NYSGA 2014
NEW YORK STATE GEOLOGICAL ASSOCIATION (NYSGA)
86th ANNUAL MEETING
86th Annual Meeting October 10-12th NYSGA 2014

A1: LOWER PALEOZOIC SEDIMENTARY SUCCESSION OF THE ST. LAWRENCE RIVER VALLEY, NEW YORK AND ONTARIO
Antun Husinec
Department of Geology
St. Lawrence University
Canton, NY 13617
J. Allan Donaldson
Department of Earth Sciences
Carleton University
Ottawa, Ontario
Canada K1S 5B6
INTRODUCTION
Our field trip to the St. Lawrence River valley in New York and Ontario will showcase the two Lower Paleozoic formations outcropping along the river (Fig. 1.), the middle to upper Cambrian Potsdam Sandstone, and lower-middle Ordovician Theresa (March in Ontario) Sandstone. We will be able to examine the non-conformable contact between Potsdam Sandstone and the underlying Proterozoic basement of the Grenville orogeny, the disconformable contact between Theresa and Potsdam Sandstones, different primary sedimentary structures, trace fossils, and microbial structures preserved within the formations.

A1: LOWER PALEOZOIC SEDIMENTARY SUCCESSION OF THE ST. LAWRENCE RIVER VALLEY, NEW YORK AND ONTARIO

STOP 1. Nonconformity Between Proterozoic Basement and Potsdam Sandstone at Alexandria Bay, New York
One of the best exposures of the non-conformable contact between the Potsdam Formation and the underlying Proterozoic basement in northwestern New York is in a road cut located approximately 1.5 miles northeast of Alexandria Bay, where New York State route 12 (NY-12), a two-lane undivided roadway cuts into a hill some 360 m (1,180 ft) west of the Cranberry Creek bridge (Fig. 5). The basal Potsdam Formation is in sharp contact with the Proterozoic gneiss that shows signs of alteration (illite, Fe-chlorite, and siderite; Selleck, 1993). The basal ∼2 meters of Potsdam Sandstone weather into thin and friable slabs that are composed of low-angle cross-laminated (Fig. 6), non-arkosic and non-conglomeratic quartz arenite that is overall moderately sorted and contains some coarse grains within the predominantly fine- to medium-grained framework. Sorting increases up-section, and the sandstone becomes more massive. No body fossils or trace fossils are present in the Potsdam Formation at this stop.

STOP 2. Primary Sedimentary Structures in Basal Potsdam Sandstone, Goose Bay, NY
The northwest face of this road cut nicely exposes lowermost approximately 2.5 meters of Potsdam Sandstone (Fig. 7). The basal ∼1.5 meters is a cross-laminated medium- and fine-grained quartz arenite showing tabular cross-bedding with curved bases and sharp erosive tops. The overlying ∼40-cm-thick tabular quartz-arenite bed is characterized by planar to very low-angle cross lamination. The topmost set exhibits discontinuous, faint wavy (possible erosive bases) and parallel stratification, and becomes more massive up-dip. The exposed section is barren of body and trace fossils and contains no obvious microbially formed structures.

STOP 3. Nonconformity Between Proterozoic Basement and Potsdam Sandstone at Goose Bay, NY
The road cut on the southeast side of the road exposes the nonconformity between Potsdam Sandstone and the underlying Grenville basement rock (Fig. 8). The contact is sharp but irregular, and displays heavily weathered, friable Proterozoic basement rock below the unconformity. Basal Potsdam Sandstone is a quartz arenite composed of poorly sorted and angular grains without any fossils. The lower 50-60 cm is white to light gray in color; the color changes in the upper part of the outcrop to pink and red. The reddish color around detrital grains and within secondary silica and illite cements is due to presence of finely crystalline hematite, goethite and anatase that formed by breakdown of detrital magnetite and ilmenite grains (Selleck, 1993). This basal, low-angle cross-laminated sandstone weathers more easily than the overlying, more massive quartz arenite that we observed at Stop 1, but is missing at this outcrop.

STOP 4. Potsdam Sandstone at Schermerhorn Harbor, Hammond, NY
Note that between this stop and Alexandria Bay, all the outcrops are either Proterozoic Grenville basement or Potsdam Sandstone, indicating that Potsdam Formation blanketed topographic lows of the Proterozoic surface. Topographically higher areas likely were not sites of Potsdam deposition, or alternatively, sandstone was subsequently eroded from these areas. At this an unconformity within Potsdam Sandstone is exposed, (Fig. 9). The basal 3 meters above the unconformity is characterized by cross bedding (Fig. 10), with alternating poorly sorted, coarse- (up to very coarse in places) to medium-grained quartz arenite. This basal sandstone characteristically weathers into thin slabs. Upward in the section, the sandstone becomes more massive.

STOP 5. Disconformity Between Potsdam Sandstone and Theresa Formation at Chippewa Bay, NY
The disconformity between the Theresa Formation and Potsdam Sandstone is spectacularly exposed near Chippewa Bay, New York (Fig. 11). An unconformity without angular discordance (dashed line in Figure 11) is marked by a thin, vegetated interval on top of the Potsdam Formation. The contact between two formations is sharp and indicated by a change in color from white to pale tan silica-cemented Potsdam below, to carbonate-cemented dark gray Theresa sandstone above. Unlike the lower part of the Potsdam Formation that on previous stops is devoid of any biogenous structures, its uppermost part exposed at this stop contains abundant Diplocraterion and Skolithos trace fossils (Fig. 12), clearly indicating a shallow-marine setting. The basal part of Theresa is extensively burrowed, and most primary sedimentary structures are completely obliterated. In addition to biogenous structures, Selleck (1993) recorded numerous fragments of the brachiopod Lingulepis in both the upper Potsdam and lower Theresa at this location.

STOP 6. Lower Theresa Formation North of Chippewa Bay, NY
The road cut on the northeast side of the NY-12 exposes extensively burrowed lower Theresa Formation (Fig. 13). Individual burrows, and possible pseudomorphs after evaporite minerals, are clearly visible as molds on weathered face of the outcrop (Fig. 14), and contribute to overall high porosity of this facies. Medium-grained quartz arenite is cemented by carbonate and is moderately to well sorted. The dark gray color is typical for the lower Theresa, with rare thin yellowish intervals that follow the bedding.

STOP 7. Microbial Structures in Theresa Formation North of Chippewa Bay, New York
The road cuts at this stop expose microbial structures preserved within the white, medium-grained quartz arenites of the middle Theresa Formation. These structures are exposed on both sides of NY-12, and are distinguished by wavy-laminated stromatolite growth structures that are common in the white quartz sandstones of the middle Theresa Formation. Vertical sections of stromatolites exhibit predominantly space-linked hemispheroids with close-linked hemispheroids as a microstructure in the constituent laminae (Fig. 15). Hemispheroids vary both in amplitude and in shape, i.e. from low-amplitude (5-10 cm) and gently convex, to higher-amplitude (up to 20 cm), steeply convex to slightly rectangular, vertically stacked hemispheroids. Sub-circular, concentrically stacked spheroids up to 30 cm in diameter, with laminae composed of close-linked hemispheroids are observed in plan view (Fig. 16). Their shape closely resemble intertidal forms from Shark Bay (cf. Tucker & Wright 1990, p. 150, fig. 4.50B).

STOP 8. Upper Theresa Formation at Morristown, NY
The upper Theresa is nicely exposed in road cuts on both sides of NY-37 immediately south of its intersection with High Street at Morristown, New York. The formation is characterized by thin beds of dark gray and white-yellowish calcareous quartz arenite that irregularly alternate, inter-finger and pinch out. Dark gray sandstone is thoroughly burrowed; extensive bioturbation resulted in complete obliteration of primary structures in this sub-tidal facies. On the contrary, although locally burrowed, the white to yellowish sandstone is characterized by well-preserved sedimentary structures indicating intertidal sand-flat setting with channel fills. Sedimentary structures include ripple cross-lamination, erosional (reactivation) surfaces, herringbone cross bedding, and local conglomerate-filled scours.

STOP 9. Nonconformity Between Precambrian Basement and Potsdam Sandstone
Lyn Valley Conservation Area, Ontario
From the parking area follow the gravel road northward to the start of a footpath that leads eastward through bushes and trees to the base of a 20-meter vertical section of Potsdam sandstone (Fig. 20). This cliff displays almost flat-lying beds with pebble interbeds (clasts of quartz sandstone), ripple marks and low-angle crossbedding that reflect deposition in water rather than by wind. Elsewhere within the Potsdam, large-scale steeply inclined crossbedding in pebble-free Potsdam sandstone in areas to the west reflects eolian transport. Trace fossils plus a wide range of directions of cross-bed inclination indicate a shallow-water marine environment. Curled biofilm structures (Fig. 21) plus desiccation cracks reflect occasional subaerial exposure. Surface patches of gravel and sand cemented to the outcrop surface mark places where cross-bedded Pleistocene outwash was once in contact with this Paleozoic cliff, reflecting the rapid rate at which unconsolidated deposits can be locally lithified.
Distinctive roll-up structures, especially prominent near the western end of this outcrop (Fig. 21), record the curling of sheets of sand grains bound by tenacious films of cyanobacteria in response to desiccation. Although the sands were deposited in water, the presence of such structures indicates intermittent episodes of subaerial exposure sufficiently long to allow drying, shrinking and curling of such coherent thin sheets of biofilm-bound sand. Elsewhere in eastern Ontario and northern New York State, similar curled biomats are overlain by sequences of continuous biomat layers that trapped successive interlayers of quartz sand to form distinctive domal stromatolites, indicating more persistent subaqueous conditions. Impressive examples occur within the city limits of Ottawa, Ontario (Hilowle et al., 2000, Donaldson et al., 2002)
On the return trip to the bus, check a few of the numerous boulders and blocks (Fig. 22) arrayed along the west side of the gravel roadway. Extracted from nearby outwash deposits, they provide examples of Precambrian igneous and metamorphic rocks from the Canadian Shield in addition to samples from the Paleozoic cover. The Precambrian boulders display pegmatite dykes, crosscutting relationships, inclusions, foliation, lineation and fractures (some with minor fault offsets).
Between the margin of the gravel road and the telephone pole in Figure 22, immediately north of the largest boulder, a cluster of small glacially polished and striated outcrops of Grenville basement rock displays near-vertical foliation (Fig. 23). Distinct lithologic interlayers suggest a sedimentary provenance for these Precambrian rocks, with the compositional layering probably marking original bedding.

STOP 10. Heritage Stone Building on West Main St., Lyn Village, Ontario
Its walls of Potsdam sandstone are replete with an impressive array of trace fossils, with Diplocraterion and Skolithos predominating. A few blocks, especially those in the north wall, show syndepositional deformation of these trace fossils in vertical section, comparable to the syndepositional overturn of cross-beds.

STOP 11. Paleozoic-Precambrian Unconformity, Quabbin Hill
The rock cut on the north side of the highway on Quabbin Hill directly exposes the regional nonconformity between Potsdam sandstone and the underlying Grenville basement rock. The latter displays excellent spheroidal weathering of the basement (Fig. 26, 27), which here is a massive intrusive mafic rock. The Potsdam contains trace fossils perpendicular to bedding, just like those seen at STOPS 9 and 10, but the underside of beds on the south side of the highway display a different variety of trace fossils: sinuous trails parallel to bedding (Fig. 28).
Of particular interest in the south-side road-cut are two independent indications of repeated subaerial exposure: bedding-parallel nodular zones within bioturbated silty layers that probably represent tropical soil horizons (Fig. 29), and desiccation patterns on several bedding surfaces. The unconformity surface marks a gap in the geological record of at least half a billion years. During this time, the land may have been elevated so that no rock record accumulated, although it is possible that post-Grenville Proterozoic strata were indeed deposited, but subsequently removed by erosion before the onset of Paleozoic sedimentation.

STOP 12. Precambrian Basement Exposure Within Community at Rockport Boat Line
This outcrop displays lit-par-lit sills of pink to white granite pegmatite and aplite that were injected parallel to the foliation of fine-grained gray to black gneiss and schist of presumed sedimentary provenance (Figure 30). Similar magmatic fluids were injected along crosscutting dykes during multiple episodes that can be put in sequence by observing crosscutting relationships. The metamorphic foliation of the host rock for these intrusions may have developed along the bedding of an original clay-rich sedimentary rock, such as mudstone or siltstone. Although foliation is typically vertical or dips steeply in most exposures or Grenville metamorphic rocks, it here is almost horizontal, perhaps due to location along the hinge of a fold, or to low-angle thrust faulting. The pinch-and-swell morphology of some lit-par-lit injections indicates lateral extension that was locally sufficient to form several boudinage (Figure 31). Note also the prominent sub-ice hydraulic scouring that created sculpted and polished benches along the outcrop during deglaciation.

A2: MILITARY GEOLOGY OF THE BATTLE OF SACKETT’S HARBOR (28 MAY 1813), LAKE ONTARIO, NY
Alexander K. Stewart
Department of Geology
St. Lawrence University
Canton, New York 13617
Ross W. Klepetko
Storefront Academy Harlem
New York, New York 10035
INTRODUCTION
Geology has shaped warfare since its inception; however, it was not until the end of the 18th century that geologists shaped warfare (cf., Guth, 1998; Kiersch and Underwood, 1998). For thousands of years, the idea that a hilltop position or fortification would provide the tactical advantage was considered “common sense.” Those who held the high ground were more likely to survive the fight and promulgate their ideals and to prosper. With the inception of the discipline of geology, these common-sense notions were turned into well-thought and planned military-geology strategies. There are many examples of the science of geology dictating tactical and strategic outcomes in war, from Napolean’s invasion of Egypt (1798) to recent counterinsurgency operations in Afghanistan (Stewart, 2014).
For the Battle of Sackett’s Harbor (1813), engineer siting of Fort Tompkins and anticipated defense of the mainland by Brigadier General Brown, New York Militia were “common-sense” approaches to dealing with the tactics of the day. For this field trip, we propose to highlight these “common-sense” notions as geological controls on the battle. If it were not for the inadvertent recognition of geological processes at the main naval base on Lake Ontario, this United States tactical victory would have allowed the United Kingdom complete control of Lake Ontario and their primary supply route, the St. Lawrence River.
Portions of this field trip are the result of Ross Klepetkos’ senior project (2011) at St. Lawrence University where he synthesized his love of history with geology. The purpose of this trip is to provide an enjoyable recounting of the Battle of Sackett’s Harbor (1813) by following the path of the advancing British Regulars while highlighting the local geology and its unintended controls on the battle. Had the geology of Sackett’s Harbor been different, there would have been no rebounded cliff, spit and protected harbor, tombolo and island–no naval stronghold and no battles–no Sackett’s Harbor.

A2: MILITARY GEOLOGY OF THE BATTLE OF SACKETT’S HARBOR (28 MAY 1813), LAKE ONTARIO, NY

STOP 1. UK Amphibious Assault Location-Windward Side of Horse Island
Joint-Controlled “Biostrome” (Ordovician)
At this location you are in the most concealed position with respect to Forts Tompkins and Pike near Sackett’s harbor. This windward side of Horse Island was a “natural” place for Commanders Yeo and Prevost to make an amphibious assault (Figure 2). General Brown, NYM, however, had set up defensive positions in the tree line along this edge of the island. Upon dismounting, the UK troops were under fire from 6-pound field guns and the 32-pound gun, firing long range, from Fort Tompkins. Because the NYM were untrained, completely untested and feared the regular British troops, they retreated back across the shoal (tombolo) to the mainland where they set up another defensive position behind the storm-deposited beach berm (Figures 3 and 9).
“The militia soon began to assemble, and as fast as they arrived they were armed and sent to Horse Island, which was the point at which the enemy was expected to bind.” (Rodgers, 1897)
Geologically, Horse Island is a joint-bounded (Figure 10), biostrome-like, coarse-grained limestone with well-exposed orthocone nautaloids at lake level. This coarse unit is overlain by approximately 0.5 meters of nodular calcisiltites in 5-10-centimeter beds interbedded with silts and shales. As you walk the perimeter of the island to this location you will notice an excellent assortment of glacial erratics, which have been storm-wave concentrated into piles and mantling the northern shore. These erratics, gabbros, gneisses, Potsdam sandstone, pegmatic granites and porphyritic gabbros, likely rolled out of the glacial drift covering the island. On the western-most lake-level exposure are a series of glacial striations with a mean vector of 196° (n=8, σ=3.8°; Figure 11).

STOP 2. Horse Island (Sackets Harbor) Lighthouse (C. 1870)
Initial USA Defensive Positions
The Horse Island lighthouse (Sackets Harbor, ARLHS USA-380) was originally emplaced in 1831 approximately 100 meters northwest of this location after Congress appropriated $4,000 for the building of a 1.5-story keeper’s dwelling and tower (LHF, 2014). By 1869, the lack of upkeep resulted in conditions beyond repair; therefore, Congress appropriated more funds for a new Queen Anne-Italianete brick lighthouse and keeper’s dwelling (Figure 12). In 1899, the tower received a ten-foot addition raising its height to 55.5 feet (LHF, 2014). This light was decommissioned in 1957 and replaced by a modern, steel tower.
Because Commanders Yeo and Prevost decided to await calmer waters after their arrival to the Sackett’s area on the evening of 27 May (and a bungled attempt at attack that day), Brigadier General Brown, NYM was able to set up hasty defensive positions in and around this location. Most likely, positions would have been naturally made beach-berm positions during the initial amphibious assault. As the NYM retreated through the woods at this location, the British Regulars were able to reform and be supported by their naval comrades. To the east of this location, there is a thought-to-be entrenchment used by the NYM during the assault; however, it is coincident with the jointing in the area and is most likely mass wasting in progress (similar to Stop #4).
“…separated from the mainland by a shallow strait, which is always fordable and sometimes almost dry. This strait, which, with the approach, formed a causeway 400 yards in length, had to be traversed by the attacking column…” (Hannay, 1903)
The military-geology of this island is limited to its ability to provide concealment to the amphibious landing party. This concealment was provided by trees growing atop a joint-controlled limestone “biostrome” that has withstood millennia of direct impingement of waves and storm waves. This energy was maximal on the western side of the island and wave refraction wrapped around the island providing sediment for deposition due to lack of competence in the leeward side of the island making the tombolo used for the mainland assault. Along the eastern edge of the island there is an elongated “entrenchment,” thought to be made/used by the NYM during the battle (and thereafter). Unfortunately, the structure is not man-made, but the intermediate phase of a block slump (similar to what can be seen at Stop #4). The strike of the feature is the same as the northwest-southeast joint set affecting the area. As far as its use by NYM during the War of 1812, that too is unlikely, for its orientation and location is on the Fort Tompkins-garrison side providing little defensive or offensive purpose.

STOP 3. Tombolo-Mainland-USA-Held Kill-zone for Advancing UK Force
Storm-Deposited Beach Berm
At this point, NYM had retreated after a perfect opportunity to “shoot fish in a barrel” on the island. After their splashy retreat across the tombolo, what militia remained were rallied by Brown, NYM to take defensive positions behind these storm-deposited beach berms. This defensive action is best summarized by a retreating NYM, “I fear we shall be compelled to retreat … I know we shall, and as I am a little lame, I’ll start now” (quoted in Wilder, 1994). Despite this natural, 250-meter, amphibious “kill-zone,” the British regulars were able to cross the tombolo and make the mainland with little resistance and move northward along the shoreline to Fort Tompkins; CPT Viger of the British Voltigeurs (skirmish unit) flanked to the east trying to get around the US defenses (Figures 3 and 4).
“The beach is of very limited extent, and bounded on the land side by a bank of sand several feet high, which had been thrown up by the current, as it rushed through the narrow channel separating Horse Island from the main. Behind this natural embankment, General Brown had formed a battalion of militia of 500 men” (Brown, 1827)
This tombolo is the result of both wave refraction around the island converging at this location and longshore drift, which transports material around the island (Figures 13 and 14a,b). Its current geomorphology suggests that it is active only during conditions of greater than approximately 10-knot winds and 1-meter waves (pers. obsv.) with major storms being the most productive in re-working and re-organizing the tombolo. General composition of the tombolo is fine-coarse-grained clastics and shelly material. The erratic-rich locations, however, are probably the result of glacier-drift reworking and mantling along with human input. The island currently has electricity (run along the easterly side of the tombolo) and access to/from the island is commonly with vehicular traffic. As a result, concentrations of erratic boulders are likely accentuated due to reorganization for human purposes (such as the duck blinds on the west side of the island). Based on lake levels before the emplacement of dams along the St. Lawrence, the late-May battle date and historical accounts (e.g., Brown, 1827) the water level at the time of the assault was likely about knee-deep or “fordable depth” (Lossing, 1869)(Figures 14b).

STOP 4a. Oil-Loading Facility Exposure - Western Edge of UK Advance
Cliff Retreat and Joint Exposures
Along this stretch of the lakeshore, advancing British troops were weaving through the dense underbrush of a wooded landscape. In addition, they were beginning to tackle the man-made defenses, abatis or an obstacle made of trees/sticks with sharpened end toward the enemy (Figure 5).
“The American artillerists toiling in Fort Tompkins repeatedly discharged the 32-pounder into the woods toward the oncoming British. A deluge of cannon balls, grapeshot, and bullets lopped off treetops and branches, which flew in every direction. One nearly killed Captain Jacques Adhimar [British Army].” (Wilder, 1994)
The prominent structure at this location is the oil-loading mooring “dolphin,” which was part of a Mobil Corporation petroleum bulk-storage facility that was emplaced during the 1920’s (Figure 15). During WWII, tanks were added on the northern side of Ambrose Street expanding storage. Petroleum storage continued at this site until 1988 reaching 589 Mbbl capacity with storage tanks being dismantled in 1989. Several environmental investigations have been completed since the 1980’s evaluating the nature and extent of contamination; current results suggest that the site no longer contains any threat to human health or the environment (NYDEC, 2013).

STOP 4b. Memorial Cross Vista - Western Edge of UK Advance
Cliff Retreat and Joint Exposures and Lake “Stack”
Continuing north along the cliff face, the advancing British soldiers were under heavy fire from Fort Tompkins. In this approximate location, they were limited by the cliff and were mired in abatis or a man-felled brush-and-tree obstacle, but because the NYM were in retreat, much of this advance was unimpeded until they reached the cantonment clearing 100 meters or so farther north (Figure 5).
Geologically, this area shows pronounced weathering of the northwest-southeast joint sets, which are coincident with impinging wave fronts. The notched cliff face to the south (Horse Island way), begins to give way to a more vertical cliff that begins to show joint weathering deep into the cliff to reach the antithetic east-west joint sets. As a result, “lake stacks” are forming and collapsing from this point north to the harbor. From the cross location, you can see this transition from notched cliff face to sapped and retreating, vertical cliff face.

