeirogeny uplifted the strata, resulting in deep erosion. This leads to the surprising conclusion that isostatic unroofing following uplift has stripped off thick sections of strata whose presence was previously unsuspected.

Cross original Erie Canal, built in 1822.
Amos Eaton surveyed this route at the request of Stephen Van Rensselaer (1764-1839); after this survey Amos Eaton (1776-1842) and Van Rensselaer decided to found a school for surveying,

Fulton-Montgomery Community College

STOP 4. North Tribes Hill Quarry
Route of Walk
Take the trail towards old abandoned crusher, but instead of heading towards the quarry move uphill to the first rock exposures. The rocks to be examined are near the edge of steep cliff.
Description and Discussion
At this stop microbial reef mounds are exposed (Friedman, 1996). Ordovician domal thrombolites, termed here microbial reef mounds, occupied the basal part of meter-scale shallowing-upward cycles (Fig. 10). They are part of a high-energy facies that a sharp transgressive surface separates from an underlying low-energy peritidal setting. This erosional surface served as the surface on which one of the reef mounds established itself during initial transgression before further deepening. The others overlie a floor of skeletal grainstone reflecting a high-stand sea-level facies tract. Skeletal grainstone composes the fill between the mounds. A channel and several aggrading hummocks occupy inter-reef mound areas resulting from storm events in a subtidal setting.
Reef mounds formed at or near the base of upward-shallowing parasequences. They are part of a high-energy facies which a parasequence surface of emergence or near emergence separates from an underlying parasequence which terminated a low-energy peritidal setting. This erosional surface between the two parasequences is interpreted as a partly lithified hard ground. As elsewhere in the Cambro-Ordovician of North America, microbial reef mounds commonly occur near the bases of upward-shallowing cycles and most rest directly on underlying cycle caps (Osleger and Montanez 1996).
Mound-Foundation Facies
Braun and Friedman (1969) designated the facies underlying the reef mounds as Lithofacies 7: Mottled Dolomitic Micrite and Biomicrite of the Wolf Hollow member (Fig. 8). This lithofacies is made up of a well-bedded, mottled limestone in which the mottles are composed of irregular patches of dolomite. On weathered surfaces, the limestone is whitish and the dolomite buff, but on freshly broken surfaces both limestone and dolomite are very light gray with the limestone somewhat darker and the dolomite of granular appearance. The outlines of large gastropods and cephalopods stand out, in places, on weathered bedding planes. A list of fossils found in this lithofacies was given by Fisher (1954, p. 88-89). Bird's eye structures are present in some beds as are pyrite patches. The limestone contains abundant dolomite-filled burrows, most of which are horizontal (or sub-horizontal) to the bedding plane, but some burrows have oblique to perpendicular orientations with respect to bedding. Many gastropods, especially Ophileta and Ecculiomphalous are found with the dolomite-filled burrows suggesting that these burrows may have been made by gastropods rather than by worms. However, the morphology of the shells suggests that these gastropods were not burrowers. Hence, worms or other soft-bodied organisms must have been abundant and produced the burrows.
This facies which is part of the Wolf Hollow Member (see Fig. 8), represents a low-energy deposit of micrite of a shallowing-upward cycle terminating in a sharp, planar erosion surface. This erosional parasequence surface is a marine flooding surface and represents a transgressive event for the next high-energy cycle in which the reef mounds formed. Below one of the reef mounds the underlying micrite of mound-foundation facies compacted (Fig. 11).
Differential compaction of this former lime mud, as the solid reef grew, suggests that the lime mud had not yet fully lithified (Friedman 1996). This observation differs somewhat from that of reef mounds in Virginia and Argentina where a solid hard ground served as foundation for the reef mounds (Read and Grover 1977; Canas and Carrera 1993). However, the original lime mud of the mound-foundation facies of the Tribes Hill mounds was sufficiently lithified to support the growing mounds (Friedman 1996).
Microbial Reef-Mound Facies
Reef mounds are prominent in the Wolf Hollow Member of the Tribes Hill Formation (Figs. 10,11). They occur as isolated mounds (Friedman 1996). These mounds are approximately one meter in length (measurements vary from 95 cm to 135 cm) and 60 to 70 cm in thickness. These measurements are approximate dimensions because the mounds do not stand out freely, and disappear in the enveloping facies. Moreover in the exposure, it is difficult to differentiate between short and long axes of the mounds.
The reef mounds are composed of clotted and peloidal microcrystalline matrix which was microbially precipitated. comparable to that in modern reefs (Friedman et al. 1974). In modern reef settings, peloids display the pattern of calcified coccoid cells which cyanobacteria or chemo-organo-trophic bacteria degrading the cyanobacterial organic matter precipitate (Friedman et al. 1985; Krumbein 1983). These textural features are products precipitated by the micro-environment of cyanobacteria (Nadson 1903; Kalkowsky 1908; Pia 1927; Johnson 1954; Endo 1961; Friedman et al. 1973). Peloids have been described as calcified algal filaments (Friedman et al. 1974); such calcification may be the result of precipitation of calcium carbonate on cyano-bacterial filaments in the presence of live bacteria (Chafetz and Buczynski 1992).
Because they are composed of microcrystalline carbonate matrix, the microbial reef mounds have a texture similar to that of the mound-foundation facies lithofacies 7 (see Fig. 8), a micrite. The term matrix has been commonly misapplied as a synonym of micrite, but the mounds are not composed of micrite. Mechanically deposited lime mud, following lithification, is known as micrite (Folk 1959). Identifying even modern reef rock is an experience in frustration: submarine microcrystalline or cryptocrystalline matrix that is biologically precipitated within millimeters to centimeters of the surfaces of reef rock is identical in appearance to micrite of mechanical origin. Case histories abound in which unwary geologists have misidentified the reef rock for low-energy facies composed of micrite (Friedman 1985, 1994). Since micrite of low-energy origin and microcrystalline or cryptocrystalline matrix of reefs are indistinguishable it is easy to confuse high-energy reef facies for low-energy lime- mud facies (Friedman 1985, 1994). This textural similarity led initially to an interpretation that mounds may be blocks of lithofacies 7 (micrite and biomicrite) that foundered and became lodged in channels (Braun and Friedman 1969). These reef mounds resemble blocks of micrite in tidal channels of the Bahamas that are derived by undercutting of the banks of the channels (Braun and Friedman 1969).
The reef mounds were included with lithofacies 8 of Braun and Friedman (1969) designated intrasparite and biosparite; the lithology is for the most part a skeletal grainstone. This lithofacies was referred to as channel fill, comparable to the Lower Ordovician reef mounds of western Argentina, of which Cafias and Carrera (1993, p. 169) noted "the reef mounds are dissected by conspicuous channels filled with coarse crinoidal grainstone and litho clastic rudstone" (Fig. 11). This same observation applies in part to the setting of the Tribes Hill Formation. As in western Argentina, the reef mounds of the Tribes Hill Formation formed in part on a previously lithified or partly lithified sediment surface (Cafias and Carrera 1993, p. 169); and as in other places in North America, they occur near the base of an upward-shallowing cycle. The surface on which one of the mounds developed is a sharp transgressive parasequence surface separating the underlying peritidal lithofacies 7 (micrite and biomicrite) from the overlying subtidal reef-mound facies. The other reef mounds nucleated near the transgressive parasequence surface, but on top of underlying skeletal grainstone, during the initial transgression before rapid deepening occurred, a setting which is similar to that of comparable facies in the Great Basin, U.S.A. (Osleger and Montanez 1996).
Inter-Reef Mound Facies
Skeletal grainstone composes the fill between the reef mounds. One channel and several hummocks occupy the inter-reef mound areas (Fig. 12). The top of one of the hummocks rolls into the channel fIll. The grainstones form lenses that build on top of one another. The channel displays the typical asymmetric profile of a tidal channel with a steep cut-bank and a low-angle slip-off slope (Fig. 12). However, the channel is entirely within grainstone, hence it is not a normal tidal channel which would be fine-grained on the steep side and coarse-grained on the opposing side. In normal tidal channels, as the channel shifts it leaves behind a layer of coarse debris at the bottom of the channel (Friedman, et al. 1992). No such channel-floor lag layer is present in the inter-reef mound facies. Hence the channel must be related to storm deposition of the grainstone since filling is not the result of the shifting of the channel. Truncation by the channel and aggradation of the hummocks occurred at the same time. The channel and hummocks formed as a result of storm events in a subtidal setting. Following transgression, storm tides and currents generated this channel between which reef mounds and inter-reef mound facies accumulated. This channel, which was incised down to 30 cm into the underlying grainstone, is conspicuous and displays sharp margins (Fig. 12).
Of the various reef mounds one rests directly on the mound-foundation surface; grainstone of inter-reef mound facies makes up the floor of all the others.
The reef mounds formed in shallow-subtidal to possibly low intertidal settings in an agitated environment devoid of lime mud.
This stop was very difficult to locate. We are told "Going east on Route 67, after the intersection with Route 147, five roadcuts on either side of Route 67 expose outcrops of the Lower Ordovician Gailor Formation." However, if we are east of 147 we are no longer in the Lower Ordovician. Furthermore, stop 8 is said to be near Waite Road, and stop 9, presumably further east, is at Manny's Corners. Yet Manny's Corners is definitely west of the intersection of 67 and Waite Road. My conclusion is that the author had confused field trips, and these stops were from a trip in which participants went clockwise, instead of counterclockwise, around the loop. Hence I locate stop 5 just west of the intersection with 147, and move to the west from there. Since the author gave no guidance on mileage, I guess on locations - a lot!

