NYSGA 2002
New England Intercollegiate Geological Conference
94th annual meeting
New York State Geological Association
74th annual meeting
Guidebook for Fieldtrips in New York and Vermont
Lake George, New York
Organized and Edited by
James McLelland and Paul Karabinos
Meeting hosted by Colgate University and Williams College
THIS GUIDEBOOK AND MEETING ARE DEDICATED TO
W. YNGVAR ISACHSEN - (1920 - 2001)
GEOLOGIST, FRIEND, AND TEACHER EXTRAORDINAIRE
For many years, Yngvar served as Senior Scientist at the New York Geological Survey in Albany. During that time, he investigated, and became familiar with, the nature and problems of rocks in his beloved Adirondacks. He also was keenly attuned to the Proterozoic of the Hudson Highlands and the remarkable tectonic history of the Taconic belt of New York. His efforts were crucial in compiling the 1970 edition of the Geological Map of the State of New York that all of us still rely on to this day. At about the same time, he and Jim Olmsted organized an immensely successful and important conference on the "Origin of Anorthosite and Related Rocks". Yngvar served as editor for the volume generated by that conference and bearing the same name Today the "Origin of Anorthosite and Related Rocks" (New York State Museum and Science Service Memoir 18, 1969) continues to serve as a centerpiece for investigators in this very significant field.
Throughout his career Yngvar was deeply interested in brittle fracturing and produced important studies on "zero-displacement crackle zones in the Adirondacks and mylonitization along the Carthage-Colton Zone separating the Highland and Lowland terranes. In his later years he also conducted research on Catskill fracture patterns that may record a meteorite impact. One of his most important pieces of research was that concerning the age of uplift of the Adirondack Dome. He provided compelling evidence for a recent (20 Ma) date for this event.
Besides his scientific contributions, Yngvar was one of the great teachers of his time. He was blessed with the ability to transmit ideas in remarkably clear terms and to generate interest as he did so. Those of us who now spend our time investigating Adirondack problems are forever grateful to this Teacher of Teachers for showing us the way, pointing out problems, making astute suggestions, and ALWAYS encouraging us. Whatever advances come out of current and future Adirondack research, they will forever carry the stamp of Yngvar's influence.
Finally, we note that Yngvar was a dear and important friend to many of us. Although we might not always agree, we could always exchange ideas in constructive and cooperative ways. A day in the field with Yngvar was a day spent generating both ideas and excitement - and a great deal of warm laughter. When the day was done, it was time to join in good food and good drink, and there was never a better companion in these endeavors - many of them at his lovely Brant Lake home with his gracious wife, Stasia. The times were many; the memories are grand.
We miss him; we think of him; and we honor him. Hail Yngvar!

A1: GEOLOGY AND GEOCHRONOLOGY OF THE SOUTHERN ADIRONDACKS
James McLelland
Dept. Geology
Colgate University (Emeritus) and Skidmore College
Saratoga Springs, NY
Lara Storm and Frank Spear
Dept Earth and Environmental Studies
Rensselaer Polytechnic Institute
Troy, NY
INTRODUCTION
The location of the Adirondacks within the larger Grenville Province is shown in Fig. 1.Topographically, the Adirondacks are divided into Highland (H, Fig 2) and Lowland (L, Fig. 2) sectors. The former is underlain principally by orthogneiss, the latter by paragneiss rich in marble, and the Carthage-Colton Mylonite Zone separates the two regions (Fig 2a,b). The region has experienced multiple metamorphic and intrusive events and large-scale ductile structures are common (McLelland, 1984;McLelland et al., 1996). U-Pb zircon geochronology (Table 2) indicates that the oldest igneous rocks exposed are ca 1350-1300 Ma tonalitic, arc-related plutons (Fig.1a) intrusive into older Highland paragneisses of uncertain age (McLelland et al., 1996). The oldest igneous rocks in the Lowlands consist are ca. 1200 Ma granodiorites (Fig.1a) intrusive into older paragneisses of uncertain age (Wasteneys et al., 1999). In both the Highlands and Lowlands metapelitic migmatites 1200-1300 Ma contain anatectic material of -1170 Ma age. Following intrusion ca. 1207 Ma granodiorites in the Lowlands, leucogranitic and tonalitic rocks were emplaced at ca 1172 Ma (Fig.la,b) and were accompanied by deformation and metamorphism (Wasteneys et al., 1999) assigned to the latest, culminating phase (ca. 1220-1160 Ma) of the Elzevirian Orogeny of Moore and Thompson (1980).
From ca 1160-1150 Ma the entire Adirondack-Frontenac region was intruded by anorthosite-charnockite- mangerite-granite (AMCG, 1155 Ma, Fig. 1b) magmas that are associated with four anorthosite massifs (Marcy, Oregon, Snowy, Carthage) recently dated directly at 1155 ± 10 Ma by SHRIMP II zircon techniques. Early age determinations of the anorthosite (McLelland and Chiarenzelli, 1990; Silver, 1969) were based on multigrain dating of MCG granitoids that exhibit mutually crosscutting relationships with the Marcy Anorthosite Massif (MM, Fig; 2b) and are interpreted as coeval with it. The absence of direct dating of the anorthosite was due to the sparse igneous zircon populations in rocks of this composition, a condition that posed a serious obstacle to the study of anorthosites prior to the advent of single grain TIMS and SHRIMP II methods. Current SHRIMP II direct dating of the Adirondack anorthosite massifs demonstrates that both they and their associated ferrodiorites and granitoids were emplaced at 1155 ± 10 Ma and that the entire complex represents a classic AMCG suite.
The final major events of Adirondack evolution comprise: 1) the emplacement of the Hawkeye granite suite at ca 1095 Ma (Fig 2b) followed almost immediately by 2) high-grade metamorphism (Storm and Spear, Appendix to this article; Spear and Markussen, 1997; Bohlen et al., 1985; Valley et al., 1990) resulting in vapor-absent, peak: granUlite facies conditions in the Highlands (T - 750°-800° C, P - 6-8 kbar) and associated with widespread recumbent, isoclinal folding and the development of intense penetrative fabrics. This granulite facies metamorphism and deformation are assigned to the collisional Ottawan Orogeny of Moore and Thompson (1980), evidence of which occurs throughout the Grenville Province (cf. Rivers, 1997; McLelland (et al., 2001 b). Toward the end of this major orogenic event much of the Adirondack region was intruded by late- to post-tectonic leucogranites (ca. 1055 Ma, Fig 1a) belonging to the Lyon Mt. Granite (LMG) and thought to be related to delamination and extensional collapse of the orogen (McLelland et ai, 2001). The Elzevirian and Ottawan Orogenies, taken together, comprise the Grenville Orogenic Cycle (ca. 1350-950 Ma) of Moore and Thompson (1980).

A1: GEOLOGY AND GEOCHRONOLOGY OF THE SOUTHERN ADIRONDACKS

STOP 1. PECK LAKE MIGMATITIC METAPELITE. (30 MINUTES).
This exposure along Rt 29A just north of Peck Lake is the type locality of the Peck Lake migmatitic metapelites that consist of restitic sillimanite-garnet-biotite-quartz-oligoclase melanosome and quartz and two-feldspar leucosome of approximately minimum melt composition (McLelland and Husain 1986). Small red garnets are common in the leucosome and, in most cases, grow across foliation. The leucosomes occur as irregular, elongate bodies generally parallel to foliation but quite commonly exhibiting crosscutting relationships with respect to the restite and to one another.
Based upon these compositional and crosscutting relations, the Leucosomes are interpreted as anatectites and a reasonable metapelitic source rock can be prescribed by reintegrating their composition with that of the restite (Table 3) to yield a greywacke-slate precursor. The restriction of these partial melts to within the migmatite is thought to be the result of near-solidus, hydrous melting that would cause ascending melts to intersect the solidus as they began to rise. Close inspection of the leucosomes reveals that, in their most pristine configuration, they consist of coarse granite and pegmatite. In low strain zones elsewhere in the Adirondacks it is manifestly clear that the leucosomes originally formed an anastomosing arrays of veins, dikes, sheets, and pods. Subsequent high strain resulted in rotation into psuedoparallelism and commonly produced disruption that caused separate grains of white feldspar some of which show elongate tails - ie, the rock was on its way to becoming a mylonitic "straight gneiss". At the last stop (Stop 9) of this trip we shall see the end result of this process in a series of platy, stretched, and highly grain size reduced and mylonitic equivalents of the rocks seen here. The anatectic origin of these units is further suggested by the uncommon, but not rare, occurrence of plagioclase- and/or garnet-rimmed hercynitic spinel in the restite. Recently, Bickford and McLelland have run a U/Pb zircon pilot study on the age of the anatectites and have found that they contain 1220-1250 Ma cores, 1020-1050 Ma metamorphic rims, and relatively thick, nicely zoned mantles that fall into the interval 1190-1170 Ma. Some of the rims show zoning, but this is minor compared to the mantling zircon. These results demonstrate that majority of the anatectites were produced by partial melting during the culminating Elzevirian Orogeny dated at ca 1210-1170 Ma (Wasteneys and McLelland, 1999) and are not of Ottawan origin. Currently, we are continuing this research by utilizing both zircon and monazite geochronology.
This outcrop provides an excellent mesoscale example of Adirondack structure (Fig. 7a). The overall strike of foliation is N60-70W and dips vary from north to south and pass through the vertical demonstrating that the roadcut defines a large minor isoclinal, recumbent fold. Smaller, meter-scale, isoclinal folds are also well exposed and are aligned paralle1 to the roadcut-scale fold. All of these are similar in style, and parallel to, the very large, regional Canada Lake isocline or nappe (F2 - Fig 7a). It is clear that the Canada Lake nappe stage of deformation rotates an earlier foliation (due to F1) and has a moderately dipping axial planar foliation that intersects the earlier foliation high angles. The time interval between these foliations remains unspecified, but it is clearly post-anatexis and probably all Ottawan (ca 1090-1030 Ma).1t appears that the Ottawan bas successfully obliterated most Elzevirian fabrics. Note that some of the smaller minor folds contain apparently terminated compositional layers that could very well represent Elzevirian isoclinal noses. If so we are looking at isoclinally refolded isoclines.
Migmatitic metapelites of this sort occur throughout the southern and western Adirondacks as well as in the Adirondack Lowlands. We interpret them as flysch sequences of shale and greywacke that were being shed from the Elzevirian magmatic arcs that dominated the region from ca 1400-1300 Ma. Tonalites and granodiorites of the arcs locally crosscut the metapelites and provide a minimum age for them. In a broad. sense, they are thought to be coeval. The metapelites of the Adirondack Lowlands are compositionally similar, but are thought to represent a different, younger, and uncorrelative arc environment.

Nick Stoner Inn
Nick Stoner Inn on west side and Nick Stoner Golf Course on east side of Rt29A-10.

STOP 2. IRVING POND QUARTZITE. (20 Minutes).
The Irving Pond quartzite unit cores the Canada Lake isocline and is folded back on itself. At map scale it is exposed across strike for over 3000m but its "true?" thickness is on the order of 1000m. Notwithstanding, it represents an enormous volume of orthoquartzites with minor pelitic intercalations. Its minimum age is unknown but is being investigated by zircon geochronology. Similar quartzites along the St. Lawrence River contain zircons as young as 1300 Ma, and the same may prove to be true for the Irving Pond. It is suggested that these quartzites may have been deposited along the present eastern margin of the Adirondack-Green Mt block between ca 1250 and 1200 Ma (see Fig. 4). During that interval, this margin was passive as the block moved westward towards the Elzevirian subduction zone that dipped westward beneath Laurentia and its leading margin, ie, the Adirondack Lowlands (Fig. 4b,c). Upon collision, (ca 1200-1170 Ma), the accumulated sandstones were deformed and metamorphosed. Although this scheme is speculative, it is consistent with the little that is known about Elzevirian events.
Within the clearing there are three small, but informative outcrops. The first consists of 30-40 cm scale layers of pure quartzite together with 2-3 cm scale metapelitic 1ayers of. These dip gently to the southeast. The bulk of the Irving Pond quartzite consists of layer upon layer of the pure quartzite with little, if any, intervening metapelite. The second exposure is located a few tens of feet to the west where the rocks form a low ledge running uphill. The southern termination of the ledge that faces the clearing shows steeply dipping layers that are discordant with underlying layers. The discordance is due to some sort of ductile shearing in the outcrop but does not affect the following interpretation. Looking back to the first outcrop, it is clear that the gently dipping layers seen there can be projected up above the ground surface so that they must have been situated overhead at this locality; however, they must also have suddenly dipped steeply and rotated through the vertical in order to have their present configuration, i.e., they form recumbent isoclinal folds of the F2 set. The trend of this recumbent, isoclinal fold axis is ˜N70W and the plunge is ˜15 deg. southwest, i.e., it is a minor fold of the Canada Lake isocline family and of Ottawan age. At the base of the southernmost outcropping, and below a small ledge, there is preserved a foot-scale isoclinal nose. Inspection of the geometry makes it clear that the outcrop preserves an isoclinally refolded F1 isocline -perhaps of Elzevirian age. Farther down the clearing towards the highway a clean outcrop consists of pure, glassy quartzite dip slopes exposed together with dark layer-like bodies of fine-grained pyroxene-plagioclase granulite. The chemistry of the granulites approximates that of diabase and locally exhibits microscopic diabasic texture (Mclelland and Husain, 1986). The igneous nature of the granulites is further manifested by the xenoliths of quartzite present in them. Elsewhere the metadiabases are isoclinally folded and the limbs of these folds clearly truncate 10-15 cm-scale intrafolial isoclinal folds (F1) that are rotated by the isocline. We interpret this older set of isoclinal minor folds to be of Elzevirian age. A recently blasted F2 fold in metadiabases is shown in Fig 7h. The metadiabases have not been dated, but they are similar to some mafic AMCG rocks and are tentatively assigned a ca 1150 Ma age.

STOP 3. CANADA LAKE CHARNOCKITE. (20 MINUTES).
Large roadcuts expose the type outcrops of the Canada Lake charnockite (Figs. 8-11). This highly deformed orthogneiss has a granodioritic composition, and consists of 20-30% quartz, 40-50% mesoperiliite, 20-30% oligoclase, and 5-10% mafics including small, sporadic grains of orthopyroxene. The exposures exhibit the drab olive color typical of charnockites around the world. In the woods these rocks tend to weather pink and exhibit a maple sugar brown weathering rind. The unit is ˜500m thick and consists throughout of relatively homogeneous granitoid with pegmatites and minor amphibolitic layers. A multigrain U-Pb zircon age of 1251 ± 33 Ma (McLelland and Chiarenzelli, 1990) indicating that this unit belongs with other calc alkaline rocks of broadly Elzevirian age. Mapping along its contact for a total distance of ˜300 km, has not revealed any crosscutting features, but xenoliths of country rock have been recognized and substantiate an intrusive origin. The apparent conformity is attributed primarily to extreme and ductile tectonism. Tn addition, an original conformable, sheet-like form is quite possible. Rocks of similar age and composition are found in the Green Mts. of Vermont (Ratcliffe and Aleinikoff, 1991).

STOP 4. ROYAL MT. TONALITE (30 MINUTES).
Steep roadcuts exposed across from the Canada Lake Store expose typical tonalitic rocks that are relatively common within the southern and eastern Adirondacks, the Green Mts of Vermont, and the Elzevir terrain of the Central Metasedimentary Belt of the Canadian Grenville. In all of these occurrences, the tonalites manifest the presence of magmatic arcs that existed along the southeastern margin of Laurentia during the interval ca 1400-1200 Ma diagnostic of the Elzevirian. Within the Adirondacks, multi- and single-grain TIMS U-Pb zircon geochronology indicate emplacement of the tonalitic magmas at ca 1350-1300 Ma. The present outcrop has been dated by both TIMS methods and the single grain age is constrained at 1307 ± 2 Ma (Aleinikoff, pers comm., 1991). Well-documented examples of these arc- terrains extend to the Llano uplift and Van Horn areas of Texas (Mosher, 1999; Patchett and Ruiz, 1990; Roback, 1996) and attest to a global-scale system that may have been an ancient analogue of the present day East Indies arc. During the Elzevirian, various arcs must have collided and amalgamated, and continental arcs may have come into existence as well. These details remain to be unraveled, but in the meantime we note that the Elzevirian came to a close at ∼1210-1170 Ma with the collision of the Adirondack Highlands-Green Mt block with the Andean-style arc then existent along the southeastern edge of Laurentia (Fig 4c). McLelland, in Wasteneys et al (1991), refers to this collisional event as the "culminating Elzevirian Orogeny" that closed out the Elzevirian interval of arc magmatism and amalgamation in the area.
The whole rock chemistry of the Adirondack calcalkaline rocks are shown in Figs (9,10,11) where their calcalkaline affinities are clearly visible. In addition, eNd characteristics are presented in Fig. 8 and demonstrate that the tonalites represent juvenile additions to the crust from the mantle, and partial melting of the older rocks can produce younger Adirondack granitoids. Neither the Sm-Nd data, nor any other isotopic data, give any hint of pre-1350 Ma crust in the Adirondacks and suggest that the original arc must have been of ensimatic origin.
Disrupted amphibolitic sheets are present here and within Adirondack tonalites in general. Their origin is enigmatic but they do not appear to have been derived locally. Their elongate, sheet-like character suggests that they may represent coeval mafic dikes of the sort commonly observed in tonalites. It is also possible that they represent enclaves incorporated from an amphibolitic source region. Isotopic studies are needed here.
In a few places the tonalite crosscuts the migmatitic metapelites seen at Stop 1. This fixes a minimum age for the latter at ca 1300 Ma. As indicated in the discussion at Stop 1, the tonalites and meta pelitic rocks may represent an original, essentially coeval, magmatic arc-flysch system. .

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

Low roadcut in migmatitic metapelites.

Avery's Hotel on left (west) side of Rt 10.

Roadcut through granite, gabbro, and tonalite.

Pink AMCG granitic rocks, gabbro, and small exposure of fine-grained anorthosite.
Pink AMCG granitic rocks, gabbro, and small exposure of fine-grained anorthosite. A large inclusion of calcsilicate in pink granite is exposed and interpreted as a xenolith.

Roadcut of metasedimentary quartzites and pelitic rocks together with anatectites

Fault breccias in charnockite.

Pink granitic rocks of the Piseco anticline.
Long roadcut of mylonitic ribbon gneiss in pink granitic rocks of the Piseco anticline. Ribbons trend N70W and plunge 10-15 SE parallel to both recumbent F1 isoclines and upright F2 such as the Piseco anticline. A multigrain U-Pb zircon date from this outcrop gives an age of 1155 ± 10 that is interpreted as the age of magmatic emplacement.

STOP 6. RIBBON GNEISS IN THE CORE OF THE PISECO ANTICLINE. (30 MINUTES).
A long, low roadcut on the left (northwest) side of Rt 8 exposes outstanding examples of ribbon gneiss developed in a megacrystic facies of the Piseco core rocks. The overall composition here is similar to that of the Rooster Hill megacrystic charnockite examined at Stop 5, but here the original feldspars and quartz have been extended into ribbons on the order of 60cm x 0.25mm x 0.05mm (McLelland, 1984). Assuming interstitial quartz aggregates of ˜1 cm diameter, the current dimensions indicate an extension of 6000%, i.e., if this strain were to be equally distributed throughout a 1 km thick block of crust that block would be 60km long,25 km wide, and 0.005 km thick. Of course, such high ductile strain cannot be integrated through the entire crust, but the numbers provide some flavor for the strain involved.
The ribbon gneisses exposed here are folded into a minor anticline that rotates the foliation of these L»S tectonites. The axis of this minor fold trends N70W and plunges 10-15 SE. This orientation is parallel to the axis of the F3 Piseco anticline, to the F2 recumbent isoclines of the Adirondacks (Figs. 6,7), and to the axis-parallel elongation lineations associated with these isoclines. The fact that all of these structural elements are parallel to one another indicates that they all shared some common kinematic experience during their tectonic evolution. In order to identify the common experience, we first note that all of the deformation in these ca 11S0 Ma rocks must be of Ottawan age. Next we note that minor isoclinal folds in the outcrop lie with their axial planes in the foliation and their axes defining rod-like structures parallel to the N70W extension lineation. The low dip of the lineation suggests that the mechanism responsible for it was likely to have been thrust faulting out of the southeast and towards the northwest. This would have resulted in the stretching that elongated the quartz and feldspar. The thrusting could also have rotated earlier fold axes into the thrust plane and parallel to the direction of tectonic vergence. However, rather than rigid rotation of the axes, it is proposed that the early folds became too ductile to buckle and their axes began to "flow" as passive markers in the direction of tectonic transport, i.e., the recumbent fold axes began to undergo increasing curvature in response to the velocity gradients in the ductile flow field. The ultimate result of this process is sheath folds, and it is suggested that the isoclinal folds with axes parallel to the ribbon lineation represent parts of sheath folds. It is further suggested that the Canada Lake isocline may itself be a large sheath fold whose southern closure is obscured by Paleozoic overburden. The large F3 folds such as the Piseco anticline and the Gloversville syncline are interpreted as structures formed by NS constrictional forces that arose in response to the EW elongation associated with the large scale, ductile thrusting of the Ottawan Orogeny. Later NNE upright folds of the F4 set were superimposed on the F2 sheath folds and the F3 corrugations by continued, but waning, NW-SE contraction associated with the NNE suture situated somewhere beneath, or beyond, the present day Coastal Plain. Finally, it should be remarked that the ribbon lineations and upright lineation-parallel corrugations described above are similar to those encountered in core complexes. The problem with this alternative is that the late- to post-tectonic Lyon Mt Granite emplaced at ca 1050 Ma is unaffected by these fabrics. Accordingly, we associate the fabrics with ductile Ottawan thrusting. Orogen collapse similar to core complex kinematics is thought to have taken place at ca 1050-1030 Ma.

STOP 7. MARBLE AND CALCSILICATE. (30 Minutes).
Exposed in roadcuts on both sides of the highway are examples of typical Adirondack marbles and their associated lithologies, i.e., garnetiferous amphibolite, calcsilicates, and various disrupted blocks of both internal and external members. The disruption and boudinage attests to the extremely ductile behavior of the marbles. Also exposed are vertical, crosscutting, and undeformed veins of tourmaline-quartz symplectite. Besides calcite, the marbles contain diopside, tourmaline, sulfides, and graphite. The graphite has a biogenic carbon signature, and the marbles are thought to have formed inorganically via stromatolite accretion in an evaporitic environment. Some calcsilicate layers consist of almost monomineralic white, Mg-rich diopside crystals up to 10 cm in length. These are probably the result of metasomatism by fluid phases. The presence of the approximately Mg-pure assemblage tremolite=>enstatite+diopside+quartz+H2O allowed Valley et a/(l983) to calculate a fluid with X(H2O) = .11-.14 for the reaction. Elsewhere in the outcrop localized occurrences of wollastonite reflect the presence of H2O as an agent lowering CO2 activity. In both cases, the metamorphic conditions were T=71°C and P = 7 kbar. These cases provide excellent examples of how the dilution of a fluid phase by a second constituent lowers activities and allows reactions to run to the right at P, T conditions below that for the pure fluid.
In the High Peaks region there exist many well-exposed examples of places where marble and calcsilicate xenoliths occur within anorthosite. This relationship documents that the marbles are older than ca 1150 Ma. It is likely that they formed during the same shelf sequence event that was posited for the Irving Pond quartzite precursor sands, ie, ca 1220-1200 Ma.

Roadcuts of steeply dipping metasediments.