STOP 5. Fort Kentucky (1814) - NYS Militia Retreat
Earthworks and 32# Cannon
This 1960’s reproduction 32# gun represents the firepower available atop Fort Tompkins. Just after the repulse of the British, Sackett’s Harbor defenses were augmented with Forts Kentucky (palisaded earthwork), Chauncey (stone tower), Stark (earthworks) and a ring of earthworks connecting them to Forts Virginia and Volunteer (NAF, 2014). As the British advanced along the lakeshore, dealing with the woods and abatis, the NYM were making their way through the same along this path to the barracks, Fort Tompkins and Volunteer.

STOP 6. Barracks - In-Line Engagement
This is the approximate location of the Smith (Basswood) Cantonment (palisaded barracks complex with four blockhouses) where the soldiers were garrisoned (officers were quartered in village homes). At this point, the Americans were in line and prepared for the mired British advance. When the first “red jacket” was visible, the Americans fired from the line. Sending a staggering blow to the advancing British, the well-trained British responded in kind and were able to overcome the Americans with rapid and precise fire and advance, pushing the Americans back to the protection of the cantonment (Figures 5 and 6).

STOP 7. Navy Point and Sackett’s Harbor
Dry-Docked USS General Pike and Fort Pike - Harbor and Storm-Deposited Spit
From this location, the “Navy Point” spit is visible (Figures 16 and 17). In this relatively deep-water harbor (3-4 m) the recently “laid down” USS General Pike, a 28-gun (24# each) corvette, named for the recently killed explorer/general Zebulon Pike (of Pike’s Peak fame) was being constructed. This was the strategic key to the battle - completion of the USS General Pike would promise lake superiority to the US before the end of the sailing season, leading to control of Lake Ontario. This, however, was not the case, during the melee a signal was confused and sailors began setting the stores on Navy Point and the USS General Pike ablaze (Figures 6 and 7). Luckily, such hasty creations during wartime were made of green wood and the damage was limited, yet significant enough to guarantee British control of Lake Ontario until the following sailing season (1814).
Geologically, Navy Point is a storm-deposited spit resulting from the longshore transport of sediments from the southwest between here and Horse Island. At the time of the battle, this spit was likely a relatively fine-grained mass that would be altered during significantly stormy periods. With increased human intervention, the spit has become mantled and a more-or-less permanent structure for the protection of the harbor. Contemporaneous maps of the harbor (e.g., Wilkinson, 1816) show the “recurve” or waning wave-refracted, longshore drift appendage at the terminal end of the spit. This portion of the landform, too, has been mantled and has been turned into a docking structure.

STOP 8. Fort Tompkins - Primary Harbor Defense
Joint-Controlled Mass-Wasting Block
This location was the center of the primary, blockhouse fortification for Sackett’s Harbor. Mounted with a 32# cannon (like the replica at Fort Kentucky, Stop 5), it was the deciding factor in the success of the Americans and their downfall! Well-aimed shots from this gun kept the British at bay; however, because of confusion and lack of communication, over-shots from the British Naval ships landing in/around Navy Point, were interpreted as enemy-controlled Fort Tompkins (Figure 6). Regardless, Fort Tompkins remained in American hands and was key to defending the harbor.
In this approximate location there is a block of harbor cliff that is in a medial state of mass-wasting failure. This intermediate failure is recognized by the joint-controlled gradient changes in the lawn adjacent the boathouse. These gradient changes (or escarpments) are coincident with the overall joint sets in the area (Figure 10). This portion of the cliff face is significantly larger than the other recently failed cliff sections (seen from Stops 4a,b) and, as a result, was “too large” to fully fail.

STOP 9. Lake Ontario Shoreline - Defensive Cliff Face
Shoreham Interval Exposure and Wave-Cut Cliff
Naval support of the British invasion at Sackett’s Harbor was nil. Primarily, relegated to an amphibious support role, the calm winds rendered their support useless. Of the ships near Horse Island, only one was equipped with “sweeps,” or long oars, the HMS Beresford, which was able to slowly maneuver itself into a supporting role. This supporting role, however, was hampered by the incessant pounding of the the 32-pound Fort Tompkins gun and the inadvertent, yet careful placement of the Fort. The 7-meter cliff face below the Fort was a serendipitous protector (Figure 18). Because of the cliff and the placement of the Fort, at ranges naval firing ranges of 500 meters or so, accurate fire required an approximately 0.4-degree-bore-sight elevation precision to place a shot through an approximately 25 m2 window of “opportunity.” Although the British seamen were excellent shots and well trained the task proved too difficult.
“Most of the enemy’s shot fell against the rocks [cliff] below the battery. One of these (a thirty-two-pound ball) came over the bluff, struck the earth not far from Sackett's mansion (then occupied by Vaughan's family), and plowed a deep furrow into the door-yard.” (Lossing, 1868)
Geologically, this cliff represents about seven meters of Trenton Group, “Shoreham” interval limestones of the Sugar River Formation (Table 2). From this location toward Horse Island, the overall topography of the land dips approximately 8 m/km along the strike of the lakeshore (Wallach and Rheault, 2010), which is due to the Lake Ontario Homocline (Kay, 1942). This homoclinal structure is likely the result of Adirondack mountain orogenesis to the east with relatively recent glacial isostatic rebound. Along the lakeshore, from this location, there are excellent opportunities for fossil collection and observation of mass-wasting processes due to shoreline weathering and erosion.

A3: GEOCHEMISTRY OF TOURMALINE FROM SOME ADIRONDACKS LOCATIONS - INDICATOR OF THE HOST ENVIRONMENT
Marian v. Lupulescu
Research and Collections
New York State Museum
Albany NY
Jeffrey R. Chiarenzelli
Department of Geology
St. Lawrence University
Canton NY
“With its plethora of chemical constituents, its wide range of stability from conditions near the Earth’s surface to the pressures and temperatures of the upper mantle, and its extremely low rates of volume diffusion, tourmaline can acquire a chemical signature from the rock in which it develops and can retain that signature through geologic time.” (Dutrow and Henry 2011)
INTRODUCTION
Tourmaline is a complex borosilicate and it can record the chemistry of its host environment and the chemical changes in the composition of the generating fluids or associated minerals.
Tourmaline-supergroup minerals are found as a primary accessory phase in granites and pegmatites, breccia, and hydrothermal quartz veins (London and Manning 1995), metamorphic rocks ranging from zeolite to granulite facies (Henry and Dutrow 1992; Ertl et al. 2010), volcanogenic massive sulfide deposits (Slack et al. 1996), and authigenic overgrowths in sedimentary rocks (Henry et al. 1994; Lupulescu et al. 2010). Variation in the chemical composition of tourmaline species allows them to form in a wide range of environments. This feature makes them an excellent indicator of the environmental conditions of their host and is further enhanced by their very low rate of diffusion at high temperature and resistance to weathering (Hinsberg et al. 2011).
Due to its hardness and lack of cleavage, the tourmaline in sedimentary rocks can be used as a provenance indicator (Henry and Guidotti 1985; Henry and Dutrow 1992; Morton et al. 2005; Marschall et al. 2008) and the varied colors contribute to make it a desired gemstone (Pezzota and Laurs 2011). It also has pyroelectric, piezoelectric, and optical (polarizing light) properties. The ability to polarize light adds credibility to the legend of the “Viking Sunstone”, but it was used in the modern times to make polarizers for petrographic microscopes.
The goal of this trip is to visit and discuss the significance of the chemical, major and trace elements, of tourmaline from different locations and environments within the Adirondack Lowlands, such as tourmalinites, talc-tremolite schists, quartz veins, simple pegmatites, Balmat SEDEX-type deposits, and authigenic tourmaline. In addition, some new geochronological information is present to help put the tourmaline occurrences in geologic context. The quantity and variability of tourmaline occurrences in the Adirondack Lowlands is still poorly understood and this trip hopes to bring attention to their diversity.

STOP 1. Tourmaline Bearing Gneisses at Richville (Rt. 11)
Outcrop on north side of busy State Highway 11, exercise extreme caution
Exposed in this outcrop is a layered sequence of quarzto-feldspathic gneisses containing up to 50%, or more, black tourmaline as small equant grains (Figure 3). These tourmaline-rich rocks and tourmalinites can be traced for 50 kilometers or more north of Route 11 in the Lowlands. They occur as a northeast-trending belt of northwest-dipping tourmaline-bearing metasedimentary rocks in the lower marble and were originally studied by Brown and Ayuso (1984). Spectacular breccias comprised of angular blocks of these rocks help define the trace of a large Paleozoic wrench fault north of Goveurneur (Selleck, 2005).
Well defined layering within the outcrop is suggestive of a detrital origin for the rocks. Despite deformation and upper Amphibolite facies metamorphism layers are defined by changes in grain-size, layer thickness, color and staining, and modal mineralogy, including variations in quartz, k-spar, tourmaline, and pyrite. Note the iron staining related to sulfide-rich areas in the outcrop. A metasedimentary origin is also indicated by detrital zircon studies which are summarized below.
The meter-scale layer quartzose layer dipping slightly across the face of the outcrop to the west was sampled for U-Pb zircon analysis (Figure 4). The rock is a feldspathic quartzite with abundant detrital? tourmaline grains. Approximately one kilogram of rock was collected and a small number of zircons, ranging in size from 25-150 microns were separated by standard techniques at the Arizona Laserchron Center at the University of Arizona in Tucson. The zircons separated were imaged by scanning electron microscope including both back scattered electron (BSE) and cathode-luminescence (CL) modes (Figure 5).
While zoning and other internal features were common, metamorphic rims are restricted to 1-2 micron euhedral overgrowths on a small population of the zircon present prohibiting their analysis. The shape of grains ranges considerably from nearly perfect rounded grains in two dimensional cross-sections to a wide variety of euhedral and subhedral to subrounded shapes. The grains have substantially difference BSE and CL signatures suggestive of a wide variation in trace element chemistry and multiple sources.
Zircons from the quartzite layer have a complex age distribution (Figure 6). Most grains appear to have been derived from Proterozoic terranes to the west, yielding ages corresponding to the Mid-continent Granite-rhyolite province and Mazatzal, Yavapai, and Penokean orogens, rather than a more limited source to the north in the Grenville Province. Particularly telling are the lack of ca. 2.65-2.75 Ga zircons from the adjacent Southern Superior Province and the relative abundance of 1.80-2.00 Ga zircons compatible with Penokean or Trans-Hudson sourcing. A possible southern source is ruled out by the presence of Early Proterozoic and Archean components and the scarcity of 1.30-1.40 Ga detrital zircons indicative of the Southern Adirondack tonalitic rocks.
The youngest zircons in the rock yield an age of 1257.6±16 Ma. This age yields maximum age of deposition for the Lower Marble between ca. 1241-1273 Ma and establishes its deposition prior to the Elzevirian Orogeny. This result is consistent with the Lu-Hf dating of apatite in the Lower Marble (1274+/-9 Ma; Barfod et al., 2005) and the presence of large isoclinal folds in the Lower Marble crosscut by the Antwerp-Rossie granitoids (1203±13.6 Ma; Chiarenzelli et al., 2010).
While the ultimate origin of these tourmaline-bearing gneisses is still equivocal; here they are undisputedly a reflection of a sedimentary environment yielding copious amounts of tourmaline. This environment punctuated the carbonate dominated shelf sedimentation of the Lower Marble. Here we interpret the rocks as arkosic to mud-rich playa deposits deposited within isolated fault blocks and rift basins during the opening of the Trans-Adirondack Back-arc Basin (Chiarenzelli et al. 2011).

STOP 2. Danburite Location at Russell (The Old Van Buskirk Farm).
Private property, please ask permission from the landowners to visit and collect. $5/person collection fee.
At Russell, danburite is found as compact masses, alternating layers of danburite and diopside and vugs with terminated prismatic crystals in a diopside-rich rock. It is associated with diopside, very rare datolite, dravite, phlogopite, quartz, pyrite, calcite (rare), fluorapatite, marialite, microcline, rare chalcopyrite and bornite, titanite, and tremolite.
Danburite formed as a prograde mineral in the assemblage danburite – diopside – microcline – quartz – scapolite. Dravite developed as a late mineral, probably at the expense of danburite. The metamorphic peak for the area seems to be T 650ºC to 700ºC and P 6 to 7 kbars (Edwards and Essene 1988). Grew (2002) considered that the layered danburite-diopside assemblage probably represents the highest-temperature danburite association ever reported. Danburite veins, cutting the host rock, formed after the metamorphic peak at lower temperature.

STOP 3. Bower Powers Farm, Pierrepont, New York
Private property, please ask permission from the landowners to visit and collect. $5/person collection fee.
The quartz-tourmaline veins from the Bower Powers Farm location occur cross-cutting metasedimentary rocks of Grenville-age. The main minerals found in different trenches and pits in this area are uvite-dravite, quartz, calcite, tremolite, diopside, fluorapatite, phlogopite. A very detailed mineral composition of the veins can be found in Chamberlain and Robinson (2013).
Fluid inclusion microthermometry on quartz from the quartz-tourmaline vein yielded a temperature of 438º – 500ºC and P 3.1 to 3.5 kbars (Kelson et al. 2013). A sample of coarse-grained tourmaline-quartz-calcite pegmatite yielded a substantial number of zoned euhedral zircon crystals (Figure 7) that gave an U-Pb zircon age of 1158.3±3 Ma Figure 8). This age is interpreted as the timing of intrusion of pegmatitic veins cross-cutting the country rock as can observed in the trenches along the stream. Given the lack of fabric in the pegmatites, this result suggests intrusion after, or during, the waning stages of the Shawinigan Orogeny.
Tourmaline crystals (some up to tens of cm in diameter) are black and shortened on the “c” axis. A few comments on the composition of some species and how this relates to the species name could be easily exemplified by the tourmaline from the Bower Powers Farm in Pierrepont. This is a well-known location and the name of the species varied in time from one author to another – uvite (Dunn et al. 1977) or dravite (Grice and Ercit 1993; Dyar et al. 1998; Lupulescu 2008). Some of our recent analyses (unpublished data) indicates dravite is present. The crystals from this location are large and zoned. Our analyses show very little difference between Na+ and Ca2+ in terms of their atoms per formula units at the X-site; some analyses have more Ca2+, others with more Na+. It is very difficult to assign the name uvite or dravite to the tourmaline from this location and would require multiple analyses on different parts of the large crystals from thorough the mineralized area. We feel that for the moment the best choice is not to use the species name, but to utilize the solid solution name uvite-dravite.

STOP 4. Wintergreen Mine Waste Rock Pile, Talcville
Our goal at this locality is to take a look at some mineral associations that are similar with those occurring in the Arnold Pit in Balmat. The Arnold Pit is no longer available for study and we will focus our observations and discussions on the rocks found on the dump of the Wintergreen mine. This location is part of the Balmat-Edwards talc belt, a metasedimentary sequence of rocks containing mainly talc and tremolite, mangancummingtonite, and serpentine. Locally, in the talc-tremolite sequences, some Mn-bearing minerals occur. The Mn-bearing-dravite or -uvite are common and in places they form centimeter wide layers containing more than 50% tourmaline.
From the tailings dump along the Oswegatchie River one can collect pink to purple tremolite (hexagonite), brown, acicular crystals of mangancummingtonite, talc, and Mn-bearing tourmaline. Talc often shows a dark green to bluish color which gave the mine its name. Tourmaline from this location is dravite and contains up to 3% MnO. An uvite crystal collected from the Arnold Pit (Ayuso and Brown 1984) contained 5% MnO.

A4: MESOPROTEROZOIC MAGMATISM OF THE ADIRONDACK LOWLANDS: THE RESULT OF CLOSURE OF THE TRANS-ADIRONDACK BACK-ARC BASIN
Sean P. Regan
Department of Geosciences
University of Massachusetts
W.H.Peck, B.W.Selleck, and M.S.Wong
Colgate University
Hamilton, NY 13346
J.R.Chiarenzelli
Canton NY 13617
INTRODUCTION AND REGIONAL BACKGROUND
The Mesoproterozoic Grenville orogenic belt is one of the world’s largest preserved Precambrian convergent margins, with a North American extent from Labrador to Texas. It records several tectonic, magmatic, and accretionary pulses typically referred to as the Grenville orogenic cycle (Ca. 1.30 – 1.0 Ga; Rivers, 2008). This sequence of punctuated crustal growth is thought to have culminated during a ca. 1.08 Ga terminal collision, which resulted in the final assembly of the supercontinent Rodinia (McLelland et al., 2001). The Adirondack Mountains in New York are a domical uplift forming the southern extension of the Canadian Shield into New York, and expose rocks that formed and evolved during this orogenic cycle. The Adirondack Highlands and Adirondack Lowlands are separated by the Carthage-Colton shear zone (Fig. 1; Selleck et al., 2005). The Adirondacks record sedimentation, plutonism, high-grade metamorphism and deformation during the Elzevirian and Shawinigan orogenies of the Grenville orogenic cycle, and are variably affected by terminal, Ottawan, deformation (Fig. 1b; McLelland et al., 1992; Chiarenzelli et al., 2011c;).
The Adirondack Lowlands are dominated by supracrustal, amphibolite-facies metasedimentary rocks arrayed in NE-SW striking belts (Wong et al., 2011; Baird and Shrady, 2012). The Lowlands contain relatively subordinate, variably deformed igneous intrusives (Carl and deLorraine, 1997; Wasteneys et al., 1999; Peck et al., 2013). In contrast, the Adirondack Highlands are dominated by granulite facies meta-igneous rocks, with minor supracrustal rocks (McLelland et al., 2004). Published data suggest that unlike the Highlands, the Adirondack Lowlands did not experience penetrative Ottawan deformation, and that the rocks of the Lowlands were likely at a higher crustal position during the Ottawan (McLelland et al., 1992; Selleck et al., 2005; Heumann et al., 2006). Thus, the Adirondack Lowlands are an ideal location to study earlier phases of Grenvillian orogenesis because of a lack of pervasive overprinting, such as is seen in the Highlands.
This field trip focuses on the major igneous rock suites that were emplaced into the metasedimentary units of the Adirondack Lowlands. Recent work on these rocks is presented in several contributions (Chiarenzelli et al., 2010b; Wong et al., 2011; Peck et al., 2013), which grew from a Keck Geology Consortium Project led by Colgate University during the summer of 2008. This project produced a new igneous whole rock geochemical dataset from many of the major intrusive suites exposed in the Lowlands. Accompanying these data, Sm-Nd isotopic analyses were acquired along with SHRIMP-RG U-Th-Pb analyses of zircon and titanite (Chiarenzelli et al., 2010b; Wong et al., 2011; Peck et al., 2013). This trip will visit localities where igneous bodies are in contact with surrounding meta-sediments, providing an opportunity to address the deformation history of the metasedimentary rocks. However, the central goal is to document how magmatism evolved during Shawinigan orogenesis, and the relationships between igneous rock geochemistry and regional tectonic history of the Grenville Province.
The Black Lake shear zone represents a major lithotectonic boundary within the Mesoproterozoic basement of New York and Ontario marking the NE extent of ca. 1180 – 1200 Ma subduction and the >1250 Ma ophiolite in the Lowlands (Chiarenzelli et al., 2010b; 2011a; Wong et al., 2011; Peck et al., 2013). This shear zone has recently been interpreted to represent magmatism and the approximate location of the rifted margin of Laurentia prior to Shawinigan orogenesis (Fig. 3; Chiarenzelli et al., 2010b). This interpretation suggests that stratigraphic correlations across the shear zone are possible, and would predict shallower facies assemblages (ie quartzites, psammites) consisting of clastic material to the northwest, grading into carbonate (limestone) and pelitic (mud) assemblages to the southeast. The gross stratigraphy preserved within the Adirondack Lowlands is interpreted to record a period of protracted opening and subsequent basin closure and restriction consistent with a back arc model (Carl et al., 1990; Chiarenzelli et al., 2010b; 2011b)
A maximum age on rifting is not well constrained, despite well understood stratigraphic and spatial relationships. Detrital zircon analyses from quartzites in the Frontenac terrane yield a maximum deposition age of 1306 Ma (Sager-Kinsman and Parrish, 1993). Detrital zircon geochronology three samples from the metasedimentary sequence indicates deposit prior to ca. 1250 Ma (Chiarenzelli et al., 2013), and likely represents silicilastic deposition during basin evolution from opening to closure. Deposition of carbonate and evaporitic rocks of the upper Marble likely occurred in isolated remnants of the collapsing basin restricted from the open ocean immediately prior to Elzevirian orogeneisis. Cessation of extension likely occurred during the onset of the Elzevirian orogeny within the Central Metasedimentary Belt to the NW (ca. 1220-1240 Ma) and stepped to the southeast into the Adirondack Lowlands (Chiarenzelli et al., 2011b).

STOP 1. Route 11 east of Antwerp, NY
This outcrop is one of many exposures of the Antwerp-Rossie suite, and surrounding marble, exposed along this section of US-11. Interestingly, the granitoid appears to cross-cut an isoclinal, recumbent fold in the marble, and contains quartzite xenoliths. However, Shawinigan deformation is generally accepted to postdate 1200 Ma. An important question to consider at this outcrop is the origin of this layering. Is the folding related to deformation around the more competent granitoid, or does the granitoid cross cut an earlier phase of deformation?

STOP 2. The Steers Head locality
The steers head locality is one of the many classic field trip localities in the Adirondack Lowlands. The marble contains graphite-rich layers and dismembered calc-silicate gneisses with well-developed reaction rims along their margins. The reaction rims contain tremolite, diopside, phlogopite, quartz, and calcite, among others. Do you see any other minerals in the reaction rims? Also, try to find the folded tourmaline vein on the north end of the outcrop within the marble.
It is important to note that the ARS has a wide range of SiO2 contents. The rocks seen at Stop 1 and 2 are representative of the more felsic members of the ARS. Samples from this belt have relatively juvenile Sm-Nd systematics with no systematic relationship between εNd with SiO2, suggesting a minimal, if any, signal of crustal contamination.
This outcrop was host to an important stable isotope study by Cartwright and Valley (1991). Stable isotope ratios were measured in both metaigneous and metasedimentary country rocks across the intrusive contact to assess the importance of metamorphic fluid flow; essentially looking for a hydrologic ‘breakthrough curve’ caused by the difference in δ18O between these lithologies. No breakthrough curve is present, and the gradient in δ18O between lithologies contains an inflection point at the contact, with a several meter equilibration gradient from high δ18O values in the marble, to lower δ18O values in the granite. The preservation of these steep isotopic gradients and lack of offset of the inflection point suggest that minimal fluid flow occurred during regional metamorphism (Cartwright and Valley, 1991).