STOP 5. Exposures of Gailor Formation
On the north and south sides of the road are exposed a massive dolostone unit overlain by a bedded dolostone unit. Roughly 10' of section are exposed here. The massive unit is dark gray in color while the upper bedded unit is lighter gray and coarser as well. Large clasts in a variety of shapes and sizes, composed of micrite and medium-textured dolostones are scattered all over the outcrop and concentrated in the basal massive unit. Pods and lenses of chert colored black and white are profuse. Calcite mineralization is also a common feature, occurring in patches and veins in orange, white and black. Stromatolites are observed in the section on the north side of the road.
The breccia observed in the section perhaps represents dissolution collapse. The massive unit in the lower part of the section may be a microbial build up. The stromatolites suggest a peritidal origin for these dolostones.
This stop was very difficult to locate. We are told "Going east on Route 67, after the intersection with Route 147, five roadcuts on either side of Route 67 expose outcrops of the Lower Ordovician Gailor Formation." However, if we are east of 147 we are no longer in the Lower Ordovician. Furthermore, one stop is said to be near Waite Road, and the next stop, presumably further east, is at Manny's Corners. Yet Manny's Corners is definitely west of the intersection of 67 and Waite Road. My conclusion is that the author had confused field trips, and these stops were from a trip in which participants went clockwise, instead of counterclockwise, around the loop. This is the first roadcut I could find west of the intersection with 147.

STOP 6.
Farther east on the north and south side of 67 are exposed roughly 27' of section. The following features are observed:
-wavy beds,
-massive units alternating with bedded dolostone units,
-the massive units appear to have a lenticular mound-like form, perhaps representing former microbial buildups,
-the beds overlying the mounds display dips on flanks of the mounds perhaps indication fore-reef slopes,
-pervasive dolomitization seems to have obliterated original depositional features,
-alternatively the wavy bedding may represent tidal channels or hummocky cross stratification,
-other features observed in these units are fine laminae in the dolostones, burrow mottling, stylolites, presence of pyrite and white calcite mineralization in vugs and fractures.
This stop was very difficult to locate. We are told "Going east on Route 67, after the intersection with Route 147, five roadcuts on either side of Route 67 expose outcrops of the Lower Ordovician Gailor Formation." However, if we are east of 147 we are no longer in the Lower Ordovician. Furthermore, one stop is said to be near Waite Road, and the next stop, presumably further east, is at Manny's Corners. Yet Manny's Corners is definitely west of the intersection of 67 and Waite Road. My conclusion is that the author had confused field trips, and these stops were from a trip in which participants went clockwise, instead of counterclockwise, around the loop. Although there is another outcrop to the west of Stop 5, it is small, and only on the south side of the road. This is one looks 27' tall and is on both sides of the road.

STOP 7.
Farther east are exposed roughly 15' of bedded dolostones displaying planar stromatolites, intraclasts, and bird's eye structures indicating peritidal environments of deposition.
This stop was very difficult to locate. We are told "Going east on Route 67, after the intersection with Route 147, five roadcuts on either side of Route 67 expose outcrops of the Lower Ordovician Gailor Formation." However, if we are east of 147 we are no longer in the Lower Ordovician. Furthermore, one stop is said to be near Waite Road, and the next stop, presumably further east, is at Manny's Corners. Yet Manny's Corners is definitely west of the intersection of 67 and Waite Road. My conclusion is that the author had confused field trips, and these stops were from a trip in which participants went clockwise, instead of counterclockwise, around the loop. The author gives no mileage guidance

STOP 8. Small section on the north side of 67
This is a small section on the north side of 67 near Waite Road. The section displays intraclasts and stromatolites.

STOP 9.
Farther east on Route 67 near Manny Corners is a section exposing roughly 26' of bedded dolostones. The following features are observed here:
- planar and domal stromatolites,
- intraclasts ranging in size from 1 to 2 inches,
- vugs and fractures partially or completely filed with black and white calcite,
- bird's eye structures,
-bedded and nodular chert,
-breccia representing dissolution collapse?
The above mentioned features displayed in these fine- to medium-textured dolostones point to a peritidal environment of deposition.
This stop was very difficult to locate. We are told "Going east on Route 67, after the intersection with Route 147, five roadcuts on either side of Route 67 expose outcrops of the Lower Ordovician Gailor Formation." However, if we are east of 147 we are no longer in the Lower Ordovician. Furthermore, one stop is said to be near Waite Road, and the next stop, presumably further east, is at Manny's Corners. Yet Manny's Corners is definitely west of the intersection of 67 and Waite Road. My conclusion is that the author had confused field trips, and these stops were from a trip in which participants went clockwise, instead of counterclockwise, around the loop. The author gives no mileage guidance. This outcrop is close to Manny's Corners, but it doesn't look like 26' of section...

A6: GEOLOGIC FIELD GUIDE TO THE BALMAT ZINC MINE, ST. LAWRENCE COUNTY, NY
W.F. deLorraine and J. Johnson
Mining Division
Zinc Corporation of America Mines, Inc.
Balmat, NY 13609
Introduction and Geologic Setting
Orebodies of the Balmat-Edwards-Pierrepont Zinc mining district in the Northwest Adirondacks of New York State rank as world-class deposits with over 45 million tons of past production plus total reserves. Orebodies occur in "clean" Proterozoic siliceous dolomitic marble that was metamorphosed and poly-deformed during the Grenvillian Orogeny about 1.1 billion years ago. The NW Adirondack Lowlands are underlain by metasedimentary rocks, predominantly marble, calc-silicate and subordinate paragneiss and metavolcanic rocks. The Adirondack Highlands to the east are predominantly underlain by anorthosite-mangerite-charnockite- granite suite (AMCG) rocks metamorphosed to granulite facies. The two provinces are separated by the Carthage-Colton line or "Highlands-Lowlands" boundary, a narrow northeasterly-trending, northwest-dipping zone of mylonitization separating the upper amphibolite-facies Lowlands from the granulite-facies Highlands to the east. Flat-lying lower Paleozoic rocks surround and unconformably overlie the crystalline rocks of the Adirondack dome.

Zinc Mine
The presently known zinc orebodies of the Balmat mine lie within one mile of the central No.4 shaft, in an area of approximately two square miles. A majority of the ore shoots are arranged in clusters customarily known as the Balmat No.2, 3, and 4 mines, after the shafts which service them. In this guide, the ore clusters will be described as belonging to No.2, 3, and 4 mines, although these are more properly interconnected mining units now centralized around No.4 shaft, from which Balmat ore is hoisted to the adjacent mill (Fig. 15).

A7: GEOLOGY, LIMNOLOGY AND PALEOCLIMATOLOGY OF GREEN LAKES STATE PARK, NY
Martin F. Hilfinger IV and Henry T. Mullins
Department of Earth Sciences
Heroy Geology Laboratory
Syracuse University
Syracuse, New York 13244
Green Lakes State Park, located ~15 km east of Syracuse, New York near the village of Fayetteville, is a stunning geologic and limnologic wonderland (Fig. 1). Located along the axis of a narrow glacial meltwater channel, the Park consists of two small, moderately deep (maximum water depth of ~53 m) lake basins (Green Lake and Round Lake), protected along their flanks by steep walls of Silurian bedrock and separated from each other by a small modern wetland (Figs. 2 and 3).
These lakes are unusual, not only for their geologic setting, but also their water characteristics which have attracted the attention of limnologists for decades. In fact, limnologically, Green Lake is one of the best studied lakes in the world (Torgersen et al., 1981; Thompson et al., 1990). There are two aspects of the limnology of Green Lake which make it so interesting. Foremost is the fact that it is a meromictic lake (as is Round Lake) which means that when the water column turns over seasonally, it does so only above a strong chemically-controlled density boundary or chemocline at a water depth of ~18 m (Fig. 4). Because of this, bottom waters >18 m deep are permanently stratified and devoid of oxygen (anoxic). The second unusual limnologic feature of these lakes is that they annually precipitate large amounts of calcium carbonate (calcite), both from the open, wind-mixed surface waters of the lake (epilimnion) and along their shallow, near shore (littoral zone) margins in the form of "reefs" or bioherms.
The annual precipitation of calcite from the epilimnion coupled with the fact that deep bottom waters are anoxic, results in the deposition and preservation (no bioturbation) of annual sediment laminations or true varves. Because of this, Green Lake and Round Lakes have enormous potential from unraveling decadal to annual scale climate change (paleoclimatology) for at least the entire Holocene (past 10,000 years). With our current interest and concerns about possible global warming, the meromictic lakes of Green Lakes State Park may well provide a longterm, high-resolution template of natural climatic variability in which to view future environmental change. The purpose of this field trip is to acquaint you with the basic geology, limnology, and preliminary paleoclimatology of Green Lakes State Park. We will draw upon an extensive published literature to discuss: (1) the geologic origin of the lakes; (2) their fundamental limnology; (3) the annual, open water precipitation of calcium carbonate which will explain why "Green Lake is green"; and, (4) the origin of their "fringing reefs"; we will then draw upon new, unpublished data that will: (5) document the natural history of relative lake level fluctuations, and thus climate change, throughout the Holocene; and finally, (6) examine in high resolution, the natural and anthropogenic environmental changes which have occurred over the past 2,500 years.
This fieldtrip will be a leisurely half-day walking tour around both Green and Round Lakes. Scientific information will be provided via a series of large, colorful posters supplemented by on site examination of sediment cores. There will be ample time for discussion, and if the weather cooperates, we will be treated to a spectacular feast of fall foliage against the backdrop of aqua lakes; so, bring an open mind and your cameras!