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

STOP 8. MASSIF ANORTHOSITE AND FERRODIORITE OF THE OREGON DOME. (30 MINUTES).
This small, but instructive roadcut has outstanding examples of two important facies of anorthosite as well as a typical ferrodiorite dike associated with massif anorthosite. The best vantage point for examining the anorthosite and ferrodiorite is on top of the roadcut. Whole rock analyses of these rocks are given in Table 1.
On climbing to the top of the outcrop at its south end, one immediately sees a distinctive dark dike of ferrodiorite filled with plagioclase grains (white, -An.3) and andesine xenocrysts (blue-gray, ∼An43 note reaction rims on the blue-gray andesine. The dike exhibits somewhat soft contacts and irregular veinlets shoot off in a fashion suggesting that the anorthosite was not yet wholly solidified when the dike intruded. Farther up the outcrop two ˜30cm-scale xenoliths occur in a ∼5-10 m-scale ferrodiorite; one of these is fine grained and the other coarse. The eastern contact of the ferrodiorite is situated just to the highway-side of the coarse xenolith. Calcsilicate xenoliths also occur in the ferrodiorite and are characterized by sulfidic staining near road level.
Farther up the outcrop the exposed rock consists entirely of anorthosite. A fine-grained, leucocratic facies forms a matrix to large (5-20 cm long) crystals of blue-gray, iridescent andesine that is typical of massif anorthosite. In places these crystals appear broken as if disrupted from a larger mass. An example of a still intact mass is found at the far end of the outcrop where a raft of very coarse grained Marcy-type anorthosite with ophitic to sub-ophitic orthopyroxene sits in a matrix of fine-grained leucoanorthosite that clearly disrupts the raft at its edges. Close inspection of the matrix reveals that its small pyroxenes exhibit subophitic texture; hence the finer grained facies must be magmatic. Workers in Adirondack anorthosite have always noted that grain size reduction produces leucoanorthosite similar to this this fine-grained facies, and examples of this may be seen at this stop. However, it is possible to distinguish between the grain size reduced and the fine-grained igneous varieties by noting that plagioclase in the former has the same composition as the coarse plagioclase, whereas the igneous variety is invariably 5-10% less anorthitic than the coarse plagioclase; in this case 52% versus 44% (Boone et at, 1969). Genetic interpretations and further details of the anorthositic suite are given in the main text.
The smooth, upper surface of the outcrop affords excellent opportunities to examine the garnet coronas or "necklaces" that are found throughout the Adirondack anorthosites. These are accounted for by the reaction Orthopyroxene + Plagioclase + Fe, Ti-Oxide => Garnet + Clinopyroxene ± SiO2. Spear and Markussen (1997) have studied this, and other garnet-producing reactions, and determined that they took place during isobaric cooling (P=∼6-7 kbar) from ∼700°-630°C during waning stages of the Ottawan Orogeny (1090-1030 Ma). McLelland et al (2001) have suggested that this reaction took place at ca 1050 Ma, which is the age of small, equant metamorphic zircons associated with the coronites. The zirconium was provided by Fe-Ti oxide, which accepts significant Zr into the Ti lattice site.
Recent SHRIMP II U-Pb dating of zircons from the ferrodiorite documents its age as 1155 ± 9 Ma, and this sets a minimum age for the anorthosite. Thirty kilometers to the northeast at Gore Mt, charnockite enveloping, and showing mutually crosscutting relationships with the Oregon Dome anorthosite (Lettney, 1969), yields a SHRIMP II age of 1155 + 6 Ma. These results document that the anorthosite series of the Oregon Dome was emplaced at ca 1150 Ma. This age is indistinguishable from emplacement ages of the Marcy massif and marks the time of emplacement of the Adirondack AMCG suite.

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

Charnockite on the north limb of the Glens Falls syncline.

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

A2: BEST KEPT GEOLOGIC SECRETS OF THE ADIRONDACKS AND CHAMPLAIN VALLEY
Mary K. Roden-Tice
Center for Earth and Environmental Science
Plattsburgh State University of New York,
Plattsburgh, NY 12901
INTRODUCTION
Finding clear, well-exposed field locations to help illustrate important aspects of petrology and structural geology to undergraduate students can be a challenge for a professor. In this field trip, we will visit the following four field stops in the Adirondack Mountains and Champlain Valley that are routinely used in my classes: the Cannon Point intrusive south of Essex; the Craig Harbor fault-line scarp in Port Henry; the intrusion breccia in Roaring Brook on Giant Mt.; and Split Rock Falls crackle zone south of New Russia. In structural geology, the class works on field exercises that illustrate normal faulting at Craig Harbor and different orientations of joint sets at Split Rock Falls. Petrology students compare Proterozoic anorthosite containing xenoliths of metasedimentary rocks from the Roaring Brook intrusion breccia on Giant Mt. with a Cretaceous laccolith of trachyte porphyry on the shore of Lake Champlain. All of these outcrops are easily accessible and provide excellent opportunities for experiential learning.
REGIONAL GEOLOGY
Summary
The Adirondack Mountains are a regionally elevated exposure of Grenville age (ca. 1.0-1.35 Ga; McLelland et al., 1988; McLelland and Chiarenzelli, 1990; Mezger et al., 1991; McLelland et al., 1996) high-grade metamorphic rocks in northern New York state. This field trip will examine granulite-facies metaplutonic rocks (McLelland and Isachsen, 1986) in the Adirondack Highlands which comprise the central and eastern portions of the massif. An abrupt transition in topography occurs between the Adirondack Highlands and the Lake Champlain Valley between New York and Vermont with local relief along Lake Champlain varying from ∼170 to 370 m. The flatter terrain of the Champlain Valley in New York is composed of Paleozoic sedimentary rocks ranging in age from Cambrian Potsdam Sandstone through Ordovician carbonates of the Trenton Group (Fisher, 1968; Isachsen and Fisher, 1970). Both the Proterozoic metamorphic rocks and Paleozoic sediments are intruded by Early Cretaceous (105-146 Ma) alkalic plutons and dikes throughout northeastern New York, northern New England and southern Quebec (McHone and McHone, 1999). In this field trip, outcrops of Precambrian anorthosite, gneiss, metagabbro and marble, Paleozoic carbonates and a Cretaceous trachyte porphyry will be visited (Figure 1).

A2: BEST KEPT GEOLOGIC SECRETS OF THE ADIRONDACKS AND CHAMPLAIN VALLEY

STOP 1. CANNON POINT LACCOLITH. (45 Minutes)
This is the north end of the intrusion which is located on private property owned by the Crater Club. We will be walking south along beach at the low water mark. The intrusion is exposed along the shore and beneath private homes. If visiting this outcrop during May through October, permission to access the intrusion should be obtained from the Crater Club and homeowners. We will walk south along the lakeshore through the exposure of the Cannon Point intrusion and observe its contact relationships with the host Canajoharie Shale.

STOP 2. CRAIG HARBOR FAULTLINE SCARP. (1.5 Hours).
This is private property but permission to visit the site is easily obtained. The name and address of the owner are given on the posted sign at the entrance. We will begin by walking south along Lake Champlain and the railroad tracks to examine the Paleozoic carbonates. These tracks are used several times a day by both Amtrak passenger and freight trains. Caution and listening for trains is strongly advised. There is generally room to stand safely on the side of the tracks.
After studying the Paleozoic dolostones, we will proceed north along the tracks across Craig Harbor to see the faultline scarp, Proterozoic marble, gneiss, and metagabbro. The faultline scarp is visible just north of Craig Harbor. From the tracks, one can determine that the marble in the cliff has an unusual structure. We will get to look at the marble up close later. Orange-stained, strongly foliated, Proterozoic gneiss is the first rock type encountered along the railroad tracks north of Craig Harbor. A few feet further north, the fault contact between the gneiss and coarse grained, iron-oxide stained, metagabbro is visible. Further north across from a small railroad cut on the lakeside of the tracks, a rough path disappears uphill into the woods. We will scramble up it and view the magnetite prospect.
Be careful of the poison ivy!
After returning to the cars, we will drive a short distance within the campground to one of the campsites that borders the faultline scarp. At this location, there are fantastic exposures of Proterozoic marble with abundant and diverse deformed inclusions. A contact between the Proterozoic marble and the Ordovician Whitehall Dolostone can be seen along the faultline at the west end of the outcrop.
If permission to visit the campsite cannot be obtained, similar spectacular outcrops of marble with deformed inclusions can be seen along the lakeshore by walking north along the railroad tracks from the magnetite mine for about 0.5 mi. The entire site can be accessed from the south by walking along the railroad tracks from the marina south of the dolostone outcrops. Parking is available at the public boat launch in Port Henry. McHone (1987) gives a good description of this route to Craig

STOP 3. ROARING BROOK FALLS AND INTRtJSION BRECCIA. (2.5 Hours).
The trail to Roaring Brook Falls exits at the north end of the parking lot. After a few hundred feet, the trail forks. Take the uphill (left) trail that leads to the top of the falls (sign says Giant). The lower trail goes to the base of the falls. We will climb for ∼0.3 mi. (15 minutes) to reach the overlook at Roaring Brook Falls. The climb is generally an easy grade. At the trail junction sign to Giant/Nubble, Washbowl, turn right and follow trail to the top of the falls. Here, a beautiful view of the Great Range of the Adirondack High Peaks, Including the Wolf Jaws, Armstrong, Gothics and Sawteeth, as well as the Ausable Club in the valley, can be obtained. We will examine the gabbroic anorthosite exposed in the stream bed and the overlook and look for an elusive syenite dike with blue microperthite described by Jaffee and Jaffee (1986) which I have never been able to find.
After returning to Roaring Brook trail, continue climbing for ∼0.4 mi. (12 minutes) to where the trail crosses over Roaring Brook. Here, we will leave the trail and bushwhack up Roaring Brook ∼0.2-0.3 mi. to reach the intrusion breccia outcrop. The best route alternates between the stream bed and the adjacent wooded area. After 15 minutes of bushwhacking, large pavement outcrops of gabbroic anorthosite occur in the stream bed (44°09.152', 73°44.928'). These exposures contain the inclusions of Proterozoic metasedimentary rocks which are quite spectacular.

STOP 4. SPLIT ROCK FALLS-ZERO DISPLACEMENT CRACKLE ZONE. (40 minutes).
The large outcrop across Rt. 9 (west side) contains some interesting features including large coarsely-crystalline inclusions of anorthosite with plagioclase megacrysts in a strongly foliated gabbroic anorthosite matrix, slickensides and two thin basaltic dikes. On the east side of the road is Split Rock Falls. We will examine the crackle zone at both the upper and lower falls. The dominant fracture trends are easily seen in both areas. The lower falls can be reached by climbing down through the woods.

A3: EARLY PALEOZOIC CONTINENTAL SHELF TO BASIN TRANSITION ROCKS: SELECTED CLASSIC LOCALITIES IN THE LAKE CHAMPLAIN VALLEY OF NEW YORK STATE
James C. Dawson
Distinguished Service Professor
Center for Earth and Environmental Science
State University of New York
Plattsburgh, NY, 12901
james.dawson@plattsburgh.edu
TRIP ABSTRACT
Well-exposed classic localities in the Lake Champlain Valley of New York provide opportunities to examine the transition from Early Cambrian to Early and Medial Ordovician continental clastic and carbonate shelf sequences (e.g. Upper Cambrian Potsdam Group, Lower Ordovician Theresa Formation and Beekmantown Group, and Middle Ordovician Chazy Group) to the Medial Ordovician foreland basin carbonates and calcareous shales (Middle Ordovician Trenton Group).
INTRODUCTION
This field trip is designed to begin at the SUNY Plattsburgh campus at 8:00 a.m., and end near Plattsburgh, NY about 4;00 p.m., in time for participants to drive to Lake George, NY to make their lodging arrangements, have dinner, and attend the conference welcoming party. This trip visits many of the localities that are commonly visited in the introductory geology classes at SUNY Plattsburgh. The field trip description is largely based on the work of Fisher (1968) and Isachsen, et al. (2000) and is not based on any significant research by this author. The trip is included among the package of four trips originating in Plattsburgh, NY, in order to round out the offerings so that they include an opportunity to see a portion of the Lower Paleozoic section. Other faculty at SUNY Plattsburgh are involved in offering opportunities to see Adirondack bedrock geology, glacial geology, and hydrology and geomorphology localities in the area.

A3: EARLY PALEOZOIC CONTINENTAL SHELF TO BASIN TRANSITION ROCKS: SELECTED CLASSIC LOCALITIES IN THE LAKE CHAMPLAIN VALLEY OF NEW YORK STATE

Note the 'old base' stone buildings
Note the 'old base' stone buildings constructed in the 19th century of Potsdam Sandstone.

STOP 1. Montreal Member of the Glens Falls Limestone and Dike
Montreal Member of the Glens Falls Limestone of the Trenton Group and Lamprophyre (Monchiquite Dike). You are parked on the apron of the former Plattsburgh Air Force Base marina and boat launch. To reach the outcrop hike, approximately 2000' north, along the cobble beach shore of Lake Champlain. The Pleistocene sediments forming the bluff between the railroad track and Lake Champlain along the way include a contact between earlier Lake Vermont varved and later Champlain Sea sediments. The Montreal Member and a cross cutting Lamprophyre Dike are exposed to varying degrees depending on the water level of Lake Champlain. The Montreal Member is a medium-bedded limestone with some shale partings, that weathers to medium dark gray. Crab Island (the smaller island that can be seen in Lake Champlain two miles southeast of this stop) takes its name from the trilobite, Isotelus, that is commonly found in exposures of the the Montreal Member on the island. Fisher (1968) maps the Lamprophyre Dike as a monchiquite dike, containing phenocrysts of titan augite with some biotite and barkevikite (a monoclinic amphibole), phenocrysts of olivine and titanomagnetite.

Pass the Fort Brown historic site markers on your left/west.
Pass the Fort Brown historic site markers on your left/west. Fort Brown was the left/west flank of the American defense during the land portion of the Battle of Plattsburgh on September 6 - 11th, 1814.

STOP 2. Cumberland Head Argillite of the Trenton Group.
The Cumberland Head Argillite is exposed in this road cut on both sides of Route 9, although the exposures are better on the left/west side. The roadcut (Fisher, 1968 - Figure 27) is typical of the Cumberland Head Argillite. The thin bedding (it looks like varves) of fairly distal turbidite deposition can clearly be seen as more calcareous layers alternate with the thinner, pale buff-colored quartz-silt bearing layers being slightly more resistant to erosion. Fresh surfaces of the rock are uniformly black; however, the thin layers can still be distinguished. The rocks in the vicinity are folded; but, the strike is generally N35E with a dip of 23 degrees N. Be sure to look north along Route 9 and see the slight dip that the highway makes north of the road cut. STOP 3 is just beyond the dip.

STOP 3. Providence Island Dolostone of the Beekmantown Group.
The Providence Island Dolostone is exposed in this road cut on both sides of Route 9, although the exposures are better on the right/east side (Fisher, 1968 - Figure 17). Before examining the outcrop, be sure to look south along Route 9 and recognize that you can see Stop 2. The Rocks at STOP 3 are nearly horizontal, compared to the steeper dips at STOP 2 and the entire Chazy Group and lower portion of the Trenton Group (some 400' or more of rock section) are missing. Fisher (1968) has mapped the edge of the Cumberland Head Allochthon, a Taconic thrust fault, as passing between these two trip stops on the basis of missing section at both the surface and in a 200' deep water well on Cumberland Head. The Providence Island Dolostone is a thick bedded, massive unit that weathers to the buff color characteristic of dolostones. With careful examination, horizons that display significant soft sediment rip ups and other features characteristic of supra-tidal environments. The dolostone is sometime vuggy and fracture zones that have been filled with secondary crystalline calcite are common.

STOP 4. Cumberland Head Argillite at Point au Roche State Park.
Walk the short distance from the parking lot to Lake Champlain and examine the large outcrop of Cumberland Head Argillite along the shore. The folding and jointing associated with the Taconic thrusting is very obvious. The rock is essentially the same as that seen at Stop 2 and again the thinning bedding of the fairly distal turbidite deposition can be seen as the more calcareous layers alternate with the thinner, pale buff-colored, quartz silt bearing layers being more resistant to erosion.

A hump in the road will indicate that you are crossing the Ingraham Esker (Denny. 1972).

A series of sand pits can be seen to the right/east
A series of sand pits can be seen to the right/east after you first cross the Esker. Route 9 crosses the Ingraham Esker again, just before the hamlet of Ingraham on the left/west.

STOP 5. Day Point Limestone of the Chazy Group.
(THIS IS NOT A HAMMER STOP.)
This road cut is part of the upper part of The Day Point Limeston, probably the Fleury Member of Oxley and Kay (1959). It is a medium gray calcarenite that is both cross bedded and regular bedded in this small exposure. Fisher (1968) mentions this outcrop as containing pelmatozoan debris.

STOP 6. Valcour Limestone of the Chazy Group at Sheldon Lane.
(THIS IS NOT A HAMMER STOP.)
This is the first of two opportunities that we will have (see also STOP 10) to see the bioherms in the Valcour Limestone.

Road to old Quarry
Just past a red barn on your right/east there is an amateurish 'stone henge' and an old road. The road leads to the abandoned International Lime and Stone quarry where the Pamelia Dolostone, the Lowville Limestone and the Isle LaMotte Limestone formations of the Black River Group (Fisher 1968 - Figure 25) were once exposed. A few years ago the remainder of this outcrop was mined and the quarry floor today consists of Valcour Limestone. None of the Black River Group remains.

STOP 7. Crown Point Limestone of the Chazy Group at the Northway in Chazy
Maclurites magnus Le Sueur, Death Assemblage
(THIS IS NOT A HAMMER STOP.)
To reach the outcrop cross the open area and enter the woods on a vague trail behind and to the right of the obvious transmission pole. The vague trail heads north a few yards in a white cedar forest and then bends left/west toward the Northway. The outcrop is close to the Northway, a short distance north of the Fiske Road Bridge and is illustrated in Fisher. 1968 (Figure 20).

STOP 8. Lacolle Melange at the Chazy Railroad Bridge.
The road cut is located along the west side of Route 9 as you walk toward the railroad bridge and is figured in Fisher, 1968 (Figure 30). Fisher, 1968 maps this locality and the STOP 9 locality as part of the Tracy Brook (normal) Fault. The wide range of angular clasts are composed of angular sandstone fragments in a sandstone matrix, presumably a tectonically crushed portion of the Keeseville Sandstone.

Cross Corbeau Creek.
Corbeau Creek was studied, as a monitored watershed, by a succession of students at SUNY Plattsburgh in the 1970s and early 1980s. In recent years SUNY Plattsburgh students have been engaged in watershed studies on the Altona 'Flat Rock' lands owned by the William H. Miner Institute and some of this work is described in Field Trip A6 of this field conference.

STOP 9. Lacolle Melange at Coopersville.
The small outcrop is located adjacent to the south side of Route 9B and is figured in Fisher, 1968 (Figure 29). The breccia clasts include a wide range of sizes and the lithology of the clasts is predominately made of carbonate fragments that appear to be of Black River group and Trenton Group origin.

STOP 10. Valcour Limestone at the Bechard Quarry, Near Kings Bay.
To reach the quarry walk across the agricultural field to the higher vegetated are that is visible to the west. Once you reach the top of the vegetated mound the quarry can be seen as illustrated in Fisher, 1968 (Figure 22). The bioherms consist intertidal stromatolites (algae), stromatoporoids, and bryozoans with the various colonies being completely compatible (Pitcher, 1964b) and building their colonies on top of one another.

STOP 11. Keeseville ss of the Potsdam ss at the Clinton Farm Supply, Champlain.
The outcrop is exposed along the north side of the northernmost of the two driveways and in the flat area behind the Farm Supply building. In a series of articles (Erickson, 1993a. Erickson, 1993b. Erickson and Bjerstdt, 1993. Erickson. Connett, and Fetterman, 1993) the stratigraphy and trace fossils of the Keeseville Sandstone and Theresa Dolostone/Sandstone have been described. STOP 11 (Erickson, 19938, and Erickson, Connett and Fetterman, 1993) consists of medium bedded, cream colored Keeseville Sandstone. At this locality large scale ripple marks, and an unusual abundance of trace fossils can be seen.

STOP 12. Theresa Dolostone/Sandstone at Champlain.
The Theresa Dolostone/Sandstone consists of an inter-layered, medium and thick bedded, quartz dolostone. Bjerstedt and Erickson (1989) and Erickson and Bjerstedt (1993) have described the trace fossils of the Theresa in detail and Skolithos can be located on the south wall of this quarry.

STOP 13. Keeseville ss of the Potsdam ss at the Gadway Preserve.
This stop requires driving in to the Gadway Preserve of the Adirondack Nature Conservancy and consists of three short STOPS within the STOP. Vehicles with low clearance should not attempt this and we will consolidate riders at the entrance.

STOP 13a. Large scale ripples
Continue west on the Gadway Road and stop and park at a bare rock rock wash. Carefully cross over a barbed wire fence to examine a pavement of large scale ripples in the Keeseville Sandstone.

STOP 13b.
Continue west on the Gadway Road and STOP and park at the next significant open area on the right/north and stop and park. At the north end of the open area near a red posted sign there is a 4 meter long (10 cm wide) Climachtichnites track. Yochelson and Fedonkin (1993) illustrate several similar examples of this trace fossil in the Keeseville Sandstone of northern New York. Additional trace fossils can be seen near the Adirondack Nature Conservancy sign.

STOP 13c.
Continue west on the Gadway Road and STOP and park at the next significant open area on the right/north and stop and park. At the north end of the open area near a red posted sign there is a 4 meter long (10 cm wide) Climachtichnites track. Yochelson and Fedonkin (1993) illustrate several similar examples of this trace fossil in the Keeseville Sandstone of northern New York. Additional trace fossils can be seen near the Adirondack Nature Conservancy sign.

STOP 14. Trough Cross Stratification in the Keeseville ss of the Potsdam ss
The trough cross stratification is located on the east side of the road, within what goes for a right of way. It is a small outcrop of well exposed cross stratification in the Keeseville Sandstone

Pass the historic Robinson Tavern on the left/east.

STOP 15. Ausable Arkose of the Potsdam ss on the Military Turnpike.
Cross to the ditch on the west side of Route 190 Military Turnpike and walk north. Excellent examples of the 'Basal' Member of the Potsdam Sandstone, a hematitic, feldspathic and micaceous sandstone can be found over much of the distance to the Blue Chip Farm driveway.

Pass the Rand Hill dikes on both sides of the road
Pass the Rand Hill dikes on both sides of the road with the best exposure on the left/east side of Route 190.

STOP 16. Crown Point ls with Lamprophyre Dikes
This extensive road cut along both sides of Route 374/Cadyville expressway contains two lamprophyre dikes on the south side of the road that intrude the Crown point Limestone with a strike of N70W and a dip 85 degrees N. Only the easternmost of the two dikes appears on the north side of the road. Interesting features, including baked zones, xenoliths and partial xenoliths, can be seen associated with the intrusions. The Crown Point Limestone is mainly a medium to dark gray calcilutite. Some of the large planispiral gastropod, Maculurites magnus Le Sueur are present along with trilobite and brachiopod fragments.