STOP 3. County Road 17 north of De Kalb Junction
The megacrystic Hermon Granite gneiss is a distinctive rock unit throughout the Adirondacks, with coarse K-spar augen. Here, the Hermon gneiss was emplaced into marble. With an age of 1182 Ma (Heumann et al., 2006), all deformation within the Hermon Granite gneiss is likely associated with Shawinigan tectonism. As noted above, this suite of rock shares many similarities to the ARS, but lacks a major Pb anomaly when normalized to primitive mantle and typically plots on field boundaries on tectonic discrimination diagrams (Pearce et al., 1984). Note the strong fabric development at this outcrop. Do the different minerals show varying amounts of deformation or recrystallization? What is the strike and dip of the gneissic layering?

STOP 4. South of Beaver Creek, County Route 7. Park on Childs Road.
BE CAREFUL!!! SMALL SHOULDER!
This outcrop contains meter to sub-meter concordant layers of the Hermon granite gneiss (HGG) within a fine grained sulfidic, pelitic gneiss. Both contain a well-developed, upright foliation, and a moderately plunging stretching lineation. Near the margin of the Hermon granite gneiss, there are sub-meter scale pelitic xenoliths aligned with the tectonic fabric (Figure 10). What is the strike and dip of the major fabric in the metapelite? The granitic gneiss? The foliation is fairly representative of the major structural fabrics throughout the Adirondack Lowlands and is well developed within the Hermon granite gneiss, suggesting that the NE-SW structural grain, predominate throughout the Adirondack Lowlands occurred after the emplacement of the Hermon granite gneiss.
Similarly to the ARS, the Sm-Nd systematics for the HGG form a fairly steep evolutionary curve. They plot between CHUR and contemporaneous depleted mantle at time of crystallization suggesting a juvenile source. The sample analyzed plots entirely within the evolutionary envelope for the ARS.

STOP 5. County Road 76, north of Cooper Rd, south of Turner Rd.
BEWARE ANTS
This outcrop is representative of the Hyde School Gneiss (HSG). The HSG has been interpreted alternatively as having an intrusive, or an extrusive (ash flow) protolith. However, substantial work has confirmed its plutonic origin (McLelland et al., 1992; Wasteneys et al., 1999).This outcrop displays rhythmic layering characteristic of the HSG, with amphibolite lenses to continuous layers that are relatively thin here. On a larger scale, the HSG forms km-scale domical lenses within surrounding metasedimentary packages, and contains a composite igneous/metamorphic fabric that parallels its margin. Medium grained in this locality, the HSG ranges from coarse leucogranite to tonalite, to finer grained aplitic layers. Geochemically, it shares many of the same characteristics as the older magmatic suites. It contains minor enrichments with respect to its LILEs and minor depletions in its HFSE when normalized to primitive mantle (Sun and McDonough, 1989), and plots from alkali-calcic to calc-alkalic with constant MALI with increasing SiO2. Furthermore, unlike the ARS and Hermon Gneiss, the HSG samples plot as largely ferroan with regard to their Fe-index, with a subset plotting in the magnesian field (Peck et al., 2013; Frost and Frost, 2008).

STOP 6. Route 58 north of junction with Acres Road outside Pope Mills, NY.
The youngest igneous suite within the Adirondack Lowlands is a series of ferroan rocks collectively referred to as the Edwardsville pluton (Buddington, 1929). Ranging from syenitic to subordinate gabbro, this suite of rocks is both texturally and compositionally heterogeneous. Recent single grain SHRIMP analysis (Peck et al., 2013) has corroborated older multi-grain TIMS work (1164 Ma; McLelland et al., 1993), which yield an age of 1149 ±22 and 1161±16 Ma from a monzonitic member of the Edwardsville pluton and a granite from the Honey Hill syenite, respectively (Peck et al., 2013). These ages correspond to anorthosite-mangerite-charnockite-granite (AMCG) magmatism within the Adirondack Highlands (McLelland et al., 2004; Hamilton et al., 2004), and the similarly ferroan magmatic event in the Frontenac terrane to the northwest (1180-1160 Ma).
Although texturally similar to the Hermon Gneiss, can you see a mineralogical difference? Is there a strong tectonic fabric? Lastly, is this outcrop homogeneous, or are there texturally distinct rocks types, and, if so, what is the relationship between them?

STOP 7. Rossie Diorite, Split Rock Road.
Very close to private property, please be courteous!
The Rossie diorite represents a more mafic component of the ARS suite. It is important to note that we are very close to the Black Lake shear zone, an important structural and lithologic discontinuity separating the Adirondack Lowlands from the Frontenac Terrane (Wong et al., 2011). Although slightly more mafic than its southern counterparts, the Rossie diorite displays similar geochemical signatures. However, despite these similarities, the belt of dioritic rocks exposed along the eastern margin of the Black Lake shear zone contain the lowest εNd(T) values, and yield the oldest TDM ages, approximately as much as 400 My older than ARS samples analyzed further south. These data suggest a major source of contamination by, or interaction with, older more evolved continental lithosphere (n=3 samples along the northwestern extent of the ARS).

STOP 8. Rockport Granite; US 12.
These exposures are on the NW side of the Black Lake shear zone (Wong et al., 2011). We are now in the Frontenac Terrane. This outcrop exposes the voluminous Rockport granite (ca. 1170 Ma; Wasteneys et al., 1998), which extends across the Thousand Islands region and into Ontario, Canada. Here, the Rockport Granite is in contact with strongly deformed calc-silicate (metasedimentary) gneisses. There is a faint, near-vertical gneissic layering within the Rockport Granite that is shared by the adjacent calc-silicate gneisses. This suite of granitic rocks is time equivalent to the domical Hyde School gneisses in the Adirondack Lowlands. However, the Rockport Granite has a much different intrusive style, and is hosted most commonly by quartzite and quartzite-calc-silicate country rock. The Rockport granite is generally slightly to moderately deformed, and doesn’t occur as strongly deformed domical bodies like its Hyde School gneiss counterpart. Furthermore, no exposures of ARS or Hermon Granite gneiss have been identified on this side of the Black Lake shear zone.

A5: GEOLOGY OF THE BALMAT ZINC REGION
William deLorraine
St. Lawrence Zinc Co. LLC,
408 Sylvia Lake Rd.
Gouverneur, NY 13642
The Balmat-Edwards district zinc orebodies are located in the northwest Adirondack Lowlands of northern New York State (Fig. 1). The Adirondacks are an extension of the Canadian Shield into New York State via the Frontenac Axis/ Thousand Islands region located along the St. Lawrence River and are underlain by poly-deformed rocks metamorphosed to a high grade during the Grenvillian orogeny about 1.1 bya. Zinc orebodies of the Balmat district occur in an eight mile long strip of rocks comprised of dolomitic marbles and calc-silicate rocks referred to as the Balmat-Edwards marble belt. The Balmat-Edwards marble belt cores a nappe-scale, doubly plunging, recumbent fold referred to as the Sylvia Lake syncline (Fig. 1). Lithologically similar silicated dolomitic marble host rocks, sulfide ore, and structure reappear 28 miles/45 km to the northeast of Balmat at Pierrepont, NY (Fig.1).
On this field trip we will make a NW-SE transect across the Lowlands beginning in Gouverneur and ending at Balmat, NY. We will examine, in a general sense, the structural architecture and tectono-stratigraphic setting of what has been referred to as the “Trans Adirondack Back-Arc Basin” (TABB) comprising the NW Adirondack Lowlands between the Highlands and the Black Lake region (deLorraine and Carl, 1993; Chiarenzelli et al., 2012). As we progress from Stop 1 to Stops 2 and 3 we’ll gradually move up section.
Stops 1 and 2 in calcitic lower marble record rift and drift phase chemogenic sedimentation in the (TABB) whereas up section at Stop 3 Popple Hill gneiss records basin fill turbiditic sedimentation initiated by convergence (Chiarenzelli et al., 2012). Stop 4 takes us to dolomitic upper marble at the top of the section. Host rocks record cyclic dolomitized and silicified algal mat carbonate and evaporite deposition in shallow restricted marine seas during the final stages of back-arc basin fill. Overall objective will be to take a look at the different high grade metamorphic rock types and mineral assemblages, styles of folding, relate minor folds to major folds, and to gain an appreciation for why the tectono-stratigraphic setting of the NW Adirondack Lowlands as a back-arc basin might be the ideal setting for the formation of SEDEX ore deposits such as the Balmat-Edwards-Pierrepont deposits
Zinc orebodies are believed to have originated as Mesoproterozoic seafloor exhalations that accumulated as conformable sedimentary massive sulfide lenses and layers deposited some 1.25 bya in the TABB within a sequence of dolostones, thinly layered stromatolitic cherty dolostones, and evaporites (Fig. 4). Sphalerite and pyrite accompanied by very minor galena are the principal sulfide minerals. Host rocks and orebodies attained burial depths of 20 km or more during Grenvillian orogenesis when they underwent polyphase deformation. Upper Amphibolite grade peak metamorphic conditions generated migmatites and anatectic melts in surrounding rocks and mobilized significant sulfide fractions within and away from the orebodies. In a few places, sheet-like, aerially extensive, nearly pure sphalerite “dikes” segregated from their parent lenses and intruded macrofracture surfaces where these large cracks intersected parent sulfide masses. Macrofractures evolved into thrust faults at a high metamorphic grade during deep-seated regional ductile shearing. Sheetlike “daughter” orebodies thus formed are found on three major ductile fault surfaces that are refolded by the Sylvia Lake syncline. Most of the intruded sphalerite segregated along the fault surfaces into tabular lozenges elongated in the direction of associated fold hinges. A typical daughter orebody “footprint” is several hundred feet wide by 6000 ft/1800 m long by perhaps 1.8-2.5 m thick on average. Tonnages average up to 6 million.
Daughter orebodies are inherently lower grade and finer grained because of their locations on fault surfaces making them inherently more susceptible to recurrent deformational strains. Ductile faulting abraded and rounded admixed adjacent wall rock inclusions in the ore thereby imparting their signature deformational texture referred to as “Durchbewegung” texture. “Parent” ore lenses from which the daughter orebodies are derived can be 1500 ft/450 m or more in length, 450ft/150 m wide, and up to 80ft/25 m or more thick. Parent ore is coarse grained and twice the tenor of daughter ore. Presently there are several daughter orebodies in production or have been drilled whose predicted source bed or parent massive body awaits discovery.
Parent orebodies originally deposited in more ductile carbonate lithologies that flowed rather than faulted during deformation did not necessarily give rise to daughter offshoots. Rather, those strata-bound sulfides remobilized within and parallel to fold hinges associated with the Sylvia Lake syncline. The largest of these orebodies, the Main and Hanging Wall, are located at No. 2 mine within the core of the Sylvia Lake syncline in Unit 14 (Fig. 3).
Individual orebodies exhibit good continuity and are quite predictable in the down-plunge dimension, though the details of host structure change and evolve from section to section in the mine. This leads to a variety of dips of ore shoots from flattish to vertical hence requiring an array of different mining methods, from inclined room and pillar to long-hole extraction.
Multiple lines of evidence suggest differentiation of sphalerite from stratiform massive sulfide ore lenses occurred during the Upper Amphibolite grade Shawinigan (Grenville) Orogeny (ca. 1.20-1.15 Ga) prior to development of the host Sylvia Lake Syncline. Syntectonic macrofractures impinged upon massive lenticular, polymetallic stratiform orebodies forming subsidiary ores as sheet-like radial dikes. Several of the monomineralic sphalerite daughter orebodies remain linked to parents by thin, refolded, “durchbewegt” ore sheets that cross-cut stratigraphy and metamorphic fabrics (Fig. 3, daughter orebodies L and lg linked to UG near No. 4 mine). Long after intrusion of daughter sphalerite dikes, admixed xenolithic fragments of wall rock contained within the initially massive, coarse-grained, sphalerite became rounded by tectonic milling resulting from recurrent strains partitioned along macrofracture surfaces. Daughter orebodies thereby acquired distinctive durchbewegung textures while nearly doubling their thicknesses but halving their tenors as macrofractures evolved into tectonic slides -- some with km-scale displacements.
Polyphase deformation culminated with superposition of the nappe-scale Sylvia Lake syncline accompanied by refolding and inversion of parent-daughter ore complexes on its upper limb. This complex fold displays elements of both passive flow and concentric geometries due to extremes of ductility contrast induced by intercalation of marble, exceedingly ductile anhydrite strata, and competent quartz-diopside-rich units. Long thought to have been the driving force or “tectonic engine” responsible for large scale sulfide mobilization to dilatant sites within its major and minor fold hinges, the Sylvia Lake syncline is now known to refold durchbewegt “daughter” ore sheets and thus is a late stage feature in the evolution of orebodies of the district. Mobilization of originally disseminated sulfides from throughout the Upper Marble section to dilatant structural sites either within fold hinges or along syn-kinematic faults via metahydrothermal fluid phases is precluded by direct linkage of daughter ore to parent source beds. As well, the emplacement of daughter mineralization along synmetamorphic faults that transgress metamorphic foliations and recrystallized bedding suggests emplacement after cessation of major devolatization reactions under “dry” ambient metamorphic conditions.
Geochemical groupings of orebodies are consistent with three “SEDEX” horizons (Fig. 4). Significantly, three major epochs of evaporite deposition in that interval coincide with the three epochs of strata-bound sulfide deposition such that anhydrite sedimentation shortly precedes sulfide deposition in each case. Resulting from this are three “evaporite—ore” couplets within units 6, 11, and 14 (Fig. 4). Stratiform massive sulfide source bed lenses or layers specifically occur in those units. The association of evaporites and ore is an indirect, but important one that is common to other important sulfide deposits.
Deciphering the origin of the Balmat zinc deposits is made somewhat problematic by the effects of high grade metamorphism, polydeformation, and post-metamorphic annealing recrystallization. Nevertheless, the following scenario can be pieced together from petrological, geochemical and isotopic data, underground mapping and photographs, deposit-wide and regional structural geology, and deposit and regional stratigraphy. Orebodies fall into three geochemical groupings, each of which appears in the section not long after major evaporative epochs. Orebodies lowest in section are highest in Hg and vice versa, consistent with the interpretation that earliest SEDEX orebodies precipitated from saline brines that stripped most of the available Hg from the volcano-sedimentary pile in the Back-Arc basin repository during the first SEDEX cycle. In so doing less and less Hg was left for successive ore-forming cycles. This interpretation places evaporites immediately available to SEDEX ore fluids as a source of saline brines capable of leaching metals and presumes depletion of halide reservoirs after each SEDEX depositional epoch. Consequently SEDEX ore deposition is precluded until after the end of the next evaporative sedimentation epoch pending restoration of anhydrite and halite reservoirs.
An interpretation consistent with available data is that Balmat Zn deposits are ultimately of late Proterozoic, vent-distal, carbonate-hosted, SEDEX origin and that metamorphogenic “Daughter” orebodies spawned from massive “Parent” orebodies during Upper Amphibolite grade metamorphism as differentiated offshoots in the form of sheetlike sphalerite “dikes”. Local preservation of flow banding, sulfur isotope gradients, color gradations in sphalerite and mineralogic differentiation ZnS+Py>ZnS>PbS in the direction of transport suggests that plastic flow or fluid-aided plastic flow was the dominant mechanism of mass transfer. Distances of migration are truly epic with readily demonstrable, km-scale, cross-stratal flow.

STOP 1. Scapolite marble roadcut, lower marble.
Dips to south at end of Reservoir Hill leucogneiss dome. Note black, N60ºE trending, late dike. Near the base of the lower marble section. Scapolite may be evidence for early evaporative chemogenic sedimentation in the Trans-Adirondack Back-Arc Basin (TABB).

STOP 2. “Train Wreck” road cut, lower marble.
“Graph-phlogo-cal” marble with brown tourmaline. Stratigraphically overlies the scapolitic marble. “Rift and Drift” phase of development of the TABB. Large blocks represent dismembered mafic dike fragments rotated and “strewn” about in the marble like a derailed train. This is the marble quarried extensively in the late 1800’s to early 1900’s making Gouverneur famous as “Marble City”. It is a prominent local building stone which is a major component of churches, post offices, and other buildings in the region. 1274 ± 9 Lu-Hf .

STOP 3. Popple Hill migmatite gneiss road cut.
Plag-qtz-biot gneiss; ca 1220. Migmatite-agmatite roadcut. Look for refolded migmatitic leucosomes, ptygmatic folds, folds with “S” asymmetric sense appropriate for structural position on upper limb of the Sylvia Lake Syncline and many that aren’t... ? Basin Fill stage metasediments; here a poly-deformed migmatite-agmatite exposure. 1180-1160ma anatectic zircons. Local sill-gar-biot gneiss layers; rare carbonate seams. Currently interpreted as immature turbiditic sediments, plag-rich wackes, qtz-arenites to Al-rich shales fining upward toward contact with upper marble as the back arc basin evolved into a foreland basin. Sourced from the Southern Adirondack terrane (Chiarenzelli et al, 2012).

STOP 4. Units 4/5 Contact, No. 4 Mine.
Upper marble Ca. 1210 ma. Upper limb of the Sylvia Lake syncline. Contact between silicated Unit 4 and overlying dolomitic marble, Unit 5. Unit 4 is stromatolitic. How are they identified? Look for laminated qtz-diopside rocks locally retrograded to serp-talcose-diop. What is stratigraphic “tops”? How do you think stroms formed and what might be their mode of preservation? How did they survive Upper Amphibolite grade metamorphism and polydeformation? “Footprint” of one inverted strom slightly elongated N10º W parallel to major and minor fold hinges.

STOP 5. Langevin Lane (private drive).
Turn right, drive up paved driveway to top of hill.
Private residence. Permission required. South end of Sylvia Lake, viewing N. Sweeping northerly vista of the lake with Balmat No. 4 mine headframe in right background. Topographic expression of the lake is a reflection of the geometry of the hourglass-shaped core of the Sylvia Lake syncline. Filled with Unit 11A anhydrite, the core of the fold was easily eroded so that the shape and great depth of the lake of over 150 feet in places is attributable to this unusual lithology.

STOP 6. Turn left onto short, dead-end paved road segment
Leading downhill toward abandoned open pit mine.
Permission Required. Gouverneur Minerals Co. (formerly referred to as Gouverneur Talc Co); Vanderbilt Corp. Exposed in the east sidewall of the pit facing Rte. 812 is a spectacular isoclinal fold hinge in Unit 15 “Mica Hanging Wall” in the core of the Sylvia Lake syncline. Look for axial planar schistosity and transposed bedding in this tightly appressed hinge. Good examples of boudinage. Are the long axes of the boudins parallel to the fold axes?

STOP 7. Gate to No. 2 mine.
(Gate is locked, access by appointment only St. Lawrence Zinc Co.).
We’ll proceed into the No. 2 mine grounds, past the inclined shaft and hoist house, to the No. 2 open pit. The shaft and hoist are kept in good running condition on a stand-by basis serving as a second way out, escape route for the No. 4 mine. High tenor zinc ore is exposed in the pit along with numerous examples of high temperature cross-cutting relationships with host rocks. Numerous minor folds, examples of surficial oxidation of the massive sulfide ore resulting in “gossan”, examples of bird’s eye alteration texture, numerous different rock types and mineral parageneses. A suite of high grade carbonate metamorphic minerals for the student mineralogist and retrograde assemblages as well.

STOP 8. St. Lawrence Zinc Co., No. 4 Mine entrance.
Turn left into parking lot. Stop to view mine models, maps, core logging facilities, diamond drill core displayed on core racks, ore specimens, recap trip and discuss genetic models.

A6: THREE AND A HALF SKARNS
George W. Robinson
Research Associate
Canton, NY 13617
Steven C. Chamberlain
Center for Mineralogy
3140 CEC, New York State Museum,
Albany, NY 12230
The Central Metasedimentary Belt (CMB) of the Precambrian Grenville series rocks of St. Lawrence, Jefferson and Lewis Counties, New York has been the source of a wide range of high-quality mineral specimens for over 150 years. Most of these localities occur in the Grenville Marble and associated calc-silicate rocks, as well as in skarns or skarn-like assemblages, with or without an obvious intrusive igneous source that supplied the heat, pressure and fluids required to make the skarn. Some of the latter have been called “vein-dikes,” and their origin is not well understood. Recent research suggests their calcite cores have both a mantle-derived carbonatitic component as well as a crustal derived marble component (Sinaei-Esfahani, 2013). This trip will visit sites hosting both types, and includes four stops:
Stop 1 – Bradley Farr property, Natural Bridge, Lewis Co.
Stop 2 – Gouverneur Minerals’ open pit wollastonite mine, Lake Bonaparte, Lewis Co.
Stop 3 – Rose Road Skarn (Purple Diopside Mound) adjacent to Rose Road skarn
Stop 4 – Rose Road Skarn (Mulvaney property) near Pitcairn, St. Lawrence Co.
Stops 1, 3 and 4 are best known to mineral collectors for having produced high-quality specimens of titanite, diopside, scapolite, zircon, albite, microcline, and other minerals. Wollastonite is present at three of the four sites, and is commercially mined at Stop 2. At stops 1 and 4 wollastonite occurs as euhedral crystals, some of which have been replaced by quartz, calcite and pyroxene, resulting in sharp pseudomorphs. Descriptions of these sites, their collecting histories, and major minerals of interest are summarized below and treated in greater detailed by Chamberlain et al. (1987, 1999, and 2013).