A7: Geology, Limnology and Paleoclimatology of Green Lakes State Park, New York

STOP 1. Geologic Setting and Origin of the Lakes
Our discussion of the geology of Green Lakes State Park will take place on the swimming beach at the north end of Green Lake. From here we will have an unobstructed view to the south along the axis of the glacial meltwater channel in which Green and Round Lakes reside.
As we look to the south, steep bedrock walls can be seen rising from the surface of Green Lake at an elevation of 418 ft. (127 m) to at least 550 ft. (168 m) above sea level, although the highest summit (South Hill) adjacent to the park extends to an elevation of 788 ft. (240 m) (Fig. 2)(Muller, 1967). Total bedrock relief in the area, as measured from the top of South Hill to the bottom of Green Lake is 543 ft. (166 m). However this value should be considered a minimum because the thickness of sediments beneath the floor of Green Lake is not known, although Muller (1967) reported that drilling "just north" of Green Lake penetrated 138 ft. (42 m) of unconsolidated sediment without reaching bedrock.
The rocks that make up the walls of the meltwater channel in Green Lakes State Park are part of the Salina Group of Silurian (∼440 to 410 million years ago) age (Muller, 1967). The oldest stratigraphic unit in the Park is the Vernon Shale which consists largely of easily eroded red and green shale as well as gypsum (CaSO4·H2O). The top of the Vernon Shale extends upward to within ∼40 ft. (12 m) of the surface of Green Lake (Fig. 4; Muller, 1967). Overlying the Vernon Shale is more shale and argillaceous dolostone of the Syracuse Formation which contains evidence ("hopper casts") of evaporite minerals such as halite (NaCl) and gypsum. This middle unit extends upward to about the present day surface elevation of Green Lake (Fig. 4). The remainder of Silurian stratigraphy here consists of a relatively resistant dolostone member of the Syracuse Formation (Fig. 4) and the overlying Camillus Shale capped by a dolostone of the Bertie Formation which forms the summits of surrounding hills (Muller, 1967). The lake basins, however, rest within shales and evaporites of the Syracuse Formation and underlying Vernon Shale (Fig. 4).
Both Green Lake and Round Lake are located in an east-northeast "cross channel" (Fairchild, 1909; Sissons, 1960) which cuts obliquely across the Onondaga Escarpment at the northern edge of the Appalachian Plateau (Fig. 1 and 2). Today, this cross channel is drained by a "markedly underfit stream" (Muller, 1967) and near its southwestern end there is an abrupt topographic drop of ∼100 ft. (30.5 m) from a dolostone member of the Syracuse Formation to the surface of Round Lake (Fig. 2).
As noted by Muller (1967), Green Lake (which was originally known as Lake Sodom) was first depth sounded by Vanuxem (1839) and subsequently by Miner (1933), who also sounded (via lead-line) Round Lake. Both lakes consist of steep-sided, flat-floored, semi-circular basins with maximum water depths slightly >50 m. However, Green Lake also contains an elongated "neck" or "handle" which shoals to the north (Fig. 2). As also noted by Muller (1967), the surface elevations of Green Lake (418 ft; 127 m) and Round Lake (421 ft.; 128 m) are significantly (∼6 m) below the highest stand of pro-glacial Lake Iroquois which existed from ∼12.5 to 11.4 ka 14C (Muller and Prest, 1985; Anderson and Lewis, 1985). The surface of Green Lake is also ∼6 ft. (2 m) lower in elevation then the water surface of the nearby Erie Canal (Fig. 2).
Green Lakes State Park is located north of the Valley Heads Moraine which was deposited ∼14.4 ka 14C (Muller and Calkin, 1993) during a brief phase of Laurentide ice sheet instability coincident with Heinrich event H-l (Mullins et al., 1996). Thus, modern Green Lake and Round Lake must be less than 14,400 14C years in age (Fig. 5). The meltwater channel in which both lakes reside connects in the village of Fayetteville with Pompey Hollow, a large glacial trough whose origin is likely linked with the nearby Finger Lakes (Mullins et al., 1996). The Green Lakes meltwater channel is but one of many east-west orientated glacial drainages in the Syracuse-Oneida region (Sissons, 1960). The origin of the Green Lake and Round Lake basins is still uncertain although a number of hypotheses have been proposed (Muller, 1967), including: (1) plunge-pools; (2) solution sinkholes; and, (3) subglacial meltwater "potholes". Native American legend of "Kayahao" originally suggested that the lakes were volcanic in origin although there is no evidence to support it. This legend speaks of a Native American woman who infuriated the spirit of the lake which then " ... spit forth fire and flame, (and) caused a great commotion of the waters ... " (Gibson, 1925, p. 21).
The most widely accepted hypothesis is that these lake basins were scoured out at the base of large, late-glacial waterfalls. Based on stratigraphic relationships and morphology, Miner (1933) developed the hypothesis that the lakes are glacial plunge-pool basins (Fig. 6) which followed Vanuxem's (1839) earlier suggestion that they were eroded by a "whirlpool of great magnitude". However, problems with this hypothesis have centered around the source of water for the waterfalls plus the fact that there are two ("twin") lake basins.
Fairchild (1899) suggested that water impounded in high-standing proglacial lakes to the south of the Park may have cut the Green Lakes meltwater channel, whereas Sissons (1960) argued that water flowed from the Great Lakes in a complex system of marginal, subglacial and englacial conduits. If Fairchild's (1899) view of a proglacial source of water is correct, then Pompey Hollow, which connects with the Green Lakes meltwater channel at an elevation of ∼530 ft. (162 m), is a likely source. Pompey Hollow, a 1 km wide, flat-floored (∼780 ft., 238 m) trough, with as much as 520 ft. (159 m) of relief, extends south of the Park for ∼20 km (Fig. 1). As ice receded from the Valley Heads Moraine at Delphi Falls shortly after 14.4 ka 14C proglacial lake waters would likely have been impounded in Pompey Hollow, dammed by the Valley Heads Moraine to the south and the Laurentide ice sheet to the north (Fig. 5). In the nearby Finger Lakes, high-standing proglacial lakes existed until -13.9 ka 14C when the ice sheet retreated north of the Onondaga Escarpment opening up lower elevation outlets to the east which resulted in catastrophic drops of proglacial lake levels, cutting of gorges (glens), and a drainage reversal (Mullins et al., 1996). Muller (1967) has argued that the Green Lakes Channel originated as a pre-late Wisconsin feature. If correct, then as the northward retreating Laurentide Ice Sheet uncovered the precursor Green Lakes Channel impounded proglacial lake waters in Pompey Hollow (hundreds of feet higher in elevation) may have catastrophically poured through the Green Lakes meltwater channel (Fig. 5). Today, unlike the Finger Lakes, Pompey Hollow is a dry, flat-floored lake valley.
Such a scenario may have been similar to that documented for the Syracuse meltwater channels which experienced catastrophic floods with waterfalls the size of Niagara Falls when a 600 ft. (183 m) deep proglacial lake in the Onondaga-Tully trough drained to the east (Hand and Muller, 1972; Hand, 1978) creating a plunge-pool basin at Clark Reservation. Hand (1978) estimated that such catastrophic floods may have had water discharge as high as 7,000,000 ft.3/sec. (∼2.1x106 m3/sec.) and flow velocities exceeding 50 ft/sec. (∼15 m/sec.)! Similar catastrophic floods from Pompey Hollow, flowing over relatively resistant dolostone caprock as a waterfall in Green Lakes Channel may have eroded underlying shales to produce plunge pool basins which today we know as Green and Round Lakes.
The second problem which must be overcome for the plunge pool hypothesis to be viable, is the origin of "twin basins" which has previously been discussed by Muller (1967). Because simple headward erosion of a waterfall should result in a single gorge, Fairchild (1909) proposed a simple solution; that the Green and Round Lake basins do occupy a single bedrock trough but have been isolated by post-glacial sediment accumulation (Fig. 6). If correct, then the depth to bedrock beneath the wetland between Green and Round Lakes should be similar to that beneath the basins (i.e. >50 m). However, we have collected sediment cores only up to 11 m in length from this wetland (see Stop #4) which represent ∼9,100 14C years. Also, one of our cores from the center of the wetland penetrated only 4.8 m and appeared to bottom on bedrock. Thus, there may be a bedrock high or sill separating Green and Round Lakes. A land-based seismic reflection/refraction profile along the axis of this wetland would likely resolve the question of whether or not the lakes are separated by a bedrock sill.
A second possibility is that the lake basins formed sequentially, with Green Lake having formed first. This would necessitate a phase of rapid waterfall retreat to form the elongated "neck" of Green Lake followed by a pause in the rate of retreat to form the main basin of Green Lake. A second phase of rapid waterfall retreat would then be required followed by a second pause to excavate the Round Lake basin. Such a scenario could be reconciled if the outpouring of pro glacial lake waters from Pompey Hollow occurred in the two discrete phases, perhaps as a result of glacial oscillation (Muller, 1967).
A third hypothesis, advocated by Muller, 1967), is that there were twin waterfalls which migrated headward simultaneously. For this mechanism to work there must have been some stratigraphic (or structural) control to initiate twin falls during a single meltwater episode. Muller (1967) suggests that both basins were controlled by water falling over dolostone caprocks; Green Lake basin by the "Middle Dolomite" of the Syracuse Formation and Round Lake basin by the "Upper Dolomite". The fact that there is no headwall scarp at the southwest end of Green Lake (as there is at Round Lake) would have to be explained by either incision and erosion of the lower caprock or by "burial" of the scarp by subsequent impounding of waters in Green Lake (Muller, 1967).
It is apparent from the above discussion that the origin of Green and Round Lakes continues to be uncertain. However, morphologically, they do appear to be plunge-pool basins but satisfactory explanations for the source of the water and why there are "twin basins" must be acceptable. Most useful would be the acquisition of subsurface data, both geophysical and drill core.