A5:LATE GLACIAL WATER BODIES IN THE CHAMPLAIN AND ST. LAWRENCE LOWLANDS AND THEIR PALEOCLIMATIC IMPLICATIONS
>
David A. Franzi. Plattsburgh State
University, Plattsburgh, NY 12901
John A. Rayburn, Binghamton
University, Binghamton, NY 13902
Catherine H. Yansa, Michigan State
University. East Lansing. MI 48824
Peter L.K. Knuepfer
Binghamton University
Binghamton, NY 13902
The late glacial legacy of the Champlain and St. Lawrence Lowlands and northeastern Adirondack Upland region (Figures 1 and 2) is recorded in the deposits and landforms associated with proglacial and marine water bodies that formed during deglaciation. The largest and most persistent proglacial lake in the region was Lake Vermont, which occupied the central Champlain Lowland and was dammed at its northern margin by the receding ice front (Chapman, 1937; Connally and Sirkin, 1969, 1973; Parrott and Stone, 1972; Wagner, 1972; Denny. 1974; DeSimone and LaFleur, 1985, 1986). Lake Vermont expanded northward until ice receded north of the St. Lawrence Lowland and allowed marine water to inundate the isostatically depressed St. Lawrence and Champlain Lowlands, thus forming the Champlain Sea (Chapman, 1937; Occhietti et al., 2001). Deglaciation of the region postdates deglaciation in the upper Hudson Valley (ca. 13.2 14C kaB.P.; Connally and Sirkin, 1971) and was completed prior to the Champlain Sea marine incursion (ca. 12.0 14C ka B.P. to 11.5 14C ka B.P.; Clark and Karrow, 1984; Fulton et aI., 1987; Anderson, 1988; Rodriguez, 1988; Occhiett
The chronology of lake phases in the Lake Champlain basin provides insight into the style and timing of Late Wisconsinan deglaciation, but may also provide information that is relevant to global paleoclimate studies. Broecker et a1. (1989) discussed the possibility that freshwater drainage from proglacial lakes within continental North America during the last deglaciation affected North Atlantic ocean circulation and thereby altered global climate. The Champlain Lowland occupied a strategic position during the Late Quaternary deglaciation of the northeastern United States. The north-south trending lowland served both as a source of cold meltwater and as a corridor fur the transmission of proglacial lake discharges from the Great Lakes Region to the North Atlantic. The region is also located at the juncture of two freshwater discharge routes. Discharges from Lake Vermont in the Champlain Valley and proglacial lakes in the eastern and central Great Lakes basin were initially routed southward through the Hudson Valley. Ice recession eventually opened the lower St. Lawrence Valley allowing proglacial lakes in the Champlain Valley and Great Lakes basin to drain northeastward to the Gulf of St. Lawrence. The Champlain Valley region is thus a key fur recognizing when and where large freshwater discharge events entered the North Atlantic.
On this trip we shan discuss the preliminary results of our on-going investigations of the geomorphic and stratigraphic record of the late glacial water bodies in the Champlain Lowland and the significance of meltwater outflow and through flow from the lowland to the North Atlantic during deglaciation.

ALATE GLACIAL WATER BODIES IN THE CHAMPLAIN AND ST. LAWRENCE LOWLANDS AND THEIR PALEOCLIMATIC IMPLICATIONS

Rae Brook exposure
The Rae Brook exposure lies on the east side of the small stream valley to your right. The base of the section consists of dark gray, calcareous diamicton (till). The diamicton is overlain by 1.0 to 1.3 meters of thinly laminated rhythmites, which are in turn overlain by marine clays. The faunal assemblages described by Cronin (1977,1979,1981) represent the late glacial transition from lacustrine to marine environments in the Champlain and St. Lawrence Lowlands approximately 11.6 to 12.0 14C ka. B.P. The bottom water temperatures and salinities at this time probably ranged from -2°C to 10°C and 0 to 18 ppt, respectively (Franzi and Cronin, 1988).

Low-relief marine beach ridges
Denny (1970, 1974) mapped low-relief marine beach ridges on the right (east) of Barnaby Road. The hummocky topography for the next 0.5 miles was mapped by Denny (1974) as recessional moraine.

STOP 1. COBBLESTONE HILL BEACHES. (20 minutes)
The beaches at this location were first described by Woodworth (1905a) and later by Denny (1974). The deposits consist predominantly of moderately rounded to well rounded, cobble gravel in multiple, low relief ridges or terraces that extend along the northern and eastern flanks of Cobblestone Hill at elevations between 206 and 17S meters above sea level. The highest ridges lie near the projected highest shoreline of the Upper Lake Fort Ann. Individual ridges are typically 1 to 2 meters high and 10 to 20 meters wide, and often extend laterally for more than 400 meters (Denny, 1974). The gravel is almost exclusively composed of Potsdam Sandstone that was presumably derived from the alluvial cobble to boulder gravel that composes Cobblestone Hill.
The large (0.2 to 1.4 meter diameter), angular boulders that comprise the core of Cobblestone Hill can be seen along the road a short distance above the highest beach ridge. The boulders of Cobblestone Hill represent material washed into Lake Vermont from the sandstone pavements by ice-marginal streams from the breakout of glacial Lake Iroquois (Woodworth, 1905a; Denny, 1974; Clark and Karrow, 1984; Pair et al., 1988). Reworking of these alluvial deposits by wave action with relatively little longshore transport probably formed the beach deposits (Denny, 1974).

Scarpit Road at Miner ("Million Dollar") Dam
Bear right and onto a concrete road (Scarpit Road) at Miner Dam. The Scarpit Road presents many hazards, especially for those driving it for the first time. Please drive slowly and cautiously. The road lies on the southwest flank of Cobblestone Hill following the abandoned shoreline of the former reservoir behind Miner Dam.
Miner Dam was part of a failed hydroelectric project initiated by William Miner in 1910 (Gooley, 1980). By the time of its completion in March, 1913, the concrete dam, known locally as the "Million-Dollar Dam", had a maximum height of over 10 meters and stretched more than 700 meters across the Little Chazy River valley. The design capacity of the reservoir was more than 3.S million cubic meters.
The inadequate flow of the Little Chazy River and ground water seepage through Cobblestone Hill, which formed the eastern flank of the reservoir, proved to be major design flaws for the project. A 10 to 15 cm layer of concrete grout was spread over more than 100,000 m2 along the flank of Cobblestone Hill (the Scarpit) to mitigate the seepage loss. A deep trench was excavated at the base of Cobblestone Hill behind the dam for the purpose of pouring a grout curtain to the underlying sandstone and thereby, presumably, sealing the northeastern flank of the reservoir. The dam and generating station were completed in 1913 but it took almost two years to fill the reservoir to capacity. The grouting effort was partially successful and the power generating plant began operation on January 21, 1915, more than four years from the beginning of the project (Gooley, 1980). The power plant produced electricity intermittently for seven years before mechanical problems forced the abandonment of the project.
Construction of a second dam, the Skeleton Dam (Gooley, 1980), approximately 1.5 km upstream was begun in 1920 to provide supplemental flow to the main impoundment The Skeleton Dam project, however, ended with the failure of the Miner Dam generating station and was never completed.

STOP 2. COBBLESTONE HILL ICE MARGINAL DEPOSITS. (40 minutes)
The Cobblestone Hill boulder deposits occur at two distinct elevations at this location. The upper level lies between 225 and 232 meters above sea level and may correspond to cobble and boulder deposits at a similar elevation near Bear Hollow, on the southwestern side of the Little Chazy River valley. The elevation of these deposits is close to projected elevation of the Coveville Stage if the Coveville shoreline is extended northward from where Chapman (1937) and Denny (1974) mapped the northernmost Coveville shoreline deposits in the Saranac River valley, assuming a northward isobase gradient of approximately 1.2 m/km. The lower level lies between 206 and 215 meters above sea level and corresponds to the boulder deposits observed at Stop 1. The lower level boulder deposits lie close to the elevation of the Upper Lake Fort Ann high stand shoreline (Chapman, 1937; Denny, 1970.1974).
The northeastern flank of Cobblestone Hill contains several large depressions that we interpret to be kettle holes. The northeastern ends of the kettles rise onto a broad terrace composed of beach deposits (Denny, 1970, 1974) at elevations between 201 and 204 meters above sea level. These beach deposits correspond closely to the elevation of Lower Lake Fort Ann.
We believe that these data indicate that the ice margin stood at Cobblestone Hill at the time of the Lake Iroquois breakout and that proglacial water levels in the Champlain Lowland dropped during deposition of the Cobblestone Hill boulder deposits deposition of the Cobblestone Hill boulder deposits.

Note the outcrop of Potsdam Sandstone on your left.
The largest boulders on Cobblestone Hill have long dimensions that exceed 3 m.

Entering Altona Flat Rock
The Scarpit Road makes a sharp right tum and the concrete pavement ends. The road emerges onto Altona Flat Rock within 30 meters of the tum. The transition from the northern hardwood forest on Cobblestone Hill to the jack pine barrens on Altona Flat Rock is abrupt at this location.

STOP 3. ALTONA FLAT ROCK SANDSTONE PAVEMENT AND JACK PINE BARRENS.
(20 minutes) The large areas of sandstone pavement provide habitat for some of the largest jack pine (Pinus banksiana) barrens in the eastern United States (Woehr, 1980; Reschke, 1990). Jack pine is a relatively short-lived «150 years), shade-intolerant, boreal species that maintains communities on the sandstone pavements because of its adaptations to flJ"e and ability to survive in an area with thin (or absent), nutrient-poor soils.
A large proportion of the pine barrens in northeastern New York are owned by a few public and private sector organizations. The William H. Miner Agricultural Research Institute is the largest landowner of pine barrens with almost 1000 ha (hectares) of jack and pitch pine barrens on Altona Flat Rock. New York State owns an additional 600 ha of the Altona Flat Rock barrens, approximately 100 ha of the Gadway barrens and 200 ha of pine barrens at The Gulf near Covey Hill. The Adirondack Nature Conservancy owns 222 ha of the Gadway jack pine barrens at Blackman Rock.
Plattsburgh State University and the William H. Miner Agricultural Research Institute have collaborated in research and teaching initiatives in the Altona Flat Rock pine barrens for more than 30 years. The hydrogeological equipment and instrumentation at Stops 2 and 3 are part of the Ecosystem Studies Field Laboratory (ESFL), a field station dedicated to undergraduate teaching and research in geology and environmental science. The field site offers an excellent geological. hydrological and ecological setting for illustrating the interdependence of normal processes and the effects of human activities on natural ecosystems. For the past three years the ESFL site has been the focus of the Plattsburgh Research Experiences for Undergraduates program, which is funded by the National Science Foundation and the William H. Miner Agricultural Research Institute. The reader is referred to Franzi and Adams (1993,1999) for a more detailed description of the Altona Flat Rock pine barrens and the Ecosystem Studies Field Laboratory Project.

A marine beach ridge
A marine beach ridge can be seen in the field on the right (south) near the intersection with Vassar Rd.

STOP 4. INGRAHAM ESKER AT THE KALVAITIS GRAVEL PIT. (40 minutes).
The Ingraham Esker is one of the most conspicuous glacial landforms in the northern Champlain Lowland. This pit contains esker fan deposits, such as described by Diemer (1988), and deposits resedimented by wave action in the Champlain Sea as described by Denny (1972. 1974). Most of the pit is cut into proximal to medial subaqueous fan gravel and sand. The resedimented deposits consist primarily of fossiliferous gravel that occur as dipping bedsets on the western flank of the esker. Individual beds are several centimeters to a few decimeters thick and are laterally continuous for several meters.

A gravel pit containing esker tunnel
A gravel pit containing esker tunnel and proximal subaqueous fan gravel and sand can be seen on the right (west) side of the road.

STOP 5. INGRAHAM ESKER AT ESKER ROAD. (10 minutes)
This location is the "West Pit" section of Diemer (1988), Most of the sediment consists of coarse-grained channel fill deposits. These deposits generally occur in lenticular beds that are meters thick, tens of meters wide and may be traceable for tens of meters in the flow direction (Diemer, 1988). Marine reworking of the esker deposits at this location is restricted to a thin layer (1 to 2 meter) of interbedded sand and gravel near the top of the pit.

Remains of esker to south
The esker was removed in the I-87 corridor but the ridge can be seen on the right (south) side of Stratton Hill Road east of I-87.

The esker ridge parallels Rte. 9
The esker ridge parallels Rte. 9 on the right-hand (east) side of the road. The ridge crosses the Rte 9 at 26.2 miles and continues its southward trend on the left-hand (west) side of the road.

STOP 6. MARINA SECTION AND LUNCH STOP. (80 minutes)
The bluffs along the shore of Lake Champlain extend for more than 1 km north from the former Plattsburgh Air Force Base marina. The bluffs probably contain a complete late glacial stratigraphic section, however, no single location contains all of the stratigraphic units. A massive gray diamicton lies at the base of the glacial section. The diamicton is exposed at the north end of the bluffs where it overlies striated bedrock. The upper contact is not exposed.
The base of the section near the marina consists of more than 3 m of dark gray clayey rhythmites, which were probably deposited as varves in glacial Lake Vermont. The rhythmites occur as clay and silty clay couplets that range from a few centimeters thick in the lower part of the section to thin couplets that rarely exceed a few millimeters in thickness near the top of the unit. Soft-sediment deformation structures are common. Rock and sediment clasts are distributed throughout the unit as individual clasts and in discrete layers along bedding planes. A deformed bed of medium sand that is 0 to 0.2 m thick occurs near the base of the exposed section. The lateral extent of this unit is not known. A thick reddish brown clay lamina occurs near the top of the rhythmite unit. This lamina is similar in nature and stratigraphic position to a red clay bed observed in cores taken north of Montreal, Canada (Jan Aylsworth, pers. comm.).
The rhythmites are conformably overlain by 1.5 to 2.0 m of laminated to thinly bedded fossiliferous marine mud. The mud facies coarsens upward to horizontally bedded silt and fine sand. The silt and sand unit is approximately 7 m thick and the unit coarsens upward. Individual beds range from a few centimeters to a decimeter or two thick and are generally normally graded. These deposits probably record the incursion and gradual regression of the Champlain Sea in the region.

STOP 7. AUSABLE CHASM DISCUSSION AND PHOTO OP.
(15 minutes) AuSable Chasm is one of the most unique scenic spots in the Champlain lowland. The AuSable River has carved a spectacular gorge that exposes a 135 m thickness of the Keeseville Member of the Potsdam Sandstone. The AuSable River also cuts through the upper marine delta noted in the road log, and thus, the cutting of the chasm postdates the Champlain Sea interval.

Lower Lake Fort Ann delta
Rte. 9 rises out of the AuSable River Valley and onto the surface of a delta that was built into Lower Lake Fort Ann. The delta surface elevation is approximately 156 meters above sea level.

STOP 8. KEESEVILLE INDUSTRIAL PARK EXPOSURE. (40 minutes)
The Keeseville Industrial Park section is exposed in a landslide scar on the south flank of the AuSable River. The river is deeply incised into a deltaic terrace graded to Lower Lake Fort Ann. The surface elevation of the delta surface is approximately 155 m.
A massive to crudely bedded, dark gray diamicton forms the base of the section. The diamicton is overlain by approximately 2 m of rhythmically laminated silt and clay couplets. Clay laminae are generally 1 cm or less thick and the silt laminae or beds range from about 0.5 to 4 cm thick. The silt beds are commonly internally laminated. The rhythmite section contains about 67 couplets. The rhythmite section is conformably overlain by approximately 7 m of deltaic silt and sand that coarsen upward to sand and gravel.
The sediments at Keeseville Industrial Park record ice recession from the AuSable Valley. The basal diamicton is interpreted to be a till and thus represents ice cover. The rhythmites are probably varves and thus record inundation of the lower AuSable Valley by proglacial Lake Coveville. Assuming that the entire varve sequence represents proglacial Lake Coveville and the overlying silt and sand record the drop of proglacial lake level to Upper Lake Fort Ann, then Coveville occupied the lower AuSable Valley for approximately 67 varve years before proglacial lake levels dropped to the Upper Lake Fort Ann level. The ice front may have receded about 30 km north to the Cobblestone Hill Ice margin over this time interval at an average retreat rate of approximately 0.45 km/yr.

Highlands Road offers spectacular views of Lake Champlain.
Burlington, Vermont lies directly across the lake at this point and is visible on a clear day.

A series of marine terraces lie to the left (east) side of the road.

A prominent Upper Lake Fort Ann sand and gravel spit
A prominent Upper Lake Fort Ann sand and gravel spit parallels the right (west) side of the road. Chapman (1937) identified this feature and measured its surface elevation as 161 meters above sea level.

STOP 9. LONG POND CORE. (40 minutes)
The Long Pond basin was a deep embayment in Lake Coveville. The drop to the Upper Lake Fort Ann level, however, left proglacial water levels in the Champlain Lowland below the threshold of Long Pond. The bottom sediment of Long Pond was vibracored in March, 2002. Wood obtained 1 cm above a horizon that may represent the drainage of Lake Coveville at this site yielded a date of 10.9 ± 76 14C ka B.P. (Wk - 10957). These discussions will continue at the Pok-O-MacCready Outdoor Education Center 0.5 miles (63.8 miles) west of this location where a portion of the Long Pond Core will be displayed.

Upper Lake Fort Ann delta
The Deerhead Road traverses an Upper Lake Fort Ann delta built by the North Branch of the Boquet River.

Top of a large Lake Coveville delta, The Plains
The road rises to the top of a large Lake Coveville delta, known locally as "The Plains". The "The Plains" delta was built by North Branch flow that was augmented in its early stages by outflow from proglacial lakes in the AuSable River basin to the west via the "Channel Belt" (Kemp and Alling, 1925; Diemer and Franzi, 1988).

There is a good view of "The Plains" delta to the left (south).

Channeled kame terraces
The road crosses a series of channeled kame terraces and eventually rises to the quarry.

OPTIONAL STOP 10. NYCO WOLLASTONITE QUARRY. (40 minutes)
The quarry operators have excavated deeply into ice-marginal stratified drift and diamicton on the flank of Oak Hill. The composition, stratification and texture of these deposits is highly variable. The stratified sediment probably represents sedimentation by ice-marginal streams flowing from the "Channel Belt" (Kemp and Alling, 1925) and local impoundments. The diamicton are probably till or sediment flow deposits.

A7:THE CHAMPLAIN THRUST SYSTEM IN THE WHITEHALL-SHOREHAM AREA -- INFLUENCE OF PRE- AND POST-THRUST NORMAL FAULTS ON THE PRESENT THRUST GEOMETRY AND LITHOFACIES DISTRIBUTION.
Nicholas W. Hayman
Department of Earth and Space Sciences
Box 351310
University of Washington
Seattle,WA 98195
nickh@u.washington.edu
W.S.F. Kidd
Department of Earth and Atmospheric Sciences
University at Albany
Albany, NY, 12222
Introduction
The Taconic foreland is an exemplar of collisional tectonics and the Champlain Valley of West-Central Vermont is a pivotal region for mapping and reconstructing this system (figure 1). This is because between the latitudes of Middlebury, VT and Whitehall, NY, the shelf sequence gives way to the fore-deep and rise sequences along the strike of the eastward dipping range (Stanley, 1987; Kidd et al., 1995) (figure 2). Although many significant contacts between the sequences have been recognized as faults, the accrued transport on these faults, and the relationship between them in time and space, is ambiguous. This is partly due to the difficulty in determining the map trace of significant faults across regions of poor outcrop and sharp contrasts in stratigraphic units and sedimentary facies.
This field trip visits outcrops along the Champlain thrust system (CTS). including the Mettawee (River) fault, a late normal fault that significantly truncates thrusts of the CTS. These outcrops are from the areas discussed in Hayman and Kidd (2002), a synthesis of our work in the region. Our mapping focused on the CTS, which was responsible for large (>80 km) transport of the shelf section (Rowley, 1982). A cornerstone of our interpretation is that many abrupt changes in map units along-strike, and problems in restoration of cross-sections across the map area, are the consequence of reactivation of preexisting faults. The early (pre-thrust) faulting occurred in response to extension along the synconvergent flexural forebulge. Other significant lithic changes are the result of late (post-thrust) normal faulting associated with one of the several phases of extension that affected the region. The thrust system projects to depth and thus is dynamical1y related to the overlying Taconic thrusts, and collectively forms a decollement beneath the Green Mountain crystalline core of the Taconic orogen (Rowley, 1982).

A7:THE CHAMPLAIN THRUST SYSTEM IN THE WHITEHALL-SHOREHAM AREA -- INFLUENCE OF PRE- AND POST-THRUST NORMAL FAULTS ON THE PRESENT THRUST GEOMETRY AND LITHOFACIES DISTRIBUTION.7

STOP 1. Orwell/Main Champlain Thrust at Shoreham
ASK PERMISSION AT THE FARMHOUSE to visit this outcrop. Unplanned encounters with an unfriendly bull may result if you ignore this instruction.
From Slateledge Farm house, walk east along the road passing the track to the barn nearest the house, and take the next track which heads toward the low ridge extending north from the second barn. About 100 meters from the road, small exposures of dark grey shales/slates occur at the side of the track. After passing a fence line (and maybe crossing an electric fence across the track) small exposures of steeply east-dipping calcareous dark shale, and more extensive loose material occur on the western slope of the ridge [43°: 53.740'N, 73° 19.723'W). Based on lithic correlation, these are mid-Ordovician in age, termed Stony Point Formation in this region. They represent the transition to deep-water deposition from the previous shallow passive margin carbonate sedimentation as the Laurentian margin subsided rapidly on entering the Taconic subduction system. On top of the ridge here are outcrops of pale tan-weathering, grey quartzite, gently east-dipping at this position, becoming flat and then moderately west-dipping in a small ramp anticline on going north along the ridge. This quartzite is identified as Potsdam Fm., of mid-Cambrian age. The westernmost major thrust of the Champlain Thrust System at this latitude is located between the quartzite and the dark shale. This fault to the north of the Shoreham area is termed the Champlain Thrust; in the area of Shoreham, where several major splays occur, this one has been called the Orwell Thrust, although we have placed the name "Main Champlain Thrust" on it. The fault surface itself is not exposed here, although the position can be constrained to an interval less than 10 meters wide.
Looking east from the ridge towards Shoreham village, no outcrops occur in the lowlands, but extrapolation from outcrops along strike imply that this belt is underlain by Cambrian-early Ordovician Beekmantown carbonates, largely dolostones, which are stratigraphically above the Potsdam. In Shoreham Village, and farther east, imbricate slices of the Shoreham and Pinnacle Thrusts occur, which are examined in subsequent stops. To the west, the Champlain lowlands, under the blanket of Quaternary lake clays, contain a few outcrops of mid-Ordovician shaly strata which are the Taconic trench/foredeep sediments. Walk back to the road, and slates, siltstones and fine-grained greywacke arenites of these mid-Ordovician flysch sediments can be seen in outcrops by the road opposite the farmhouse, and in the small quarry adjacent [43° 53.639'N, 73° 19.832'W]. A well-developed moderately east-dipping slaty cleavage cuts moderately NW-dipping bedding, and excellent slate pencils formed by this oblique intersection weather out. The bedding strikes obliquely towards the N-S trace of the Main Champlain Thrust. We infer from the strong cleavage and down-plunge lineation that this is probably an imbricate thrust slice attached to and considerably transported by the Main Champlain Thrust; more autochthonous flysch west of the Champlain Thrust in this region tends not to have prominent cleavage.

STOP 2 - Potsdam/Ticonderoga quartzites/arenites at base of Shoreham Thrust Duplex
This stop has DANGEROUS TRAFFIC CONDITIONS (high speeds and limited sight distance) - please park as far off the paved road as possible, be extremely careful crossing the road. and do not stand on or step into the pavement when contemplating the outcrop.
From the parking place specified, cross the road and walk north to the roadcut [43° 55.241'N, 73° 18.743'W]. Pale tan-weathering, grey, well-indurated quartzites and local dolomitic arenites dip gently east. These can be identified as either the upper part of the Potsdam and/or the arenite-rich part of the overlying unit, the Ticonderoga Fm. Small-scale cross-bedding occurs locally. Mapping shows that they form an imbricate slice on the Shoreham Thrust. In this area near Shoreham. the Shoreham Thrust Slice is a complex imbricated stack of slices, forming the Shoreham Duplex.