A6: THREE AND A HALF SKARNS

STOP 1. Classic Dana mineral locality for wollastonite, titanite, scapolite, diopside, etc.
This stop is famous for its well-formed crystals of wollastonite that occur in calcite-filled veins in syenite, associated with crystals of titanite, meionite, microcline, diopside and zircon. Known prior to 1840, it is among the earliest of American mineral localities, and has been referred to as the Cleveland, Ashmore, or Whitestone farm, depending on when reference is made. It is also commonly referred to as simply Natural Bridge, Diana, Lewis County, NY, and is the type locality for the two now discredited species lederite (= titanite) and nuttallite (= meionite).
GEOLOGY
The two principal rock types present at the locality are Grenville Marble and a gneissic pyroxene syenite of the Diana Complex. Previous investigators all conclude that the observed skarn assemblage formed when the syenite intruded the marble (Smyth and Buddington, 1926; Agar, 1923). However, what first appears to be a simple contact metamorphic skarn, upon closer examination also shows characteristics of calcite vein-dikes (Mills, 2014; Moyd, 1990). There can be little doubt that mineralogically, the observed assemblage is typical of a skarn, but if it formed by the syenite intruding the marble, then one should expect to see additional skarn development at the syenite – marble contacts in the immediate area, rather than being confined solely to the “veins.” The syenite in contact with the skarn assemblage developed on the walls of the veins shows a marked cataclastic texture that would have provided a preferential pathway for the mineralizing fluids to enter and form the minerals observed. Similar skarn-like deposits occur throughout the CMB of the Grenville in Canada, and have recently been interpreted as having formed from a mixed crustal and mantle-derived fluid based on carbon and oxygen isotope studies (Sinaei-Esfahani, 2013). Preliminary carbon – oxygen isotope analyses of the calcite from the Natural Bridge locality yield similar results (δ13Cpdb = -1.1 ‰, δ18Osmow = +16‰), but additional isotopic and trace element analyses, as well as radiometric age determinations of both the skarn and syenite will be required to resolve which (if either) model is correct. Unfortunately, detailed mapping of the area is greatly hampered by copious glacial overburden and vegetation.
MINERALS
The brief descriptions that follow are for the more important species only. For more complete information on these as well as additional minerals present, see Chamberlain et al. (1987).
Diopside occurs as dark green to black prismatic crystals to 10 cm associated with white microcline, titanite, wollastonite and meionite. While sometimes referred to as “augite” microprobe analyses show the dark pyroxene crystals are a diopside – hedenbergite solid solution series with approximately Di75% - Hd25%.
Meionite forms equant gray-white crystals to 5 cm associated with diopside and microcline. Sometimes the meionite has a distinctive gray-blue surface coating of a K-Al-silicate (probably microcline) that was originally described as the now discredited “species” nuttallite.
Microcline is perhaps the most abundant of the well-crystallized minerals in the skarn assemblage, and occurs as blocky, somewhat rounded, white crystals several cm across.
Titanite occurs in sharp, dark brown, wedge-shaped crystals to 7 cm, though most are much smaller. One of the crystal habits of titanite from this locality was originally described as the species lederite (Shepard, 1840), but has since been shown to be titanite.
Wollastonite forms elongated, pinacoidal, cream-white crystals to nearly 30 cm in length, making them among the largest and best examples of the species known.
Zircon forms small, pink, prismatic crystals seldom exceeding 2 cm in length. They are not common.
Retrograde minerals. Several of the species described above show evidence of partial to nearly complete replacement by other minerals, resulting in a variety of pseudomorphs. These are thought to have formed as a result of retrograde metamorphism, and include diopside coated by ferroactinolite, titanite by rutile and amorphous Ti-oxide, and wollastonite by quartz, calcite, diopside and hedenbergite.

STOP 2. Gouverneur Minerals No. 4 (Valentine) wollastonite mine.
Originally known as the Valentine property, this active wollastonite mine commenced operations in 1977, and is today operated by the Gouverneur Minerals Division of Vanderbilt Minerals, LLC, producing commercial grade wollastonite which is used primarily in the ceramics and plastics industries. In addition to the wollastonite, other minerals of interest include abundant blue calcite, graphite, prehnite, quartz, and the rare species orientite and brewsterite-Ba, the latter for which it is the type locality.
GEOLOGY
The main orebody at the mine consists of a wollastonite skarn developed along the contact of Grenville Marble with a metasyenite of the Diana complex. The wollastonite occurs as coarse interlocking creamy white crystals 1- 6 inches long, locally cut by veinlets of fine-grained, fibrous wollastonite and veins or pods of massive, pale yellow prehnite. It is thought that the skarn resulted from silica-rich hydrothermal solutions reacting with the silica-poor marble under relatively low pressure, forming wollastonite (CaSiO3). Subsequent retrograde alteration is responsible for epidote and quartz replacing the metasyenite and wollastonite, the prehnite and secondary wollastonite veins noted above, and calcite veins cutting the metasyenite and marble. Open cavities in some of the prehnite and quartz veins host a variety of crystallized minerals, the more interesting of which are described below. For an in-depth discussion of the geology of this deposit, as well as a more detailed coverage of the 30 some mineral species that have been found here, the reader is referred to Chamberlain et al. (1999), Gerdes (1991), Petersen and Totten (1993) and Gerdes and Valley (1994).
MINERALS FOUND IN THE MARBLE
Calcite. Ton quantities of pale-to-dark blue intergrown cleavages of calcite constitute much of the marble, particularly along the eastern edge of the skarn. The blue color is thought to be due to the development of radiation-induced lattice defects in sheared calcite.
Graphite is found throughout the marble units, where it occurs in a variety of habits varying from small euhedral hexagonal plates to unusual spherical aggregates several millimeters across. The hexagonal crystal plates occasionally show remarkable spiral growth features on their {0001} pinacoids.
Wollastonite is uncommon as a “secondary” mineral, yet it has been found here clearly as a late-stage species in dense, fine-grained fibrous veinlets in the earlier-formed wollastonite and marble skarn. This material may resemble a fine-grained quartzite in hand specimens, and fluoresces a powder blue color with an intense, and long-lived phosphorescence (afterglow) in long-wave UV light.
MINERALS FOUND IN CAVITIES IN MASSIVE PREHNITE
Apophyllite group minerals occur as chalky white tapered pyramidal crystals generally less than 2 mm in length. Electron microprobe scans show they are compositionally zoned with respect to fluorine, and both fluorapophyllite and hydroxyapophylite may be present within a single small crystal. Rarely fluorapophyllite cleavages up to 2 cm across have been found.
Brewsterite-Ba is a rare member of the zeolite group that was first noted here in the late 1980s, but not formally described until 1993 (Robinson and Grice, 1993). At the same time the mineral was also found and described from Italy (Cabella, et al. 1993). Both occurrences have been referred to as “type localities.” At the No. 4 mine it occurs sparingly in some cavities in the massive prehnite as aggregates of colorless, tabular crystals 1-3 mm across. It is easily found using a shortwave UV lamp, since it fluoresces yellow-green.
Datolite is an uncommon mineral in some of the prehnite veins, but has been found in glassy, yellow-green crystals to 1 cm.
Hedenbergite is relatively common in some of the prehnite veins, where it forms elongated dark green prismatic crystals1-2 mm long. Diopside is also present, but tends to be much paler in color.
Microcline is a common accessory mineral in the prehnite cavities. It’s most common habit is small, white, adularia-habit crystals, often with serrated edges.
Orientite is a rare species, and is included here solely because of that fact. It is known from only six other localities, worldwide. At the No. 4 quarry it forms brush-like aggregates of acicular brown crystals to 1 mm as well as smaller, granular, rounded red-brown coating prehnite and babingtonite. To date, only two specimens are known (see Chamberlain et al., (1999) for a more complete description).
Pectolite occurs as micro tufts of white acicular crystals, which are practically indistinguishable from the more common wollastonite without X-ray or chemical data.
Prehnite crystals 1-5 mm across occur as pseudo-octahedral, dipyramidal, pale yellow-green crystals lining cavities in the massive prehnite.
MINERALS OF THE NORTHWEST ALTERATION ZONE
Albite. In June, 1995, a large brecciated quartz vein was encountered while building a ramp along the northwest boundary of the skarn. While quartz was the predominant mineral present (q.v.), there were others worthy of mention, among them albite, which occurred as twinned, white, crystals to 1 cm associated with acicular microcline and reddish brown, botryoidal andradite on quartz.
Andradite occurs as coatings of red-brown microcrystals on quartz.
Calcite occurs as gray scalenohedral crystals up to 2 inches long, frequently with red and black inclusions of hematite and goethite, making interesting and attractive specimens.
Epidote crystals form aggregates up to a foot across in some of the larger pockets encountered.
Quartz is the mineral of primary interest in the northwest alteration zone, where it forms colorless to milky prismatic crystals several inches in length, and crystal groups several times that size. Some of the crystals have a dusting of hematite, giving them a pinkish hue, while others have a wide variety of hematite and goethite inclusions, resulting in an equally wide variety of appearance. Amethyst is also present, but less common.

STOP 3. Rose Road - Purple Diopside Mound near Pitcairn, NY
Although the mineralization at this stop was only recently discovered, an interesting variety of specimens has been recovered in the past two years. Despite being a few hundred yards from the classic wollastonite skarn, the purple diopside mound has a very distinctive suite of minerals of different age and origin from the skarn. Large quantities of well-crystallized scapolite, fine-grained masses and crystals of purple diopside, black titanite, and transparent phlogopite are noteworthy. The discovery of lesser amounts of pink corundum and pink spinel, secondary natrolite and prehnite, and still unidentified “gieseckite” pseudomorphs, have all added interest.
GEOLOGY
Excavations have revealed an occurrence of calcite-filled “vein dikes” typical of regional metamorphism in the Grenville Central Metasedimentary Belt. The zircon age of the vein-dike mineralization is 1170 Ma (Chamberlain et al., 2014), making it slightly older than the nearby wollastonite skarn. The principal veins are lined with scapolite crystals facing into a coarse calcite center. At places black titanite is an accessory mineral. Locally significant masses of acicular natrolite crystals coat the scapolite. Other parts of the mound have silicate-marble contacts with purple diopside crystals and “gieseckite” crystals developed at the contact. Phlogopite crystals are developed in the marble near the silicate-marble contacts. Spinel occurs as a rare accessory with the phlogopite in marble. Pink corundum occurs as anhedral masses in scapolite several centimeters away from the scapolite crystals developed facing into the vein. The association of corundum with fluorescent scapolite is noteworthy and appears paragenetically similar to the recently discovered gem sapphire occurrences in the Kimmirut area, Baffin Island, Nunavut, Canada (Cade et al., 2005; Lepage and Rohtert, 2006).
MINERALS Since no detailed description of the minerals identified from this occurrence has yet been published, a summary of the complete list follows.
Albite occurs as white euhedral crystals to 1 cm and as white masses in purple diopside. Corundum occurs as bright purplish-pink, anhedral masses to 25 mm in meionite on the margins of the scapolite vein dike.
Diopside occurs as grayish purple prismatic crystals to several cm and as fine-grained masses. The composition includes about 1 wt % FeO and 1 wt % TiO2, which may account for its unusual color. The dark green diopside crystals, the yellow acicular diopside formed as an alteration of wollastonite, and the dark blue-green actinolite overgrowths at the adjacent wollastonite skarn all contain significantly more iron and no detectable titanium (by SEM/EDS).
Fluorapatite forms bright blue-green transparent prisms to 8 cm in the marble.
“Gieseckite” occurs as sharp euhedral pseudomorphs to several cm associated with graphite rosettes. These pseudomorphs have a complex internal replacement texture and variable composition. Some have zones of fine-grained plagioclase (labradorite). The original mineral has not been found nor inferred.
Graphite occurs as sharply developed crystals and rosettes to several mm, most often associated with “gieseckite”.
Natrolite is present as radial aggregates of coarse, white, crystals to 5 cm and also as masses of acicular crystals in localized portions of the scapolite vein. This is a sodium zeolite with little detectable calcium that appears to have formed from the breakdown of scapolite, similar to other zeolite minerals found in similar “skarn” assemblages in the Grenville CMB (VanVelthuizen, et al., 2006).
Phlogopite occurs as honey-brown, transparent crystals to 1 cm and as masses of 20 cm or more in marble. Rarely, pseudomorphs of phlogopite after scapolite occur in the scapolite vein.
Prehnite occurs sparingly as brown botryoidal masses to 4 cm on a matrix of calcite, graphite, and pyrite. Like natrolite, it probably formed from the breakdown of scapolite.
Pyrite occurs as modified cubes to 5 mm, and as minute crystals often altered to goethite in “gieseckite”, prehnite and scapolite. It also forms oriented arrays of acicular crystals embedded in meionite.
Scapolite occurs as sharp crystals to 20 cm on the walls of the scapolite vein. These crystals have intermediate compositions close to the middle of the meionite (calcium-rich)-marialite (sodium-rich) isomorphous series. Generally glassy gray scapolite is slightly enriched in calcium and is meionite. Sometimes a rind of white, translucent scapolite coats the crystals and is slightly enriched in sodium (marialite). Most of the volume of scapolite is meionite.
Spinel occurs as pink octahedral crystals to 1 mm in marble associated with phlogopite.
Titanite occurs as tabular, almost black crystals to 5 cm associated with scapolite and also as minute crystals of complex morphology embedded in albite.
Zircon occurs as minute euhedral crystals embedded in meionite.

STOP 4. Rose Road - Wollastonite Skarn near Pitcairn, NY.
This classic locality (also known as the Mulvaney sugar bush) has been a source for well-formed crystals of wollastonite, titanite, diopside and albite for about 125 years. Williams (1885), Luquer (1893), and Ries (1893-94, 1895) all published reports on various aspects of the mineralogy of this site. All of the various species reported from this site (Chamberlain and Robinson, 2013) are still collectible today
The locality is a wollastonite skarn formed by the emplacement of the Diana Complex (estimated at ∼ 1164 ± 11 Ma (Hamilton et al. 2001)). It is a secondary skarn in the sense that it was apparently formed by solutions from the Diana Complex invading fractures along a contact between the Grenville Marble and a metamorphic albite-diopside rock. The age relationships between the skarn formation and the regional metamorphism of the Adirondack Lowlands are somewhat confusing, given the zircon age we report for the nearby vein dikes at Stop 3. Significant alteration of the wollastonite and minor uralitization of the diopside have occurred subsequent to the original mineralization.
MINERALS
A summary description of the minerals is provided below. For more details see Walter et al. (2009) and Chamberlain and Robinson (2013).
Actinolite occurs relatively rarely as a blue-green coating on diopside crystals. It contains 7.5 wt % FeO which is significantly more iron than is found in the diopside.
Albite occurs as equant or elongated white blocky crystals to 8 cm. Both Carlsbad and Manebach twins have been found (Chamberlain, 1985; Walter, 2009; Chamberlain and Robinson, 2013). The perthitic texture is due to minor exsolution lamellae of microcline making this an antiperthite.
Diopside is perhaps the best known species from this locality, and forms lustrous dark green prismatic crystals to 20 cm. It contains about 2.5 wt % FeO. Diopside also occurs as a granular replacement of wollastonite.
Fluorapatite occurs as prismatic crystals to 8 cm in colors ranging from dark blue to bright blue to blue-green.
Goethite usually replaces pyrite, forms brown stains on the adjacent albite and occasionally forms small clusters of microscopic needles.
Graphite rarely forms small tabular crystals to 2 mm. It is more common as gray inclusions in albite.
Pyrite occurs as microscopic crystals on albite. Most have completely altered to goethite.
Quartz occurs as granular masses and white to colorless rounded crystals to 1 cm replacing wollastonite.
Titanite occurs as tabular, clove-brown crystals to 12 cm. Some of the finest titanite crystals found in New York have come from this skarn. Most of the larger crystals have somewhat rounded faces and prominent parting.
Wollastonite occurs as white well-developed, euhedral crystals to 35 cm. These are almost always altered, at least on the surface. Pale brown alteration consists of acicular yellow diopside crystals containing 5 wt % FeO intermixed with calcite formed from the breakdown of wollastonite. Greenish, granular alteration is mostly diopside. White granular alteration is quartz. Green and white patchy altered wollastonite crystals are typical (Chamberlain et al. 2009).

B1: BLACK RIVER AND TRENTON GROUPS, NORTHWESTERN NEW YORK STATE
Bruce W. Selleck
Colgate University
Hamilton, NY 13346
Black River and Trenton Group strata in eastern North America record the last phases of deposition on the “Great American Bank” (Sloss, 1963), the long-lived Cambrian through late Ordovician cratonic-passive margin carbonate depositional system. Carbonate deposition ended with tectonically-driven foreland basin deepening, coupled with the influx of terrigenous clastics from Taconic Orogenic source lands to the current east of the continental interior. The change from carbonates to clastics was diachronous across the basin in the New York State region, with Utica Formation muds derived from the orogenic hinterlands progressively flooding the platform from east to west. This facies transition, and the structural framework of basin development linked to collisional mountain-building, were critical to the plate tectonic model elucidated by Bird and Dewey (1970), which was among the very first studies to document the link between plate collision and basin development on ancient continental margins.
The Black River and Trenton Groups have also been central in the historical development of our understanding of the relationships of stratigraphic patterns such as lateral facies changes, stratal continuity and chronostratigraphic correlation, and abrupt vertical contacts, to basin dynamics, including tectonic, eustatic and climate forcing factors. Following the early descriptive work of Vanuxem (1838, 1842) and Hall (1847), subsequent paleontological (Young 1943a, b) and stratigraphic analyses (Kay, 1939, 1960; Fisher, 1962, 1977, for example) documented large-scale facies relationships, biostratigraphic correlations, and established a workable stratigraphic nomenclature. Sedimentological analyses and actualistic facies models were also developed, including, for example, the work of Winder (1960), Textoris (1968), and Walker (1973). Integrated paleoecological-stratigraphic studies (Cameron and Mangion, 1977; Titus and Cameron, 1976; Brookfield, 1988; Brookfield and Brett, 1988), and continued refinement of graptolite and bentonite-based regional correlations (e.g. Mitchell, et al, 1994, 2004) provided important new understandings of the Black River-Trenton-Utica system. Within the last 20 years, revision of the global time scale has resulted in assignment of the Black River-Trenton-Utica sequence to the Upper Ordovician, whereas in older literature, these rocks were considered Middle Ordovician. This sequence is placed within the Caradoc Stage, with deposition spanning the interval 457-449 million years ago.
In the last two decades, the framework established formed the basis of studies that further documented regional temporal and spatial depositional patterns, and placed these within a modern depositional systems framework. In a series of excellent papers published by Carleton Brett and Gordon Baird, and co-workers (e.g. Baird, et al, 1992; Brett and Baird, 2002, Baird and Brett, 2002; Brett et al, 2004), a modern depositional systems model for the Black River-Trenton-Utica system was developed. Additionally, the fine work of Cornell (2001, 2008, Cornell, et al 2005) deserves special mention, as much of the information in this guidebook article is derived from those studies.
The Black River-Trenton-Utica interval has also received attention as an economic resource system. Important natural gas reservoirs hosted by Black River/Trenton carbonates – so-called hydrothermal dolomite (HTD) reservoirs – were developed in the southern Finger Lakes region of New York State. The origin of these high-permeability systems is linked to burial diagenetic processes, with saline, higher-temperature fluids derived from underlying rocks, associated with basement faulting (Smith, 2006). Some of these reservoirs host exceptionally productive wells; however, there have been few new discoveries in the New York HTD play in the last decade. Application of horizontal well and hydraulic fracturing technology has promoted development of gas and gas-condensate extraction in the Trenton-Utica interval in western Pennsylvania, Ohio, and West Virginia. The Point Pleasant Formation, for example, is age-equivalent to the upper Trenton Group in New York, and has seen significant recent hydrocarbon production in Ohio (Wickstrom, et al, 2012). Organic carbon content and thermal maturation play key roles in determining the production potential for these unconventional resource intervals.
This field guide focuses on the Black River Group and basal Trenton Group exposed in northwestern New York State, in the area south of the St. Lawrence River, and east of Lake Ontario, extending south to the Black River Valley. This guide provides only a brief overview of these units. Interested readers are referred to the bibliography, and the extensive resources available online at

B1: BLACK RIVER AND TRENTON GROUPS, NORTHWESTERN NEW YORK STATE

STOP 1. Lowville Formation at Perch River Game Management Area
The basal ∼2 meters of the outcrop on the NE side of Rt. 12 exposes the contact interval between the uppermost Pamelia Formation and lower Lowville Formation of the Black River Group. This unit, known as the Pittsburgh Quarry Bed (Cornell, et al 2005: Cornell, 2008) is now considered to form the base of the Lowville Formation. It consist of weak-weathering dolomitic mudstone, overlain by thick laminated and occasionally cross-laminated, relatively massive brown to yellow weathering quartz sandy dolostone. Angular to subrounded polycrystalline quartz grains in the massive dolostone approach 4 mm in diameter, and must have been derived from exposed Proterozoic basement nearby. Overlying the Pittsburgh Quarry Bed, shaley dolostones and medium to thick bedded lime mudstones comprise the lower (unnamed) member of Lowville Formation.
Key sedimentary features to observe are cryptalgal/microbial laminations, bentonite (volcanic ash) beds, mudcracks, evaporite (gypsum/anhydrite) mineral casts, and storm event beds represented by mud rip-ups clasts, intraclast breccias, stylolites and solution voids. Grey limestone beds contain vertical “Phytopsis” burrows that are backfilled with grainy sediment or occasionally filled with calcite spare cement. Body fossils are relatively rare, but some beds contain abundance ostracodes; scattered gastropod and trilobite debris is present in some beds. The SE end of the outcrop is capped by a massive, tan-weathering laminated dolostone with abundant large mineralized voids. Immediately beneath the dolostone, beds of grey limestone are arrayed in a low amplitude fold with an axial trend approximately N80W, parallel to the highway. The basal beds of the overlying House Creek Member of the Lowville Formation are poorly exposed and partially hidden by brush at the summit of the exposure on the west side of the highway.
Thin (mm-cm thick) dolomite+calcite+pyrite+barite veins trending N70E are present in the limestone, and cm-scale calcite spar-filled voids are common in the dolostone and some limestone beds. The mineralized fractures here are small-scale hydrothermal features that may be related to the dolomitization and related mineralization. Note also that some of the larger voids have dark, hydrocarbon-stained calcite as the void fill.
Examine the limestone - dolostone contact and see if you can trace mineralized fractures into the dolostone. Is the dolomitization here ‘early’ (preburial – related to Mg-rich brines derived from seawater that had precipitated gypsum) or is the dolomitization related to the later mineralized fractures? What is the relationship between dolomitization and stylolitization? Why is some rock dolomitized whereas other rock remains nearly pure calcite limestone? Was there significant porosity in this rock at any time during its diagenetic history?
Is the folding related to the minor mineralization or a later phenomenon? When during the burial history were hydrocarbons generated?