STOP 2. Limnologic Setting of Green Lake
From our vantage point at Stop 1 we will progress clockwise and walk a few hundred meters along the lake path to an open area just south of the diving platform. Here we will continue to have an open view of Green Lake and have an opportunity to present and discuss it's basic limnology (characteristics of the water column). Limnologically Green Lake is probably the best studied meromictic lakes in the world (Thompson et al., 1990); thus, there are considerable data from which to draw upon.
According to Thompson et al. (1990), Clark (1849) was the first to report (quoting L.C. Beck) that there was something unusual about the limnology of Green Lake: "Water drawn from a depth of one hundred and sixty-eight feet (51 m) was found to be strongly charged with sulphurated hydrogen ... (which) blackened silver powerfully, and gave copious precipitates ... ". In 1931 Eggleton also noted the "foul odor of deeper mud deposits" and speculated that Green Lake might be stagnate, characterized by "incomplete spring and fall overturns" and perhaps a complete lack of whole-lake circulation. Eggleton's (1931) account of Green Lake is believed to be the first of a meromictic lake anywhere in North America (Thompson et al., 1990). Following his early assessment, Eggleton (1956) also reported results of a 25-year study of Green Lake concluding that "for all practical purposes, there is no significant amount of dissolved oxygen in the water below 19-20 m at any time of the year" (Thompson et al., 1990).
More recent studies have confirmed Eggleton's pioneering work on Green Lake, while enhancing our understanding of why Green Lake is meromictic. Of particular importance are publications by Deevey et al. (1963), Turano and Rand (1967), Takahashi et al. (1968), Brunskill and Ludlam (1969) and Torgensen et al. (1981). Based on a study of sulfur and carbon isotopes in Green Lake, Deevey et al. (1963) noted unusual quantities of sulphate which they suggested were derived from nearby outcrops of Silurian gypsum entering the lake as surface waters and perhaps "gypsiferous springs". Deevey et al. (1963) concluded that geochemically, Green Lake is more akin to Gulf Coast salt dome systems than to most temperate lakes.
In 1965, Turano and Rand (1967) conducted a series of measurements on the temperature and water chemistry of Green Lake from late winter (February 25) to early spring (April 18). They found that waters of the mixolimnion are colder (3-5°C) than waters below the chemocline (∼7°C). However, surface water temperature of Green Lake varies greatly (3-22°C) as a function of season whereas the monimolimnion maintains a relatively constant temperature of ∼7°C (Fig. 7; Brunskill and Ludlam, 1969).
Turano and Rand (1967) also found that dissolved oxygen concentrations are highest in the mixolimnion >e;10 ppm) of the lake, decrease abruptly between 16 and 19 m and then remain relatively stable to the bottom of the lake (Fig. 7), consistent with previous results. However, unlike Eggleton (1956) who reported that the monimolimnion of Green Lake was completely devoid of oxygen, Turano and Rand (1967) did detect very low levels of oxygen (∼1 ppm). Based on these data they concluded that the monimolimnion of Green Lake is not completely stagnant but rather must have some exchange with other waters, perhaps groundwater. However, Turano and Rand (1967) are the only investigations who have reported some oxygen below the chemocline.
Turano and Rand (1967) also found that both sulfate (oxidized) and sulphide (reduced) increase below the chemocline (Fig. 8) but that only sulphate is present above the chemocline. In the monimolimnion, sulphide concentrations are high (up to 50 ppm) representing a "very considerable quantity" of H2S (Turano and Rand, 1967). Maximum concentrations of sulphide occur near the bottom of the lake suggesting a source via the degradation of organic matter in the sediment primarily by sulfur reducing bacteria (Turano and Rand, 1967). However, sulphide is also oxidized by purple, photosynthetic bacteria that inhabit the chemocline in Green Lake (Culver and Brunskill, 1969). The abundance of sulphide in the monimolimnion and absence in the mixolimnion, of Green Lake, though, is directly related to oxygen concentrations.
Carbon dioxide and pH values in Green Lake are mirror images of each other (Fig. 9): CO2 concentrations are relatively low in the mixolimnion and high in the monimolimnion whereas just the opposite is true for pH (Turano and Rand, 1967). CO2 concentrations are controlled largely by photosynthesis above the chemocline and decomposition of organic matter below the chemocline. pH (a negative measure of hydrogen ion concentration) is inversely controlled by CO2 concentrations. However, pH values are neutral or above (∼7.2-8.0) throughout the entire water column of Green Lake (Fig. 9).
Turano and Rand (1967) also used data on water hardness (Ca, Mg concentrations), alkalinity, pH and temperature to calculate the saturation state of Green Lake with respect to calcium carbonate, both above and below the chemocline. They found that both water layers of the lake are "considerably supersaturated" with respect to calcium carbonate which should promote the precipitation of large quantities of marl.
In 1968 Takahashi et al. published their geochemical study of Green Lake. Although Green Lake had traditionally been believed to be meromictic because of "its great depth ... and small area ... " (Deevey et al., 1963), Takahashi et al. (1968) concluded that Green Lake's meromixis is a result of two different water types entering the lake. Below ∼18 m groundwater entering the basin must be nearly twice as saline as surface waters, regardless of temperature (Takahashi et al., 1968). They found that surface waters are 4x10-4 g/cm3 less dense than deep water. Thus, the permanent stratification of Green Lakes can be fully explained by salinity differences, which is contrary to most lakes which stratify due to temperature-controlled differences in density. This is verified by conductivity (a gross measure of salinity) measurements in Green Lake which display an abrupt increase at ∼18 m water depth (Fig. 10; Brunskill and Ludlam, 1969). Thus, the meromixis of Green Lake is the consequence of a chemically-controlled density boundary or chemocline. However, the salinity of Green Lake is controlled not by Na and Cl as in the oceans, but rather by Ca and S04 derived from gypsum in surrounding bedrock (Takahashi et al., 1968).
That the chemocline in Green Lake is the result of the inflow of saline groundwater is further supported by chemical studies of interstitial waters recovered from bottom sediments. Brunskill and Harriss (1969) found that interstitial waters in a core from the deep (>50 m) central basin of Green Lake show continuous increases of Na, Mg, Sr, Ca, Cl and Br with depth indicating the mixing of monimolimnetic waters with more saline groundwater diffusing upwards into the bottom sediment.
Takahashi et al. (1968) also estimated the residence time (the amount of time a water molecule resides in the lake before being removed) of waters in Green Lake. It had previously been known that the surface outflow of Green Lake exceeds the surface inflow by about an order of magnitude (Turano and Rand, 1967); thus, there must be appreciable inflow of groundwaters. Because of this, plus the meromictic nature of the lake, considerable differences in residence times might be expected. Takahashi et al. (1968) indicated a residence time on the order of only 2 years for surface waters whereas deep waters are replaced no less than once every 35 years. Brunskill and Ludlam (1969) concurred with the short surface water residence times estimated by Takahashi et al. (1968), and further argued that surface inflow supplies <50% of water flow into Green Lake. Based on 3H-3He water mass ages in Green Lake, Torgensen et al. (1981) have confirmed the short residence times of the mixolimnion. However, they concluded that the residence time of water in the monimolimnion is considerably less than previously reported, ranging up to only ∼7 years. Torgensen et al. (1981) also present evidence for a secondary chemocline at ∼33m where there is a significant increase in the concentration of S, CH4, CO2, Na, and Cl. They explain both the short residence time of the monimolimnion and the secondary chemocline as a consequence of diffuse groundwater input.