STOP 3 - Middlebury Limestone in the Shoreham Duplex
ASK PERMISSION to park and view this outcrop AT THE GARAGE, OR AT THE HOUSE ON THE CORNER OF THE 74/22A INTERSECTION.
From the east side of the parking area, walk to the north-east corner, then go east to the nearest outcrop. This is massive limestone, showing a fairly well-developed steep east-dipping cleavage in places. About 50 meters east of the first outcrop, another ridge of limestone [43° 54. 144'N. 73° 18.377W] contains locally prominent large coiled gastropod fossils, probably Maclurites, which is common in autochthonous sections of Chazy Group mid-Ordovician limestones. In these transported rocks this limestone is usually termed Middlebury Limestone. These outcrops map as imbricate slices within the Shoreham Duplex. Recent expansion of the parking area has unfortunately resulted in the best outcrop being buried by bulldozed trees, soil and trash. If time is short, or permission is refused. a smaller roadcut outcrop of this limestone can be viewed just east of the Mobil Station parking lot on Route 74 east, but the gastropods are not clearly seen there.

STOP 4A. Potsdam quartzites of the Pinnacle Thrust slice adjacent to the Lemon Fair Fault
From the parking place, cross the road, and view tan-weathering, grey indurated quartzites, locally dolomitic, of the Potsdam Fm. East of the intersection [43° 52.386'N, 73° 16.321'W] they dip steeply southwest; in the outcrop west of the intersection, they dip gently southwest. The monoclinal fold outlined by this dip change, and the oblique strike, we interpret as due to proximity to the Lemon Fair Fault, and the lateral ramp it defines in the thrust geometry. These quartzites are regionally part of the Pinnacle Thrust Slice; this thrust is not exposed here, but its position can be constrained by mapping to the south. The Lemon Fair Fault runs here along the river; looking across from the parking place, limestones on the other side of the fault can be seen at the base of the old bridge abutment.

STOP 4B - Carbonates adjacent to the Lemon Fair Fault, and just beneath the Pinnacle Thrust
Roadcut outcrop [43° 52.398'N, 73° 16.280'W] of limestones, dolomitic limestones, and minor dolostones which are adjacent to the Lemon Fair Fault. The generally rather fractured appearance, and the common presence of calcite veins, are indicative of this proximity. The dominant fractures and veins strike NW, dipping steeply north to vertical, subparallel with the Lemon Fair Fault. One vein shows prominent steeply pitching slickensides. These rocks are part of a slice of early-mid Ordovician limestone-dominated carbonates localized immediately under the Pinnacle Thrust in the area north of the Lemon Fair Fault, but not present to the south. This and other substantial changes in structural geometry that occur across the Lemon Fair Fault show that it was used as a major lateral ramp structure by the thrusts of the Champlain System. Because facies and stratigraphic thickness variations only occur in mid-Ordovician units across this structure, we infer that it was generated as a mid-Ordovician cross-strike "flexural" normal fault, and then converted to use as a lateral ramp in the Champlain Thrust System.

STOP 5 - Champlain/Orwell Thrust at Stevens Orchard
Walk up the road from the parking place by the house and barn/garage. On the east side, an outcrop of calcareous dark shale/slate, with a local phacoidally cleaved fault zone, is mid-Ordovician Stony Point Formation. Farther up on the same side just above a driveway there is a smaller outcrop and loose material of the same unit. About 20 feet above this, seen in the roadcut and the ridge going south from it, are dolostones of the Beekmantown Group forming the base of the Main Champlain (or Orwell) Thrust here. A little farther up on the north side of the road [43° 47.077'N, 73° 20.586'W], similar dolostones, showing not very prominent medium to thick-bedding and homogeneous brown-weathering, are exposed. This outcrop also shows the characteristic "fretted" weathering pattern produced by differential removal of abundant narrow calcite veins that are ubiquitously and abundantly developed in these transported dolostones, but are not usually seen in autochthonous dolostone sections in areas to the south. Dolostones like these form all of the body of both the Main Champlain (Orwell) and Shoreham Thrust Slices at this latitude; the Potsdam quartzites are absent because the thrusts have ramped up laterally to within the late Cambrian-early Ordovician carbonate section. Rocks like these are conventionally termed "Providence Island Formation", but we have little confidence that this is stratigraphically meaningful in the sense of strict correlation to the type section on Providence Island. The map pattern of the Main Champlain Thrust in this local area shows that it is very gently east-dipping, or almost flat.

STOP 6 - Shoreham Thrust near Orwell
On the corner with the roadcut outcrop [43° 47.439'N. 73° 19.236'W], east of the parking spot specified, deformed calcareous shales ("Stony Point Fm") make up most of the exposure, although at the western end a contact can be observed with fractured tan-weathering Beekmantown (late Cambrian/early Ordovician) dolostone above the shales. This dolostone forms a low ridge in the woods that parallels the road westwards to beyond the Bascom Road junction and, back along the road in this direction, smaller outcrops of both calcareous and non-calcareous deformed shales can be found locally, with dolostones near or in contact above the shales. The thrust fault defined by this contact is the Shoreham Thrust; the local map pattern demonstrates that this fault also is near flat-lying here. As for the Main Champlain Thrust, the absence here of Potsdam Fm is due to the thrust climbing section southwards, in part across the lateral ramp structure of the Lemon Fair Fault. A much better, even excellent outcrop of this fault exists 20 meters west into the woods near the end of Wilcox Road, about 1.0 mile SW of this outcrop, but the current landowner is most unwelcoming and is quick to display a gun.

STOP 7. Orwell/Main Champlain thrust sheet at Benson Bay/Blue Ledge
STOP 7. Orwell/Main Champlain thrust sheet at Benson Bay/Blue Ledge ASK PERMISSION AT THE HOUSE.
The main part of the outcrop visible from the parking area is now only accessible through the pig mud wallow, unfortunately. We recommend viewing it from the road, from where the nearly flat-lying well-defined planar bedding in the dolostones is clearly seen. To see rocks closer up, walk east along the road, to a large oak tree on the corner [43° 45.889'N, 73° 20.399'W], and then go into the bushes a short distance where the same well-bedded sugary-textured dolostones continue to outcrop. Here the existence of common dark chert as nodules and patches in the dolostones can be seen, both in the outcrop and in the talus fragments. This unit is clearly identifiable in the Beekmantown stratigraphy of Fisher (1985) as the Whitehall Formation. Nearby, on White Ledge, the cliff visible to the east of the road, more Whitehall Formation is exposed; above it one can find in the woods good outcrop that shows overlying Beekmantown units matching those of Fisher (1985), including the unmistakeable Winchell Creek cross-bedded arenites. Structurally, we are back down in the Main Champlain Thrust slice (the fault is well~ constrained by outcrop on Blue Ledge a few hundred meters to the northwest of this outcrop), yet the whole section carried by the Thrust here is utterly different from the homogeneous, poorly-bedded, veined dolostones at and just south of Stop 5, no more than 1-2 km. away. We identify this change to occur across a NW-trending valley just to the northeast of this stop location, and interpret this to be a lateral ramp in the thrust system inherited from a flexural Ordovician normal fault. The change in the shelf facies that occurs in the "Main Champlain" Thrust sheet here is by far the most pronounced anywhere, and it must represent a very substantial relative distance (we think many tens of kilometers) between the sites of detachment of the stratigraphic sequences north and south of this change. Because the sequence in the MCT slice from here south so closely matches the autochthonous stratigraphic sequence defined by Fisher (1985), we think it is unlikely to have been displaced very far <10 km?). Because it is likely the MCT to the north has significantly larger displacement (at least several tens, and perhaps 80 km - Rowley, 1982) we think there must be a large displacement transfer here from the MCT north of this site, along the lateral ramp fault up to the next thrust, the Shoreham Thrust. Not only is there a substantial change in the MCT hanging wall here, but the footwall also contains a similar (and we think closely related) feature about 1 km SW of this stop. There is an abrupt change, from mid-Ordovician shales and calcareous shales of the Stony Point Fm as the footwall strata, to an autochthonous shelf carbonate/clastic section of the "Whitehall" sequence. This change takes place across a mapped cross-fault which affects the MCT as a small-offset lateral ramp/tear fault, but which has a much larger stratigraphically-defined offset in the footwall. We think this is also a mid-Ordovician normal fault used soon after to localise a ramp in the lowest part of the thrust system.

STOP 8 - Shoreham Thrust and end of Orwell/Main Champlain Thrust near Benson Landing
GET PERMISSION AT FARM ON N. LAKE ROAD 0.4 MILES WEST OF FRAZIER HILL ROAD JUNCTION
This stop may be omitted if time becomes short.
Enter field on south side of road [43° 43.781'N, 73° 20.562'W] through gate [OPEN AND CLOSE IT BEHIND YOU - PLEASE DON'T CLIMB OVER IT]
About 50 meters from gate [43° 43.751 'N, 73° 20.599'W] dark calcareous shales of Stony Point Fm. Follow the track on the east side of the valley for another 50 meters or so; small outcrops of calcareous shales occur at the foot of the ridge to the east, while massive fractured dolostones occur up the slope and at the crest of the ridge. This defines the Shoreham Thrust here (to the south this has been called the Shaw Mountain Thrust). On the hillsides and tableland to the west of this valley, on a longer walk than we have time for, flat-lying shelf strata of "Whitehall facies" that map in the MCT slice are sparingly exposed; there has to be a normal fault bounding these against the Stony Point Fm. shales which continue down this valley in and near the stream. The Main Champlain Thrust terminates against this normal fault, and is not present farther south. We think that this normal fault originated as an Ordovician "flexural" normal fault, and was subsequently reactivated with reverse displacement sense and used by the Champlain Thrust System as a "terminal ramp" including the transfer of the residual displacement on this southernmost part of the MCT up to the Shoreham Thrust.
If desired by those on independent schedules, continue walking down the valley until the stream enters a continuous section exposing dark calcareous shales [43° 43.575'N. 73° 20.934'W]. Outcrops of near flat-lying dolostones (Fort Edward) or cross-bedded arenites (Winchell Creek) are found on the northwestern slope of the valley. e.g. dolostones at [43° 43.636'N. 73° 20.895'W], at an elevation significantly above the shales in the stream bed.

STOP 9A. Shoreham Thrust at West Haven
From the parking place [430 38.727'N, 73° 20.888'W] walk about 120 meters south along Book Road [43° 38.663'N. 73° 20.929'W] to roadcut on east side. In the center portion of this exposure. black non-calcareous shale (mid-Ordovician) up to 1 meter thick, with a pronounced phacoidal cleavage indicative of large shear strain. occurs below a sharp contact with highly fractured massive dolostones ("Providence Island" lithology). This is the Shoreham Thrust, and mapping of the surrounding area shows unequivocally that it is here the westernmost thrust in the Champlain System - in other words that the "Main Champlain Thrust" farther north has disappeared. To the south across the Vermont-New York border at the Poultney River, the Shoreham Thrust too disappears as a discrete mappable structure, although minor folding and incipient ramps can be detected in shelf rocks at its expected position as far south as the latitude of Whitehall. A major WSW-ENE cross-fault passing north of this outcrop cuts and displaces the autochthonous shelf section and the Champlain Thrust System stack, crossing Book Road at the parking spot. The Shoreham Thrust is offset by this structure with an apparent left lateral displacement of about 2 kilometers. Two other faults like this one pass through the next valley north near West Haven hamlet. These cross faults are discussed in the description of Stop 9B below.

STOP 9B.Mettawee River Fault at WestHaven
ASK PERMISSION AT BOOK FARM BEFORE ENYERING - FIND BOOK FARM ALONG BOOK ROAD 1.1 MILES SOUTH OF STOP
From parking place [43° 38.727'N. 73° 20.888'W]. enter field on east by walking between the two fences, opening the gate if necessary. Follow the grassed-over track into the edge of the woods (about 100 meters), then along the edge of the woods (about another 200 meters), then angle down the crest of a low ridge in the meadow (about another 100 meters) to the edge of the woods crossing the low ridge [43° 38.850'N. 73° 20.513'W]. Outcrop of fractured tan-weathering dolostone ("Providence Island" lithology) is found just inside the woods at this point. Use the sketch map below to navigate and understand the geology from. this point on. About 200 feet east of the first dolostone outcrop. in the bed of the stream near the south end of the small ravine, dark slates ("Hortonville"), of presumed mid-Ordovician age, are exposed in contact with the dolostones [43° 38.853'N, 73° 20.473'W]. This contact is clearly a steep (60°) east-dipping fault, truncating the moderately east-dipping thick bedding in the dolostones, and the cleavage in the slates. Given the regional context, with no thick dark slates occurring west of this point, but rather the carbonates of the Shoreham Thrust slice, then the underlying autochthonous Cambro-Ordovician "Whitehall facies" sequence, with Potsdam sandstone/quartzite at its base lying unconformably on Grenville Basement, this fault must be a normal fault, and is in fact the Mettawee River Fault first proposed by Fisher (1985) in New York. It is the only complete exposure of this fault, although our mapping demonstrates that it runs continuously from the Mettawee River (to be seen at stop 12) north through this locality at least as far north as the latitude of Shoreham, and all along this length it regionally truncates the thrust stack of the Champlain Thrust System. We suspect that it extends significantly farther north and south than this, but this falls outside our mapping area, and there is a lack of clear truncation of components of the Champlain Thrust System beyond these extents. One reason for suspecting a larger along-strike extent is the minimum displacement, which can be constrained from accurately drawn cross-sections, and which demand at least 1000 meters of throw, and perhaps more than 1500 meters. This fault cuts out very significant parts of the Champlain Thrust stack, and is one of the main reasons why previous attempts to trace the major thrust faults of the System failed. What age is this normal fault? We have no direct younger age limit in the mapped area, other than the observation that it is cut and displaced by the WSW-ENE faults which also cut the Shoreham Thrust. In fact this set of local outcrops (see sketch map) shows that relationship directly, with a smaller splay of the larger WSW -ENE fault displacing in an apparent left-lateral sense the dolostone-slate contact (the Mettawee River Fault). Despite the prominence of the WSW -ENE faults, we have seen no direct evidence in outcrops near the fault valley indicating whether they are really left-lateral strike slip faults or whether the displacements are created by dip slip movement (in this case, it would require north side down) working on consistently east-dipping strata and faults. Thus these ENE-WSW faults, and the Mettawee River Fault, could be Taconic in age, although they might be younger, e.g. Acadian, or even Mesozoic, and they do not necessarily have to belong to the same orogenic episode.

STOP 10. Pinnacle/Comstock Thrust, and Mettawee River Fault, at Fish Hill
OBTAIN PERMISSION FROM HARMONY HILL FARM [0.5 MILE DOWN BUCKLEY ROAD FROM JUNCTION]
Enter field to north of parking point through gateway/over electric fence. Go round pond on western side, loop round north end, crossing dry bouldery bed of stream, and head NE up slope toward outcrops just into the woods. [Cross electric fence at edge of woods]. Find contact between grey indurated quartzites above and tan-weathering crudely bedded dolostones below [43° 33.942'N, 73° 21.398'W]. all dipping gently east. About 1 meter below the base of the quartzites, there is a projecting ledge exposing a bedding-parallel surface of fractured and veined dolostones, smears of limestone, and loose fragments of dark shale coming from the 10-30 cm unexposed interval between the dolostone above and the ledge with limestone smears below. This is the Comstock Thrust, which we maintain corresponds with the Pinnacle Thrust to the north of Orwell (or more precisely here it is the lowest fault surface in a zone which places imbricated Potsdam and Ticonderoga/Whitehall quartzites and dolostones over Fort Ann and Providence Island dolostones). Walk east along the base of the hillside outcrop; an ∼10 meter section mostly of quartzites is conformably overlain by dolostones which extend to near the west end of the barn, where there are some dark cherty patches in the dolostones. Climb up on the outcrop to avoid the manure piles and to find another imbricate fault contact placing the quartzites back on top of the dolostones, with truncation of the dolostone bedding in the footwall as a thrust ramp. Near the east end of the barn [43° 33.920'N, 73° 21.305'W], the quartzite outcrop ends, with locally steep and folded bedding probably reflecting accommodation to the ramp/imbricate geometry underneath. Walk north up into the field over quartzite outcrop, then curve to the east across exposure gap of about 40 meters to outcrop of dark and strongly deformed, melangy mid-Ordovician shales [43° 33.957'N. 73° 21.281'W]. You have just crossed the Mettawee River Fault. which again must have substantial normal, down to east displacement. You may find the evidence here for the existence of the Comstock/Pinnacle Thrust not wholly compelling, but raise your eyes to the view towards the southwest, where, in the hills south of Whitehall, the map of Fisher (1985) unquestionably demonstrates a) the existence of the Pinnacle/Comstock Thrust duplicating most of the Cambrian-Ordovician shelf stratigraphy from the Potsdam upward, and b) the unquestionable fact that this thrust is truncated on a large scale by the Mettawee River Fault.
To the east of the shale outcrop, up the slope of Fish Hill, limestones and dolostones of early-mid Ordovician age form another of several thrust slices involving only the upper part of the outer Laurentian shelf and the overlying mid-Ordovician dark shales. These slices rest below the westernmost fault bounding the Taconic Allochthon, which outcrops only about a kilometer to the east.

STOP 11. Mettawee River Fault at the type locality
NO HAMMERS OR SAMPUNG PLEASE - NY DEC REGULATIONS
From parking area [43° 28.259'N. 73° 22.153'W], walk from northeast corner along path which quickly descends to a bouldery small stream bed. If wet, the path down can be slippery since it passes through well-rounded fluvial pebble gravel and underlying lake clays; the sloping bedding surfaces of the carbonate outcrops can also be treacherous in places. Cross the stream bed near/at the Mettawee River bank and get onto the outcrop along the west· bank of the River [43° 28.283'N. 73° 22.103'W]. These carbonates are mapped by Fisher as Providence Island Formation, and dip about 10-20 degrees east. The lowest part of the section here consists of dolostones; the upper 4 meters contains substantial limestones; we think it is possible that the part of the section containing limestones is either basal Chazy or Black River. This shows some bedding surfaces with well-developed ripples, and small mudcracks, and selectively dolomitized burrows. Local meter-scale folding of periclinal type affects these beds near this end of the section. Walk north along the west bank on the outcrop; mostly dolostones are exposed farther north. About 70 meters north, dark highly deformed mid-Ordovician shales extend from the east bank across the river to an almost complete section at the small rapids. Depending on the streamflow, it may or may not not be possible to examine the shale closely. The contact of the deformed shale with the dolostone beds at the west side of the stream is not quite fully exposed (even at the lowest water we have seen, there is a gap of 10 cm or so), but is clearly sharp, and parallel with bedding in the dolostones. This is somewhat surprising, for a contact which maps out a short distance to the north as a sharp fault that unquestionably truncates strata of the carbonate shelf stratigraphy, as well as the major thrust fault which duplicates this sequence (see the map of Fisher, 1985). Also surprising, for a fault that must have substantial normal sense displacement, is the fact that most or all quartz fiber slickensides in the shale here give thrust sense of displacement!
We infer either that the normal fault mapped to the north passes east of this outcrop within the shale belt, or that the motion is confined to a surprisingly narrow zone along the contact; the full exposure at Stop 9B suggests the latter possibility is more likely than one might be inclined to think based on this outcrop at Stop 11 alone.

STOP 12 - Pre-thrust normal fault in Comstock/Pinnacle Thrust slice at Tyler Farm
ASK PERMISSION at Tyler Farm house [43° 26.849'N, 73° 24.432'W] before going into fields to north of road. If visiting the western part of the ridge, also ask at the old stone house about 0.2 mi west of Tyler farmhouse.
This stop may be skipped if time runs short.
Refer to the outcrop map to navigate. Outcrops on the slope above the Tyler farmhouse, and above the old stone house 0.2 miles west of this, begin with mid-Ordovician limestones of Black River and/or Trenton age; the thin-bedded shaly ones are termed Glens Falls Fm, the more massive ones stratigraphically below are Orwell Fm. They strike about E-W and dip steeply south. Units of the upper Beekmantown (Providence Island, Sciota Limestone, and Ward Siltstone) are found on and near the crest of the first ridge, with decreasing dips towards the valley to the north in which a fault is located. The zone of steep dips is interpreted as a bend fold ["drag fold"] against the hanging wall of (inferred from its straight trace, and the limestone maximum dips) a steeply south-dipping normal fault, and perhaps also is partly an accommodation to the lateral ramp this fault has provided to the Comstock Thrust. North of this ∼E-W striking fault, various units of the older parts of the Beekmantown are found forming mostly a moderately east-dipping panel truncated at the normal fault. These strata are in the hanging wall of the Comstock Thrust (=Pinnacle Thrust), which here is the last and only thrust fault in the exposed part of the Laurentian shelf sequence between the autochthonous Grenville basement and mid-Ordovician shales of the Taconic foredeep. The Comstock Thrust is included on the map of Fisher (1985), although he shows its continuation by joining it to the E-W fault referred to above, which we think is geometrically and kinematically unlikely. Our map (based on mapping by Y. Pan and W. Kidd), instead interprets the E-W norma1 fault to intersect the Comstock thrust and to be truncated by it. In addition, we interpret the Comstock Thrust to continue south across the point of this truncation, and the map pattern requires that it abruptly climb stratigraphic level from the Ticonderoga or Potsdam sandstone to near the base of the mid-Ordovician limestones despite no change in its structural level. This requires the thrust to have truncated an existing normal fault that had a throw of several hundred meters, nearly the full thickness of the shelf sequence. Since there is no indication approaching the E-W fault of any thickness or facies change in the pre-Black River section, we infer that this normal fault too is a mid-Ordovician age structure, generated during flexure of the Laurentian continental margin as it approached the Taconic trench and subduction system. (As a additional note on the structure of the loca1 faults, we find no evidence of the existence of the rootless klippen of Potsdam-Whitehall Fm strata shown not far southwest of this locality on Fisher's 1985 map. We think he confused gently dipping autochthonous Winchell Creek and stratigraphically conformable dolostones with the lower units here; there is in our view no stratigraphic or structural evidence requiring this peculiar object to exist).

B1: GEOLOGY AND GEOCHRONOLOGY OF THE EASTERN ADIRONDACKS
[
James McLelland
Dept. Geology, Colgate Univ.
and
Skidmore College, Saratoga Springs, New York, 12866
M. E. Bickford
Department Earth Sciences, Syracuse University
Syracuse, New York, 13244-1070
Frank Spear and Lara Storm
Department Earth and Environmental Sciences
RPI, Troy, New York, 12180
Geologic background and detail for the Adirondack region is summarized in the introduction section of trip A-I, "Geology and geochronology of the southern Adirondacks" by Mclelland, Storm, and Spear. The reader is referred to that section for relevant background information. Here we summarize broad considerations pertaining to the eastern Adirondacks.
The eastern Adirondacks contain a wide variety of rock types, and this trip aims to visit representatives of the major lithologies. Stops include units recently dated by U-Pb geochronology including multi- and single-grain TIMS and SHRIMP II analyses. Results of these investigations are presented in the text of trip A-I, to which the reader is referred. Broadly, the various lithologies fall into the following groups: 1300-1350 Ma tonalites and granodiorites; 1150-1160 Ma anorthosites, mangerites, charnockites, and granites (AMCG suite); 1090-1100 Hawkeye granite; and 1035-1060 Ma Lyon Mt Granite (LMG). The two major orogenic events associated with these are the Elzevirian and Ottawan Orogenies (Moore and Thompson, 1980). The former falls into the interval ca 1350-1170 Ma and involves protracted arc magmatism and accretion capped by a culminating collisional orogeny from ca 1200-1170 Ma (Wasteneys et al., 1999). The latter refers to the Himalayan-type collision of Laurentia with Amazonia during the interval ca 1090-1000 Ma. Most of the metamorphic and structural effects present in the Adirondacks are the result of the Ottawan Orogeny, but Elzevirian features can be recognized locally. Both the AMCG suite and the LMG are thought to be late- to post-tectonic manifestations of delamination of over thickened orogens.
Structurally, the eastern Adirondacks are dominated by the same large, recumbent fold-nappe structures (F2) as found in the southern Adirondacks, and with fold axes oriented dominantly ∼E-W parallel to lineation. As in the southern Adirondacks, the fold-nappes are thought to have sheared-out lower limbs, but this has yet to be demonstrated on a map scale. At least two distinguishable upright fold events are superimposed on the nappes: F3 with shallow plunging ∼E-W axes and F4 with shallow-plunging NNE axes. All of three fold sets affect Hawkeye and older units, and thus must be of Ottawan age. This is also the case with the strongly penetrative rock fabric, including strong ribbon lineations, that are present in these rocks and are associated with the large fold-nappes. Intense fabric of this sort is largely absent from the Lyon Mt Granite and this is interpreted to reflect its intrusion in late, post-nappe stages of the Ottawan. In the northern portion of the eastern Adirondacks the NNE, F4, folds become quite tight and have a strong lineation associated with them. This may be the result of rock sequences being squeezed between 1arge, domical prongs of anorthosite. Finally, we note that there exists abundant local evidence of small isoclinal F1 folds that pre-date F2. These, and their associated fabrics, are thought to be Elzevirian in origin, and in a few cases this can be shown to be the case.
A dominant feature of the eastern Adirondacks is the great Marcy anorthosite massif that underlies almost all the High Peaks. Over two-dozen zircon age determinations demonstrate that the anorthosite was emplaced at 1150 ± 10 Ma (Table 2, Trip A-I) and that associated granitoids and coronitic metagabbros are coeval with it. These relationships make it clear that thermal energy from the anorthosite had nothing, whatever, to do with the granulite facies that characterize the Adirondack Highlands and post-date the AMCG suite by 60-100 million years. The specifics of Adirondack P,T relationships and uplift history are discussed trip A-I and its appendix AA-l. It is likely that most of the region experienced peak temperatures on the order of ∼800° C and pressures of ∼8 Kbar. Based upon extensive isotope work by John Valley and his students, it appears most likely that this metamorphism proceeded under fluid-absent conditions (Valley et al, 1983). Note, however, that this does not exclude the presence of late, post peak, fluids associated with the emplacement of Lyon Mt Granite.