STOP 2. Upper Lowville Formation and Chaumont Formation at Black River Gorge, Glen Park, NY.
Glen Park entrance to parking area for hydroelectric fishing access. Walk to gate entrance, crossing canal. Follow signs to fishing access site, descending path to river level. Watch for poison ivy along path. Be aware that warning sirens provide only 5 minutes before river level rises due to water releases. Be cautious – rocks are slippery.
The Black River Gorge from Great Bend to Brownville, NY exposes Black River and Trenton Group strata in a series of spectacular narrow river channel walls, waterfalls, and challenging rapids favored by rafters and kayakers. The river has been partially diverted and dammed for hydropower, and here at Glen Park the control structures provide access to riverside exposures along a foot path which descends to the natural channel south of the diversion canal. The strata exposed in the river gorge in the immediate vicinity of the access area include the upper portion of the House Creek Member of the Lowville Formation, including the Weaver Road Beds, the entire Watertown Formation, and the overlying Selby and Nappanee Formations of the basal Trenton Group.
At the river bank where the access trail ends, the outcrops consist of massive to thick-bedded bioclastic packstone/grainstone of the Watertown Member of the Chaumont Formation. The gorge wall across the channel exposes the Weaver Road Beds (House Creek Member of the Lowville Formation), and LeRay and Glenburnie Members of the Chaumont Formation, succeed by the massive Watertown. The Selby and Nappanee Formations of the Trenton Group overlie the Watertown. Looking west, downstream, it appears that the strata are downdropped to the north, although the perspective here is somewhat problematic. This could represent a tight monoclinal fold or high-angle fault. The course of the river here may be in part controlled by this structure. Coarsely crystalline ferroan dolomite in the Watertown here suggests hydrothermal fluids were active, perhaps related to the faulting/folding.

STOP 3. House Creek Member of Lowville Formation Overlain by Chaumont Formation on NYS Route 283 Near Calcium, NY.
The massive Watertown Member of the Chaumont Formation forms the resistant upper ledge at this road cutting. A prominent reentrant at the base of the Watertown is likely a bentonite horizon, with the underlying House Creek Member of the Lowville Formation forming the base of the exposure. Notably, the LeRay and Glenburnie Members of the Chaumont do not appear to be represented here (compare with Stop 2).
The House Creek Member beds here include typical Tetradium-rich wackestone with scattered, poorly-preserved ostracodes, gastropod and trilobite fragments. The moderately weathered faces of the overlying Watertown Member provide a good opportunity to examine this facies in some detail. Good examples of colonial tabulate corals (‘Hexagonaria”) and the small solitary rugose coral Lambeophyllum are relatively common. Brachiopods, crinoid plates, bryozoans and rare trilobites are also present. This fauna signals fully open marine conditions with normal salinity and well-oxygenated water.
Typical of the Watertown Member, chert nodules and certified bioclastic material is common in this outcrop. Nodular cherts are generally arrayed in laterally discontinuous horizons, perhaps related to the abundance of original siliceous skeletal material (sponge spicules, radiolarian) or possibly volcanic ash. Biogenic silica and volcanic ash are chemically unstable in seawater, and would provide dissolved silica for certification. Are some fossil types here more likely to be chert-replaced than others? Use a hand lens to examine the void filling minerals in the tabulate coral fossils. What diagenetic minerals are present?

STOP 4. House Creek Member of Lowville Formation on NYS Route 12 Near Lowville, NY.
The road cut on the west side of NYS Route 12 is developed in the House Creek Member of the Lowville. Only the uppermost House Creek Member was seen at Stop 3, and here a fuller range of lithologies is exposed. The subtidal facies in this exposure are typical of the House Creek, with grey-weathering Tetradium wackestones forming the basal portion of the interval, succeeded by a series of thick beds of bioclastic and intraclastic packstones and grainstones. There are relatively few large whole shell fossils in these units, but abundant fragmental ostracode, trilobite, bryozoan and coral material is present.
Note the wispy dolomite developed in the upper beds. Is the dolomite fabric-selective, that is selectively replacing a particular primary carbonate material, or are these wisps of stylocumulate origin? Carefully examine the intraclastic packstone bed. Is the fabric indicative of simple rip-up clasts, the result of bioturbation, or are these algal ‘clots’?

STOP 5. Kings Falls and Napanee Formations at Sugar River on NYS Route 12.
The road cuts on Route 12 and adjacent stream sections in the Sugar River at this locality allow observation of more or less continuous exposures of the lower and middle Trenton units. The Barrett Corporation quarry on the west side of Route 12 is developed in the upper Black River Group (upper Lowville and Chaumont Formations). On this trip we will limit our observations to the section exposed north of the Sugar River on the west side of Route 12.
Medium to thin bedded lime mudstones and silty shales of the Napanee Formation are overlain sharply by thick bedded, coarse bioclastic grainstones with thin interbedded shales of the basal Kings Falls Formation. The Napanee mudstones and shales are relatively organic-rich, and represent highstand facies. The succeeding wave and storm-dominated Kings Falls facies overlie a sharp, possibly erosional contact with the Napanee. Coarse, coquina-like brachiopod and crinoid grainstones are characteristic of the Kings Falls Formation. Many bedding surfaces show very fine preservation of brachiopods, crinoids, bryozoans and rare trilobites. Horizontal veins within shale horizons in the Napanee Formation document bed-parallel thrusting, perhaps related to high fluid pressure during burial diagenesis. Fluid inclusion temperatures in vein carbonate suggest burial heating to >145°C.

B2: ULTRAMAFIC/MAFIC ROCKS OF THE PYRITES COMPLEX
Jeffrey Chiarenzelli
Canton, New York 13676
Marian Lupulescu
Research and Collections
New York State Museum
Cultural Education Center
222 Madison Avenue
Albany, NY 12230
David Bailey
Department of Geosciences
Hamilton College
198 College Hill Road
Clinton, NY 13323
General overviews of the geology of the Adirondack Lowlands incorporating new results can be found in Chiarenzelli et al. (2010a; 2010b; 2011a; 2012), Lupulescu et al. (2010), Peck et al. (2013), and Wong et al. (2011), among others. The purpose of this trip is to provide additional insight into the tectonics of the Adirondack Lowlands by highlighting mafic and the ultramafic rocks exposed in the region. The trip will leave from Alexandria Bay and end at Pyrites, where ultramafic rocks of the Pyrites Complex occur (Figure 1).
The Adirondack Lowlands are a small part of the greater Grenville Province (Figure 2), a Mesoproterozoic orogenic belt of global proportions. The tectonic assembly and collisional orogenesis recorded here, and elsewhere, resulted in the formation of Rodinia, one of the Earth’s Supercontinents. In fact, breakup of Rodinia set the stage for the deposition of the widespread blanket of Potsdam sandstone along the rifted margin of Laurentia, to our east in the Champlain Valley, and exposed nearby on the fringes of the Ottawa Embayment.
The Adirondack Lowlands are delineated from the Adirondack Highlands by a long-lived shear zone that displays evidence for both a ductile and brittle history (Figure 3). This structure, named the Carthage-Colton Shear Zone, marks a number of distinct geologic changes such as the proportion of metasedimentary to metaigneous lithologies, differences in metamorphic grade, variation in the timing of metamorphism, the orientation of structural grain, and a change in topographic relief and elevation. The Lowlands is composed mostly of metasedimentary rocks, predominantly marble, and was metamorphosed to Upper Amphibolite facies during the Shawinigan Orogeny (ca. 1150-1200 Ma), leading to widespread melting in some pelitic gneisses and the growth of metamorphic zircon at 1160-1180 Ma (Heumann et al., 2006). Its position higher in the crust has led to a lack of widespread (any?) effects of the Ottawan Orogeny (ca.1020-1090 Ma), in contrast to the Highlands where metamorphic zircon of this age and resetting of zircon is widespread (Chiarenzelli and McLelland, 1993; Chiarenzelli et al., 2011b).
The oldest known rocks in the Lowlands consist of metasedimentary rocks of Grenville Supergroup, first noted in the literature by Sir William Logan in 1863 to describe the sequence of metasedimentary rocks that occur across large areas of southern Ontario and adjacent parts of Quebec. In the Lowlands these rocks provide evidence of shallow water, marine deposition and include, among other lithologies, voluminous marbles, calc-silicate gneisses, tourmaline-rich quarztofeldspathic gneisses, the sedex sulfide exhalatives of the Balmat zinc mines, and metaevaporites. A tripartite stratigraphy and/or structural stratigraphy has been recognized and consists of the lower marble, Popple Hill Gneiss, and the upper marble, which has been further subdivided into 16 lithologic units (deLorraine, 2001; deLorraine and Sangster, 1997). In contrast to the proposed setting of the carbonate rocks, a deep water turbidite origin has been proposed for the Popple Hill pelitic and psammitic gneisses and equivalents (Chiarenzelli et al., 2012). However, all of these rocks may be allochthonous, as older basement rocks have yet to be identified. The maximum depositional age for the Grenville Supergroup in the Lowlands has recently been constrained between <1258 and <1284 Ma (Chiarenzelli et al., in review).
Recent investigation has revealed that ultramafic-mafic igneous rocks occur within northeast-trending linear belts (Figure 4) in the Lowlands and may be older or equivalent in age to the Grenville Supergroup. Named the Pyrites Complex for exposures along the Grasse River near Pyrites, New York, these rocks have required rethinking of the depositional environment of not only the Grenville Supergroup equivalents in the Adirondack Region but the origin of the Zn-Pb sedimentary exhalatives of the Balmat-West Pierrepont Belt. A model incorporating the discovery of ultramafic rocks suggests that the rocks of the Grenville Supergroup were deposited after 1300 Ma in a back-arc basin which closed at about 1240 Ma during the Elzevirian Orogeny (Chiarenzelli et al., 2011a; 2012). This back-arc basin has been named the Trans-Adirondack Back-arc Basin and is thought to be temporally equivalent to a similar basin in the Central Metasedimentary Belt of Ontario and Quebec (Dickin and McNutt, 2007). This implies that a broad region of the southeast margin of Laurentia underwent rifting and extension at this time.
During this trip we will see what is thought to be the remnants of a highly disrupted ophiolite complex and deep water sedimentary rocks including hydrothermal altered ultramafic rocks (Stops 4 and 5), amphibolite (Stop 1), and possible turbidites (Stop 4). We will observe a lamprophyre dike cutting the ultramafic rocks at Stop 5. In addition, two stops will be made to investigate candidates for inclusion in the the Pyrites Complex including a large block of serpentine (Stop 2) and a hornblendite unit (Stop 3), both which require further investigation. Isotopic and geochemical studies indicate widespread enrichment in incompatible elements suggests that the Pyrites ultramafic complex and associated rocks formed in a suprasubduction zone setting. Although the surface exposure of the ultramafic rocks is small, the largest gravity anomaly in the Lowlands extends from Pyrites to the southwest for over 10 km parallel to the trend of the Carthage-Colton Shear Zone, indicating substantial amounts of high density rocks at depth.

B2: ULTRAMAFIC-MAFIC ROCKS OF THE PYRITES COMPLEX

STOP 1. Amphibolite and Pyritic Gneisses near Antwerp
We have driven south from Alexandria Bay to the southern belt of pyritic gneisses, amphibolites, metagabbros, and ultramafic rocks exposed in the Lowlands. Here near Antwerp, New York we will see some of the amphibolites, associated pyrite-rich metasedimentary rocks, and igneous rocks which intrude them (Figure 4). This belt can be traced nearly 50 kilometers to Pyrites and beyond. The large outcrop in the foreground shows a sequence of rocks dipping shallowly to the northwest. The darker rocks are amphibolites with MORB-like geochemistry. Those with rusty staining are of metasedimentary origin and contain detrital zircons, in addition to pyrite. Unfortunately these zircons are metamict and discordant. However, similar rocks at Pyrites have yielded intriguing results that will be discussed later at Stop 4. Both of these rocks are cut by a variety of intrusive pegmatites which are correlative in age to the Antwerp-Rossie Suite (ca. 1200 Ma).
The intrusive age of the Antwerp-Rossie Suite provides a minimum age for rocks of the Grenville Supergroup and ultramafic mafic rocks of the Pyrites Complex. Further east along Route 11 large banded outcrops of marble, and isoclinal folds within it, are cut by rocks of the Antwerp-Rossie Suite (Figure 5) dated at ca. 1200 Ma. The Antwerp-Rossie Suite is calc-alkaline and the product of northward subduction which preceded the Shawinigan Orogeny (ca. 1150-1200 Ma). The deformation preserved in the marble must pre-date this subduction event and subsequent Shawinigan orogenesis (Figure 5). This implies that the Grenville Supergroup and Pyrites Complex where deformed during the Elzevirian Orogeny (ca. 1220-1245 Ma), consistent with the minimum age for the Grenville Supergroup in the region constrained by detrital zircons (Chiarenzelli et al., in review).
The amphibolite outcropping here has been characterized by geochemistry and Nd-systematics. The amphibolites and metagabbros of this belt yield unrealistically old NdTDM ages of 1508-1624 Ma, as do ultramafic rocks at Pyrites, which in some cases are even older (Figure 6). This suggests disturbance of the Sm-Nd isotopic system and the influence of a metasomatically altered mantle wedge in their origin. Elevated oxygen isotopic ratios have been found zircon in plutonic rocks extending across the Lowlands into the Frontenac Terrane by Peck et al. (2004). Peck et al. (2004) interpreted these elevated ratios as indicative of the influence of subducted hydrothermally altered oceanic crust and/or sedimentary material beneath the leading edge of Laurentia. As a group, rocks of the Pyrites Complex plot between the Nd isotopic arrays of Quebecia ca. 1570 Ma (Dickin and Higgins, 1992) and Ontario Juvenile Crust ca. 1277 Ma (Dickin and McNutt, 2007) indicating an age in excess of 1277 Ma (Figure 7) but younger than 1570 Ma.
An age greater than 1277 Ma for the Pyrites Complex is compatible with several lines of evidence. Barfod et al (2005) obtained a 1274±9 Ma Lu-Hf date on apatite from the lower marble and interpreted it as the time of diagenetic apatite growth in the marble. Sager-Kinsman and Parrish (1993) obtained a maximum depositional age of <1306±16 Ma from the youngest detrital zircon in a quartzite unit of the Grenville Supergroup on Wellesley Island. And as will be discussed below, Chiarenzelli et al. (in review) found a maximum age of 1284±7 on detrital zircons from turbiditic rocks overlying ultramafic rocks at Pyrites. A model for the generation of oceanic crust within a series of back-arc basins along the leading edge of Laurentia has been proposed by Chiarenzelli et al. (2011a). The opening of the Trans-Adirondack Back-arc Basin is believed to have been synchronous with the opening of a similar basin in the Central Metasedimentary Belt of Ontario and Quebec (Dickin and McNutt, 2007).

STOP 2. Yellow Lake Serpentinite
This stop is on private land and permission was obtained for this field trip to visit the area. Please respect the landowner’s wishes and do not revisit this location without permission. Respecting land owners rights insures that future geologists will have access to key scientific sites. If in doubt, ask first.
While serpentine is a common mineral throughout much of the Lowlands, massive serpentinite, like that exposed here, is unusual. While this exposure may be a large serpentinized block of metasedimentary rock, the location of this body puts it within a linear trend of mafic and ultramafic rocks approximately 10 kilometers north and west of the Pyrites belt which we will visit again in the near future. These and other linear belts may represent separate thrust sheets, or perhaps, opposing limbs of a major nappe-like structure. In any event, the uniqueness of this exposure makes it an excellent place to visit regardless of its origin.
The Yellow Lake serpentinite is an unusually large body (∼100 m) of massive serpentine hosted within marble of the Adirondack Lowlands metasedimentary sequence. The only detailed geologic map of the area shows a thin band of lower marble surrounded by the "Hermon granite-amphibolite" unit (Buddington, 1934). On the Adirondack Sheet this unit is denoted as amg and consists of highly disrupted and intruded amphibolite and metagabbro engulfed within the Hermon granitic gneiss. While localized areas of disseminated serpentine within the marble were noted in the report and are particularly common in the upper marble which is found to the south, this large mass of serpentine was apparently not recognized or described. Intriguingly about 10 kilometers to the southwest near Theresa a small stream is named Soapstone Creek, suggesting more massive serpentinite or talc is yet to be rediscovered.
It appears that the Yellow Lake serpentinite first came to the attention of geologists in the mid-1980s when it was discovered by George Robinson and Michel Picard while on a St. Lawrence County mineral collecting excursion (G. Robinson, pers. comm.). Since that time, local mineral collectors have explored the outcrop and found nice sky blue apatite crystals and a few nicely formed pseudomorphs of serpentine after forsterite (S. Chamberlain, pers. comm.).
The outcrop is extensively fractured and incoherent, and because of this property, has been used locally as a source of rock aggregate and fill. Most of the exposure is composed of a dull yellow-green massive serpentine with abundant radiating dendrites of an unknown black oxide/hydroxide mineral coating fracture surfaces and disseminated throughout the massive serpentine (Figure 8). Small (2-20 cm wide) irregular veins of coarsely crystalline calcite cut through the unit, and it is within these veins that the blue apatite crystals (up to 3 cm long) are found. Pseudomorphs after forsterite occur along the sides of the veins (S. Chamberlain, pers. comm.). Pyrite, talc, and an unidentified Ca-Mg-silicate (possibly harkerite?) have also been found within these veins (Chamberlain and Bailey, unpub. data). Table 1 gives the composition of select mineral phases in the serpentinite, including Mn-rich dendritic growths and serpentine.
The occurrence of a relatively large body of ultramafic rock within the metasedimentary sequence of the Adirondack Lowlands is enigmatic; possible explanations for its origin include a localized accumulation of komatiitic ash, or an unusual clay-rich evaporite horizon within the carbonate sequence. Alternatively it may prove to be another block of highly altered ultramafic metaigneous rock such as that exposed at Pyrites. Additional field, mineralogical, and geochemical studies at this location and along this belt are needed to better understand this unusual ultramafic body and place it in its proper geologic context.

STOP 3. Hornblendite near Elmdale
This outcrop consists largely of hornblendite whose age and origin is unknown. The rock falls along the general trend of serpentinite (last stop), and within a 40 kilometer linear, but highly disrupted, belt of amphibolite, pyritic gneisses, and metagabbro extending to the northeast from Yellow Lake to near Kendrew Corners, north of Dekalb. It is composed 1-2 centimeter blocky, equant black crystals of hornblende surrounded by k-feldspar, titanite, and apatite (Table 1 and Figures 10 and 11). The rock shows relatively little evidence of wholesale deformation but is cut by pegmatite, which is also undeformed.
Hornblende-rich ultramafic rocks are relatively rare and can have a variety of origins. Detailed petrographic, geochemical, and geochronological study will hopefully provide further constraints on the origin of this rock. However, it may represent the one of the magmatic products of subduction in the Lowlands. Previous work has identified a strong subduction signature in many of the plutonic rocks exposed in the Lowlands (Figure 12; Chiarenzelli et al., 2010b). In particular, the lamprophyre dike we will see later in the day is indicative of this enriched mantle signature prevalent in mafic and ultramafic rocks of the Pyrites Complex as well.
Shawinigan plutonic suites in the Lowlands crosscut the fabric and isoclinal folds (Figure 5) in metasedimentary rocks of the Grenville Supergroup (Chiarenzelli et al., 2010a; Peck et al., 2013). The geochemistry and Nd signature of these plutonic suites (Antwerp-Rossie Suite, Hermon Granitic Gneiss, and Hyde School Gneiss), intruded between 1170-1200 Ma, mark the evolution of arc magmatism to granitic magmatism of the Anorthosite-Mangerite-Charnockite-Granite suite prevalent in the Adirondack Highlands but also found in isolated plutons in the Lowlands (Peck et al., 2013). These rocks track the tectonic evolution of the region and underlying mantle dynamics.
Based on its state of deformation and geochemical affinities (alkali-rich), the hornblendite unit exposed here may be an ultramafic variant of the Antwerp-Rossie Suite. If this is verified, it would be the only ultramafic body yet recognized as part of the Antwerp-Rossie suite. Intrusive gabbros and diorites associated with the Antwerp-Rossie suite occur nearby at Pleasant Lake and Split Rock, and to the south near Balmat and Harrisville (see Carl, 2000; Chiarenzelli et al., 2011). If it is indeed part of the Antwerp-Rossie Suite, the hornblendite’s occurrence along the linear trend of mafic and ultramafic rocks would be coincidental. However, like the serpentine body just investigated, the hornblendite occurs within a large belt of the Lower Marble. It may be an older tectonically emplaced block as will be demonstrated for the ultramafic rocks at Pyrites later in the trip. Regardless of its origin, this ultramafic rock is unique and has much to tell us about the geologic history of the Adirondack Highlands.