STOP 3. Precipitation of Calcium Carbonate
From Stop 2, near the dive platform of Green Lake, we will continue our clockwise walk around the lake to Deadman's Point, located at the eastern junction of the "neck" and main basin of Green Lake. Here we will first discuss the seasonal, open water precipitation of calcium carbonate ("whitings") from Green Lake, followed by a discussion of littoral zone precipitation and the formation of the prominent "reef" or bioherm that we will see at Deadman's Point.
Whiting Events
One of the most commonly asked questions about Green Lake is "Why is it green"? Well, the truth is that the lake is not always green. During winter (when not frozen) and early spring months, open waters of the lake tend to be blue, whereas they become green from late spring through fall, when most people visit the Park.
The color of water in a lake varies largely as a function of the amount and type of particulate matter in the water column. Oligotrophic (low nutrient) lakes with little particulate matter, like Skaneateles Lake, tend to display a blue color, whereas those lakes with high concentrations of particulate matter from primary biological productivity, like Oneida Lake, tend to be green, especially during the summer. The reason for this is that clear water scatters light predominantly from the blue portion of the electromagnetic spectrum, whereas when the amount of particulate matter increases in the water column longer wavelength radiation is preferentially scattered, and the lake appears green (Wetzel, 1975).
Green Lake is an oligotrophic lake (Thompson et al., 1997) with annual primary productivity estimated at ∼290 g C/m2 (Culver and Brunskill, 1969). However, ∼83% of this annual primary production in Green Lake occurs within the primary chemocline (∼18 m water depth) due to the photosynthetic activity of purple sulfide oxidizing bacteria. Obviously, an alternative explanation for the seasonal green color of Green Lake is required. A hint comes from Wetzel's (1975) statement that lake sediments rich in light-colored calcium carbonate (marl) reflect more light than non-marl sediments, and that lakes with large amounts of suspended calcium carbonate backscatter energy from the green spectrum of visible light.
The abundance of calcium carbonate deposits in Green Lake has long been known (Clark, 1849; Bradley, 1929, 1963; Howe, 1932; Eggleton, 1956; Brunskill, 1969). Takahashi et al. (1968) as well as Brunskill (1969) documented that the entire water column of Green Lake is supersaturated with respect to calcium carbonate. However, based on δ13C data, Takahashi et al. (1968) concluded that calcite precipitates only from the surface waters of Green Lake. They showed (via data from Deevey et al., 1963) that calcite crystals from the sediment of Green Lake have a δ13C (a measure of the ratio of 13C/12C) value of -4 ‰, which is much more similar to that of surface waters (-7 ‰) than to waters below the chemocline (-18 ‰). By independent means Brunskill (1969) also showed that a 3-4 fold increase in calcite supersaturation occurs in the mixolimnion of Green Lake from May through August, which is the time of massive precipitation of calcite at rates as high as 2 g/m2/day. From sediment trap studies, Brunskill (1969) found calcite crystal loads in the water column of 35 g/m2 in June and July of 1967 (Fig. 11). He argued that the precipitation of calcite in Green Lake is an inorganic process controlled largely by seasonal temperature change- " ... temperature is the direct causal factor in the initiation of calcite precipitation ... with photosynthesis playing only a secondary, and probably minor, role" (Brunskill, 1969, p. 844).
However, this concept of a purely inorganic origin for calcite precipitation in Green Lake has recently been challenged by Thompson and Ferris (1990) and Thompson et al. (1997). Based on experimental and transmission electron microscopy results of waters collected from Green Lake, Thompson and Ferris (1990) concluded that calcite precipitation is the direct result of the photosynthetic activity of the cyanobacterial pico plankton (0.2-2.0 µm) Synechococcus which occurs at concentrations of ∼105 cells/ml of water in the mixolimnion of Green Lake. They convincingly demonstrated that calcite (as well as a gypsum and perhaps magnesite) precipitates epicelluarly around Synechococcus cells (Fig. 12). The reason for this is that the photosynthetic metabolism of Synechococcus directly results in the alkalinization (increase pH) of the microenvironment immediately surrounding the cell because of its ability to use bicarbonate (HCO3-) as a carbon source which results in the production of hydroxyl (OH-) ions (Thompson and Ferris, 1990). Calcium (Ca2+) ions which concentrate on the surface of Synechococcus cells serve as nucleation sites for the precipitation of calcite and gypsum (Fig. 13).
This new view of Green Lake "whiting events" (direct precipitation of fine-grained calcite from the open water column) was further tested by direct monthly field measurements from May 1989 to April 1990 (Thompson et al., 1997). Secchi disk data (Fig. 14) clearly document the onset of the annual Green Lake whiting in May, when water clarity diminishes dramatically (Secchi disk readings decrease from 18 to 8 m). Minimum water clarity (Secchi disk readings of 4.5 m) occurred during July and August and then gradually improved through fall, winter and early spring (Thompson et al., 1997). During this whiting event the color of Green Lake changed from bluish-green to greenish-white (Thompson et al., 1997).
The onset of this whiting event was directly related to the spring bloom of Synechococcus triggered by increasing temperatures and light intensities (Fig. 15; Thompson et al., 1997). That Synechococcus is directly responsible for the precipitation of calcite is supported by δ13C data which reveal that whiting calcite (as well as sediment calcite) in Green Lake is enriched by -3-5 %0 relative to the δ13C of dissolved inorganic carbon (DIC) of summer surface waters (Thompson et al., 1997). The reason for this isotopic enrichment is the preferential uptake of relatively light 12C by the photosynthetic picoplankton (Fig. 16). According to Thompson et al. (1997) individual cells of Synechococcus can precipitate calcite crystals a number of times per season because they periodically shed their outer "S-layer" once it is mineralized.
To summarize, the reason that Green Lake is green (particularly during spring and summer) is because of the widespread precipitation of calcium carbonate. Calcite crystals precipitated around picoplankton result in an increase of particulate matter in the water column which preferentially scatters back green light. In addition, the accumulation of calcium carbonate in littoral zone waters of Green Lake provides a permanent surface of reflection for wavelengths within the green spectrum of visible light.
Bioherms
Here at Stop #3 we can literally stand on a massive, in situ accumulation of calcium carbonate or bioherm. Although the best developed bioherms in Green Lake are here at Deadman's Point, they are also found across the lake along the south-facing shoreline (Fig. 17). These unusual littoral zone accumulations of massive carbonate in Green Lake have been known for many decades (Walcott, 1914; Bradley, 1929; Howe, 1932; Eggleton, 1956). The bioherms of Green Lake have traditionally been viewed as algal in origin, a concept most recently advocated by Dean and Eggleston (1975), Eggleston and Dean (1976), Dean (1981), and Dean and Fouch (1983). In fact, Dean and Eggleston (1975) compared these freshwater bioherms of Green Lake to marine algal "cup" reefs of Bermuda. However, Thompson et al. (1990) have reinterpreted the "algal bioherms" of Green Lake as thrombolitic microbialites which simply means a massive carbonate accumulation precipitated by microbes (rather than algae). SCUBA observations here at Deadman's Point revealed that the microbialites have grown up and out from dolostone beds within the Syracuse Formation. They are lobate features that protrude as overhangs into Green Lake and are as much as 10 m thick (Fig. 18). The undersides of the overhangs are colonized by aquatic mosses as well as sponges, and have laminated stromatolitic caps (Thompson et al., 1990). The internal structure of these features is characterized by a non-laminar, clotted texture. The active, outer growing portion of these bioherms is heavily colonized by a benthic variety of Synechococcus (Thompson et al., 1990). They believe that this benthic species of cyanobacteria, rather than algae, is responsible for the in situ development of the bioherms in Green Lake. Synechococcus is specifically adapted for life in oligotrophic, hard water lakes like Green Lake because of its ability to directly use bicarbonate (HCO3-) as a carbon source and its very low nutrient requirements (Thompson et al., 1990).
Benthic forms of Synechococcus colonize most any hard surface in the lake as evidenced by the accumulation of calcite on human artifacts such as bottles and cans (Fig. 19; Dean and Fouch, 1983). It is also common to see a white calcite coating on trees that have fallen into the lake. However, if one assumes that the bioherms at Deadman's Point originated shortly after Green Lake formed (∼11-12 ka) their growth rate would only be ∼1 mm/year (Thompson et al., 1990).
Before leaving Deadman's Point for our next stop, please take a few minutes to examine the marl deposit which extends to ∼1 m above present day lake level. This calcite-rich deposit contains an abundance of lacustrine, littoral zone gastropods, thus providing clear evidence for a previous highstand of Green Lake (Eggleton, 1956). Harmon (1970) assumed that the level of Green Lake dropped when the Erie Canal (which is located just north of Green Lake; Fig. 1) was built in the early 1800's. However, as we will see at our next stop, there is a well preserved record of natural lake level fluctuations beneath the wetland which connects Green and Round Lakes.