B1: GEOLOGY AND GEOCHRONOLOGY OF THE EASTERN ADIRONDACKS

Sulfidic staining on NNE fault in granite.

Exposure of hornblende granite on left.

Pegmatite on right.

Pass metagabbro

Hornblende granite on both sides.

Hornblende granite

Hornblende granite with metagabbro

Metagabbro, pegmatite, and metasediments on right.

STOP 1. MEGA-GARNET AMPHIBOLITE. (20 MINUTES).
This small roadcut at the intersection of Schroon River Road and Wall Street is an example of what is referred to as a "Poor Man's Gore Mt.". The roadcut contains large (up to 8" across) almandine-rich garnets rimmed by black hornblende and set into an amphibolite matrix that can be shown to have formed from metagabbro. The assemblage is reminiscent of the famous deposits at the Barton Garnet Mine on Gore Mt where garnet growth has been dated at ca 1050 Ma (Mezger et al., 1990) using Sm-Nd techniques. The northern contact of the present exposure is characterized by a deformed hornblende-biotite-quartz-oligoclase pegmatite that also crosscuts the interior of the garnetiferous amphibolite. Several other examples of "Poor Man's Gore Mt" occurrences are present in the Adirondacks, and all - including Gore Mt - share in common proximity to high angle faults that host pegmatites, quartz-veins, etc. We propose that the faults served as plumbing systems for hydrous magmas that provided water to the amphibolites at temperatures of ∼700°-800° C. This anomalous circumstance resulted in elevated diffusion rates that enhanced garnet growth at a small number of nucleation sites. The hornblende rims represent the excess of hornblende constituents in the area of local garnet growth. Post peak decline in the activity of H2O, or continued increase of T, caused reactions between garnet and hornblende to produce calcic plagioclase (AN80) and orthopyroxene.

STOP 2. ca 1300 Ma TONALITE. (20 MINUTES).
These outcrops consist of tonalitic rocks dated elsewhere as ca 1300-1350 Ma by U/Pb zircon. Their chemistry is calcalkaline, and, together with associated granodiorites, they are interpreted as having formed in one, or more, magmatic arcs during the Elzevirian events of ca 1400-1220 Ma. Locally these units contain xenoliths of, or crosscut, metasedimentary rocks, thus establishing a minimum age for most, if not all of these units in the Adirondacks. (See Stop 4, Trip A 1-1)

STOP 3. ca 1050 Ma LYON MT LEUCOGRANITE. (30 MINUTES).
The medium grained, pink leucogranite is characteristic of the main quartz-microperthite unit of Lyon Mt Granite that occurs in broad tracts across the Adirondack Highlands. The classical low-Ti Kiruna-type magnetite deposits of the Adirondacks are closely associated with the granite, either occurring within it or in rocks (usually metacarbonates) immediately adjacent to it. Magnetite mineralization is commonly accompanied by sodic alteration manifested by checkerboard albite and by quartz-albite (Ab98) replacements of the original quartz-microperthite rocks (McLelland et al., 2002). A total of six occurrences of Lyon Mt Granite from across the Adirondacks have been dated by U/Pb techniques including singIe and multigrain TIMS and SHRIMP IT (Fig. 2). All of these fall into the interval 1050 ± 10 Ma (McLelland et al., 2001). The present roadcut was dated by single grain TIMS methods and yields an age of 1049 ±2 Ma (Fig. 3). Staining of the roadcut reveals that it contains several sheets and layers of quartz-albite one of which is associated with a irregular veinlet of magnetite. Locally, quartz-sillimanite nodules are present in these rocks and are attributed to leaching by high temperature hydrothermal fluids (McLelland et al., 2002). Pegmatites and quartz veins are common.

Quartz-sillimanite nodules in Lyon Mt Granite included by Walton and deWaard in their "Basement Complex".

Metapelites assigned by Walton and deWaard to their "Older Paragneiss".

Marble, calcsilicate of Walton's "Paradox Lake Formation in near continuous outcrop for 2.6 miles.

Long roadcut of migmatitic metapelite
Long roadcut of migmatitic metapelite assigned by Walton to the "Treadway Mt Formation".

STOP 4. HIGHLY DEFORMED MIGMATITIC METAPELlTES (> 1300 Ma). (30 minutes),
The southern and eastern Adirondack Highlands, as well as the Northwest Lowlands, contain a significant thickness of dark garnet-biotite-quartz-oligoclase + sillimanite metapelite accompanied by a garnet-bearing white quartz-feldspar leucosome. McLelland and Husain (1986) interpret the leucosomes as in situ anatectic products of the Original rock and the dark fraction as restite. Greywacke is a likely precursor. The presence of armored spinel, and rarely corundum, are consistent with this interpretation. In regions of low strain the leucosomes are present as crosscutting, anastomosing, irregular sheets, dikes, and veins of clearly coarse-grained pegmatitic material. As strain increases, the pegmatitic material gets pulled into pseudo-conformity, disrupted, and grain size reduced so as to yield porcellanous layers that are commonly parallel but locally retain their original crosscutting configurations. In short, the apparent layering in these rocks is almost wholly tectonic and has nothing to do with original stratigraphic superposition. They are, in fact, mylonitized migmatites, of the variety referred to as "straight gneiss". As indicated in stop 2, these metapelites are crosscut by the ca 1300 Ma tonalitic suite, and therefore post-date these rocks. Our interpretation is that their current reintegrated compositions are best accounted for by a greywacke-shale precursor, and that these were arc-related, flysch-type sediments approximately coeval with the Elzevirian tonalites. This would also help to explain the relative proximity between these lithologies. The age of anatexis is currently being investigated, but a pilot SHRIMP II study on a leucosome from near Speculator reveals a complicated pattern of zoning and inheritance with cores of age ca 1240 Ma and metamorphic overgrowths of 1010-1040 Ma. Sandwiched between these are mantles with oscillatory zoning and ages of 1170-1180 Ma. We interpret these mantles as zircon grown during anatexis and dating this event as Elzevirian. The absence of further significant anatexis in the granulite facies Ottawan Orogeny (1090-1030 Ma) is thought to be the result of earlier dehydration of the migmatites during the Elzevirian anatexis. Both mafic and granitic sheets can be traced across the roadcut face and reveal the presence of fault offsets. Note that the name "Treadway Mt Formation" assigned to this unit by Walton suggests that it has stratigraphic characteristics and continuity. This assumption is no longer considered to be correct and is, at best, a lithotectonic assignation.

Granite and gabbro.

Large isoclinal fold in quartzites and metapelitic rocks.

Undeformed gabbro.

Marble

Elephant rock.
Consists of leucocratic sillimanite-garnet-quartz-feldspar gneiss (khondalite).

STOP 5. SWEDE MT SEQUENCE. (60 MINUTES INCLUDING LUNCH).
The metasediments exposed along the Rt 8 at Swede Mt provide an exceptional opportunity to closely examine both the rocks and structure. At the southwestern end of the sequence (mile 48.9) a well exposed isoclinal fold can be seen in the large road-cut on the south side of Rt 8. Folded units include marble, quartzite, sillimanite-garnet-quartz-feldspar (khondalite) gneiss, and rusty sulfidic Dixon Schist. The latter is thought to represent a sheared, altered, and graphitic variety of the khondalite. The fold axis trends ∼E-W at a low angle of plunge about the horizontal, and its axial plane, which dips from ∼50 degrees SW to horizontal, has been folded about upright E-W axial planes. The fold is typical of refolded isoclines in the Adirondacks. Similar rocks at Dresden Station, on the east side of Lake George, are intruded by a metagabbro dated at 1144 ± 7 Ma. At this locality the metagabbro truncates foliation and even individual garnet grains in the khondalite. These relations indicate that the khondalite was deposited and first metamorphosed in Elzevirian times. Early workers considered the khondalite to be a stratigraphic unit (Hague Gneiss or Spring Hill Pond Formation), but regional mapping suggests that individual units are continuous over distances of miles. Both Dixon schist and the khondalite were mined for flake graphite during the early part of the century. The region around Swede Mt is known as the Dixon National Forest and the former mining hamlet of Graphite is located at mile 49.8. As Rt 8 is followed eastward, an undeformed dolerite dike is encountered. The dike strikes parallel to the highway and has caused brecciation and alteration of the country rock that consists principally of marble at this locality. At the top of the hill, and across from Swede Pond, a long roadcut exposes typical Adirondack marbles. At the eastern end of these exposures, a thick bed of quartzite is wrapped around the nose of an isoclinal fold. Note the linear features along its axis.

STOP 6. TYPICAL AMCG CHARNOCKITE OF TICONDEROGA DOME
(15 MINUTES).
These outcrops were dated at 1113 ±10 Ma by Silver (l969) but are certainly ca 1150 Ma member of AMCG suite. Silver's "young" reflects the fact that air abrasion was not yet in use when he did the analyses. Note the strong fabric that is typical of AMCG granitoids.

Complex roadcut of marble, calcsilicate, and gabbro.

STOP 7. LOW-TI MAGNETITE ORE IN LYON MT GRANlTE WITH QUARTZ-ALBITEFACIES.
(30 MINUTES). Old magnetite mine workings are located a few hundred feet uphill from the highway and are defined by the excavation of magnetite ore along meter-scale layers dipping ∼45 degrees north. A meter-thick pillar of magnetite-apatite ore was left in place to support the hanging wall. The area below is strewn with waste rock from the 19th century mining operation. The hanging wall of the mine consists of the quartz-microperthite facies of Lyon Mt Granite. Thin section study reveals that for about 150 ft uphill the perthite is progressively replaced by albite with replacement, and checkerboard albite, increasing in the direction of the mine opening. As the mine opening is approached the amount of K2O in the rock decreases, and below the mine pillar, the rock consists of quartz-a1bite rock with magnetite and small quantities of aegerine. Mclelland et al. (2002) have interpreted both the magnetite and quartz-albite rock to be the result of metasomatic replacement due to the percolation of regional fluids at high temperature. It is thought that the fluids were derived from evolved surface brines rich in Na, CL, and scavenged iron. Heat from Lyon Mt Granite drove these hydrothermal cells as they penetrated into the outer margins of Lyon Mt Granite plutons shortly after crystallization. This is consistent with oxygen isotope results (Mclelland et al., 2002). SHRIMP II geochronology demonstrates that the quartz-albite rock contains cored zircons with mantles showing good oscillatory zoning and ages of ca 1050 Ma. These grains are interpreted as unreplaced zircons remaining from the original quartz-microperthite precursor.

Marble then granitic gneiss.

Marble for next 2.3 miles.
Note disrupted amphibolite forming "tectonic fish").

STOP 8. MARCY ANORTHOSITE MASSIF DATED AT ca. 1155 Ma. (30 MINUTES).
This outcrop has yielded a relatively large volume of remarkably well-formed euhedral zircon grains exhibiting excellent oscillatory zoning. Two of these grains are shown on the guidebook cover. A summary of all Adirondack zircon dating appears in Fig 3 and Table 2 of trip A-I where it may be seen that that 14 anorthosite suite samples from across the Marcy massif have now been directly dated by SHRIMP II techniques. These cluster tightly about 1155 Ma, which corresponds, almost exactly to the 15 ages for AMCG mangerites and charnockites also shown in Table 2. These results leave little doubt that the Marcy anorthosite and its granitoid envelope were emplaced at ca 1155 Ma. The results also emphasize that the other AMCG complexes in the Adirondacks were emplaced at ca 1155 Ma as well. Attempts to assign a ca 1040 Ma age to the Marcy anorthosite are inconsistent with this geochronology and with field evidence as well. The characteristics of Adirondack massif anorthosite, and a model for its origin and evolution, are presented in Trip A-l to which the reader is referred for these discussions. The present outcrop is typical of much of the Marcy facies and much of the Marcy massif. Large blue gray andesine grains account for most of the outcrop, but a finer grained matrix is also present. Whether this represents grain size reduction or a distinct magma has not yet been investigated (see Stop 8, Trip A-I). The outcrop contains numerous narrow veinlets of ferrodioritic material that are interpreted as filter-pressed late interstitial magmas from the anorthosite (McLelland et ai, 1994, see Trip A-1).

Roadcuts in Whiteface facies of the Marcy massif.

STOP 9. CORONITIC OLIVINE METAAGABBRO DATED AT 1150 ± 10 Ma.
(20 MINUTES).
This roadcut is typical of Adirondack coronitic metagabbros that have been described by McLelland and Whitney (1977, see Trip A-1). Under the microscope these rocks display remarkably well-preserved igneous textures most of which are sub-ophitic. Superimposed on these are coronas of garnet and clinopyroxene formed between olivine or pyroxene and plagioclase. During the reaction, the olivine is transformed to orthopyroxene and therefore may be totally exhausted. Excess Fe, Mg, and Na from the reaction site migrates out into surrounding plagioclase to form green hercynitic spinel while, at the same time, Ca displaced from the plagioclase migrates back to the reaction site to form garnet The effect is to cloud and impart a distinctive pistachio green color to the plagioclase and to reduce its anorthite content. As a consequence, most plagioclase in these rocks has An∼32% although normative plagioclase is An∼65% (rarely preserved in cores of large grains). The reaction appears to halt at An∼28, presumably because plagioclase this sodic no longer participates in the garnet-forming reaction. As of late-August, 2002, this metagabbro had not yet been dated by SHRIMP II. However, an identical metagabbro near the intersection of the Northway and the Blue Ridge Highway yielded abundant zircons many of which were of excellent igneous morphology and showed good zoning. These gave a well-constrained age of 1150 ±14 Mil, and we expect that the same age will be found for the metagabbro at this stop. Coronitic metagabbros of this type occur throughout the Adirondacks but are notably concentrated near the anorthosite massifs. This, coupled with the similarity in age, indicate that these bodies represent batches of the magma that ponded at the core-mantle interface to form plagioclase-rich crystal mushes that ascended to form anorthosite (See trip A-I). The difference is that these gabbroic magmas did not pond but broke through to the upper crust at an early stage. A spectrum of continuous compositions does exist among the metagabbros and suggests that tapping off of these magmas took place at different stages of differentiation.

B2: GEOLOGY AND MINERAL DEPOSITS OF THE NORTHEASTERN ADIRONDACK HIGHLANDS
Philip R. Whitney
New York State Museum
3140 CEC
Albany NY 12230
James F. Olmsted
48 Haynes Road
Plattsburgh, NY 12901
On this trip we will examine three of the most notable features of the Adirondack Highlands. The first stop is at the NYCO wollastonite mine at Oak Hill near Lewis, where wollastonite-garnet-pyroxene skarns preserve a record of a giant Proterozoic hydrothermal system that was fed by meteoric waters and driven by heat from the nearby Westport Dome anorthosite intrusion. The following two stops, at Arnold Hill and Palmer Hill near Ausable Forks, are in the granitoids and felsic metavolcanics of the enigmatic Lyon Mountain Gneiss, host to numerous low titanium magnetite deposits that were the basis of a flourishing iron mining industry in the late nineteenth century. The remaining three stops, one at Jay and two near Elizabethtown, are in metamorphosed anorthosite and its mafic derivatives, and illustrate the structural complexity and lithologic diversity of these rocks which ordinarily appear on maps as undifferentiated blobs.

B2: GEOLOGY AND MINERAL DEPOSITS OF THE NORTHEASTERN ADIRONDACK HIGHLANDS

This is NYCO property.
This is NYCO property; if following this trip on your own, be sure to get permission at the NYCO office in Willsboro.

STOP 1. OAK HILL WOLLASTONITE MINE.
At the first small quarry on L, park on the R well off the road. From here we will walk through the entire mine complex observing and describing the rock units. As you can see the mine is under development so there may be dangerous situations for which care must be exercised. Please use caution and follow instructions with care.
The Oak Hill Mine was discovered by accident during a logging operation, when a skidder exposed white rock which the logger recognized as wollastonite. It is located at a sharp bend in the OBZ, where the structural trend changes from NNE with a gentle W dip to nearly EW with a gentle to moderate S dip. As of June 2002, mine is being actively developed. Two test pits have been blasted in ore, and work has begun removing overlying rock and glacial deposits in order to begin open-pit mining. Additional exposures may be available at the time of the trip, but others may have been removed or covered. Figure 9, constructed from drill core data, shows the complex nature of the ore body.
The lowermost exposures are knobs of skarn consisting largely of orange-red grossularitic garnet, locally with abundant green, diopsidic clinopyroxene and/or veins and patches of sodic plagioclase and quartz. The garnetite encloses blocks of gabbroic anorthosite gneiss, metagabbro, and pyroxene granulite. Some of these blocks have narrow alteration haloes but otherwise are in sharp contact with the garnetite. They appear to be fragments of disrupted layers in what was probably a marble-hosted tectonic breccia, since metasomatically transformed to garnetite (other interpretations welcome!). Garnetite is much more abundant here than at either Lewis or Willsboro. Drill core data indicate that the the proportion of garnetite to ore increases toward the NE, eventually forming a layer of nearly 200', without ore.
Proceeding up-section, several layers of ore are exposed, separated by septa of mafic gneiss and/or garnetite. Garnetite layers are discontinuous; some appear to be megaboudins.
The hanging wall of the exposed ore consists of strongly foliated and lineated gabbroic anorthosite gneiss. This is overlain in turn by a chaotic marble melange with detached blocks of intricately folded calcsilicates. The marble also contains masses of crumbly-weathered skarn consisting largely of black (dark orange-brown in transmitted light) garnet and dark green clinopyroxene. This skarn resembles those associated with magnetite deposits elsewhere in the Adirondack Highlands rather than the garnet-pyroxene layers in the ore. The fact that this marble has apparently escaped extensive metasomatism suggests that the metasomatizing fluids that formed the ore were strongly channelized.
Overlying the marble is mafic gneiss with interlayers of charnockite and coarse, leucocratic augen gneiss, the latter an extreme L-tectonite. Above the gneiss complex, coronitic olivine metagabbro is exposed. The section in the mine is cut by several E-W, nearly vertical faults, one of which exhibits slickensides plunging steeply E. On the N side of the mine road, at the time of writing, bedrock is concealed beneath a thick section of varved glacial lake sediments inter-layered with boulder till.

Crossing north branch of the Boquet River.

Spectacular fault breccia on L.
This is one of the numerous NNE- to NE- trending faults that occur throughout much of the eastern and central Adirondack Highlands. Movement on these faults, possibly associated with the opening of the Iapetus Ocean, began in the latest Proterozoic and continued at least into the Ordovician.

Pokomoonshine State Park.
The prominent cliff is mostly granitic gneiss of the AMCG suite, containing several layers of mafic gneiss that are probably transposed dikes.

Marble xenolith in jotunite
The prominent white spot in the roadcut on the R is a marble xenolith in jotunite and badly contaminated gabbroic anorthositic gneiss. The marble contains abundant graphite, phlogopite and diopside and lesser amounts of chondrodite, grossular, and sulfides. Wollastonite is absent.

Enter hamlet of Clintonville.
Reportedly, in the late 19th century this was a town of 10,000 miners, wood cutters, iron forge workers and their families.

Stop 2A. LYON MOUNTAIN GNEISS; Albite gneiss facies.
Park off the road on R. The recently blasted low cuts on the L are largely albite gneiss with about 30% quartz, 65% plagioclase (ca. An2Ab96Or2) with minor magnetite and/or hematite, and traces of biotite, clinopyroxene, titanite, apatite, and zircon. The fine-grained, equigranular texture is typical of this facies of the LMG. Note the small, nearly isoclinal Z-folds, N 10-20° E lineation, and numerous crosscutting quartz-albite pegmatites and quartz veins. The pegmatites locally contain masses of partially martitized magnetite, possibly remobilized from nearby ore bodies.

Stop 2B FOLDING IN LYON MOUNTAIN GNEISS.
These flat outcrops at the intersection of Arnold Hill Road and a logging road on the L show complex folding in albite gneiss similar to that at the last stop. Abrupt color changes in the rock crosscut the folding and appear to be contacts of more and less oxidized rock, suggesting a post-deformation oxidation event. This area is not far from the site of the Arnold and Nelson Bush Mines, opened in 1830 and operated sporadically until 1906 (Kemp and Alling 1925; Postel 1952). The ore at Arnold Hill occurs as three foliation-parallel layers from 3-25' thick in albite gneiss. The layers were known to the miners as the gray, black, and blue "veins". The gray and black veins consist of magnetite, quartz, plagioclase, chlorite, and hornblende; in the blue vein the ore mineral is martite (Postel 1952).

Small exposure of LMG on R side of road.

Park off road on R.
Walk around the iron gate and proceed S along the unpaved road furthest to the R.

Stop 3. PALMER HILL.
Follow the road approximately 0.3 miles to the top of Palmer Hill. Along the way are several small pavement outcrops exposing heterogeneous LMG showing complex folding and containing numerous small clots of magnetite ore. Were these formed in place, or are they disrupted fragments of a vein of ore emplaced before deformation? At the top of the hill are several large boulders of Potsdam Sandstone (middle to late Cambrian). The Potsdam unconformably overlies the Proterozoic rocks around much of the perimeter of the Adirondack Dome, and is exposed at the surface less than 6 miles to the NE. An abandoned fire tower here has been converted to a cellphone relay. We will then follow a powerline for about 0.1 mile S to the old mine workings. USE EXTREME CAUTION, especially in wet weather. Cross a narrow rock bridge over the open cut and turn L to view the mine workings. Do not go down into the cut; samples of ore will be available. The ore here occurs in a discontinuous layer up to 20 feet thick, striking NE and dipping NW, roughly parallel to foliation on the SE limb of a NE-plunging synform (Postel 1952). It consists of magnetite with quartz, feldspar, and, locally, apatite, fluorite, and andraditic garnet. The host rock is predominantly the microperthite facies of LMG, although drill cores show almost the entire range of LMG lithologies (Postel 1952). Fluorite is a common accessory in the gneiss, and andraditic garnet occurs locally. Mining at Palmer Hill began in 1825 and continued until 1890 (Newland, 1908).

"Lake Placid Granite" quarries entrance on R.
The "granite"here is actually anorthosite.