STOP 4. Contact between Pyritic gneiss and Ultramafic Rocks along the Grasse River
The rocks at Pyrites lie with the triangular core zone of a large winged structure defined by amphibolite and gabbroic rocks that extends northeast from Stellaville (former site of the largest pyrite mine in St. Lawrence County) nearly 15 km to Crary Mills (red star, Figure 6 and Figure 13). The general shape of the structure indicates it is folded and most of the amphibolitic rocks have a strong foliation and are highly disrupted by later magmatism. The belt of mafic and ultramafic rocks lies within marble. Although the actual surface exposure of the ultramafic rocks is modest (∼ 1 km2), the largest gravity anomaly in the region (Revetta and McDermott, 2003) extends from here southwest for about 10 kilometers (Figure 14). This suggests that it represents just a small part of the body and that a considerable mass of ultramafic rock, extending downward for some distance, lies parallel to the Carthage-Colton Shear Zone.
At this stop we will examine rocks associated with one of several pyrite mines that was in production between ca. 1880 to 1920 (Prucha, 1957) as a source of sulfur for industrial processes and sulfuric acid feedstock. Two adits can be seen on the side of the mustard-colored hill (Figure 15) as you approach the outcrop along the bank of the Grasse River. The adits intersect the steeply-dipping ore zone and the workings continued downward for as much as two hundred feet and extend underneath the river. On the bank, the rocks can be seen dipping steeply to the northwest and show differential weathering related to the variation in their mineral content and extent of weathering. Investigation at the top of the exposure shows the characteristic red coloration of pre-Potsdam weathering indicating that the unconformity was likely within a few feet of the current erosional level at the top of the hill.
The main ore zone consists of a ∼2m thick planar layer of pyritic breccia. Examination of the rock at the outcrop, and in thin-section, indicates the ore is composed of rounded 1-2 cm fragments of breccia indicating enrichment of the ore during repeated fault movement. A black, rounded, 5 centimeter wide fragment composed of breccia within breccia can be seen at the foot of the hill within the primary ore zone. Carefully examination of the ore and surrounding rock reveals highly weathered lath-like crystals of sillimanite. Other minerals within the sequence include quartz, feldspar, mica, chlorite, sericite, garnet, graphite (Figures 16-18), and a host of trace minerals including pyrrhotite, rutile, sphalerite, chalcopyrite, molybdenite (Figure 17), and chromite (Tiedt and Kelson, 2008). Tiedt and Kelson (2008) suggest a possible biogenic origin for the pyrite on the basis of sulfur isotopes, in line with the presence of graphite. However, a hydrothermal volcanic influence is also possible based on the variety of metals found and association with mafic and ultramafic rocks.
In addition to a strong foliation and compositional layering, the sequence is cut by shallowly dipping veins which weather in positive relief. These veins are composed mostly of pyrite, minor pyrrhotite, and silica and represent a much later event in the history of the deposit. Early deformation is defined by compositional layering which is isoclinally folded. Thin seams of breccia less than a 1 cm wide can be seen cutting the foliation and/or folded layers in several areas in the exposure. Near the ore, within white feldspathic pods, relatively large, brown rutile crystals can be seen (Figure 18). Other non-sulfide minerals observed in the ore include chlorite, graphite, quartz, rutile, and monazite (Figure 18). Note the variation in resistance of the rocks to weathering, colors, and variation in layering as you walk upstream.
Towards the end of the accessible outcrop area along the river the rocks change in character and color to a massive, knobby weathering green ultramafic rock. The actual contact between the ultramafic rocks and metasedimentary sequence is well exposed in the wall of the river bank. Whether the contact is sedimentary or tectonic, or both, can be debated here. Regardless of the nature of the contact, the layered metasedimentary rocks become progressively greener in color and more Mg-rich towards the contact. Given the occurrence of chromite noted by Tiedt and Kelson (2008) in these rocks, proximity to ultramafic rocks during deposition is likely. Deposition along a transform fault is one possibility that could account for the juxtaposition of these two rock types and lend additional credibility to a deep water turbidity origin for the pyritic gneisses.
Several meters above the contact a large isoclinal fold can be seen. About the hinge of the fold, one can see both recessive and resistant cm-scale layering (Figure 19). The recessive layers consist of quartzofeldspathic gneiss with both sillimanite and garnet, whereas the more resistant layers consist of up to 85% or more quartz. This exposure provides insight into both the original nature of the layered metasedimentary rock and its metamorphic grade. Couplets of mud and sand are suggestive of a turbiditic sequence deposited in a relatively deep water setting. At some point, likely associated with the Shawinigan Orogeny (1150-1200 Ma), these rocks were intensely deformed and metamorphosed to Upper Amphibolite facies (garnet-sillimanite).
The thin, resistant, quartz-rich layers were sampled for detrital zircon geochronology and yield a sparse, but very homogeneous, population of detrital zircons. The zircons are between 25-100 microns in length, fairly euhedral, and zoned (inset Figure 20). Limited evidence of metamorphic rims or xenocrystic cores was found, despite BSE and CL imaging of cross-sectioned grains. The zircon yielded a near unimodal population of grains with a maximum depositional age of 1284±7 Ma (Figure 20). This result is compatible with a restricted source consisting of igneous zircon, perhaps derived from a sheet of rift-related volcanic rocks.
The massive green rock at the end of the outcrop is composed of a mass of hydrous secondary minerals including Mg-rich phases such as talc, tremolite, chlorite, serpentine, and phlogopite. While the vast majority of primary minerals are replaced, isolated core fragments of augite survive in addition to chromite. Consisting of ∼40% SiO2 and nearly 33% MgO, it also contains substantial amounts of Cr and Ni (Chiarenzelli et al., 2011). Its knobby texture likely represents the pseudomorphic replacement of large oikocrysts. Further upriver the ultramafic rock varies from peridotite to pyroxenite to layered ultramafic rocks (Figure 21) all highly altered and nearly completely pseudomorphically replaced (Figure 22). Nonetheless, original layering is frequently preserved and no indication of ductile deformation is found in stark contrast to the adjacent metasedimentary lithologies, which are isoclinally folded. Classification based on normative mineralogy determined by the CIPW method indicates that most ultramafic rocks are peridotites and a few pyroxenites, consistent with field observations (Figure 23).
The lower grade assemblage and undeformed nature of the ultramafic rocks appears to be at odds with the intensely folded garnet-sillimanite gneisses they are in contact with. This can be explained either juxtaposition of the rocks after peak metamorphic conditions or the competent and hydrous nature of the ultramafic rocks which contain as much as 10% H2O.
Several attempts were made to directly determine the age of the ultramafic rocks. A range of rock types were analyzed for both Rb-Sr and Sm-Nd isotopes but unfortunately both systems appear to have undergone strong disturbance and provide isochron ages with large errors. Small zircons (∼50-200 microns) were observed within thin mm-scale, calcite-rich veins in samples of both periodite and pyroxenite (Figure 24). Attempts to recover them were successful and isotopic analysis by laser ablation – inductively coupled plasma mass spectrometry (LA-ICP-MS) and Sensitive High Resolution Ion Microprobe (SHRIMP) were conducted.
Zircons from the peridotite were dated by SHRIMP in Perth, Australia and yielded an age of 1140±7 Ma with a few older grains yielding an age of 1202±20 Ma (Figure 25). Zircon from the pyroxenite were dated by LA-ICP-MS at the Arizona Laserchron Center and yielded and age of 1197±5 Ma (Figure 25). Given the unlikely occurrence of zircon as a magmatic product in silica-undersaturated rocks, these ages are thought to represent the timing of zircon growth during metamorphic events and thus provide only a minimum age for the ultramafic rocks. The ages correspond to the known effects of the Shawinigan Orogeny in the Lowlands from previous studies (Heumann et al., 2006).

STOP 5. Enel Hydropower Station on Grasse River at Pyrites
Our final stop will be just a few hundred meters upriver from the last stop at the Pyrites Enel Hydropower Station. This area is privately owned and off-limits to the public. Special arrangements have been made to bring our group here and the owners require prior submission of a liability waiver. If you wish to visit this site in the future please respect private property laws and get permission beforehand.
Along the back wall of the parking lot of the Enel Hydrostation a layered sequence of ultramafic rocks occurs. Here the layers dip steeply to the right. Obvious and widespread alteration can be seen and the rocks are composed of a variety of Mg-rich secondary minerals and are discolored by hematitic veins. Secondary minerals included substantial amounts of serpentine and few, if any, primary phases survive (Figure 26).
A number of layer perpendicular fractures occurs here (Figure 27) and appear to define columnar jointing, a typical cooling-related phenomena in some volcanic sequences. Preservation of columnar jointing seems unlikely, but the general appearance and alteration of these rocks is distinctly different than the mostly massive periodite and pyroxenite downriver. Here layering on the decimeter to meter scale is well developed. It is conceivable that the rocks are metamorphosed lavas, but, if so, they are ultramafic in composition!
Proceed carefully down from the parking lot to river. Here near the discharge from the Hydrostation, you can again see the contact between the pyritic gneisses and ultramafic rocks exposed in the far wall of the river bank. However, of special interest is a meter-wide lamprophyre dike consisting of large, cm-scale, phlogopite phenocrysts with a black serpentine-rich groundmass cutting the other ultramafic rocks (Figure 28). Again small areas of augite are preserved but little else. Large, euhedral apatite crystals, several millimeters in size are also abundant and likely a primary phase.
The geochemistry of this rock indicates substantial enrichment in incompatible elements and suggests derivation from metasomatized mantle. Despite having just 33% SiO2 a population of zircons was separated from this rock. The zircons, nearly all concordant, gave an age of 1183±6 Ma with individual grains as old as 1213±16 Ma (Figure 29), again indicative of growth during Shawinigan metamorphism or conversely, intrusion at that time. If the zircons do indeed mark the time of intrusion they are likely related to arc magmatism associated the Antwerp-Rossie suite (ca. 1200 Ma).
Further upriver these rocks are in structural contact with a sequence of gabbroic rocks cut by finer-grained dikes. The gabbros appear exceptionally coarse-grained, baseball-sized (Figure 30), however, this appears to be the consequence of anastomosing, closely-spaced shear zones and the widespread growth of coarse phlogopite rather than the original grain-size. Some low strain domains between the sheared areas show a typical gabbroic texture. Samples of gabbro and a fine-grained dike rock cutting the gabbro share similar geochemical and isotope trends with the ultramafic rocks and are likely part of the same sequence. Given the occurrence of ultramafic rocks, gabbros, amphibolites, and sulfide deposits Chiarenzelli et al. (2011b) interpreted the Pyrites Complex as a highly dismembered and intruded ophiolitic complex. If so, it must be in thrust contact with shallow water metasedimentary marbles of the Grenville Supergroup it is often enveloped by.

B3: METAMORPHIC PETROGRAPHY BETWEEN TUPPER LAKE AND BLUE MOUNTAIN LAKE
Robert Badger and James Carl
SUNY Potsdam
Potsdam, NY 13676
Kyle Ashley
Dept. of Geology
Virginia Tech
This trip was led for the 2004 NYSGA. Since that time students have examined two additional outcrops on the trip, so we now have additional probe and petrographic data.
This trip examines the petrography and petrology of roadside outcrops along Route 30 between Tupper Lake and Blue Mountain Lake. As with most Adirondack rocks, these have been intensely metamorphosed. Lost during the metamorphism are the features which distinguish sedimentary rocks from extrusive and intrusive igneous rocks. Their intrusive igneous textures and/or content of fossils, clay minerals, layers of volcanic lava and ash, have long since been transformed. We will visit outcrops of marble and quartzite that clearly represent metamorphosed limestone and sandstone, and amphibolites that most likely were basaltic dikes. Most rocks, however, are gneisses with varying proportions of feldspars, quartz, amphibole, pyroxene and garnet. Without geochemical analysis, it is sometimes difficult to determine if their protolith was an igneous or sedimentary rock. We draw your attention to features at the outcrop that support our interpretations; color microphotographs will be provided at most stops. We anticipate lively discussion of interpretations about the outcrops.
Our understanding of the geologic history of the Adirondacks is an ever-evolving project that involves the work of hundreds of researchers. We offer a brief summary, and refer you to more detailed studies by Isachsen et al. (1991), Chiarenzelli and McLelland (1991), McLelland et al. (1996), Hamilton et al. (2004).
The oldest rocks in the Adirondacks are 1.2 to 1.3 billion year old metasedimentary rocks that include marbles and gneisses seen on this trip. Included are gneisses that carry a geochemical signature of island-arc complexes, suggesting that northern New York consisted of a marine environment fed with materials from a nearby island arc. These rocks were metamorphosed during the Elsevirian Orogeny about 1210 to 1160 million years ago (Hamilton et al., 2004). Intruded into this sequence were the anorthosites, mangerites, charnockites and granites comprising the AMCG suite of McLelland at al. (1996). The ages of many of these plutonic rocks cluster around 1155 million years; the rocks at stop 1 are younger. A second orogenic event, the Ottawan, occurred between 1090 and 1030 million years ago (Hamilton et al., 2004). It was associated with the supercontinent of Rodinia and the formation of the Grenville Mountain range, whose heights were similar in magnitude to the present-day Himalayas. By the time Rodinia broke up, between 600 and 550 million years ago, the Grenville Mountains had been eroded to a gently rolling landscape, not unlike the present-day Adirondack lowlands, except barren of vegetation. Detailed apatite fission-track thermochronology by Mary Roden-Tice (Roden-Tice et al., 2000) at SUNY Plattsburg indicates that uplift and unroofing of the central Adirondacks began during the late Jurassic.

B3: METAMORPHIC PETROGRAPHY BETWEEN TUPPER LAKE AND BLUE MOUNTAIN LAKE

STOP 1. Metamorphosed, fractured and mineralized meta-syenite.
A long roadcut of reddish-colored rock on the right (west) side of Route 30. This outcrop was the focus of a senior research project by SUNY Potsdam student Scott McDonald, whose data are presented here.
Meta-syenite. This dark reddish-brown stained rock contains abundant red to tan K-feldspar, albite, and quartz (Figure 1). The dark minerals are hornblende, but locally have been altered to chlorite. The K-feldspar shows marvelous perthitic textures. The minerals display pronounced parallel alignment. We believe the protolith was an igneous syenite and classify the rock as a meta-syenite or hornblende-alkali feldspar gneiss.
Microprobe data for albite are listed in Appendix 1, for potassium feldspar in Appendix 2, and for calcite and chlorite in Appendix 3. Data was obtained on the microprobe at SUNY Binghamton, with the help of probe guru Bill Blackburn. The data show that the plagioclase was almost completely altered to albite during metamorphism, and that the potassium feldspar contains very little sodium.
The rock probably was coarse grained to begin with. What once were randomly oriented, tabular crystals of alkali feldspar were later transformed by shearing into small, parallel-aligned grains. Some chunky feldspar crystals were deformed into lens-shaped grains or augen, (German for “eye”). On the east side of the road, the gneiss is brown and finer grained, but it grades into reddish meta-syenite and probably is the same rock. Zircon crystals, ZrSiO4, were separated from similar rock at a nearby outcrop. They have a uranium-lead age of 1098±4 million years, which is the age of crystallization of the syenite magma (McLelland et al., 1988, p. 921, their rock sample #10). The meta-syenite is typical of younger granitic rocks of the Adirondack Highlands which have not been studied in detail.
Joint planes and rockfalls. This outcrop has many prominent joint planes. Note the geometric shapes of the large blocks at the south end on the west side. Produced by breakage along sets of intersecting joints, they resemble children’s play blocks perched precariously on edge (Figure 2). A small rockfall occurred by breakage along joints on the east side of the road (Figure 3). In July, 2000, crushed green plants beneath the rubble indicated that the fall had occurred a few days prior to our visit.
Mineral deposition in breccia zones. Pace off about one hundred feet from the north end of the outcrop, west side, and locate the zones of crushed, broken and angular rock fragments (Figure 4). The breccia formed long after cooling of syenite magma and long after the period of Adirondack metamorphism, perhaps during uplift of the Adirondack dome. The breccia zones are shown by the presence of irregular patches and stringers of low magnesium, Mn-bearing calcite, CaCO3 (see microprobe data in Appendix 3).
The calcite was deposited as cement in open spaces in the breccia. However, the calcite hides a sprinkling of tiny, brassy-yellow crystals of pyrite (FeS2). The crystals occur as a coating on brecciated meta-syenite rock fragments (Figure 5). Magnified under a hand lens, the pyrite consists of millimeter-sized crystals of two types: (1) cubes with six square-shaped sides, and (2) pentagonal dodecahedrons, a geometric form consisting of twelve faces, each with the five sides of a pentagon.
The calcite could have originated in one of two ways. It could have precipitated from fluids that dissolved minerals from overlying rocks, or, less likely, it could be igneous calcite (carbonatite) with CO2 derived by mantle degassing. To test these hypotheses, a sample of the calcite was sent to Steve Howe, stable isotope magician at the mass spectrometer lab at SUNY Albany. His analysis shows a delta 18O value of –14.927 per mil (compared to PDB) or +15.474 per mil (compared to SMOW), and a delta 13C value of –4.303 per mil (compared to PDB). These stable isotope numbers indicate a meteoric origin of the water and a bedrock source for the carbonate (Steve Howe, personal communication).
The carbon and oxygen isotope data therefore indicate that the calcite came from the dissolution of overlying layers of marble, the dominant carbonate rock in the Adirondacks. Calcite was dissolved in water to produce ions of calcium (Ca+2), bicarbonate (HCO3-1), carbonate (CO3-2) and molecules of carbonic acid (H2CO3). The ions and molecules were transported downward by groundwater, perhaps traveling a considerable distance before crystallizing as calcite in the breccia. The marble source of these components, probably hundreds to perhaps thousands of feet above your head, was long ago removed by erosion.
The sulphur-bearing mineral, pyrite, is virtually insoluble in oxygenated groundwater. Reducing conditions are required for transport and deposition of iron sulfide. Such groundwater is generally deep-seated, oxygen-depleted and full of dissolved material. The point being made is that the kind of groundwater required to transport calcite is incompatible with water that transported the iron and sulphur. Separate solutions from different sources must have formed the minerals, the pyrite was first and the calcite later on.

STOP 2. Calc-silicate metasediments
Park in the large, paved pullover on the right (west) side of Route 30. We are located on South Bay at the south end of Tupper Lake. Examine the roadcut opposite the middle of the parking area (Figure 6). This outcrop was discussed by Jim McLelland in the 1992 NYSGA Guidebook.
We have left metamorphosed igneous rocks and entered a domain of metamorphosed sedimentary rocks. This outcrop is a slice through the steep west side of a bedrock hill, and the steeply inclined rocks include pale green gneisses with abundant diopside, CaMgSi2O6 (Figure 7). These metasediments must be older than the meta-syenite and other igneous rocks that intruded them.
The mineral assemblage of diopside, quartz, K-feldspar, plagioclase feldspar and sphene formed during the metamorphism of impure limestone. A chemical reaction occurred among minerals that commonly occur in limestone—calcite, CaCO3, dolomite, CaMg(CO3)2, and the so-called impurity—a little quartz sand, SiO2. Note that the formula of diopside, CaMgSi2O6, contains elements derived from each of them.
A difficulty with the presence of diopside at this outcrop is the scarcity of carbonate minerals, calcite and dolomite, that were needed to make it. Note the presence of light colored seams and veins that help to define the gneissic banding. At first glance, they might be mistaken for calcite, but the minerals do not fizzle in acid and cannot be scratched with a knife. The light-colored mineral is potassium feldspar, not calcite. We suspect that calcite and dolomite of the impure limestone protolith were completely used up in metamorphic reactions to produce diopside.
Before leaving, examine the outcrop from a distance and note the color change where green, diopside-bearing gneiss gives way to gray hornblende-biotite-feldspar gneiss as you look from right (south) to left (north). The gray gneiss was interpreted by McLelland (1992, p. 40-41) as an igneous intrusive rock, even though its contact with the diopside-bearing gneiss appears gradational over a distance of a few feet. McLelland’s interpretation is that this is a highly sheared contact between two very different rock types, one igneous and one sedimentary. Another interpretation is that both are sedimentary and that the gray gneiss represents a facies change in the original sedimentary protolith, an impure limestone, to perhaps to a shale or mudstone that was metamorphosed to a hornblende-biotite gneiss. A few tens of feet north of the contact between diopside gneiss and the gray hornblende gneiss is another much smaller zone of diopside gneiss, interpreted as either a sheared fragment of the diopsidic gneiss or another facies change. You may form your own opinion.

The William C. Whitney Wilderness Area
Intersection with the north entrance road to the William C. Whitney Wilderness Area on Little Tupper Lake.
Called the “most significant Forest Preserve acquisition in twenty five years,” the estate of Mary Lou Whitney was purchased by the State of New York in December, 1997. The purchase consisted of 14,700 acres surrounding Little Tupper Lake, once the largest privately owned lake in the northeast. The eight mile-long lake is relatively shallow (less than twenty five feet deep), drains into the Bog River and is home to a native strain of fish called the Little Tupper brook trout. Canoeists were pleased with a land purchase that connected major canoe routes throughout a vast region of wetlands, lakes, bogs and cut-over timberland. In March, 2000, the area was classified as wilderness.
The road leads to an eighty acre site on the lake’s north shore, once used as the headquarters for Whitney Industries. Now classified for “administrative use,” the site includes a ranger station in a grove of tall pine trees, a parking area and a canoe launch. The lake has become one of the most widely used, flat water canoe routes in the East, and several tens of primitive camping sites have been constructed around its perimeter.

STOP 3. Marble and other metasediments
Drive beyond the outcrop and park beyond the guard rail at the south end. Walk north and examine three outcrops on the west side. These outcrops were studied by SUNY Potsdam student Aaron Fiaschetti for his senior research.
Here’s a good example of diopside-bearing marble, an easily recognized metamorphic rock that once was a limestone (Figure 9). The marble lies between gneissic layers of debatable origin. Separated by intervals of soil, the three parts of the outcrop include (1) garnetiferous biotite gneiss with quartzite layers at the south end, (2) marble outcrops in the middle, and (3) garnet-bearing gneiss at the north end.
South outcrop. The dark gneiss contains alkali feldspar, quartz, hornblende, seams of biotite and locally garnet (Figure 10). The mica seams are close together and locally give the gneiss a schistose appearance. Some of the biotite appears to have been altered to bronze colored phlogopite. Non-geologists have misidentified such grains as flakes of gold. Thin layers of quartzite that broke away from weathered mica seams give the rock a slab-like appearance. The hornblende gneiss contains pods of quartz with small grains of garnet (Figure 11) whose lavender color indicates a small spessartine (i.e. manganese, Mn) component.
Middle outcrop. The white marble and calc-silicate rock are easily recognized in the center outcrop (and across the road) (Figure 9). They consist of coarse calcite grains that were recrystallized from limestone during metamorphism. Most likely, the marble is much lighter in color than the original limestone. Finely dispersed coloring agents such as organic matter, clay minerals and iron oxides, were either expelled from the limestone or incorporated into the crystal structure of another mineral during metamorphism.
Glassy grains of green diopside are present, but there are fewer of them than at Stop 2. Also present are muscovite, brown sphene, shiny black flakes of graphite (pure carbon), tiny grains of pyrite, smoky to clear grains of quartz, and light colored seams of chunky alkali feldspar that, at first glance, can be mistaken for calcite. Look for veins of milky quartz. The graphite is probably the recrystallized remains of primitive algae that once lived in the warm marine environment in which the original limestone was deposited. A breccia zone at the south end of the outcrop may have formed at the same time as the breccia zone at stop 1.
North outcrop. This hornblende-biotite-garnet gneiss has more hornblende and less biotite than gneiss at the south outcrop. There are layers and lenses of pink alkali feldspar, as well as quartz veins that parallel the gneissic layering. Some layers were deformed into ragged and elongate stringers, whereas others were pulled apart into sausage-like boudins. Garnets in the boudins are iron-rich and reddish in color.
Layers of quartzite in the south outcrop and the proximity of the intervening marble lead us to interpret the gneisses at both ends of this outcrop as shales or mudstones prior to deep burial and metamorphism. There’s a certain chemical “richness” to shale, thanks to the presence of sheet-structured clay minerals such as illite, chlorite, kaolinite and montmorillonite. They became unstable and decomposed during high temperature metamorphism, and their elements were re-combined to form a variety of high-grade metamorphic minerals. Imagine the conversion of compact and drab mud into an attractive crystalline aggregate of garnet, hornblende, biotite mica and feldspars (Figure 11). Hence, the well known saying, “There’s nothing wrong with a sedimentary rock that 10 kilobars can’t cure.”