STOP 4. Wetland Stratigraphy and Natural Lake Level Fluctuations
From Deadman's Point we will continue our clockwise journey to the southern end of Green Lake. Here the path will divide and we will bear to the left. This will take us along the intervening wetland between the two lake basins(Figs. 2 and 3). We will stop at the next juncture of the path at the east end of Round Lake to present the subsurface stratigraphy of the wetland and discuss its implications for natural lake level fluctuations as well as regional paleoclimate change over the past -8,000 years. Sediment cores previously recovered from the wetland will be available for viewing and discussion at this stop.
During the summer of 1996 we collected sediment cores, up to 11.2 m long, from three sites in the wetland between Green and Round Lakes (Fig. 20). There is at least 11 m of sediment beneath both the eastern and western ends of the wetland (Fig. 20). However, at the middle site only ∼5 m of sediment was recovered before terminating on what appeared to be bedrock. If correct, there is a very shallow bedrock sill separating the Green and Round Lake basins.
Most of the sediment recovered at all three sites is fossiliferous marl consisting of 60-90% calcium carbonate (Figs. 21 and 22). The abundance of lacustrine fossils indicates deposition in relatively shallow, oxic waters (i.e.littoral zone). However, a number of organic-rich (up to 90%) peat layers, representing various wetland environments, were also recovered and radiocarbon dated (Fig. 21). Based on radiocarbon data, the cores extend back to ∼9.1 ka 14C. Average sediment accumulation rates for the long cores are on the order of ∼1.2 m / 1000 years, although intervals of marl have accumulated at rates of up to 2.6 m / 1000 years (Fig. 23).
Of particular interest is the series of marl-peat cycles (N=4) recovered at core site GL WL-3 nearest to Green Lake (Fig. 22). The four peat layers have been radiocarbon dated (University of Texas Radiocarbon Lab #'s TX-9191, TX-9136 through TX-9133), from bottom to top, at 7,570±70 years, 5,620±70 years, 4,545±65 years, and 2,169±;50 years. Calculation of an average period for these cycles yields an average of ∼1800 years (range 1075 to 2385).
The marls obviously represent open lacustrine connections between Green and Round Lakes whereas the peats represent wetland environments similar to those that exist today. The modern surface peat has developed during the past -∼2,000 years. Thus, it seems highly unlikely that the marl outcrop that we saw ∼1 m above modern lake level at Deadman's Point was stranded by an anthropogenic lowering of lake level less than 200 years ago during construction of the Erie Canal. This is supported by tree-ring counts from the wetland which indicate that living trees predate the construction of the Erie Canal (Brunskill and Ludlam, 1967).
The marl-peat cycles most likely record natural fluctuations of relative lake level approximately every 2,000 years over the past 8,000 14C years. If we interpret such natural lake level fluctuations (Fig. 24) as products of regional changes in precipitation minus evaporation, the data further imply relatively dry climatic conditions at ∼8 ka, 6 ka, 4 ka, and 2 ka with relatively wet climates in between. This is consistent with results from a recent series of cores collected from the Montezuma wetlands (north end of Cayuga Lake) which also revealed relatively dry events at ∼8 ka, 6 ka, and 4 ka (Mullins, in prep.). Thus, the relative lake level fluctuations recorded from the wetland in Green Lakes State Park are likely a response to regional climate changes rather than local drainage basin effects.
The ∼8 ka 14C relative lake lowstand at Green Lakes coincides with the end of a cool, dry period (10.1-8.2 ka 14C) defined by stable isotope data from Seneca Lake (Anderson et al., 1997). This event followed the well-known Younger Dryas cold interval (∼10.8-10.3 ka 14C), and was correlative with a rapid phase of melting of the Laurentide ice sheet which poured large volumes of cold, isotopically light meltwater into the Great Lakes (Anderson et al., 1997). It also occurred quite close (∼7.5 ka 14C) to a recently discovered North Atlantic cool/dry event defined in Greenland ice cores (Alley et al., 1997).
The ∼6 ka 14C relative lake lowering defined at Green Lakes may also be a reflection of broader scale climate changes. In her review of natural lake levels throughout eastern North America, Harrison (1989) discovered that lakes here in general reached their lowest levels ∼6 ka 14C, implying widespread drought for at least a short interval of time. Cores from Montezuma wetlands also display a marked drop in the percentage of calcium carbonate at this time, implying relatively dry and/or cool conditions (Mullins, in prep).
The ∼4 ka 14C relative lake drop at Green Lakes coincides with widespread regional peat development, which effectively ceased marl deposition in the Finger Lakes region (Mullins, in prep.) Approximately 4 ka also marks the end of the mid-Holocene Hypsithermal (∼9-4 ka 14C) in central New York State, which was mostly relatively warm and wet (Dwyer et al., 1996).
Why relative lake level fluctuations in Green Lakes should occur every -2000 years is open to speculation. Perhaps they are related to regional climate changes caused by changes in the production of North Atlantic Deep Water (and thus heat transport) which presumably caused the Younger Dryas climate reversal (Broecker et al., 1989); or, perhaps they are related to shifts in the mean position of the jet stream and atmospheric pressure systems (Dwyer et al., 1996). However, there may also be an intriguing relationship between lake level fluctuations at Green Lake and natural variations in solar irradiance caused by sunspot activity. Wigley and Kelly (1990) have developed a proxy for prehistorical solar variation based on the difference between tree-ring counts and 14C age, which they as well as Karlen and Kuylenstierna (1996), have used to explain Holocene climate changes. The peats dated from the Green Lake wetland match-up well with relatively high amounts of solar irradiance at ∼8 ka, 6 ka, 4 ka and 2 ka 14C (Fig. 25). Greater amounts of solar energy at these times may have led to increased rates of evaporation and thus, relative lake-level lowerings. Although this potential linkage between solar variability and lake levels is speculative, it is certainly deserving of further thought and investigation.