STOP 4. ANORTHOSITE OF THE JAY DOME.
Walk back across the bridge (Beware of traffic!) and descend the steep bank on the L to outcrops in the river. The river is usually quite low this time of year affording us good looks at several facies of anorthosite and a variety of structural features. Both leucocratic and gabbroic anorthosite are present, as well as a block of coarser gabbroic anorthosite. In one location, a thin layer of pyroxene- and oxide-rich ultramafic rock separates two anorthosite facies, and is offset by parallel small faults. The latter may have formed at relatively high temperatures; foliation in the anorthosite locally shows the effects of drag along the faults. They are subparallel to later, NE-trending brittle faults, some of which are occupied by unmetamorphosed diabase dikes. One such fault-cum-dike offsets a shallow-dipping mylonite zone in the anorthosite. A xenolith of calcsilicate rock in the anorthosite consists largely of diopsidic clinopyroxene. Numerous potholes are present in the outcrop surface.
Walk back across the bridge and, if time permits, up the road beyond the parking area to a glacially polished and striated outcrop showing rounded blocks of coarse gabbroic anorthosite in finer grained, more leucocratic anorthosite.

Several exposures, of calcsilicates, mafic gneiss and contaminated anorthosites
Several exposures, of calcsilicates, mafic gneiss and contaminated anorthosites may be seen along Rt. 9N between Upper Jay and Keene.

STOP 5. JOTUNITE WITH XENOLITHS.
Park on R beyond the guard rails; cautiously cross the road and walk back to a small outcrop on the N side. The rock here is a jotunite containing numerous metasedimentary xenoliths and one xenolith of anorthosite. The metasedimentary rocks are diopside-rich calcsilicates, some with fine, millimeter to submillimeter-scale layering consisting of alternating pyroxene- and plagioclase-rich layers. The layering is normally straight but locally shows irregular folding or chaotic disruption. The general appearance of the rocks is similar to that of stromatolite-bearing metasedimentary rocks in the Balmat area on the northwest Adirondack Lowlands. Is it possible that the layering here is of organic origin? What other processes might be responsible?
Carefully observe the foliation in the jotunite near the anorthosite xenolith. If this is, as it appears, a metamorphic foliation then it suggests that penetrative deformation may affect some rocks and not others. Beware of assuming that a rock that crosscuts foliation in another rock is necessarily younger!
Interestingly, jotunites in the Adirondack Highlands commonly form intrusion breccias, whereas granitoids, olivine metagabbro, and ferrogabbro do not. Why?

STOP 6: "WOOLEN MILL" GABBRO. ANORTHOSITE, SYENITE.
Park on the L (north) side. The best exposures are the stream N of the road near the remains of an old dam. This stop provides another opportunity to study the characteristics and complexity of Adirondack metanorthosite and its associated lithologies. Fine-grained mafic granulites consisting of granoblastic plagioclase, clino- and orthopyroxene, garnet, ilmenite, and magnetite are of ferrogabbroic composition. This rock has been informally called the "Woolen Mill Gabbro". Here it intrudes the anorthosite as a network of thin, discontinuous dikes. At the upstream end of the outcrop, one of these terminates in a small bulb. More extensive exposures of this lithology are found in the roadcut on the S side of 9N just across from the parking area. A slightly coarser version of the same rock is exposed upstream on the opposite bank, in sharp contact with the anorthosite.
Also present are small dikes of coarse syenite, some with rows of pyroxene crystals along their contacts with anorthosite. Both these and the ferrogabbro appear to be intruded along fractures in the anorthosite, sometimes in the same fractures! (Which came first? Or are these composite dikes formed from two immiscible liquids?).

B3: GEOLOGY AND MINING HISTORY OF THE BARTON GARNET MINE, GORE MT. AND THE NL ILMENITE MINE, TAHAWUS, NY WITH A TEMPORAL EXCURSION TO THE MACINTYRE IRON PLANTATION OF 1857
William M Kelly
New York State Museum
3140 Cultural Education Center
Albany, NY 12230
Robert S. Darling
Department of Geology
SUNY, College at Cortland
Cortland, NY 13045
This field trip examines the geology and history of mining at two types of ore deposits, one a metal and one an industrial mineral. Both are strongly identified with the Adirondack Mountains of New York and both have histories that extend for more than a century. While mining activities began at both in the nineteenth century, only the Barton garnet mining venture has remained in continuous operation. Attempts, successful and unsuccessful, to exploit the ore at Tahawus occurred sporadically through the nineteenth and twentieth century. During this time the minerals that were the target and gangue at Tahawus essentially reversed roles.

B3: GEOLOGY AND MINING HISTORY OF THE BARTON GARNET MINE, GORE MT. AND THE NL ILMENITE MINE, TAHAWUS, NY WITH A TEMPORAL EXCURSION TO THE MACINTYRE IRON PLANTATION OF 1857

STOP 1. BARTON MINES, GORE MT. (1.5-2 hours)
Note: there is a charge of $1.00 per pound of material collected, payable at the mineral shop.

STOP 2. KRONOS, INC., CHENEY POND DEPOSIT (1.5-2 hours)

STOP 3. ADIRONDACK IRON AND STEEL CO. "NEW" FURNACE (1 hour)
Please take only photographs at this stop. Do not take artifacts as souvenirs.

B4: NEW VIEWS ON FAULTING, FABRIC DEVELOPMENT, AND VOLUME STRAIN IN THE TACONIC SLATE BELT, WESTERN VERMONT AND EASTERN NEW YORK
Jean M. Crespi, Jonathan R. Gourley, and Christine M. Witkowski
Department of Geology and Geophysics
University of Connecticut
Storrs, CT 06269
Slate belts provide an important window into tectonic processes, because they preserve bedding despite the penetrative character of the deformational fabric. The reference frame given by bedding, together with information from strain markers, permits quantitative understanding of the deformation since deposition of the strata. Although the complete deformational history can rarely, if ever, be determined, approaches that involve the integration of different types of data are more likely to advance understanding.
The northernmost part of the Taconic Allochthon (Fig. 1) in western Vermont and eastern New York is one of the best studied slate belts in the world. Geological studies (Dale, 1899) accompanied the development of the quarrying industry, which began in the area in the 19th century, and the Taconic orogenic belt was one of the first ancient orogenic belts to be interpreted in the context of plate tectonics (Chapple, 1973). An arc-continent collision interpretation still stands for the origin of the orogenic belt (Chapple, 1973; Rowley and Kidd, 1981; Stanley and Ratcliffe, 1985); however, debate centers on whether the colliding volcanic arc was the Bronson Hill arc or the more inboard and older Shelburne Falls arc (Karabinos et al., 1998, 1999; Ratcliffe et al., 1998, 1999). Geochronologic and biostratigraphic data (Zen, 1967; Laird et al., 1984; Sutter et al., 1985; Ratcliffe et al., 1998) indicate that collisional orogenesis occurred during Middle to Late Ordovician time. The Taconic Allochthon formed as Cambrian?/Cambrian to Middle Ordovician slope and rise strata were thrust west onto approximately coeval strata of the carbonate platform.
The strata within the slate belt of the northernmost Taconic Allochthon were deformed into a series of west-vergent, tight to isoclinal folds during the main phase (D2) of deformation. The slaty cleavage (S2) is parallel or lies at a very low angle to the axial planes of the F2 folds. Although Zen (1961) proposed that the slate belt is characterized by recumbent and downward-facing folds, implying an earlier D1 phase of deformation, Bosworth and Rowley (1984) showed that F1 folds are relatively rare. Locally, S2 is overprinted by a weakly to moderately developed crenulation cleavage (S3). Recently obtained 40Ar39Ar ages from closely packed S3 domains are Devonian and are interpreted as consistent with Acadian rather than Taconian orogenesis (Chan et al., 2001).
This field trip guide builds on a previous NEIGC guide to the area (Goldstein et al., 1997): several outcrops from the 1997 trip are reexamined on this trip in light of new data from a variety of strain markers. In particular, we compare results from markers that record the total strain with those from markers that record different portions of the total strain. Aspects of the deformation that we focus on include evidence for prelithification tectonic deformation; volume change associated with S2 development; and variations in S2 orientation and state of strain with structural level. We also present new observations of meso- and micro-scale structures within the Bird Mountain slice. Specifically, we focus on the information that these structures provide on the movement history of the Bird Mountain fault with respect to regional deformational events.

B4: NEW VIEWS ON FAULTING, FABRIC DEVELOPMENT, AND VOLUME STRAIN IN THE TACONIC SLATE BELT, WESTERN VERMONT AND EASTERN NEW YORK

STOP 1. ABANDONED QUARRY. (45 MINUTES)
This quarry, which exposes the Cambrian Hatch Hill Formation, lies on the upright limb of the Mount Hamilton syncline, a regional-scale, west-vergent F2 fold. S0 lies at a low angle to S2, which dips about 30°-35° to the east. Fibrous calcite veins are abundant in the strata and lie in a variety of orientations.
The outcrop was included by Chan (1998) in his transect of the Taconic slate belt in the vicinity of Route 4. Strain fringes around pyrite framboids in the strata record top-to-west-northwest, non-coaxial flow. Passchier and Urai (1988) used veins in the exposure to estimate the kinematic vorticity number of the flow. They inferred that the dominant vein set formed at some point during S2 development; however, these veins and others display characteristics that suggest that they are early pre-folding veins, which were reactivated during S2 development (Crespi and Chan, 1996). Determination of the flow type required estimation of volume change, which Passchier and Urai (1988) accomplished through geochemical analysis of a sample with a variably developed S2. They determined a volume loss of 50%. Erslev (1998), in contrast, analyzed samples from the quarry with a penetrative S2 and concluded, on the basis of geochemical data, that large volume loss occurred only at the scale of individual cleavage domains.

STOP 2. CEDAR POINT QUARRY. (1 HOUR 15 MINUTES)
The hinge zone of the Cedar Mountain syncline, a regional-scale, overturned, west-vergent, isoclinal F2 fold, is exposed at this stop. The folded strata consist of the Cambrian Mettawee Formation. The dominant cleavage in the exposure is S2, which dips about 25° to the east and is parallel to the axial plane of the fold. A moderately developed S3 is also present.
Green ellipsoidal reduction spots are common in the maroon slate and can readily be seen in the abandoned blocks on the quarry floor. Irregularly shaped areas of reduction and bands of reduction subparallel to S0 are also common. Measurements of the ellipsoidal reduction spots have been made by Wood (1974), Hoak (1992), and Goldstein et al. (1995). Minimum volume losses during S2 development of 55% determined from the reduction spot data (Goldstein et at, 1995) are in conflict with geochemical data from the quarry, which are consistent with minimal nonvolatile volume loss (Erslev, 1998).

STOP 3. SCOTCH HILL SYNCLINE. (45 MINUTES)
This is a well known exposure in New England. The property owner asks that hammers not be used on the outcrop, so please leave your hammer in your vehicle. Also, poison ivy may be dense at the base of the exposure.
This exposure of the Ordovician Poultney Formation lies in the hinge zone of the Scotch Hill syncline, a regional-scale, west-vergent F2 fold. S2 dips about 30° to the east and is axial planar to the fold. En echelon, sigmoidal quartz veins provide evidence for flexural folding of the strata with pinning at the hinge. The veins are present on both limbs of the fold but are more systematic on the vertical limb and, thus, provide a better example of vein development within shear zones. (In this context, S0 acts as the shear plane.) The veins are easily found on the vertical limb of the fold. To observe the veins on the gently dipping, upright limb, go around the small shed and climb to the base of the exposure behind the shed. Veins on both limbs of the fold are folded and/or boudined. On the vertical limb, the veins have a Z-shaped asymmetry (viewed to the north), indicating east-side-down sense of shear along S0. On the gently dipping, upright limb of the fold, the veins have an S-shaped asymmetry (viewed to the north), indicating top-to-west sense of shear along S0. Pre-folding veins, which are not coaxial with the fold, can also be observed in the exposure. These appear approximately parallel to S2 on the fold profile plane on both limbs of the fold, but three-dimensional exposures show that the veins lie at an angle to S2. The pre-folding veins lie in approximately the same orientation on both limbs of the fold. This results because shearing within layering effectively cancelled the change in orientation from rigid rotation of the fold limbs.

STOP 4. STODDARD ROAD GRAPTOLITES. (1 HOUR)
The first exposure (SR1) that we will visit at this stop is located immediately east of the road. The second exposure (SR2) is located on the small hill in the middle of the drained beaver pond and, if the area is marshy, can be reached by crossing the small stream at the south end of the drained pond. Both exposures are rather smalL Please be kind to them with your hammers so that geologists on trips in the future can also enjoy the graptolites and structural features.
The two exposures are of the graptolite-bearing Stoddard Road Member of the Ordovician Mount Merino Formation. Exposures of the Mount Merino Formation in the vicinity of Stoddard Road are in the core of a regional-scale F2 syncline. SR1 is located on the overturned limb of the syncline, and S0 and S2 are essentially parallel. S0/S2 dips about 45°-50° to the east. Graptolites are readily observed in place on the underside of the overhanging S0/S2 surfaces and in numerous pieces of float. The hinge zone of a mesoscale fold is exposed at the north end of the outcrop where graptolites can be observed on an S0 surface at a very high angle to S2. Mesoscale folds, which cannot be recognized using stratigraphic relations, may be common in the generally poorly exposed slates of the region. SR2 is located in the hinge zone of the regional-scale F2 syncline. Well-preserved graptolites are present but more difficult to find at the exposure. S0 lies at very high angles to S2, which dips about 45°-50° to the east, and the exposure provides good examples of using S0-S2 relations to determine whether strata are on the upright or overturned limb of a fold. The axial-planar character of S2 can be observed in the hinge of the fold, which is exposed on the north side of the hill in a small block that is probably slightly out of place.

STOP 5. BIRD MOUNTAIN FAULT OVERLOOK. (20 MINUTES)
From this vantage point, the Bird Mountain fault is clearly expressed as the abrupt topographic transition between the valley containing Little Pond and Lake St. Catherine and the hills to the east. Good exposures of the fault can be viewed in a large road cut along the east shore of Lake St. Catherine on Route 30, but access is limited and the orientation of the exposure is parallel to strike. To the east of the fault in this region, the Mettawee (St. Catharine Formation of Shumaker (1967)) and West Castleton/Hatch Hill formations underlie the hills of the Bird Mountain slice. Slate quarries are absent from the Bird Mountain slice despite the abundance of the Mettawee Formation. Zen (1967) noted that the Mettawee within the Bird Mountain slice is siltier than Mettawee to the west and, therefore, is unfit for roofing material.

STOP 6. BIRD MOUNTAIN SLICE. (45 MINUTES)
This small outcrop records several deformational events to affect the Taconic Allochthon. The rocks are within the Mettawee Formation. S0 is folded and the fold axes trend northwest. A variably spaced, axial planar S2 cleavage cuts S0. S2 is well defined along limbs of folds but nearly absent in the hinges. A weak spaced cleavage, steeper than the pervasive S2 cleavage, is interpreted to be S3. Refer to Fig. 10 to see the local relation between S2 and S3. S2-parallel veins are present at the south end of the outcrop.

STOP 7. BIRD MOUNTAIN FAULT. (45 MINUTES)
Outcrops in the yard behind the house, in the woods above, and along Route 30 (beware poison ivy!) contain abundant veins that are parallel and subparallel to S2. This is the prominent S2 cleavage that dominates the Allochthon from here west, but is folded in this manner only near the Bird Mountain fault. In addition, the S2 surface and veins are deformed into asymmetric chevron folds. The green and purple rocks exposed here are the Mettawee Formation. The orientation of S2 is highly variable due to the intense chevron folding.

B5: MIDDLE ORDOVICIAN SECTION AT CROWN POINT PENINSULA
Charlotte Mehrtens
Department of Geology
University of Vermont
Burlington, VT 05405
Bruce Selleck
Colgate University
Hamilton, NY 13346
The Chazy. Black River and Trenton Group are a well known sequence of fossiliferous limestones and sandstones in the Champlain Valley of New York., Vermont and southern Quebec. These rocks record cyclic sedimentation in shallow water environments on the flanks of the Iapetus Ocean during the Middle Ordovician transgression. This field trip reviews the stratigraphy of these units and the evidence for their depositional environments.

STOP 1. Redoubt
Approximately 6 m of variously burrowed, slightly dolomitic, thin to medium bedded bioclastic packstone and grainstone is exposed in this section. Some beds are relatively well-sorted grainstones with sharp bases and are interpreted as storm deposits. Note the large intraclasts in the base of one unit in the low ledge at the southeast corner of the ditch. These indicate rip-up of cemented grainstone, apparently by a storm. Abundant "Girvanella" algal oncolites are present in beds approximately three meters from the base of the section. Rounded dark calcite grains (abraded gastropod fragments) form the cores of the oncolites, and are scattered in other beds. Fossils are relatively abundant and best seen on bedding surfaces. Trilobite fragments, brachiopods, bryozoans, pelmatozoan plates, nautiloids and large Maclurites magnus are present. Dolomite occurs in shaley weathering laminae and in burro fills.
The relatively high faunal diversity, general bioturbation, and storm-related sedimentation point to a 'low energy', shallow subtidal environment at depths slightly below normal wave base. The abundant algal oncolites and discrete calcareous algal fossils (e.g. Hedstromia) suggest depths well within the photic zone. A possible modern analogue is found in the mixed mud and sand shelf to the west of the emergent tidal flats of Andros Island Bahamas, as described by Bathurst (1971) and Purdy (1963).
The wavy, irregular dolomite laminae result from dolomitization of lime mud, followed by compaction and local pressure solution of calcite, that produced irregular, clay- and dolomite-rich stylocumulate seams. Preferential dolomitization of burrows is due to contrasts in porosity or permeability of burrow-fill versus burrow-matrix sediment. The burrow-fill material retained permeability longer during diagenesis and allowed more pervasive dolomitization. In similar facies exposed on Bullwagga Bay (west shore of peninsula), nodular limestone with shaley dolostone seams and stringers are present (Fig. 2). The limestone nodules appear to have been cemented prior to significant burial compaction, whereas the shaley dolostone material was compacted around the cemented limestone. The early-cemented limestones were resistant to dolomitization. This sort of fabric selective dolomitization is common in the Chazy and Black River Groups throughout the Champlain Valley. Can you find other evidence of early cementation in this outcrop. Are there hardgrounds?

STOP 2. Ridge outcrop extending NE from near entrance Gate
Cross-stratified bioclastic grainstones are well-exposed near the entrance road and adjacent ridge. Nearly three meters of section form a prominent belt parallel to strike extending from the entrance road to the main highway. Forset cross-strata show bipolar dip directions. Angular quartz and feldspar grains are concentrated in some laminae. The carbonate particles are dominantly subrounded, abraded pelmatozoan plates with gastropod and brachiopod fragments. Large Maclurites fragments and grainstone intraclasts are present on the upper bedding plane surfaces of the ledges northeast east of the entrance road.
We envision the environment of deposition of this facies as shallow subtidal wave and current reworked sand shoal. Active transport of abraded grains may have been accomplished by tidal currents (as suggested by the bipolar cross-beds) or by storm-generated currents that produced complex, anastomosing patterns of cross-strata and intervening reactivation surfaces. The lack of burrowing and well-preserved whole shell body fossils may be due to the inhospitable shifting sand substrate. This environment likely resembles the unstable sand shoal environments described from the Bahamas Platform (Bathurst, 1971; Ball, 1967). The scale and style of cross-stratification present here are similar to that predicted by Ball (1967) from study of the bedforms present on Bahamian sand bodies. Similar Chazy Group facies in the Northern Lake Champlain Valley are oolitic (Oxley and Kay, 1959).
Note that these grainstones are essentially undolomitized. Does this indicate early cementation or diagenetic stabilization prior to deeper burial?

STOP 3. Low Ledges adjacent to entrance road (Picnic Pavilion Ridge) approx. 50 meters north of 2
Brown weathering, slightly shaley dolostone exposed here contains small lenses and stringers of fossiliferous lime packstone. As at Stop 1, trilobites, small brachiopods and Maclurites fragments are common. The environment of deposition is assumed to be subtidal shelf, with less storm activity than at Stop 1.
Note that some of the fossils are almost entirely encased in dolomite, which is assumed to be of replacement origin here. Why are some of the fossils so well-preserved, and not dolomitized?

STOP 4. SE moat of Fort Crown Point
Approximately 3 meters of thickly laminated limestone and dolostone is exposed in the southeast "moat" of the British Fort. The dominant facies here consists of alternating 0.5-2 cm thick laminae of limestone and dolostone - often termed a "ribbon rock". The limestone ribbons are very fine peloid grainstones of 'calcisiltites' and appear blue-grey on slightly weathered surfaces, and are indentations on highly weathered surfaces. The dolostone ribbons weather tan-brown and consist of an interlocking mosaic of 20-300 micron dolomite crystals of replacement origin. Quartz silt grains are present in the dolostone ribbons, versus medium to fine quartz sand in the limestone, suggesting that the limestone ribbons were slightly coarser than the dolostone when originally deposited.
An erosional surface with 10-20 cm of relief is exposed near the base of the south wall. Similar erosional surfaces occur within this facies in other exposures, and appear to represent micro-karstic solution surfaces on a tidal rock platform that developed during subaerial exposure of cemented limestone (Fig. 5). Typically, the rock below the surface is mostly calcite limestone, suggesting that cementation and diagenetic stabilization of the limestone occurred prior to development of the erosional surfaces. Overlying rock contains more dolomite. Maclurites shell hash can be found in pockets on the erosional surface, suggesting wave transport of shells onto the rocky platform of the erosional surface. Dolomitized burrows cut across the limestone ribbons in some parts of the outcrop, and trough cross-strata that appear to fill low scours are also visible.
On the less-weathered prominence on the SE corner of the moat, shallow scours containing brachiopods and gastropod debris are seen. Intraclasts or pseudoclasts of limestone in dolostone are also present. Some 'clasts' appear to be cored by dolomitized burrows.
We interpret this sequence as a tidal flat to shallow subtidal facies. The alternating limestone/dolostone "ribbon" fabric may represent rhythms of slightly coarser (limestone) and slightly finer (dolostone) sediment deposited on the lower reaches of a tidal flat, similar to the bedding described by Reineck and Singh (1980) from the mud/sand tidal flats of the North Sea. These coarse-fine alternations might also reflect storm-related, ebb-surge deposition. Early cementation of the slightly coarser limestone ribbons made this lithology less susceptible to later dolomitization, which affected the finer, muddy ribbons that are now dolostone. Variations in intensity of burrowing reflect subtle differences in duration of subaerial exposure of the flat and/or reworking by tidal currents. Limited in situ faunal diversity is also expected in a stressed tidal flat setting. The absence of mudcracks and evaporite minerals may indicate that only the lower portion of a humid climate tidal flat system is preserved here.

STOP 5. Parade Grounds near barracks.
As we enter the parade ground from the southwest corner of the moat, note the array of carbonate rocks used in construction of the barracks. Chazy, Black River, and rare Trenton lithologies can be identified.
Restoration of the barracks began in 1916. More recently, starting in 1976, the N.Y. State Division for Historic Preservation undertook protection and stabilization of the ruins.
The low rock pavement just north of the barracks is within the upper part of the 'ribbon rock' unit seen at stop 4. Immediately up-section, cross-stratifiai grainstone beds are visible. Coarse quartz and feldspar sand is easily seen on weathered surfaces. Trough cross-strata and 'herringbone' co-sets of planar-tabular cross-strata are visible on the low vertical faces. Large angular clasts of slightly dolomitic grainstone and Maclurites are present on bedding surfaces. We interpret this facies as a current-dominated sand shoal environment similar to that seen at Stop 2B.
Note the polygonal pattern on some outcrop surfaces. Are these mudcracks?