STOP 4. Multiple igneous intrusive rocks?
Park beyond the guard rail and examine the rocks on the opposite (east) side of the road. This outcrop was studied by SUNY Potsdam student Kyle Ashley for his senior research and presented at the NE GSA meeting in 2009. The chemical data and interpretations come from his study and were updated this year reflecting his current graduate training at Virginia Tech.
Description
This outcrop contains a variety of metamorphic rock types, from pink, garnetiferous gneiss to gray garnet-poor gneiss to pegmatite with coarse, black hornblende crystals. Clearly, multiple protoliths are called for. From south to north, seventeen different units were differentiated which can be divided into six different rock types.
Table 1. Characteristics of rock varieties at stop 4.
Units | Description
1, 2, 15 | Hornblende gneiss with crosscutting pegmatite
10, 12, 14 | Hornblende-biotite amphibolite dike
3, 9, 16 | Garnetiferous gneiss with small pockets of qtz-gar-ksp pegmatite. Garnet is small and scattered.
6 | Hbld-ksp-qtz pegmatite with euhedral hbld
4, 13 | Garnetiferous gneiss with large scattered garnets
5, 7, 8, 11, 17 | Ksp-plag-qtz gneiss with little to no other major minerals present
The massive, pink and gray gneisses that predominate at the south end of the outcrop (Fig. 4-1) contain abundant quartz and potasssium feldspar with lesser amounts of plagioclase feldspar, biotite and a few small garnets scattered here and there. In thin section, the potassium feldspar contains a wide range of excellent examples of perthitic textures. The rock has a fairly uniform grain size. It lacks the pronounced banding and segregation of minerals that we observed at Stop 3. It also lacks dark colored minerals, but chemical reactions that produced the garnet may have used them up. Most likely this rock was originally a granite.
Garnet compositions throughout the outcrop are almandine (Table 2), and plagioclase feldspars have An compositions between 23 and 40.
Present in the light colored granitic gneiss is a distinctive 10-15 inch wide layer of amphibolite (Figure 12), composed predominantly of black hornblende and plagioclase feldspar. This rock probably was a basaltic dike that intruded the granite prior to metamorphism.
The mafic content of the gneiss increases towards the north end of the outcrop, and the rock becomes an orthopyroxene-bearing granitic gneiss, or “charnockite” (see discussion of charnockites at Stop 9). Much of the opx has been altered to chlorite and other alteration minerals, but the persistence of relict cleavage indicates that the mafic grains were originally pyroxenes (table 4). Other portions of the rock contain significant amounts of hornblende. Again, an igneous protolith is called for, but the high mafic content suggests this intrusive rock unit differs from the mafic-poor granite gneiss at the south end of the outcrop. The prescence of pyroxene suggests crystallization deep in the crust at higher temperatures. The melt is derived from orogenic collapse of an over-thickened Grenville continental crust.
The hornblende in the amphibolite of bed #10 has a composition as follows:
(n = 2)
A coarse grained pegmatite occurs near the north end of the outcrop (Figure 13). It contains coarse grains of potassium and plagioclase feldspar, hornblende and two varieties of pyroxene – one a pale green with second order interference colors of clinopyroxene, the other a pale pink with first order interference colors of orthopyroxene. The opx shows more alteration than cpx, which is typical in Adirondack rocks with co-existing pyroxenes. Apatite is also a significant component. The coarse minerals are randomly oriented as though crystallized from a silicate melt in the presence of a watery fluid. The pegmatite appears to be a late igneous intrusion that post-dates metamorphism. Or, given the compositional similarity to the surrounding rock, it could be recrystallized gneiss in the presence of hot, watery fluids of late stage metamorphism.
In thin section, quartz extensively exhibits chessboard extinction, which indicates the combined operation of prism [c] and basal <a> slip systems active in quartz, occuring at minimum deformation temperatures of ∼630 °C (Blumenfeld et al., 1986; Mainprice et al., 1986). Garnet-biotite exchange thermometry considering the effects of Mn partitioning from Holdaway (2000) constrains peak metamorphic temperatures at ∼750 °C – consistant with 2-pyroxene thermometry results from Hoffman (1979), McLelland and Whitney (1977), and Bohlen and Essene (1979) (Figure 14).

STOP 5. The granite quarry outcrop.
Park at the south end of this long roadcut, beyond the guard rail, and examine the outcrop on the right (west) side (Figure 15).
Quarries. No, they don’t quarry the granite for building stone here. In a sense, however, all road cuts are quarries because rock was removed to make room for the roadbed. The Department of Transportation blasted through the massive rock, and much breakage occurred along the natural fractures (joints). Breakage along prominent near-horizontal sheeting joints has provided some flat and narrow surfaces that resemble crude stair steps in the face of the outcrop, similar to the working face of an active stone quarry.
Metamorphosed granite. The finely foliated meta-granite or alkali feldspar-quartz-hornblende gneiss, has small grains of dark hornblende aligned as parallel stringers (Figure 16). Some of the hornblende has altered to chlorite, but most is intact. A little plagioclase feldspar is present, along with apatite and a few grains of zircon. Cutting diagonally across the outcrop is a dark amphibolite layer (Figure 15) that once was a tabular dike of basalt that intruded the granite. A small fault extends from the granite into the dike about midway up. See if you can find it. The top part of the dike was displaced a few inches to the right of the bottom part. Also note whether the dike lies parallel to, or cuts across the gneissic layering in meta-granite.
Small faults in meta-granite are often coated with fault gouge. Slickenlines in the gouge indicate the direction of fault movement (Figure 17). The fractures are coated with green hydrous alteration minerals, probably chlorite and serpentine.

Sabattis Road
Intersection on the right (west) side with the south road (Sabattis Road) to the William C. Whitney Wilderness Area.

STOP 6. The marble cake outcrop
(that only has a little marble). Evidence for igneous intrusive origin.
SUNY Potsdam student Christopher Mack studied this outcrop for his senior research, and presented his data at NEGSA in 2013.
Examine the outcrop on the left (east) side of Route 30. The rock is cream and black mottled gneiss that reminds us of a marble cake. The light portions are composed primarily of plagioclase (An36, optically determined) and perthitic alkali feldspar, with very little quartz, while the dark zones are predominantly composed of clumps of hornblende, plagioclase and diopside, with minor garnet, orthopyroxene, magnetite and scattered apatite. Some of the hornblende contains relict orthopyroxene cores. At the south end of the outcrop and on the west side of the road is a different type of rock. Thus, our “marble cake” rock may be of small volume and confined to this outcrop, but we think it is the most unusual rock on this field trip.
To make a marble cake, one takes a yellow cake batter and divides it into two parts. Chocolate is added and thoroughly blended in one part. Spoonfuls of the chocolate batter are then dropped into the remaining yellow batter, and a spatula is used to swirl the blobs about. This outcrop is the result. Or so it seems. The dark, steeply inclined layers contain abundant hornblende and pyroxene (Figure 18). The “free form” pattern of layers, ribbons, loops and swirls, however, was not produced by stirring. The pattern results from the intersection of the viewing surface with the plane of gneissic banding, whether parallel to it, perpendicular, or somewhere in between. Marble cakes and rocks that resemble them have beautifully convolute patterns, but the anticipation of eating is absent here.
An early effect of metamorphism was to segregate a coarse-grained igneous rock into light and dark layers. The rock then became “mobile” and began to flow. Some of the dark layers behaved rather stiffly. They were ruptured and shredded while the light-colored, feldspar-rich layers flowed plastically around them (Figure 19).
Evidence that the swirled gneiss was an igneous rock includes a large fragment of diopside-bearing marble in the gneiss (Figure 20). The marble is typical of marble found throughout the Adirondacks, consisting of coarse calcite with green diopside and brown sphene. The fragment is a xenolith whose presence is a powerful argument that the gneiss was once magma and that the fragment was caught up in it.
Another xenolith is composed of green diopside and red garnet (Figure 21). Note that the gneissic layering is deflected around it. The swirled gneiss also contains large feldspar crystals up to eight inches across (Figure 22). Minerals of that size generally have crystallized from low viscosity silicate melts, often in the presence of watery fluid. The question is how such large crystals survived the metamorphism, especially the shearing that produced the gneissic banding.
Our explanation is not very profound. A wide range in grain size is expected if the shearing had been a hit and miss operation. Some parts of the rock were reduced in grain size, whereas other parts were left undisturbed. Primary igneous textures are preserved elsewhere in the Adirondacks, including anorthosite outcrops near Saranac Lake which contain single plagioclase feldspar grains more than 12 inches long.
We note that angular fragments make up a breccia zone parallel to the gneissic banding (Figure 23). This post-metamorphic breccia zone may be synchronous with the one seen at Stop 1.
A change in rock type occurs at the south end of the outcrop. The banded rock gives way to a pinkish and green, faintly foliated syenitic gneiss whose minerals include pink alkali feldspar, quartz, biotite, a little hornblende, and locally significant amounts of red brown garnet (Figure 24). Some garnets are up to 4 cm in diameter.

STOP 7. Layered “gray gneiss,” a common metamorphic rock.
Examine the left (east) side of the road. Please no hammers here – it is a small outcrop and with only a couple of the large, interesting hornblendes in the pegmatite.
Welcome to the “gray gneisses,” a common and widespread rock type in Precambrian age rocks throughout the eastern United States and Canada. This medium grained example has fewer surprises than the last outcrop. It, too, contains plagioclase and alkali feldspar, hornblende, pyroxene, garnet, quartz and a little biotite mica. The gneissic banding consists of stringers and seams of cream-colored alkali feldspar. Some bands are convoluted and pinch out, and all are vertical or steeply inclined (Figure 25). The gneiss is peppered with garnets, and the large ones are prominently displayed in the light colored seams. The rocks at this outcrop lack the large feldspar crystals and marble xenoliths seen at the previous stop, and may not be intrusive in origin.
This gray gneiss may have been volcanic ash, an extrusive igneous rock. Similar-looking rock underlies a belt more than seventy miles across in the Adirondack Lowlands to the northwest. Published geological studies suggested that the original rock was sedimentary in origin, a shale or an arkose, but the chemical data were a good match for dacite volcanic rocks (Carl, 1988). We have not analyzed the rock at this outcrop but it, too, could represent a thick pile of volcanic ash.
Note the presence of a fifteen inch-thick pegmatite vein that cuts across the gneiss (Figure 26). It consists of smoky quartz and coarse, chunky grains of alkali feldspar, cream to gray in color. The vein also contains dark and tabular hornblende crystals about three quarters of an inch wide and up to four inches long (Figure 27). The crystals are randomly oriented as expected from growth within a low viscosity, water-bearing silicate magma. Such magma would promote the growth of large crystals, as well as the crystallization of the hydrous minerals hornblende and biotite observed along the vein margins.

Bridge, beach and floatplane dock at Long Lake

STOP 8. Buttermilk Falls on the Raquette River.
Park near a small sign on the right (north) side of North Point Road. Walk the short distance to the waterfall. We are located upstream less than a mile from Long Lake.
This relatively steep stretch of the Raquette River lies east of Forked Lake and southwest of Long Lake. The river drops 116 feet in five miles, and Buttermilk Falls is the third of three canoe carries from Forked Lake. Rapids here extend for about one hundred yards before ending in a waterfall, and the total drop in elevation is about forty feet.
The bedrock at the falls is a highly weathered, dark brown, metamorphosed granite. One can see pods and veins of milky quartz that crop out near the shore. The bedrock is covered with pine needles and shallow soil. Trees are shallow-rooted and unstable in high winds.
Unless many rock hammers improve the quality of fresh outcrop exposures, this is a structure stop, and the petrography of the rock is best examined at Stop 9. Two prominent directions of jointing are present here (Figure 28), with one extending across the river. The joint planes had little influence in creating the stream channel, in contrast to joints that lie parallel to stream courses. But “cross joints” like these are partly responsible for the presence of rapids. Blocks were removed from the downstream side of a joint, resulting in stair steps that decline in the downstream direction. The river spills over them. People who sit in spaces where the blocks once sat will either enjoy a shower or a bouncing ride downstream (Figure 29).
The same joint spacing, often outlined by pine needles caught in the grooves, can be observed in the bedrock surrounding the falls. The origin of jointing is unclear. Is it due to isostatic rebound following glacial retreat? Or is it much older, related to uplift of the Adirondacks and erosion of the overlying bedrock? Or is it older still, caused by tectonic forces shortly after metamorphism?
The falls and spray sparkle in the sunlight, and our geology field trips get bogged down when the college students disperse among the rocks. They sit at the falls and perch with the demeanor of poised herons, legs crossed, patiently waiting for a stray fish. They relax, get wet and cannot hear their professor’s voice over the roar of the falls. If the timing is right, this is a good place for lunch.

STOP 9. Dark brown meta-granite with pegmatite.
Park on the shoulder of Route 30, before the curve and near the center of a long outcrop on the right side.
The dark brown colored meta-granite (Figure 30) is similar to the bedrock at Buttermilk Falls. This type of rock was quarried and used as a building stone in handsome but very dark-walled churches in Saranac Lake, Lake Placid, Tupper Lake and elsewhere. The major minerals are alkali feldspar, hornblende and orthopyroxene with lesser amounts of quartz, plagioclase feldspar, clinopyroxene, garnet, magnetite and apatite. Gneissic banding is faint, steeply inclined and hard to see when the outcrop is wet.
The presence of pyroxene is unusual for granite, and geologists have a special name for orthopyroxene-bearing granite: charnockite. The name honors Job Charnock, the founder of Calcutta, India, who died in 1693. Charnock was not a geologist and may not have appreciated this particular rock, but his memory is preserved among geologists because the first described example of charnockite was his tombstone.
Charnockites are igneous rocks that were subjected to high grade metamorphism. The presence of hypersthene pyroxene, (Mg,Fe)SiO3, raises a question. Did the mineral crystallize in the magma and, therefore, is of igneous origin? Or did it form much later during metamorphism when water-bearing silicate minerals, such as biotite and hornblende, dehydrated and converted into hypersthene? If so, hypersthene would be a metamorphic mineral. Adirondack charnockites are widespread igneous rocks that make up the “C” part of the AMCG suite of intrusive rocks.
Look closely along the outcrop, before the right bend in the road, for a pegmatite vein consisting of quartz and alkali feldspar (Figure 30). Also of interest are mottled patches of light colored, slightly coarser grained rock that lack hypersthene (Figure 31). The patches are best seen on dry surfaces around the bend towards the south end of the outcrop. The pegmatite and mottled patches are generally interpreted as having formed by partial melting of the charnockite.
During metamorphism, the temperatures were hot enough to melt some of the biotite, hornblende, quartz and feldspar. Most likely, the break-down of biotite and hornblende supplied enough water to lower the melting temperatures of the other minerals. Where appreciable quantities formed, the melt migrated into fractures to form pegmatite veins, like the one seen here, that may have migrated several tens or even hundreds of feet. The patches, on the other hand, appear to be small quantities of melt that did not migrate very far. The patches are sometimes called “sweats,” a reminder of efforts by the human body to cool itself. The sweats at this outcrop solidified near where they formed.

Adirondack Museum at Blue Mountain Lake
The Museum honors the people of the Adirondacks. There are exhibits about the lives of Native Americans, early settlers, wealthy tourists who joined them in summertime, the unique architecture of Adirondack camps, and the intricate transportation networks that brought people into the woods. One learns about farmers, hunters, miners, loggers, guides, artists, writers, furniture makers and, most of all, tourists like you, drawn here for more than 150 years by the natural beauty of the mountains, lakes and forests.
The museum itself is a mountainside courtyard of large and small buildings, some serving to house a collection of artifacts, and others to illustrate a distinctive style of Adirondack architecture. All are clustered on a mountain side, the site of a former hotel and resort. The center of the grounds is a courtyard with an artificial pond and the steam engine from the Marion River Railroad. People climb on the engine, sit on Adirondack chairs, and leisurely stroll about, enjoying a spectacular view of mountains that overlook Blue Mountain Lake. The lake and its inflowing streams are the source of the Raquette River that we encountered at Tupper Lake, Long Lake and Buttermilk Falls.
Be sure to visit the log cabin of the Buck Lake Club and examine the mineral specimens mounted in the stone fireplace that include agate, garnet, rose quartz, flint, obsidian, feldspar, tourmaline and magnetite.

B4: STRATIGRAPHY AND TERRESTRIAL TO SHALLOW MARINE ENVIRONMENTS THE POTSDAM GROUP IN THE SOUTHWESTERN OTTAWA EMBAYMENT
David G. Lowe
Department of Earth Sciences
University of Ottawa
Ottawa, Ontario
Canada
This field trip showcases the stratigraphy and sedimentology of the Potsdam Group exposed along the southwestern Ottawa Embayment, south of the St. Lawrence River in Jefferson and St. Lawrence County of New York State (Figures 1, 2). The trip will highlight some important regional unconformities which have been used to correlate the Potsdam Group from this area to the north and east, as well as two contrasting fluvial styles in the Potsdam Group: braided perennial fluvial and sheet-like perennial fluvial. Refer to figure 2 for stop locations. We will begin at Wellesley Island and travel a loop that covers some classic Potsdam sections along Highways 37 and 12.

STOP 1A. Braided fluvial architecture
interactions with active basement topography and an Early Ordovician marine transgression.
This section consists of perennial fluvial deposits of allomember 3 (Chippewa Bay Mb of the Keesville Fm) overlain by tide-dominated marine strata of Early Ordovician allomember 4 (the Nepean Mb of the Keeseville Fm). This section contains good braided fluvial architectures, interesting stratigraphy and evidence of interaction with the adjacent basement ridge directly to the south (Figure 9).
Perennial fluvial deposits of allomember 3 are present at the base of the outcrop. On the west side of the road, 2 - 3 channel belt successions are present, and are differentiated by abrupt grain size and paleoflow direction changes and by sharp bounding contacts. The lower channel belt (Figure 9A) consists mainly of tractional conglomerate facies and was a ∼WNW paleoflow. This is overlain by a sharp contact and a 1 – 1.5 m thick lateral accretion bar deposit at the base of the second channel belt succession (Figure 9A). Boulder debris increases upward and these thin debrites thin to the north. Clasts are presumably derived directly from the ridge to the south (Figure 9B). A laterally-discontinuous boulder debrite characterizes the base of the third channel belt, underlain by a sharp, angular erosional contact (Figure 9C). This “progradation” of boulder debris in the upper part of the second channel belt succession and base of the third is interpreted to represent the presence (possibly active) of the basement ridge directly to the south during deposition of allomember 3 (Figure 9B).
The contact between Allomember 3 and 4 is best viewed some 50 m north, and is present on both sides of the road. Early silcrete nodules of probable groundwater origin are present 1 – 1.5 m from the top of Allomember 3. The base of Allomember 4 is marked by an erosional contact overlain by coarse-tail graded conglomerate. Approximately 1 m above this are gutter casts (visible on the west side of the road) above which fully marine strata onlaps underlying transgressive lag deposits (this onlap is very low angle and is best viewed at the next stop).

STOP 1B. Wellesley Island, I-81
SAFETY NOTE: This section gives a broader view of the previous section, and is located only ∼200 m to the west. The best view is from the west side of the I-81, looking east across the I-81.
The Allomember 3-4 contact is well displayed here (Figure 10a). The base of Allomember 3 is a north-dipping erosional contact overlain by a ∼50 – 75 cm thick conglomerate. The conglomerate is onlapped by tabular-bedded tide-dominated marine strata of Allomember 4. The lower contact and conglomerate are interpreted as a transgressive surface of erosion and transgressive lag, respectively (Figure 10A). Although marine strata appear to onlap the lag deposit here relatively passively, this contact was accompanied by erosional gutter casts at the previous stop, suggesting local erosion by waves and tides.
Some instructive braided fluvial architectures can be seen farther north that demonstrate the initiation of a new channel belt succession and subsequent bar migration (Figure 10b, c). Again, we should be looking east across the I-81. The paleoflow here is to the NW to the N, generally coming out of the outcrop. A broad but very obvious confluence scour marks the base of the second channel belt succession from the previous stop (Figure 10b-c). Low to moderate angle dune cross-stratified laterally to downstream accreting elements (DA/LA on Figure 10b-c), interpreted as the deposits of bars- are built up into a compound bar deposit. A smaller scour element is present near the top of this compound bar succession, and is interpreted as the deposit of a cross-bar channel.

STOP 2. Ephemeral and perennial fluvial contact
This stop shows the contact between ephemeral fluvial and perennial fluvial systems in Allomember 3 (Chippewa Bay Mb of the Keeseville Fm). The ephemeral fluvial section is best viewed in the northern part of this section (stop 2A), and the ephemeral–perennial contact is exposed just 100 m to the north (Stop 2B).
Stop 2A: Ephemeral fluvial strata are exposed here, and diagnostic architectural elements include UFR bars and antidune cosets. These architectures are better exposed at later stops (stops 4, 5 and 6).
Stop 2B: The contact between ephemeral and perennial fluvial systems is exposed here, 100 m to the south of 2A along route 26 (Figure 11). The contact is sharp and horizontal (with minor, cm-scale incision). The lower ephemeral fluvial deposits are characterized by their tabular bedding, whereas the overlying perennial fluvial deposits are clearly trough cross-stratified, dominated by dune-cross stratification with mean paleoflow to the north. This contact is correlated elsewhere (Figures 4, 16a and 18c) and is interpreted to record a change in climate from semi-arid to humid.

STOP 3. Allomember 2-3 contact
Near Millsite Lake on Cottage Hill Road, east of Redwood, NY
Precambrian basement as well as Allomembers 2, 3 and 4 are exposed here (Hannawa Falls Mb of the Ausable Fm; and the Chippewa Bay and Nepean Mbs of the Keeseville Fm, respectively). All of these are exposed on the south side of the road, uphill from the Millsite Lake access parking. Grenville basement and the Precambrian unconformity are visible from the road. A ∼50 cm regolith is present just above the basement. Overlying that is a 2 m section of aeolian red beds of the early Late Cambrian Allomember 2 (Hannawa Falls Mb of the Ausable Fm). An excellent exposure of the allomembers 2 – 3 unconformity is exposed on a small ridge just a small way (∼10 m) south over the ditch and into the wooded area. The unconformity is erosional and lithified rip-up clasts of allomember 2 are present in the lower few cm of Allomember 3 (Figure 12a-b). Lithification of allomember 3 is characterized by hematite rims and kaolinite cementation (Figure 13), interpreted to have formed pedogenically by the formation of a laterite paleosol. This contact is correlated to others in the Ottawa-Embayment and Quebec Basin, and represents a long period of time (possibly ca. 10 Myr) and significant changes in provenance, depositional setting and paleogeography (from basin –wide correlations, detrital zircon analysis and regional paleoflow analysis).
Allomember 2 (Latest Cambrian – Earliest Ordovician Chippewa Bay Mb of the Keeseville Fm) is relatively thick (∼10 m) and consists mainly of ephemeral fluvial facies. Specific architectures and divisions of facies associations into ephemeral and perennial fluvial deposits are not differentiated here due to poor exposure of the section.
Allomember 3 (Early Ordovician Nepean Mb of the Keeseville Fm) is exposed in the upper ∼2.5 m of the section (near the top of the hill) and consists of medium-scale tabular dune and ripple cross stratification that is locally bioturbated by a filter-feeding Skolithos ichnofacies (including Skolithos, Diplocraterion, and Arenicolites).