STOP 5. Environmental Changes During The Past 2,500 Years
Short Cores From The Deep Lake
From our previous vantage point of the wetland between Green and Round Lakes we will continue to proceed clockwise around Round Lake. During the course of this part of the fieldtrip keep a sharp eye out for littoral zone bioherms similar to (but smaller than) those at Deadman's Point as well as marl benches indicative of a previous high stand(s) of lake level. At the southwest end of Round Lake we will also be able to observe the inlet to the lake as well as the ∼100 ft. (30.5 m) high cliff over which a late glacial waterfall presumably flowed to carve out this plunge-pool basin. We will then walk, uninterrupted, around all of Round Lake, past the north-side of the wetland, and back to Green Lake before our next stop. Our final stop (#5) will be along the path on the northwest side of Green Lake (across from Deadman's Point) which will afford us an open view of the circular, deep water basin of Green Lake (Figs. 2 and 3).
Because of the seasonal precipitation of calcium carbonate (whitings) coupled with the anoxic nature of waters below the chemocline, sediments in the deep main basin (and part of the "neck") of Green Lake are characterized by distinct annual laminations or varves (Fig. 26). Using sediment trap data Brunskill (1969) was able to document the summer accumulation of a light tan calcite laminae and the winter accumulation of a dark homogenous laminae of more organic-rich sediment; thus, proving the annual nature of these laminations (i.e.-varves). Ludlam (1969) further documented that the light colored laminae contain an average of -82% calcium carbonate whereas dark laminae contain only ∼55%. Average thickness of the more recent varves is on the order of ∼0.7 mm.
Ludlam (1969) also noted that not all sediments below the chemocline are varved; there is also an abundance of massive or "unlaminated" sediment. Because some of these massive layers (mm to cm scale) are graded, have erosional bases and contain coarse littoral zone grains, they have been interpreted as turbidites (Fig. 26; Ludlam, 1969, 1974). We have also recovered deep lake sediments from Green Lake which contain coarse gastropod shells and other material which has obviously been derived and transported from much shallower depths. There is also very good photographic evidence (Fig. 27) for the deformation of varves via subaqueous slumping (Ludlam, 1974; Dean and Fouch, 1983). Thus, there is little question that mass wasting and sediment gravity flows are processes that operate in Green Lake.
However, Ludlam (1969, 1974) has interpreted all massive layers (even those ∼1 mm thick) as turbidites and estimated that 40-65% of basinal sediments in Green Lake were deposited as gravity flows. We do not agree with this assessment which evolved largely in the late 1960's and early 1970's when turbidity current theory was being developed and widely applied. Many of the massive layers we have examined display no objective evidence for deposition by turbidity currents; they are not graded, do not contain displaced grains, and do not have erosional bases. In fact our core 6 from the central basin of Green Lake contains relatively few unequivocal turbidites (Fig. 28). Although this does not mean that all the massive layers are not turbidites, it does highlight a need for additional hypotheses. An alternative interpretation is that these massive layers (particularly the thin 1-2 mm thick ones) are simply thicker than normal winter layers of annual couplets (varves). In fact, Ludlam (1981) stated that" ... thin turbidites ... can not be separated from the annual layers" (p. 85). At the time of this writing, we are conducting compositional analyses and detailed varve versus radiocarbon chronologies in an attempt to develop objective criteria for the origin of many of these "massive layers" in Green Lake.
The accumulation rate of profundal sediment in Green Lake has also been found to be variable, both temporally and spatially. Ludlam (1981) noted a 48% increase in sedimentation rates between the late 1800's and the 1970's. Ludlam (1984) also discovered that sedimentation rates in the "neck" of Green Lake are ∼27% higher than they are in the main basin, perhaps due to the reworking and focusing of sediment from shallow water.
During the summer of 1996 we collected short (<50 cm) gravity cores at seven sites in Green Lake (Fig. 20) with the objective of using the annual deposits (varves) as high-resolution recorders of relatively recent environmental change. We have focused on a pair of cores recovered at site 6 from the floor of the central basin because they contain the lowest ratio of varved sediments to gravity flow deposits (i.e.-turbidites). Because massive zones are interpreted as instantaneous events, they were subsequently subtracted from the final varved sediment record, thus yielding a total varved section of 32.5 cm. Dating of the core is based on four approaches: (1) varve counts; (2) a radiocarbon date on terrestrial organic matter; (3) geochemical correlation with a core dated by 210Pb techniques (Whalen and Lewis, 1980); and, (4) recognition of a pink varve deposited in 1963 (Ludlam, 1969; Ludlam, 1984). The base of the varved section is ∼2500 years B.P. (504 B.C.; Fig. 29) and is thus a record of both natural and anthropogenic processes. Sediment accumulation rates between 2500 years B.P. and 167 years B.P. (1830 A.D.) averaged ∼10 cm/1000 years versus ∼70 cm/1000 years over the past 167 years.
Loss-on-ignition results for total calcium carbonate content in core 6 display a marked decrease at the same depth (18.5 cm) where accumulation rates increase by nearly an order of magnitude (Fig. 30). Average total carbonate values drop from ∼80% prior to 1830 A.D. to ∼60-65% after 1830 A.D. (Fig. 30). In contrast, average total organic matter contents in core 6 remain relatively constant at ∼10% both prior to, and after, 1830 A.D. These data indicate that starting in ∼1830 A.D., the flux of terrigenous material to Green Lake increased significantly resulting in increased rates of sediment accumulation (Fig. 30). This is supported by mineralogical data (determined by X-ray diffraction) which show an abrupt increase in the relative abundance of detrital quartz and dolomite at ∼1830 A.D. (Fig. 31). Prior to 1830 A.D. dolomite is generally absent from core 6 and quartz is only a minor component. Trace element data (determined by direct current plasma emission spectrometry) of the carbonate fraction in core 6 also display a marked change at ∼18.5 cm depth. Concentrations of Pb, Mg, Mn, Cu, and Fe all increase above this depth, whereas Ca and Sr decrease. Lead (a common anthropogenic element) is of particular interest (Fig. 32). Prior to 1830 A.D. lead concentrations range from 5 to 55 ppm with an average ∼25 ppm; whereas, after 1830 A.D. Pb concentrations range from 35 to 105 ppm with an average of ∼70 ppm (Fig. 32).
Collectively, these results argue for a significant anthropogenic influence on sedimentation in Green Lake since ∼1830 AD. Whalen and Lewis (1980) have previously proposed that deforestation of the region in the early to middle 1800's increased surficial runoff and the flux of terrigenous material into Green Lake. Our results are consistent with their suggestion. Pollution of surface waters by heavy metals at this time might also explain the observed increases of Pb, Cu and Fe since 1830 AD., although these metals may have also entered the Green Lake basin via the atmosphere as a by-product of coal burning starting with the industrial revolution. We have also sampled (N=48) varves from core 6 for stable isotope analysis in order to evaluate both natural and potential anthropogenically-influenced climate changes over the past 2500 years. These new data represent the first stratigraphic stable isotope results for Green Lake since a brief mention of results by Stuiver (1968,1970). δ18O (a measure of the 18O/16O ratio) values vary by as much as ∼2 ‰ over the length of the core (Fig. 33) which is 1 ‰ greater than the variability indicated by Stuiver (1968). δ18O values also display a gradual depletion from ∼10.2 %0 to ∼10.7 %0 between 500 B.C. and 200 AD. If this entire depletion (∼0.5‰) is due solely to temperature change (0.34 ‰/1 °C; Rozanski et al., 1993; Anderson et al., 1997) it would imply regional spring/summer cooling of up to 1.5°C. 2 However, some or all of this isotopic change could also be due to shifts in atmospheric moisture sources. In this case more depleted δ18O values would imply an increase of moisture from Pacific, Canadian or North Atlantic air masses, which all tend to be cooler than the warm, isotopically heavy air masses which advect into central New York State from the Gulf of Mexico (Rozanski et al., 1993). 2 In terrestrial records, such as ice cores or lake sediments, δ18O depletions imply relative coolings due to latitudinal fractionation of 18O. This general relationship is opposite that for marine records. This isotopic evidence for a modest cooling trend between 500 B.C. and 700 A.D. is consistent with ice core data sets which indicate global cooling during the past 4-5 ka following the mid-Holocene Hypsithermal (Larsen et al., 1995; Thompson et al., 1995). Both glacier reconstructions (Nesje and Kvamme, 1991) and general circulation models (Liao et al., 1994) indicate that Hypsithermal summer temperatures were 2-3 °C warmer than pre-industrial modem values, due to precessional changes.
Following an extreme δ18O depletion event (one data point) of ∼1 ‰ at ∼1250 A.D. δ18O values remain relatively constant until the early to middle 1800's and "colder" than modem values. Since then, δ18O values become enriched (heavier) by as much as 1 %0, although there is significant variability (Fig. 33). The increased variability is due in part to a lower sampling interval afforded by the higher accumulation rates after 1830 A.D. Whereas the pre-1830 A.D. average sampling interval is 1 sample per 90 years, the interval increases to about 1 sample every 14 years after 1830 A.D. allowing resolution of decadal scale variability.
It is tempting to speculate that the enrichment (warming) of δ18O values from Green Lake after the early 1800's is related to the anthropogenic rise of atmospheric carbon dioxide and global warming since the early to middle 1800's (Keeling et al., 1989). Absolute δ180 values over the past -150 years are as heavy as ∼9.3 %0 which is unprecedented over the 2500 year record in core 6 (Fig. 33). However, a 1 %0 increase in δ180 would indicate an ∼3 °C summer temperature rise, assuming that the entire isotopic shift is temperature dependent. Mean global surface temperatures have risen by only ∼0.5 °C since the mid-1800's. However, part of this observed increase in δ18O values from Green Lake over the past ∼150 years could also be due to an increase of moisture from warm, isotopically heavy Gulf of Mexico air masses, and/or evaporative enrichment.
To elucidate the potential causes of enriched and depleted δ18O values in Green Lake throughout the last century, a calibration using regional climate data was performed. Figure 34 shows the graphical comparison of δ18O values from ∼1900 to 1980 A.D. with yearly average summer temperatures for Syracuse, N.Y. When the temperature data is smoothed (a 15% running average function was utilized to remove inter-annual variability) decadal trends are evident that display a remarkable correlation (r=.86) with δ18O values (Fig. 34). Periods of cooler temperatures (1910 to 1930 and 1950 to 1970) correlate with isotopically depleted values while warmer periods (1900's and the late 1930's to early 1940's) correlate with isotopically enriched values. However, if the δ18O values were solely recording surface water temperatures in Green Lake during these periods, the exact opposite relationship would be observed (i.e. cool periods represented by isotopically enriched values). Therefore, the dominant control on the δ18O carbonate signature must be a function of the water composition in Green Lake, which is largely controlled by the type of airmass bringing precipitation into the area. As stated above, enriched values could be due to an increase of moisture from warmer temperature, isotopically heavy Gulf of Mexico source air masses while depleted values may result from cooler temperature, isotopically depleted Canadian or Pacific source air masses. Because Green Lake lies directly below the modern day position of the jet-stream, and the position of the jet-stream controls the type of airmass coming into the region, the Green Lake δ18O record provides a very sensitive proxy for jet-stream fluctuations and climate change.

B1: STRATIGRAPHIC INCOMPLETENESS: MILANKOVITCH IN THE MANLIUS AT THE MARGIN
Peter W. Goodwin and E. J. Anderson
Temple University
Philadelphia, P A 19122
As a result of lower subsidence rates, stratigraphic sections at basin margins are less complete than those in shelf and basine positions. Determining degree of incompleteness and its specific stratigraphic location or locations have been difficult tasks dependent on the accuracy of detailed chronologic correlations from complete basinal sections to less complete marginal sections. In the Helderberg Group (Fig. 1), large-scale models of gradual stratigraphic accumulation (e.g. Rickard, 1962: Laporte, 1969) were able to recognize major group-bounding unconformities, but did not identify internal smaller-scale unconformities. Using a model of small-scale episodic accumulation of time-stratigraphic allocycles, Goodwin and Anderson (1986, 1988) recognized cryptic unconformities internal to the Helderberg Group, surfaces at which detailed correlation of cycles revealed missing cycles (Figs. 2 and 3).
One such unconformity, traced from the Hudson Valley westward to the Syracuse area (Figs. 2 and 4), was interpreted as a sequence-bounding unconformity separating tectonically distinct portions of the Helderberg Group (Goodwin and Anderson, 1988). That third-order unconformity is the Manlius-Coeymans contact in the Hudson Valley and the Thacher-Olney contact in western sections (Fig. 1). Lesser unconformities, with fewer missing cycles, were recognized in the Manlius, but not fully explained (Fig. 2). It is the purpose of this trip to examine stratigraphic discontinuities of all scales within the Manlius Formation at four localities on the western margin of the Helderberg Basin, and to explain these unconformities in the context of the hierarchic model of orbital forcing.