STOP 6. West parade grounds
Bechtel and Mehrtens (1993) suggested that the sandstone unit in the westernmost parade ground is the basal sandstone of the Black River Group, an interpretation which differs from that of Speyer and Selleck (1988) who suggested that this unit was part of the underlying Chazy Group. In thin section this sandstone is a quartz arenite in composition, very poorly sorted, containing fewer lithic fragments and phosphatic fragments than Chazy sandstones. Visible at the very easternmost portion of this ridge an overlying buff-colored dolomite bed containing pockets of quartz sand (burrow infills?) is exposed. The sandstone and dolomite lithologies are very similar to those described by Walker (1972) for the Pamelia Formation at the type locality in New York. Alternatively, placement of the sandstone within the Chazy Group is consistent with the common presence of coarse quartz and feldspar sand within the Chazy here, whereas siliciclastic material in the Black River Group at Crown Point is mainly fine silt and clay. Whatever the stratigraphic placement of this unit, it marks an interval when sands were transported from a nearby (Adirondack?) source area. This period of subaerial exposure of the shelf was followed by marine reworking of the sand, and deposition of the muddy carbonates of the basal Black River Group.

STOP 7. Northeastern moat
There is approximately 4 feet of covered interval between the parade grounds and moat exposures. The wall of the northeastern most moat exposes several feet of the lower Black River Group (Lowville Formation, House Creek Member). At the base of this exposure a series of stylolitized gastropod-bearing (Liospira) wackestone beds are overlain by thick beds of Phytopsis-burrowed aphanitic mudstones. This sequence can be interpreted as an example of a "classic" shallowing-up cycle consisting of subtidal to peritidal lagoonal muds. Examine the sharp contact of the aphanitic mudstone with the overlying wackestones, a contact which in thin section appears to be a firmground (Bechtel, 1993).
Examination of Phytopsis burrows in thin section (Fig. 6) reveals that many are filled with graded (fining-up) geopedal silt, evidence of cementation in the meteroic vadose zone as shown below.
Continuing upsection, several thick packstone beds are exposed. More detailed examination of these beds reveals that they consist of alternating one to three inch thick intraclast and oncolite-rich packstone horizons interpreted as tempestites, interbedded with fossiliferous wackestone/packstone horizons. The tempestites, or storm-generated deposits, consist of graded and crudely imbricated intraclasts and skeletal fragments. Note the nature of the upper and lower contacts of these horizons. Horizons of tempestite beds within in situ fossiliferous muds is a second motif which occurs throughout much of the Black River Group.
It is instructive to spend several moments sketching the basal six feet of the moat exposure. Your sketch could illustrate the basal SUC's as well as the upper and lower contacts and internal fabric of the tempestite-rich beds.
The uppermost third of the outcrop appears to be a massive bed of limestone, however closer examination also reveals small scale "cycles" of alternating wackestone/packstone and grainstone, the third motif of bedding in the Black River. These are characterized by a base of thinly laminated or cross laminated grainstone horizons 1 to 2 inches thick, overlain by fossiliferous wackestones and packstones. In thin section the bases of the grainstones can be identified as firmgrounds, recognizable by the truncations of allochems and cements in the underlying mud.
Figure 7 illustrates the nature of the contact between a peloid-rich grainstone of the base of a cycle and the top of the underlying packstone. The scalloped surface and truncated grain and allochems are characteristic of these contacts.

Stop 8. Ledges extending from service road to lake shore.
A black chert layer near the top of Stop 7 provides the correlation to Stop 8, the outcrop across the service road
The limestone beds on this ridge commonly consist of alternating wackestone/packstone and planar to cross laminated grainstone beds as seen at Stop 7, however bedding plane exposures permit identification of many fossils in these, the most faunally diverse beds in the Black River. Specimens of gastropods (Liospira, Lophospira, Hormotoma). Lambeophyllum, Tetnodium, stromatoporoids, the bivalve Cyrtodonta, Strophomena sp. and cephalopods are recognizable. This ridge is most notable for its bedding plane exposures of Tetrodium and Lambeophyllum and is interpreted as recording a wave baffle margin lithofacies described by Walker (1972) at the type section.
A sketch of the third type of cycle, laminated grainstone overlain by bioturbated wackestone, with special attention paid to the nature of the contacts between cycles, would be appropriate. These sequences are interpreted to represent smaller scale 4th order 'micro-cycles' that could be the result of facies mosaicking and/or small scale base level changes.
There are at least two distinct types of chert occurrences in the Black Rivet: One appears to be fabric selective: infilling horizontal burrows, for example (many other burrows are dolomitized). The other chert occurrence is less frequent and consists of broad bedding parallel sheets. The uppermost chert horizon on this ridge, traceable down to the shoreline, is of this latter variety. Clearly, there was a significant source of silica available for chert formation, perhaps a combination of silica derived from sponges (Tetrodium?) and bentonite alteration (the Ordovician sequence in the Champlain Valley is notorious for the paucity of bentonites compared to equivalent strata in central New York and Quebec). In thin section the chert cross cuts all previous cements, including late fracture-filling calcite, and is therefore the youngest, latest example of diagenesis in these rocks.

STOP 9. Quarry
Be extremely careful around the quarry - the thick algal scum in the quarry water obscures where the grass begins and the quarry wall drops off.
The older, weathered south walls of the quarry show, by color differentiation, two cycles. Closer examination of the more visible north wall reveals more occurrences of the third motif of Black River bedding: 8 to 10 inch thick beds of planar laminated skeletal and peloidal hash overlain by burrowed wackestones overlain by an intraclast-rich horizon. Interbedded with these cycles are also tempestite couplets of mudstone/wackestone and fossil hash layers in which brachiopod-rich layers are abundant (look for 'nested' pockets of shells).
Sketch the interbedding of the two cycles below:
Before leaving the quarry area, note the numerous quarried blocks stacked between the quarry and the shoreline. See if you can recognize cycles, and from these topping directions.

STOP 10. Exposures of Glens Falls ls separated by intervals of beach gravels
Uppermost horizons in the quarry can be traced down to the shoreline to the north where the Black River section continues (Chaumont Formation of the Black River Group) with small covered intervals up to the Trenton Group limestones. Shoreline bedding planes exhibit horizontal burrows of Chondites, opercula of Maclurites, and polished surfaces also reveal large intraclasts, in other words the same internal stratigraphy that could be viewed in cross section on the quarry walls. Note: exposures may only be visible at low water levels.

STOP 11. Shoreline of eastern Bulwaga Bay
Continuing up the shoreline from the uppermost outcrops of Black River a thin covered interval (4') occurs before the basal beds of the Glens Falls Formation. The Glens Falls is characterized by thin, nodular to wavy bedded wacketones, mudstones and rare grainstones, a very different type of bedding style and faunal assemblage from the underlying Black River. Beding planes along the shoreline contain mostly Chondites and Helmanthopsis burrows, however as one moves up section recognizable piece of Cryptolithus, Isotelus, orthid brachiopods, Stictopora, and Prasopora simulatrix can be found, the latter is important became it permits the correlation of the lower Glens Falls here in the Champlain Valley to the lower Denley Limestone at the Trenton type section in central New York. MacLean (1987) interpreted the lithofacies of the basal Glens Falls to represent sedimentation in a shallow subtidal environment periodically influenced by storm activity. In thin section the nodular and wavy bedded wacketones appear thoroughly bioturbated, a process which would influence and enhance subsequent differential compaction. Grainstone beds exhibit more planar bases with basal skeletal fragment lags or finely crushed debris of brachiopod, trilobite and crinoidal material capped by mud (see Fig. 10). MacLean interpreted these as tempestite deposits. Based on lithologies, fossil assemblages and sedimentary structures, or the lack thereof, the Glens Falls Formation is interpreted to represent sedimentation in a deeper water, subtidal setting relative to the Black River deposits. For those familiar with lower Trenton localities in central New York the paucity of fossiliferous bedding planes here at Crown Point is noteworthy. The overall fine grain size and ichnofauna suggest that bathymetry increased significantly and rapidly from the Black River into the Glens Falls, a transition which might reflect not only rising sea level but base level changes as well. The sedimentologic and faunal transitions from the Glens Falls to the overlying Cumberland Head Argillite and Stony Point Shale are much more gradational than that of the Black Rived / Glens Falls contact.

B6: EARLY PALEOZOIC SEA LEVELS AND CLIMATES: NEW EVIDENCE FROM THE EAST LAURENTIAN SHELF AND SLOPE
Ed Landing
New York State Museum
The State Education Department
Albany, NY 12230
Purpose
Field studies undertaken in eastern New York in the early 1800s by Amos Eaton and others led to the birth of American geology (see review by Friedman, 1979). With this early work, geology continued as an important applied discipline, but it developed into an important intellectual and cultural activity that provided the first empirical explanations of physical and biotic change through earth's "deep time." Although almost two centuries have passed since Eaton's first work, "everything isn't known and decided" either about the geology of eastern New York or its significance in opening a window into "deep time." This field trip will illustrate recent published syntheses of eastern New York geology that help reconstruct Early Paleozoic sea-level and climate changes.

B6: EARLY PALEOZOIC SEA LEVELS AND CLIMATES: NEW EVIDENCE FROM THE EAST LAURENTIAN SHELF AND SLOPE

Small road cuts in Grenvillian Proterozoic at south end of village.

First view of high ridges in Taconic allochthon
At crest of low rise at south end of Proterozoic of French Mountain, note first view of high ridges in Taconic allochthon directly in front of vehicles.

Pass Queenbury Country Club on left.
Underlying less resistant Cambrian-Lower Ordovician explains lack of relief.

Low road cuts on left (North) in Grenvillian
Low road cuts on left (North) in Grenvillian at south end Sugar Loaf Proterozoic inlier.

Low road cuts in Grenvillian inlier east of Hadlock Pond fault.

N-S-trending ridges in Taconic allochthon
Clear crest of hill and see spectacular view (if weather is clear) of N-S-trending ridges in Taconic allochthon across pastures developed on glacial outwash.

View to left (NE) of last ridge of Adirondacks
View to left (NE) of last ridge of Adirondacks east of Welch Hollow fault. East slope is nonconformity surface with lower Upper Cambrian Potsdam Formation eroded off.

Cross bridge over small creek
Cross bridge over small creek with medium-massively bedded dolomitic limestone and replacement dolostone. The locally oolitic, thrombolitic, intraclast pebble facies exposed here are more suggestive of the Upper Cambrian Little Falls Formation, rather than the Lower Ordovician Tribes Hill Formation (e.g., "Fort Edward Dolostone" as mapped by Fisher, 1984).

STOP 1. LOWER ORDOVICIAN AT KANE FALLS: AGE AND LITHOSTRATIGRAPIDC REINTERPRETATIONS. (30 MINUTES).
A stratigraphic section for Kane Falls will be distributed on this field trip; this section is available in Landing et al. (2003, fig. 5).
Mazzullo (1974) and Fisher and Mazzullo (1976), respectively, termed the upper part of the Kane Falls sequence a "reference section" for the lower part of the lowest Ordovician "Cutting Formation" and "Great Meadows Formation" (abandoned designations). A planar-laminated silty dolostone 9.55-10.75 m above the base of this 23 m-thick section (Landing et al., 2003, fig. 5) at the top of the falls was regarded as a thin or condensed "Winchell Creek Siltstone" (abandoned term, now Sprakers Member of Tribes Hill Formation; Landing et al., 2002; see Appendix). This interpretation meant that Kane Falls is near the western feather-edge of this lowest "member" of the "Great Meadows" and that a disconformity separated the silty dolostone from the underlying "Whitehall Formation" (abandoned designation, now Little Falls Formation; Landing et al., 2003, see Appendix).
Re-investigation of the Kane Falls section (Landing et al., 2003) led to its re-interpretation. 1) There is no physical evidence for disconformity at the base of the silty dolostone. 2) The medium-gray thrombolitic dolostones and dolomitic limestones of the lower 9.55 m of the section do not resemble the light gray or white thrombolites of the upper Little Falls Formation ("Whitehall") at Steves Farm (Stop 6) or elsewhere in the Lake Champlain Lowlands. These medium-gray thrombolitic facies resemble those in the middle "Fort Ann Formation." 3) Diverse conodonts in the thrombolitic facies and rarer conodont elements in the higher laminated dolostones at Kane Falls are significantly younger than those from the Tribes Hill Formation (="Cutting"/Great Meadows," designations abandoned by Landing et al., 2003). The conodonts from Kane Falls are all referable to middle Lower Ordovician conodont Fauna D of Ethington and Clark (1971), and are comparable to those of the "Fort Ann Formation" elsewhere in the Lake Champlain lowlands.
The "moral" of this stop is the following: Mapping and stratigraphic syntheses on Lower Paleozoic carbonates on the New York Promontory require a combination of a detailed familiarity of lithofacies with adequate micro- and macro-fossil investigations to allow age discrimination of broadly similar carbonate platform units.

Cross onto uplifted (Proterozoic) block
Cross onto uplifted (Proterozoic) block of Welch Hollow Fault and pass by ca. 1 mile of Grenvillian road cuts.

STOP 2. PROTEROZOIC-TERMINAL MIDDLE CAMBRIAN NONCONFORMITY
STOP 2. PROTEROZOIC-TERMINAL MIDDLE CAMBRIAN NONCONFORMITY AND THIN LOWER PALEOZOIC SHELF SUCCESSION. (10 MINUTES). Approximately 4 m of east-dipping, medium-coarse grained, slightly dolomitic quartz arenite of the Potsdam Formation nonconformably overlie Grenvillian gneiss with east-dipping exfoliation surfaces. This photogenic locality records the absence of ca. 600 million years of earth history at this planar nonconformity. Flower (1964, p. 156) reported a trilobite fauna (presently unillustrated) approximately 5-6 m above the base of the Potsdam in this area, which he said includes a possible Crepicephalus with Komaspidella, and Lonchocephalus. Following the recent decision (January 2002) of the Cambrian Subcommission to define the base of the Upper Cambrian at the lowest occurrence of the agnostoid Glyptagnostus reticularus, this traditional "Upper Cambrian" Crepicephalus Zone assemblage must now be regarded as uppermost Middle Cambrian. This lower Potsdam Formation fauna indicates that the earliest marine onlap in this part of the Lake Champlain-Hudson River lowlands was terminal Middle Cambrian.
The purpose of this stop is to emphasize that the Lower Paleozoic of the eastern New York shelf is a relatively thin succession by comparison with Great Basin and southern Canadian Rockies sequences (i.e., hundreds of meters vs. several thousand meters). The ridge crest directly to the east is underlain by the upper Little Falls Formation and indicates that only ca. 200 m of Upper Cambrian is present above the Grenvillian. Thin Lower Paleozoic successions occur in this region of New York and Vermont (e.g., Palmer, 1971) due to its location on the slowly subsiding New York Promontory rift-margin of Laurentia (Thomas, 1977; Williams, 1978). For this reason, this passive margin can be expected to record Early Paleozoic eustatic changes as type 1 erosive sequence boundaries (e.g., Landing et al., 2003). Two of these type 1 sequence boundaries will be viewed in the course of this trip.

Topographic break corresponding to the the Proterozoic-Cambrian nonconformity.
Note that Rte 4 roughly follows a topographic break corresponding to the the Proterozoic-Cambrian nonconformity.

Pass cuts in Grenville gneisses

Dip slope at 10-12 O'Clock is nonconformity surface on Proterozoic.

Hill directly in front (North) of vehicle is Skene Mountain in Whitehall.
This rock under this hill corresponds to the entire terminal Middle-Upper Cambrian of the Lake Champlain Lowlands.

Continue north through Whitehall along the foot of Skene Mountain.

STOP 3. POTSDAM FORMATION ROADCUT. (10 MINUTES).
Approximately 30 m of Potsdam Formation siliceous quartz arenites and dolomitic quartz arenite are exposed in the Whitehall area. The village of Whitehall east of the Champlain Canal lies on the Proterozoic nonconformity, and the canal is excavated in Potsdam Formation sandstones. Herring-bone cross sets, dolomitic quartz arenite pebbles at the base of Small channels/dunes, and U-shaped Diplocraterion burrows all point to tidally influenced, higher energy deposition of the Potsdam. The Potsdam is overlain by lower Upper Cambrian dolomitic quartz arenites and quartzose dolostones of the Galway Formation (="Ticonderoga Formation," designation abandoned; Landing et al., 2003, see Appendix) higher on Skene Mountain.

Cuts in Potsdam Formation immediately on the right.

STOP 4. CROSSMAN QUARRY: UPPER CAMBRIAN LIITLE FALLS FORMATION.
The eastern end of the quarry shows ca. 18 m of medium- to massively bedded oolitic, stromatolitic, and thrombolitic limestones with 4 m of overlying calcareous quartz arenite. This shallow carbonate platform facies, which represents Fisher's (1977) undefined "Warner Hill Limestone" member of the middle "Whitehall Formation" (abandoned designation), is increasingly dolomitized and neomorphic with the loss of primary features, in the western part of the quarry, where it is laterally transitional into the "Skene Dolostone" (Wheeler, 1941) sensu Fisher (1977, 1984). Although a biostratigraphically useful Late Cambrian trilobite assemblage appears ca. 3.5 m above the base of the eastern quarry wall (Taylor and Halley, 1974), snails and other mollusks are the dominant macrofauna! elements in this sequence, and indicate a restricted, relatively near-shore biofacies (e.g., Westrop et al., 1995).
The purpose of Stop 4 is to show the lithology of non-dolomitized Upper Cambrian limestones of the Little Falls Limestone ("=Whitehall Formation," abandoned term; Landing et al., 2003, and Appendix) for comparison with lithofacies of the overlying Tribes Hill Formation at Stop 5. The dolomitic limestones at the west end of Crossman Quarry and at Steves Farm and Comstock (Stops 6 and 7) were likely produced by the movement of hydrothermal fluids during burial with the Taconic orogeny (e.g., Collins-Wait and Lowenstein, 1994).

Top (Middle Ord. Providence Island Dolostone) of the carbonate platform.
The first ridge to the east is the top (Middle Ordovician Providence Island Dolostone) of the carbonate platform. A thrust relationship suggestive of the Champlain thrust further north brings the Potsdam Formation and younger Cambrian-Middle Ordovician units onto the Providence Island Dolostone. The second ridge to the east is the front of the Taconic allochthon, which overrides the synorogenic, Lower Upper Ordovician "Snake Hill Formation" (Fisher, 1984).

STOP 5. TRISTATES QUARRY: LOWEST ORDOVIClAN TRIBES HILL FORMATION.
A stratigraphic section for the Tristates Quarry will be distributed on this field trip; this section is available in Landing et al. (2003, fig. 4).
Landing et al. (2003) measured a 32 m section along the west and north walls of the old Tristates Quarry. The section is lithically identical to the middle part of the Comstock section (Stop 7) in showing: 1) thin- to medium-bedded, current cross-bedded, fine-grained dolomitic quartz arenite in the lower part of the section; 2) a thin (ca. 1.5 m) dark gray, pyritiferous, siltstone with a middle tempestite bed (note prominent rusty-weathering layer in the middle of the west wall); and 3) a higher interval with massive thrombolites to 2.9 m high. Landing et al. (2003) demonstrated that the conodonts from the Tristates Quarry indicate a precise correlation with the Tribes Hill Formation in the Mohawk River valley (see Landing et al., 1996). In addition, they described how intervals 1, 2, and 3 (above) in the Tristates Quarry correspond to the Sprakers, Van Wie, and Wolf Hollow Members of the Tribes Hill in the Mohawk valley-even down to the presence of a storm bed in the middle Van Wie Member and thrombolites in the Wolf Hollow. These data indicate that the Tribes Hill Formation is regionally extensive on the New York Promontory, where it comprises a deepening shoaling sequence (Landing, 1998) with the Van Wie indicating the deepest facies. Fisher and Mazzulo (1976) earlier referred this section to the "Great Meadows Formation" and recognized a lower "Winchell Creek Siltstone" member (units 1 and 2, above) and assigned the higher strata to Fisher's (1977) undefined "Fort Edward Dolostone." Somewhat later, Fisher (1984) separated the interval with prominent thrombolites (lower unit 3, above) as his "Kingsbury Limestone," but maintained "Fort Edward" for higher strata, thus creating an objective homonym for the undefined "Port Edward" (see Appendix).
In the course of Stop 5, note the large fallen blocks along the west wall that show sections through the massive hexagonal thrombolites. The inter-thrombolite lithology consists of echinoderm grainstone with ellesmeroceratoid cephalopods, orthid brachiopods, snails, and trilobites (Symphysurina and Hystricurus), and is the most unrestricted marine biofacies in the local Tribes Hill Formation. The abundance of snails and other mollusks increases higher through the quarry. Note also the lateral transitions along the north wall from limestones with well preserved primary fabrics and allochems into structureless, recrystallized dolostones, probably as a result of the hydrothermal activity noted at Stop 4.

Lower part (Sprakers Member) of Tribes Hill Formation.
Road cut on right (north) in fine--grained, dolomitic quartz arenites of lower part (Sprakers Member) of Tribes Hill Formation.

Road cut on right in massive dolostones of upper Little Falls Formation

STOP 6. STEVES FARM: AGE AND LITHOSTRATIGRAPHIC REVISIONS OF THE TERMINAL CAMBRIAN.
A stratigraphic section for the Steves Farm section will be distributed on this field trip; this section is available in Landing et al. (2003, fig. 4).
A carbonate-dominated, 85 m section was measured from the pasture 0.5 km west of Upper Turnpike and through gently east-dipping strata on the wooded north side of the 108 m (360 foot) hill at this stop (Landing et al., 2003). The section immediately south of the trailer has a massive white lime mudstone with lenses of high energy ooid grainstone and intraclast pebble-ooid grainstone and silicified gastropod conchs and black chert in the 40-59 m interval of the measured section. This limestone apparently corresponds to Flower's (1964) "Steves Farm Limestone" member of the Little Falls Formation (="Whitehall Formation," abandoned). A second 0.6 m-thick, massive white limestone with ellesmeroceratoid cephalopods and snails forms the highest exposed carbonate bed on the northern slope of the hill. This second limestone is Flower's (1964) Rathbunville Limestone member of the Little Falls Formation. Following a covered interval (13 m), fine-grained, dolomitic sandstone of the lowest Tribes Hill Formation appears as float near the hill crest.
The restricted marine, mollusk-rich, trilobite-poor Little Falls Formation at this section yields low diversity conodont faunas. The interval from the "Steves Farm" through Rathbunville Limestones is referable to the uppermost Cambrian Clavohamulus elongatus through C. hintzei Subzones of the Cordylodus proavus Zone s.s. (Landing et al. (2003).
With use of the lowest occurrence of the conodont lapetognathus fluctivagus Nicoll to define the base of the Ordovician globally (Webby, 1998), the conodonts from the Little Falls at Steves Farm and elsewhere in the Lake Champlain lowlands show that the Little Falls Formation does not extend into the lowest Ordovician. Earlier reports that assigned trilobite faunas with Missisquoia from the lower Little Falls Formation to the Ordovician (Taylor and Halley, 1974) or used the "Ordovician aspect" of molluscan faunas from the "Steves Farm" and Rathbunville Limestones (Flower, 1964; Fisher, 1984) are now superseded by this new global standard for definition of the base of the Ordovician. The conodonts illustrate that the ellesmeroceratoids from the "Steves Farm" and Rathbunville are Upper Cambrian and among the oldest cephalopods known in North America.
Clavohamulus hintzei Subzone (top Cordylodus proavus Zone s.s.) conodonts from the top of the Little Falls in the Lake Champlain lowlands are younger than the C. elongatus Subzone (middle C. proavus Zone s.s.) at the top of the Little Falls in the Mohawk Valley (Landing et al., 1996). In the Mohawk Valley, Mazullo and Fisher (1976) proposed a major hiatus between the Little Falls and Tribes Hill Formations and equated the lower Tribes Hill with the "Steves Farm"-Rathbunville Members of the "Whitehall" and assigned this interval to the lowest Ordovician. Correlated. However, these new conodont-based data show the top of the Little Falls is uppermost Cambrian across the New York Promontory and the sub-Tribes Hill hiatus shortens by only two conodont Subchrons in the Champlain Lowlands. The Cambrian-Ordovician boundary is further discussed at Stop 7.
The eustatic rise equated with deposition of the Tribes Hill Formation is equated with a middle-upper Tremadocian macroscale black mudstone-limestone in the Taconian allochthons of New York and Quebec (Stop 13).