STOP 4. Antidune cross-stratification
North ledge outcrop on the north side of Highway 37: between Redwood and Hammond, NY
This is a relatively small outcrop of ephemeral fluvial strata from Allomember 3 (Chippewa Bay Mb of the Keeseville Fm). Here one of the most diagnostic elements indicating sustained deposition by supercritical flows is present: antidune cosets and convex-up formsets (Figure 14). This is interbedded with UFR bar form elements, dominated by upper flow regime plane bed indicative of fast-moving and probably shallow unidirectional flows (Figure 14).

STOP 5. Aeolian and chute-and-pool stratification in an ephemeral fluvial setting
North and south of Highway 37, between Redwood and Hammond, NY ).
This is another exposure of Allomember 3 (Chippewa Bay Mb of the Keeseville Fm) that highlights some aeolian and upper flow regime facies of the ephemeral fluvial facies association. On the north side of the road is a small outcrop that contains the remains of an aeolian dune, complete with lee-side grain flow deposits and “pinstripe” wind ripple stratification (Figure 15a). On the south side of the road is a natural pavement, with rib-and furrow structures formed by the intersection of 3D dune cross-strata, showing a paleoflow of 355o (∼N). Past this and down a steep slope of rubble (beware of broken glass here! – proceed with caution) is an outcrop exposing the same interval as at the last stop. This outcrop exposes some of the best examples of up-flow dipping chute and pool and cyclic step stratification (Figure 15b-c) recording the deposition from extremely fast-moving highly-concentrated supercritical flows.

STOP 6. Perennial-Ephemeral fluvial contact (again!) and cyclic steps
North and south of Highway 37 and on Cemetery Road, near Hammond
This is another exposure of the ephemeral – perennial fluvial contact within allomember 3 (Chippewa Bay Mb of the Keeseville Mb) and one that also showcases some thick cyclic step supercritical bedform stratification. On the north and south of Highway 37, the ephemeral fluvial deposits with moderate angle upper flow regime macroforms are incised by dune cross-stratified channel elements of perennial fluvial deposits (Figure 16a). On the east branch of Cemetery Road, south of Highway 37, a ∼N-S oriented section reveals the “macroforms” actually consist of thick, up flow-dipping cyclic step stratification (Figure 16b). Incredibly, cross-stratification formed by dunes is preserved on the stoss-side of cyclic steps, climbing and dipping in the opposite direction of the upflow-dipping chute and pool backsets (Figure 16b). On the south part of this exposure, antidune cross-stratification and formsets are present (Figure 16b).

STOP 7. Potsdam-Theresa contact
Highway 12, just north of Pleasant Valley Road .
This is a classic stop, showing the Potsdam—Theresa contact. The uppermost Potsdam is the typical tide-dominated marine facies association characterizing Allomember 4 in the area. The contact at this location is interpreted as a probable conformable but drastic facies change, into deeper water, storm- and tide- dominated shelf setting. The Theresa here is largely characterized by dolomite cemented fine- to medium-grained sandstone (locally coarse) with normally-graded beds, 2 -15 cm thick, consisting of sharp to diffuse upper plane bed laminations overlain by wave and/or combined flow ripples and bioturbated fine argillaceous sandstone. These beds are interpreted as tempestites (e.g. Selleck, 1993). Hummocky cross-stratification (HCS) in the Theresa is probably largely inhibited due to a relatively strong unidirectional flow component caused by reinforcing ebb tides. Experiential work has shown that only small unidirectional velocities mixed with strong oscillatory wave-generated currents are needed to inhibit HCS (Dumas et al., 2005).

STOP 8. Marine flooding and a groundwater silcrete linked to the Black Lake Fault.
Highway 12 near Chippewa Bay.
This stop consists of two small outcrops on either side of Highway 12. These outcrops show the contact between Allomembers 3 and 4, characterized here by a sharp erosional discontinuity and a flooding surface separating ephemeral fluvial strata below from marine strata above. A massive silcrete horizon is present here at the very top of allomember 3, with conical intraclast breccias, originally described by Selleck (1978). A recent honours study focusing the genesis of this silcrete was undertaken by Ed DeSantis at the University of Ottawa. Ed’s findings were that: (a) the silcrete is of groundwater origin, rather than pedogentic origin due to the preservation of primary sedimentary structures, occurrence of nodules and lack illuviation features; (b) multiple generations of silica overgrowths are present, indicating multiple precipitation-dissolution events; (c) the silica was likely derived from feldspars in basement rocks, given that the underlying Potsdam contains no evidence of feldspar dissolution; and (d) movement of fluids along, and movement of the Black Lake fault itself contributed silica and also stresses that generated the conical breccias in these relatively well-indurated horizons. The close spatial relationship between the distribution of the pervasive silica horizon, the conical breccias and the Black Lake fault provide good evidence for this relationship (DeSantis, 2014). The stratigraphic location of the silcrete (at the top of allomember 3) although suggestive of a pedogentic origin, is interpreted to have resulted from transgressive erosion down to the well-indurated and erosionally-resistant horizon.

STOP 9. Stratigraphy, sedimentology and diagenesis- this stop has it all!
Highway 12 just south of Schermerhorn Landing
Allomembers 2, 3 and 4 are all exposed in this section. Allomember 2 is present at the westernmost part of this exposure, where it is at least ∼8 m in thickness (the base is not exposed, Figure 4), which contrasts with the thin preservation of this unit at Stop 3 and lack of preservation elsewhere, and attests to the variable preservation of allomember 2 in this area. Here it consists of large-scale cross-stratification formed by the migration of aeolian dunes. Inversely graded grain flow avalanche cross-strata are present, and can be found on the south side of the highway. Allomember 2 also has a blocky structural fabric, which appears to be truncated by the allomember 2-3 contact (Figure 17A).
Allomember 2 (∼10 m thick) consists of planar-tabular stratified ephemeral fluvial deposits, with a few ∼west dipping cyclic step deposits towards the top of the unit, and some ∼NE-dipping subaqueous dune cosets as well as aeolian sand sheet deposits. The massive silcrete at the top of allomember 3 is well-developed here. It is coincident with bedding in the western part of the exposure, but towards the east it cuts across the bedding at a low angle, dipping east (Figure 17B). Minor conical autoclastic breccias are present here within the silcrete, and appear to be spatially associated with vertical fractures in the underlying sandstone.
At the very easternmost part of this outcrop and on the south side of the road, the eastward-dipping silcrete horizon is onlapped by marine strata of allomember 4 (Figure 17B). Elsewhere, allomember 4 has been eroded, which shows that the silcrete horizon is currently more resistant to erosion as it probably was in the Latest Cambrian- Earliest Ordovician.

STOP 10. Cyclic steps and the ephemeral-perennial fluvial contact.
Highway 12, south of Goose Bay
This outcrop exposes a thick (∼12 m) section of allomember 3, including a lower section of ephemeral fluvial strata dominated by upper plane bed stratification, with some ∼SW dipping cyclic steps and chute and pool stratification, and an upper perennial fluvial section dominated by ∼N – NE dipping dune and current ripple cross-stratification with an overall “sheet-like” architecture. Thick (up to 1 m) sets of SW-dipping cyclic step cross-stratified sandstone are exposed at the western part of the outcrop on the south side of the road (Figure 18A). Farther to the east, where the outcrop thins on the south side of the road, low angle and sigmoidal cross-stratification is present, interpreted to have formed by chutes and pools (Figure 18B). Moving east again, the outcrop thickens and exposes a sharp contact between the ephemeral and perennial fluvial deposits (Figure 18C). The large-scale cross-stratification beneath the contact was originally assumed to be from an aeolian dune; however, it lacks the characteristic inversely graded grain flow strata and it shallows to the west, grading into upper flow regime facies. Possible interpretations for this feature include: (a) a large cyclic step deposit, (b) an asymmetric or oblique channel element filled with upper plane bed strata, or (c) it is in fact an aeolian dune that has been onlapped by upper flow regime strata. The contact between the ephemeral and perennial deposits is sharp, but without significant incision. The “sheet-like” nature of the perennial fluvial deposits is interpreted to record deposition and accretion in very wide and shallow channels.

STOP 11. The Great unconformity and ephemeral fluvial architecture.
Highway 12, east of Alexandria Bay
This is another “classic” Potsdam outcrop, exposing the basal unconformity with Precambrian basement rocks. At this location, all of the Potsdam strata unconformably overlying the Precambrian is correlated to allomember 3, and interpreted as ephemeral fluvial in origin. Allomember 3 appears to onlap a basement knoll at the easternmost part of the outcrop; however, some structural tilting of the lower ∼4-6 m of allomember 3 is also vaguely evident above the basement feature. The architecture is sheet-like and dominated by planar laminations with occasional isolated west-dipping cross-beds and scour fills. Sheet-like packages generally 1.5 – 2 m thick are interpreted as discrete deposits of depositional lobes in an ephemeral fluvial setting. The best way to explore this outcrop is to walk from the west to the east, essentially moving up section. Many subtle features can be observed moving up-section; including aeolian wind ripple stratification and lags, isolated aeolian dune cross-stratification, upper flow regime plane bed, minor scour fills, possible cyclic steps and chutes and pools, ripples and adhesion structures and stratification. Near the top of the section is the contact between ephemeral and perennial fluvial facies associations.

B5: ST. LAWRENCE COUNTY, NEW YORK: NO LONGER AN AREA PRODUCING QUALITY MINERAL SPECIMENS?
Michael R. Walter
Nicholville, New York 13672
It is so often said when describing the localities available to mineral collectors, “That location is exhausted”. One might get the impression that there are no longer places where fine mineral specimens can be collected. This statement could not be further from the truth when it concerns St. Lawrence County, New York. As one of the most mineralogically productive counties in our country it continues to offer field collectors ample opportunity to find exceptional mineral specimens. Numerous localities described as exhausted have produced significant finds (often better than historic specimens in museums) in contemporary time. Exhausted, “I say not”!
Participants on this field trip will be led to two sites often referred to as exhausted but still available to collectors: Powers farm in Pierrepont, St. Lawrence County, New York and the Wolf property near Edwards, St. Lawrence County, New York will both be visited. There will be approximately three hours of collecting time at each site. To participate meet Mike Walter at the Pierrepont Highway Department parking lot at 9 am on Saturday October 11th, 2014. The Highway Department is found at the intersection of county routes 24 and 29 in the town of Pierrepont. From there participants will be taken to the first site, the Powers farm approximately two miles away.

Starting Place
OHM 2019: The coordinates for Stop 1 are really for here - the starting place.

STOP 1. Powers Farm, Pierrepont, St. Lawrence, County, New York
Common Mineral Species: Dravite, Quartz, Calcite, Phlogopite, Pyrite, Fluorapatite, Talc Less Common Mineral Species: Chalcopyrite, Magnetite, Tremolite, Actinolite, Marcasite, Goethite, Diopside, Microcline, Malachite, Sphalerite, Synchysite-(Ce), Vermiculite, Gold *many unusual pseudomorphs are found at this location, as well.
The tourmaline species from this location is now being described as dravite, as per new International Mineralogical Association (I.M.A.) guidelines. Previously known as uvite these lustrous black crystals can reach impressive dimensions. Individuals of large size are uncommon but have been known to reach 30 cm. More commonly they are in the one to three cm range. Although the target species is normally the these fine dravite crystals other species can be desirable. The mica, phlogopite, can form beautiful, large, double terminated crystals to 15 cm in size. Quartz can be found in two distinct generations. First generation quartz is found in unattractive milky to gray colored tessin habit crystals. The second generation quartz is highly prized by collectors forming in prismatic crystals to 10 cm in length of alpine quality.
Many other species can be found in well-formed crystals but most are less common. Pseudomorphs are numerous and include talc after quartz and scapolite, quartz after phlogopite, pyrite after pyrrotite, and quartz after diopside, to name a few.
The reality is that land owners are being more restrictive in allowing access to their land but many locations remain accessible if collectors simple ask for access and act responsible when it is granted. This location, however, has been open to the public since the mid 1800’s as a fee site. The land owner, Bower Powers, Jr. charges five dollars and visitors may keep all the mineral specimens they find. Field trips from area colleges, high schools, and mineral clubs frequent this location often and fine specimens are collected with regularity (fig. 1). There will be no fee charged for our visit.

STOP 2. The Wolf Property (also known as the Morgan Farm)
Common Mineral Species: Diopside, Microcline,Tremolite, Mica
Less Common Mineral Species: Pyrite, Apatite, Molybdenite
It seems that this site has been seldom visited over the past 40 years. It is on private property and we will have permission for this one time collecting effort. It has the reputation as being "exhausted" but that is not true. A large feldspar vein and the road cut where enormous diopside crystals were collected in the early 1960s are now likely depleted, but other diggings at this location will produce specimens. There has been very little collector activity at the site in the past 40 years and the location holds strong potential for further finds of quality specimens.

B6: HISTORY AND GEOLOGY REVIEW OF MAGNETIC IRON MINING IN THE WESTERN ADIRONDACKS
Jerome Zaykoski C. P. G.
Jefferson Community College
Coffeen Street, Watertown, NY 13601
jzaykoski@sunyjefferson.edu
Today when people talk about mining, they generally have an adversarial opinion of this industry. In addition, many fail to recognize the importance mining provides to today’s economy. There is an old saying that goes “if it can’t be grown, it must be mined”. Today, most mining in New York State revolves around the sand, gravel and construction aggregates industry. However, New York has a rich history of mining that goes back into the early 18th century and the dominant material mined was iron. Much of this iron was mined in the Adirondacks and, not only an important commodity to the early colonists for everyday use, it also played a critical role in the fight for independence from Great Britain. The purpose of this trip is to examine the history of western Adirondack iron mining, along with an overview of the regional and site specific geology. We will visit and discuss the mine history and limiting factors that governed the life of the Clifton Iron Ore, Jayville and Benson Mines.

B6: HISTORY AND GEOLOGY REVIEW OF MAGNETIC IRON MINING IN THE WESTERN ADIRONDACKS

STOP 1. Jayville Mine
Leonard and Buddington (1964) describe eleven principal workings at the mine with the Benson No. 1 shaft (N 44.16032°, W 75.18929°) located on the northwest side of the hill believed to have supplied most of the ore shipped from Jayville (Figures 17 and 18). This shaft is inclined about 65° NW and has a slope length of 350 feet. There was no processing of the ore at the mine, other than hand sorting with low grade ore being cast aside. Ore was loaded onto rail cars and shipped to furnaces in Scranton and Bethlehem.
Leonard and Buddington (1964) describe the other workings as follows. The author has provided the Latitude and Longitude information for each;
Fuller No. 1 (Open Cut) (N 44.15938°, W 075.18979°): Located roughly 180 feet south southwest of Benson No. 1, the cut is roughly 100 feet deep with ore analyses showing 47 and 52 percent Fe.
Benson No. 2 (N 44.15960°, W 075.18960°): Located less than 100 feet southwest of Fuller No. 1, the shaft is inclined at 55° W, 60 feet deep through 12 feet of ore ranging in grade from 37 to 49 percent.
Adit (N 44.16000°, W 075.18958°): Located about 150 feet south of Benson No. 1 shaft, this adit opens up a lens of ore 60 – 70 long and 20 feet wide.
Fuller No. 2 Shafts (N 44.15813°, W 075.19044°): Located roughly 400 feet southwest of Benson No. 2 and is said to be 50 feet deep, with ore averaging 54 percent Fe.
Fuller No. 3 (N 44.15803°, W 075.19029°): Located roughly 100 feet southeast of Fuller No. 2 shafts, a 30 to 40 foot opening accessed by a drift.
New York No. 1 (N 44.15951°, W 075.18842°): This shaft is located roughly 170 feet due south of the Leary shaft, inclined to the northwest and 300 feet deep with short drifts at the bottom. Ore averaged 47 percent Fe.
Hart No. 2 (N 44.15787°, W 075.18842°): Located slightly north and west of the former train station, this shaft is 60 feet deep.
Hart No. 1 (N 44.15685°, W 075.18890°): This shaft is located just to the west of the former railroad station and is roughly 300 feet deep, with ore shoots 10 feet thick. Ore analyses averaged 55 percent Fe.
New Find (Hustler mine): This small pit is located off the map, but contained ore analyzed at 52 and 58 percent.
Leary (New York No. 2) (N 44.16016°, W 075.18910°): Located roughly 120 feet southerly of the building foundation in the gravel pit, this shaft is roughly 50 feet deep with ore analyses ranging 22 to 38 percent Fe.
The Jayville mine lacked any ore processing facilities other than simple hand separation of low grade ore. There are many old shafts and drifts around the site and caution is advised as all are flooded.
We will walk around and visit several shafts and drifts that accessed the mine, as well as look at some of the few foundation features found at the mine.

STOP 2. Clifton Furnace at Clarksboro
Clifton Iron Company was founded by Zebulon Benton and sporadic mining operations began in 1863 (Leonard and Buddington, 1964). In April, 1864, the New York State legislature passed a bill allowing Clifton Iron Company to construct a 23.5 mile railroad from the mine to the Rome, Watertown and Ogdensburg Railroad at East DeKalb. The scarcity of iron after the Civil War made it impossible to use iron rails and maple timbers were used instead. Major mining operations then began in 1865 and a charcoal furnace was constructed at Twin Falls on the Grasse River. Later in 1868, a new furnace was constructed at the mine, located roughly 2.0 miles southwest of Twin Falls. The mill at the mine succumbed to fire in September, 1869, forcing the mine to be shuttered (Palmer and Thomas, 1969).
In 1940 the M. A. Hanna Company purchased the mine and operated it until 1952 when it closed for good. All structures were torn down at the mine.
At this stop we will carefully cross the sluiceway and visit the old furnace, ore and charcoal piles. The access to the island is a hemlock log that must be shimmied across. You may or may not wish to access the island.
Today, the mine itself, located 2.0 miles south southwest of Twin Falls is owned by the Clifton Hunting Club. The club has a strict no entry policy and therefore we will not be able to visit the mine.

STOP 3. Benson Mines at Star Lake
MINING AND CONCENTRATING METHODS
The open pit mine is now flooded (Figure 19). When in operation it utilized standard quarry methods to remove ore. A standard bench used nine inch holes drilled to a depth of 55 feet. The ore required roughly one half pound of explosives per long ton. Once broken, the muck pile is excavated with large 5 cubic yard electric shovels, loading an average of 25 tons into large dump trucks. The ore is then hauled to the 54 inch primary gyratory crusher located in the pit (Figure 20). The ore is then carried via conveyor belt to two 18 inch secondary gyratory crushers and then discharged into 500 ton silos that act as surge bins to feed the fine crushers (Tillinghast, 1948).
Originally, only magnetic ore was concentrated. However methods were later developed to process the nonmagnetic ore known as martite (Figure 21). Magnetic ore travelled from storage silos to one of three rod mills, where it was ground to pass a 20-mesh screen. From here, the ore was fed to three magnetic separators, with the waste then cycled back through for additional grinding and magnetic separation. From here, the concentrate is dewatered and transferred to the sintering plant. Tailings are pumped to a waste pond (Tillinghast, 1948).
The martite was concentrated in two sections, a rod mill section and an aerofall section, with the difference being the method of reducing the ore to a liberation size. In the rod mill, ore is ground to 18 mesh and the oversize is returned to the rod mill. The 18 mesh ore is sent to a set of rougher spirals that utilize the different specific gravities. The concentrate from the rougher moves to a cleaner spiral and the middlings return to the rougher spiral feeder tank. The tailings are classified and the minus 100 mesh portion sent to a floatation section (Jones & Laughlin, n.d.).
The concentrate is then sent to storage silos at the sintering plant. Here the roughly 64 percent concentrate is mixed with five percent coal, one percent lime and 30 percent fine sinter that is obtained from screened finished sinter. The mix is ignited and passed through an oil fired chamber where suction is applied. The entire sintering process is completed in eleven minutes when the sinter is sent to be screened and cooled. Once cooled to around 350° F., sinter is loaded onto railroad cars and shipped to steel plants at Pittsburgh and Aliquippa, Pa., and Cleveland, Ohio (Jones & Laughlin, n.d.).
BENSON MINES EPILOGUE
The mine continued operations until 1978 (Lupulescu et al, 2014). However problems, from the past industrial activity began to show in 1987 (Jesmore, 2014 NYSDEC, Personal Comm.). After almost ten years of inactivity, the quarry water rose to a level that began to float spilled petroleum to the surface and an oil sheen was discovered on the Little River. New York State Department of Environmental Conservation undertook an investigation and found that the network of underground pipes that carried fuel oil to various parts of the plant had corroded to the point where large volumes of petroleum were leaking directly to the ground. In all, it is estimated that roughly 1,000,000 gallons of fuel oil was lost. A groundwater dam was erected between the spill area and the Little River. Recovery wells were sunk and pumping activity has recovered about 350,000 gallons to date. In addition, the investigation revealed a number of PCB issues as well as a solvent spill and mercury problem (Figure 19), (Jesmore, 2014, NYSDEC, Personal Comm.).
Stop 1. View of the flooded mine looking northward. Here we will discuss the overall geology of the mine.
Stop 2. Amoeba Pit. This is a brief look at the small pit located westerly of the main workings.
Stop 3. Northern waste rock piles. The waste rock in this area was removed from the northern end of the mine where pegmatites intruded. Magnetite and martite can be found around the mine. This waste rock dump is a good location to collect minerals. Some minerals such as sillimanite, calcite, garnet, bornite, hornblende, tourmaline and molybdenite can be found. Dark green fluorite cubes have been found, though these are rare.
Stop 4. Here we will stop and discuss the processing of the ore and look at the remains of the plant.
Stop 5. A continued look at what is left of the mill.
Stop 6. Remediation site.