B1: Stratigraphic Incompleteness: Milankovitch in the Manlius at the Margin

STOP 1. Manlius Formation at Munnsville Quarry
(Rickard Loc. 137)
The Munnsville Quarry Manlius section begins in the Thacher Member at the quarry entrance and continues into the Olney Mbr. on the north wall of the quarry. The upper Olney, Elmwood, Clark Reservation and Jamesville members are exposed in the west wall of the quarry.
At this locality, a third-order unconformity forms the Thacher-Olney contact (Fig. 8); extended eastward this unconformity becomes the Thacher-Ravena contact at Schoharie and in the Hudson Valley (Figs. I and 2). Between this unconformity and the fourth-order boundary at the base of the Clark Reservation Mbr. the OlneyElmwood interval (46.5 ft thick) consists of four relatively complete fifth-order sequences (A, B, C and D). Each of these fifth-order sequences consists of four or five sixth-order sequences (PACs). The fourth-order sequence below the unconformity consists of portions of two fifth-order sequences (A and B), the upper two missing by vacuity (erosion) at the unconformity.

STOP 2. Clockville Roadcut
(Thacher-Olney Interval, Manlius Fm., Rickard Loc 142) .
At this locality, the Thacher Mbr. and the lower Olney Mbr. are exposed. The described section (Fig. 9) begins near the top of the thick-bedded Thacher of Rickard (1962). The third-order unconformity (Thacher-Olney contact) is exposed near the top of the outcrop. The Olney-Elmwood portion of the columnar section was described at Perryville, 3.5 miles to the West.
As at Munnsville (Fig. 8), the fourth-order sequence below the unconformity consists of two incomplete fifth-order sequences (A and B). The smaller number of sixth-order cycles in fifth-order sequence B at Clockville relative to Munnsville suggests slightly deeper erosion at the third-order unconformity at this locality.

STOP 3. Madison Highway Products Quarry, Perryville
(Rickard Loc. 144)
At this locality, the upper Thacher, Olney, Elmwood and Clark Reservation Members are exposed in the north wall of the old pit behind the maintenance shed and scale house. The base of the section begins six feet below the Thacher-Olney third-order unconformity (Fig. 9). The Thacher portion of the column was described at Stop 3.
As at Munnsville, the Olney-Elmwood fourth-order sequence consists of four nearly complete fifth-order sequences (A, B, C and D). The slight difference in thickness (less at Perryville) and the generally more restricted facies at Perryville suggest that this locality is onshore of Munnsville. However, the similar stratigraphic structure of the two sections indicates very little topographic relief between the two localities.

STOP 4. Onondaga County Quarry, Jamesville
(Rickard Loc. 151)
At the Jamesville locality, the section (Fig. 10) begins in the Elmwood Mbr. on the road north of the quarry entrance. The upper Elmwood, Clark Reservation and Jamesville Mbrs. are exposed in the east wall of the quarry.
At Jamesville (Fig. 10), the Olney-Elmwood fourth-order sequence is less complete than at Perryville (Fig.9), consisting of portions of three fifth-order sequences (B, C and D). Sequence A was not deposited because this area remained positive until flooding during the deposition of sequence B. As yet, we have not determined which two sixth-order flooding events in sequence B are recorded at this locality. Similarly, fifth-order sequence D is less complete at Jamesville than at localities to the East. Whether this difference is the result of hiatus at the base of the sequence or erosion at its top has not been determined.

STOP 5. Interstate 81 Roadcut
(Olney, Elmwood, Clark Reservation, Jamesville Mbrs.)
This roadcut south of Syracuse exposes the upper Thacher, Olney, Elmwood, Clark Reservation and lower Jamesville Mbrs. of the Manlius Fm. The Thacher Member is completely exposed in the roadcut of the southbound lanes ofInterstate 81, a locality not visited on this trip.
At this locality (Fig. 11), the Olney-Elmwood fourth-order sequence is even less complete than at Jamesville, approximately 4 miles to the East. As at Jamesville (Fig. 10), the sequence consists of three fifth-order sequences (B, C and D), but sequence B consists of just one sixth-order cycle (PAC). Beneath the third-order unconformity (Thacher-Olney contact), the Thacher fourth-order sequence contains more sixth-order cycles than at localities to the East, indicating that erosion did not cut as deeply (vacuity is less). Still, this fourth-order sequence is far from complete, consisting of just two fifth-order sequences.

B2: HERKIMER DIAMOND MINE FIELD TRIP
Hugh Humphreys
Tibbits Road, New Hartford, NY 13413
Herkimer "diamonds" are doubly-terminated quartz crystals found near Herkimer, New York. The clarity and perfection of these crystals make them one of central New York's most collectable minerals. The crystals are found in muddy pockets, or vugs, up to a meter in diameter in the Late Cambrian Little Falls formation near Middleville, Little Falls and Herkimer, New York. As many as 1000 crystals have been found in a single pocket (Moore, 1989).
This trip will visit the Herkimer Diamond Mine, owned by the Herkimer Diamond Development Co. near Middleville, New York. The trip will be led by Hugh Humphreys. Hugh has been an avid miner and collector of Herkimer diamonds for 15 years. After a visit to the museum at the mine, he will discuss the occurrence of the crystals and will demonstrate different mining techniques. Ample time will be provided for participants to explore the mining area and search for "diamonds." Although this trip is scheduled for a half-day, you may stay as long as the mine is open. The following description is taken from information displayed with Herkimer diamonds in the mineral collection at Hamilton College, Clinton, New York.
The exact origin of the pockets and the quartz crystals is unknown; several related theories exist. One theory is that in Late Cambrian time, quartz sand grains and minor amounts of clay were deposited in a predominantly carbonate environment where there was significant algal growth. The sand, clay, and algae were incorporated into limestone, which was dolomitized by circulating waters. The formation is a rather normal dolostone or dolomitic sandstone.
After burial, entrapped sea water or water from some other source dissolved some of the sand grains, dolomite, and algae to form the pockets. Many pockets today contain a black carbonaceous material called anthraxolite, which may represent an insoluble residue from the algae. Thus, the algae may have served as sites for solution and pocket formation.
As the pH of the water in the pockets decreased, quartz crystals started to grow. The source of the silica was the dissolved sand and clay or the remains of silica-secreting animals such as radiolaria. Some quartz crystals started their growth on the pocket wall and grew into the pocket. Many crystals, however, grew freely in the pocket water or in the mud that was the insoluble residue from pocket formation, allowing the doubly-terminated crystals to form. Most of the crystals are short hexagonal prisms terminated by rhombohedrons.
Some crystals show solid or fluid inclusions. Solid inclusions are usually specks of anthraxolite. Fluid inclusions are mostly salt water with bubbles of CO2 or water vapor. Based on a homogenization temperature of 51°C, the crystals formed at a depth of approximately 3500 feet (Dunn and Fisher, 1954).

Herkimer Diamond Mines

B3: EURYPTERIDS AND ASSOCIATED FAUNA AT LITCHFIELD, A CLASSIC LOCALITY IN HERKIMER COUNTY
Victor P. Tollerton, Jr.
23 Shepherd Place
Utica, New York 13502-5417, mltolrton@borg.com
The Upper Silurian Bertie Formation (Group of D.W. Fisher, 1960; Ciurca, 1973) of New York State and Ontario, Canada, is world renowned for its spectacular eurypterids. One of these, Eurypterus remipes Dekay, 1825, holds the distinction of being the first eurypterid in the world to be described. This species is also the state fossil of New York, as signed into law by Mario Cuomo on 27 June 1984. Another species, the second one in the world to be described, Eurypterus lacustris Harlan, 1834, has been figured on a postage stamp of Canada (as Eurypterus remipes).
Central New York is fortunate to boast of four localities where eurypterids can still be collected from the Bertie (Forge Hollow, Litchfield, Lang's Quarry, and Passage Gulf). Two others mentioned in the literature (Jerusalem Hill and Crane's Corners) are not currently accessible. The Litchfield, Herkimer County locality is a classic outcrop that displays both vertical and horizontal sections amenable for paleoecological study.

B3: Eurypterids and Associated Fauna at Litchfield, A Classic Locality in Herkimer County

STOP 1. Forge Hollow, Oneida County, New York.
BE CAREFUL ... blind curves at both ends of the outcrop.
This locality is the type locality for the eurypterid Eurypteros remipes Dekay, 1825, and the scorpion Proscorpius osborn; (Whitfield), 1885. It is also the type locality for a new genus and species of eurypterid being described.
This outcrop exposes the following units (in ascending order); Camillus Shale; the Morganville, Victor, and Phelps beds of the Fiddlers Green Member, Bertie Formation; Scajaquada Member (= Rickard's 1962 Forge Hollow Member), Bertie Formation; Williamsville Member, Bertie Formation; Cobleskill Formation; and Manlius Formation.

STOP 2. Litchfield eurypterid locality.