Pass road cuts in "Fort Ann Formation."

STOP 7. COMSTOCK SECTION
COMSTOCK SECTION - BRACKETING THE CAMBRIAN-ORDOVICIAN BOUNDARY AND OVERLYING TRIBES HILL FORMATION. (40 MINUTES).
A stratigraphic section for the Comstock section will be distributed on this field trip; this section is available in Landing et al. (2003, fig. 4).
Stop 7 includes an examination of the roadcut on the north side of Rte 22, then continues along the low section in the shallow stream at the east end of the cut, crosses a long covered interval to the east across the pastures on the Scalaro farm, and then climbs the west slope of the 132 m ( 440 foot) hill to the ridge crest.
The roadcut on Rte 22 begins at the east border of Comstock village and exposes ca. 20 m of medium- to massively-bedded, gently-dipping (10° E), upper Little Falls Formation (="Whitehall") dolostone. In 1998, these sucrosic, locally evaporitic (note gypsum molds) dolostones were sharply overlain by a 30 cm bed of fine- to medium-bedded, brown-weathering, current cross-bedded, dolomitic sandstone of the basal Tribes Hill Formation (="Cutting/Great Meadows Formation") at the east end of the cut. Almost all of this dolomitic sandstone bed was collected as a conodont sample. This dolomitic sandstone-facies (Sprakers Member, ="Winchell Creek Siltstone," abandoned designation) is present as float immediately above the exposed Little Falls, is exposed in the shallow creek above the road cut, and apparently continues east under the covered interval in the pasture behind the Scalero family farmhouse. This facies re-appears at the foot of the ridge on the west side of the pasture at the base of the 132 m ridge, where it is overlain by ca. 1.5 m of gray siltstone with a middle storm-deposited sandtsone (Van Wie Member). Thicker-bedded, increasingly dolomitic sandstones continue upward to a cliff with several horizons of spectacular thrombolites with echinoderm grain- and packstone inter-head matrix (Wolf Hollow Member;= "Kingsbury Limestone"). Higher locally oolitic and overlying structureless (i.e., burrow-homogenized) white limestone (Canyon Road Member; = "Fort Edward" + "Smith Basin" Members of Fisher [1977] and Flower [1968b]) extend to the top of the Tribes Hill. The top of the Tribes Hill Formation forms the crest of the 132 m (440 foot) hill, and has a paleokarst surface with 30 cm of relief overlain infilled by arenaceous dolostone of the "Fort Ann Formation."
Conodonts from the Tribes Hill Formation comprise a low diversity, restricted marine assemblage that persists unchanged through the formation. This Rossodus manitouensis Zone assemblage is referable to the lowest Ordovician (middle-upper Tremadocian-equivalent), and shares no taxa either with the underlying Little Falls Formation or the overlying "Fort Ann Formation." The total replacement of the Upper Cambrian, upper Cordylodus proavus Zone fauna of the top Little Falls Dolostone (Stop 6) at the base of the Tribes Hill reflects the duration of the trans-Laurentian Cambrian-Ordovician boundary hiatus on the east Laurentian shelf. Similarly, the total replacement of the Tribes Hill Formation conodonts by Fauna D conodonts at the base of the "Fort Ann Formation" reflects the duration of the trans-Laurentian, intra-Lower Ordovician hiatus that occurs in the Tremadocian-Arenigian boundary interval.

Cross into Taconic allochthon.

Lower Cambrian Bomoseen Member (argillaceous lithic arenites).
Rocks immediately ahead are probably Lower Cambrian Bomoseen Member (argillaceous lithic arenites).

North Granville Slate.
Drive past slate scrap heap and quarry (on left, West) in North Granville Slate.

STOP 8. BROWNS POND
PALEO-OCEANOGRAPHIC CHANGES IN THE LATE EARLY CAMBRIAN. (15 MINUTES).
This section may be regarded as the "type locality" for the late Early Cambrian Browns Pond dysaerobic interval on the east Laurentian slope (Landing et al., 2002). The Browns Pond dysaerobic interval is recognized throughout the external thrusts of the Taconic allochthon in New York-Vermont and in Taconian Quebec (Figure 3).
This overturned section in the upper Browns Pond Formation shows several meters of dark gray, fine-grained sandstones and slates with thin (decimeter-thick), lensing, sandstone pebble debris flows on the east side of the cut. Small-scale current cross-bedding in the orange-weathering, dolomitic quartz arenites shows that the section is overturned. Although centimeter-wide grazing trails can be found on the top of some of the sandstone beds, burrows are rare, and fine lamination and other primary structures are not disturbed in these rocks. Rare burrows, absence of a shelly fauna, and dark gray (carbonaceous) sediments are all consistent with deposition under dysaerobic conditions (e.g., Sagemann et al., 1991).
The west side of the road cut is dominated by a thick limestone pebble to (rare) boulder clast debris flow with local dark argillaceous matrix. Trilobites, archaeocyathan, mollusk, and calcareous and phosphatic problematica of the lower Elliptocephala asaphoides assemblage (see Landing and Bartowski, 1996) appear in the limestone clasts at this locality (Theokritoff, 1964). None of these clasts show derivation from the carbonate platform, and they include nodular lime mudstones and bedded fossil packstones that are interpreted as allodapic clasts that accumulated as limestone on the upper slope. Pyrite-infilled and phosphate-replaced, calcareous conoidal fossils in these clasts (E. Landing, unpublished data) is consistent with the deposition/lithification of these limestones in a strongly dysaerobic environment that developed on the upper slope (see Landing and Bartowski, 1996; Landing et al.,2002).
Debris-flow conglomerate lenses are common as the cap unit of the Browns Pond Formation, and some have received local stratigraphic names (e.g., Ashley Hill Conglomerate in Landing, 1984). Locally, allodapic fossil hash packstones and decimeter-thick debris lenses are the cap unit of the Browns Pond (e.g., Landing and Bartowski, 1996). In either case, the conglomerates or bedded limestones are directly overlain by a green/green gray or locally purple or red siliciclastic mudstone unit in the Taconic allochthon. This black-green transition in the late Early Cambrian is seen here at Stop 8 immediately above the conglomerate with the abrupt transition into the lower green slates of the Middle Granville Slate. These green slates form low outcrops in the pasture to the west. Stop 10 further illustrates this interval of improved oxygenation on the east Laurentian slope and relates the improved oxygenation to the latest Early Cambrian Hawke Bay regression of Palmer and James (1979). Sea-level still-stand and progradation of the shelf margin or sea-level fall at the onset of the Hawke Bay regression are mechanisms to explain the carbonate clast debris flows and allodapic limestones at the top of Browns Pond dysaerobic facies.

STOP 9. TANNER HILL SYNCLlNE
PALEO-OCEANOGRAPHIC CHANGES IN THE LATEST EARLY CAMBRIAN-MIDDLE ORDOVICIAN. (40 MINUTES).
The superb Tanner Hill section was first described by Rowley et al. (1979). The walk up hill crosses the overturned east limb of a large syncline.
Dolomitic quartz arenites and interbedded, minor dark gray and black siltstones and shales form the lowest part of the section. The coarse, lensing (apparently channelized), conglomeratic sandstones become thinner bedded and finer grained higher in the section, and black shales become dominant. This entire interval up to an abrupt transition into green-gray mudstones of the overlying Deep Kill Formation is the Hatch Hill Formation. The Hatch Hill records a long interval of persistent dysaerobic deposition on the east Laurentian continental slope (terminal Early Cambrian-lowest Ordovician [early Tremadocian] Hatch Hill dysaerobic interval) (see Landing, 1993; Landing et al., 2002). However, the changes in relative oxygenation of slope waters through this long interval are admittedly poorly known at present. Indeed, the development of three important Upper Cambrian "Grand Cycles" on the northeastern Laurentian shelf (Chow and James, 1987) should have been accompanied by sea-level and climate fluctuations recorded by changes in relative oxygenation on the continental slope. One explanation for the lack of any apparent record for changes in relative oxygenation through this interval may be that the transport and deposition of the thick sandstones that characterize the lower Hatch Hill served to erode and obscure much of the record of relative oxygenation that is recorded elsewhere in the Taconic succession by mudstones of various colors. Even with a maximum estimated thickness of 200 m (Rowley et al., 1979). the 20 m.y. interval bracketed by the Hatch Hill Formation indicates that it is a condensed unit that may have a number of unconformities produced during the transport and deposition of thick sand sheets. These sand sheets may have been emplaced primarily during eustatic lows.
Sandstones disappear in the upper Hatch Hill in the Tanner Hill section. The upper Hatch Hill corresponds to the interval of earliest Ordovician dysarobic mudstone deposition that has been termed "Poultney A" (abandoned designation, Landing, 1988b) by Theokritoff (1959; Zen 1964; see Appendix). A sharp transition from the Hatch Hill Formation into the lowest green-gray mudstones of the Deep Kill can be observed in the drainage ditch on the north side of Tanner Hill Road. Limited outcrop of the Deep Kin Formation likely explains the apparent absence of the black mudstone-limestone mesoscale intervals characteristic elsewhere of the formation (Stops 11, 13; Figure 3).
The transition into the lowest synorogenic sediments of the Taconic allochthon is observable just west of the crest of the hill with the appearance of low outcrops of the red, thin (ca. 50 m) Indian River Slate. Fisher (1961) attributed the red color of the Indian River to off-slope transport of lateritic sediments produced on the platform during development of the Knox unconformity. However, the rapid development of bacterial films on sediment grains with their transport into marine regimes regularly leads to grayish sediment color, and an alternative explanation for the color of the Indian River must be found. Landing (1988b) noted the multiple lines of evidence (e.g., occurrence of radiolarian cherts and thin volcanic ashes undiluted by background argillaceous sediment; thorough burrow-homogenization of much of the unit; presence of large, up to 3 cm-wide burrowers) in proposing that the Indian River reflects very slow deposition on an oxygenated sea-floor and long sediment residence time at the sediment-water interface. Cherty red slates are widespread in a number of orogens (e.g., in the Taconian allochthons from New York to western Newfoundland, southern Uplands of Scot1and, Hercynian Rheinisches Schiefergebirge and Hart Mountains, and Jurassic of Japan), where they always underlie green mudstones and higher flysch. These data suggest that red, cherty, oxygenated shales in overthrust belts reflect the following history: passage of a peripheral bulge through passive margin successions; consequent flexural uplift and restriction of sedimentation on the peripheral bulge to slowly deposited pelagic muds, radiolarian cherts. and thin volcanic ashes; and final flexural down-warping and increased rates of deposition as sediment provenance changes to the emergent accretionary prism. The transition into the green-dominated, synorogenic mudstones of the overlying Mount Merino and then into Austin Glen Formation flysch are present in the core of the Tanner Hill syncline.

STOP 10. OXYGENATED MIDDLE GRANVILLE SLATE AND THE HAWKE BAY REGRESSION.
This stop serves to fill in the stratigraphic interval between the rocks of Stops 8 and 9. The large pile of brownish red and minor green slate of Middle Granville Slate on the east side of the road apparently came from the small flooded quarry on the west side of the road. If the light is good, dense Planolites traces can be seen on many of the bedding plane-parallel cleavage surfaces of the reddish slate, where they accompany large grazing traces up to 2 cm in width. The abundance of burrows, which led to the general absence of primary depositional structures in the slate, and its reddish color (produced by traces of ferric iron) are consistent with deposition of the Middle Granville Slate under a more oxygenated slope water mass than the underlying Browns Pond Formation.
Landing et al. (2002) noted that the uppermost Lower Cambrian (upper but not uppermost Olenellus Zone) on the New York and Quebec portions of the eastern Laurentian slope is composed of red and green siliciclastic mudstones. This interval of improved oxygenation of slope waters is equated with the presumed lowered sea-levels, cooler climates, and deeper circulation of oxygenated surface waters during Palmer and James' (1971) Hawke Bay regression. Subsequent sea-level rise and the re-establishment of dysaerobic slope facies (e.g., Hatch Hill Formation and Hatch Hill dysaerobic interval) took place in the terminal Early Cambrian. This Lower-Middle Cambrian boundary interval eustatic rise and climate amelioration led the actual movement of poorly oxygenated slope water onto the continental shelf in such widely separated areas as northern Vermont (Parker Slate) and eastern California (Mule Springs Formation) (Landing and Bartowski, 1996).

STOP 11. RACEVILLE: EARLY MIDDLE ORDOVICIAN DYSAEROBIC INTERVAL.
This vertically dipping section shows a black mudstone-limestone mesoscale alternation in the lower (northern) part of the section and overlying greenish-gray, siliciclastic mudstone without limestone intercalations. Most of the limestone in the lower part of the section consist of fine-grained, argillaceous mudstone that appears to have resulted from diagenetic remobilization of fine-grained carbonate and/or methanogenic production of nodular diagenetic carbonate. Acid dissolution of these limestones yields conodonts (e.g., Histiodella species) indicative of the highest (lower Middle Ordovician, lower Whiterockian) black mudstone-limestone macroscale alternation of the Deep Kill Formation (Landing et al., 1992; Figure 3). The climate maximum, eustatic high, and resultant interval of intensified dysaerobic slope water recorded by the Raceville cut is equated with earliest Ordovician, Whiterockian eustatic rise across Laurentia and with deposition of such units as the Providence Island Dolostone (Figure 2).

Western edge of the Taconian allochthon.
The ridge to the left (east) of Rte 40 is formed of rocks at the western edge of the Taconian allochthon. The SSW trend of Rte 40 from this point to just north of Troy follows the front of the overthrust. Rte 40 meanders back and forth from the edge of the allochthon and onto structurally underlying, Upper Ordovician "Snake Hill" synorogenic flysch and wild flysch, and, locally in the area of Middle Falls, onto Lower and Middle Ordovician shelf carbonates.

STOP 12. LOGAN CYCLES: MESOSCALE CYCLES IN THE MILANKOVITCH BAND.
This is one of the most accessible outcrops that shows the color alternations characteristic of Logan cycles (Landing and Benus, 1985; Landing et al., 1992) in the Lower-lower Middle Ordovician Deep Kill Formation.
The section lies at the leading edge of the Taconic allochthon and is structurally complex. It is overturned and dips ca. 70 degrees east dip. About four Logan cycles are present in this ca. 6 m-thick section. Most of the cycles are BC cycles, and lack the lower black mudstone that appears in many, but not all, Logan cycles (Figure 4). The uppermost part of the section (about at waist level) shows the highly burrowed, buff-weathering C-unit of a Logan cycle sharply overlain by black mudstones with thin sandstones (A-unit) of an overlying Logan cycle. Well preserved Logan cycles that extend through a completely exposed section through the Arenigian Series at Levis, Quebec, appear to have an average period of less than 100 k.y. (Landing et al., 1992). Logan cycles are evident as the dominant depositional motif in the deeper-water, organic-rich, Miocene petroleum source rocks on the southern margin of the Mediterranean (Landing et al., 1992).

Continue to classic outcrop at Schaghticoke gorge on the Hoosic River.
If time is available, we will continue south on Rte 40 to examine the classic outcrop at Schaghticoke gorge on the Hoosic River. This structurally complex section features the black shale and limestone interval that records middle-late Tremadocian dysaerobia on the Laurentian continental slope from eastern New York to easternmost Quebec (Landing et al., 1992).

Note prominent Logan cycles in cut in Deep Kill Formation on left side of road.

The old Bald Mountain limestone quarry
The old Bald Mountain limestone quarry was developed in a giant block of white weathering, massive limestone with a mollusk-dominated fauna (e.g., Cushing and Ruedemann, 1914, p. 75-80). This macrofauna and conodonts (E. Landing. unpub. data) indicate a late Early Ordovician age, and allows interpretation of the Bald Mountain limestone as a massively bedded block of Fort Cassin Formation torn off the outer carbonate platform by Taconian thrusting. The block lies under the Taconic master thrust in early Late Ordovician wildflysch (see Zen, 1967, p. 31, 32, 36).

Early Ordovician carbonate platform units
For the next several miles, Early Ordovician carbonate platform units under the Taconic allochthon are exposed in low cuts along the east side of Rte 40.

This is a good place to bring class trips
Note cut in vertical beds of lower Trenton Group limestones along east side of road and opposite wide dirt pulloff on west side of road. This is a good place to bring class trips, as it provides a panorama that includes the Devonian of the Catskill high peaks to the south, the Proterozoic Grenvillian hills of the southernmost Adirondacks to the west, and the Lower Cambrian of the Taconic allochthon on the hill crest immediately east. In addition, the Upper Ordovician "Knox unconformity" can be investigated at the very north end of the outcrop. The "Knox unconformity" is seen at the contact of relatively proximal, brachiopod-dominated limestones of the Trenton Group with underlying, brown-weathering dolostones (age presently undetermined but possibly early Middle Ordovician and referable to the Providence Island Dolostone).

STOP 13. SCHAGHTICOKE GORGE
LATE TREMADOCIAN DYSAEROBIC INTERVAL. ( 40 MIUTES).
This section appears to be a tectonically isolated slice at the leading edge of the Taconic allochthon. This type section of the lowest Ordovician "Schaghticoke Shale" (term abandoned; Landing, 1988b) features the structural duplication of a thin (ca. 4 m) black mudstone-limestone mesoscale alternation in greenish-gray slaty mudstones. These facies are referable to the lower Deep Kill Formation. Graptolites from the black mudstones (Ruedemann, 1903) are representative of the lower Ordovician Matane faunas known along the leading edges of the Taconian allochthons in Quebec, but conodonts are limited to stratigraphically long-ranging protoconodonts (Landing, 1976; E. Landing unpublished data). The lithology of this middle-upper Tremadocian black shale-limestone macroscale alternation is remarkably similar to the coeval black shale-limestone macroscale alternation exposed for 400 km along the south shore of the St.Lawrence River in Quebec (Landing et al., 1986). This interval of climate maximum and intensified dysaerobia on the Taconic continental slope is equated with the eustatic rise that led to deposition of the Tribes Hill Formation (Stops 5, 6).

B7: DACRYOCONARID BIOEVENTS OF THE ONONDAGA FORMATION AND THE MARCELLUS SUBGROUP, CHERRY VALLEY, NEW YORK
R.H.Lindemann
Department of Geosciences
Skidmore College
Saratoga Springs, New York 12866
The Order Dacryoconarida Fisher 1962 is a taxon of extinct and enigmatic conical microfossils that occur in marine strata of the Devonian System. Although the taxonomic status of the group and its phylogenetic affinities are not universally agreed upon, unequivocal dacryoconarids (dacs) range from upper Lochovian strata to the Frasnian-Famennian mass extinction boundary. Regardless of their biologic affinities, the organisms themselves are understood to have been marine animals, which were predominantly planktic in habit and pelagic in habitat. Within the region of the Old Word Realm, the dacs are represented by more than thirty genera and scores of species from which sets of biozones have been developed rivaling those of the conodonts in their detail and temporal resolution (see Lutke, 1979, 1985; Alberti, 1993; Oliver and Chlupac, 1991; Weddige, 1998). Old World dacs are also diagnostic of several Devonian bioevents, most notably for present purposes the lower Eifelian Chotec Event and the upper Eifelian Kacak-otmari Event(s). Such is not the case in the Devonian of eastern North America where the dacs have been afforded little attention, and the stratigraphic ranges of the few known species are themselves incompletely documented. It is the purpose of this field trip to introduce participants to dacryoconarid genera and bioevents as they are currently understood to be recorded in the Onondaga Formation and the Marcellus Subgroup at Cherry Valley, New York (CVNY).

B7: DACRYOCONARID BIOEVENTS OF THE ONONDAGA FORMATION AND THE MARCELLUS SUBGROUP, CHERRY VALLEY, NEW YORK

STOP 1. DACRYOCONARID EVENT I.
Pull on to the right shoulder of the road just beyond a large road sign above the top of the road cut. Walk down hill to the base of the roadcut.
The base of this outcrop is in the upper Carlisle Center, which is characterized by the ichnofossil Zoophycos. The sub-Onondaga disconformity is marked by abundant phosphate nodules at the Edgecliff-Carlisle Center contact. Much of the Onondaga Formation's Edgecliff Member is a massive to thick-bedded, coraliferous limestone. The upper Edgecliff contains nodules of light-gray chert and the top of the member is easily recognized the red staining of oxidized pyrite nodules. The base of the Dacryoconarid Event I interval is coincident with the Edgecliff-Nedrow disconformity, which is at the contact of this pyriteiferous Edgecliff bed and base of the lowermost argillaceous Nedrow bed. The base of the Nedrow is also the base of the Viriatellina and Tentaculites scalariformis local range zones. At this locality, the fauna of Dacryoconarid Event I appears to be restricted to the two argillaceous intervals of the Nedrow. The top of this interval corresponds to the top of the T. scalariformis local range zone, and to the uppermost known occurrence of Viriatellina in the Onondaga Formation.
The arbitrarily chosen Nedrow-Moorehouse contact is the bed of limestone that contains the lowermost occurrence of dark-gray chert nodules. Immediately above this horizon, there are two argillaceous intervals, which may be correlative with the two uppermost Nedrow "black beds" of central New York. If this correlation proves to be correct, strata referred to the lower Moorehouse at this locality are correlative with the upper Nedrow of central New York and with the Chotec Event. In any case, the base of the Dacryoconarid Event I interval is either at, or in proximity to, the base of the Eifelian Stage.

Cherry Valley and Berne Member of Oatka Creek Fm.
Drive uphill a short distance and park in the highway department staging area. The Cherry Valley Member and the lower beds of the Berne Member of the Oatka Creek Formation are exposed in the road cut immediately above the staging area. As it is more convenient to examine and collect from this interval here than it is at Stop 3, we will do so.

Talus blocks of the Carlisle Center
Drive in the right hand lane and decelerate immediately after passing beneath the old railroad overpass. If the fates have been kind, there will be large talus blocks of the of the Carlisle Center, which may afford excellent bedding plane exposures of Zoophycos. If this is the case, we will stop briefly for photographs. Please do not damage, or collect from, these large blocks. There are plenty of small blocks to be had that do not require the use of a hammer.

STOP 2. DACRYOCONARID EVENT II
The 2 m exposure of massive grainstone and packstone on the south side of Rt 20 is the easternmost exposure of the Seneca Limestone. The Tioga B Bentonite occurs in the reentrant at the base of the exposure. The low exposure of packstone on the north side of Rt. 20 is in the uppermost Moorehouse. At this locality, both members were deposited toward the culmination of a regressive phase, the beginning of which is recorded in the Moorehouse at Stop I. The top of the Seneca is an erosional truncation surface. At other localities in central and eastern New York, this surface is a classic Devonian bone bed.
The base of the Bakoven Shale is not exposed here, though it has been excavated in a stream gully within sight of this outcrop and is (sometimes) exposed nearby at Optional Stop A. The Bakoven was deposited during a eustatic flooding phase which appears to have breached the transcontinental arch, permitting the immigration of dacs from the equatorial Old World Realm, of what is now Arctic Canada, into the Eastern Americas Realm. Striatostyliolina was the first new dac to access the northern Appalachian Basin during this highstand. Its first occurrence is coincident with the base of the Bakoven Shale and is the diagnostic of Dacryoconarid Event II.

STOP 3. DACRYOCONARID EVENTS III AND IV.
The base of the Chestnut Street road cut is in the upper concretionary part of the Bakoven Shale. The upper part of the outcrop, which is visible from the road, is in the Berne Member of Oatka Creek Formation. The massive limestone bed between them is the Cherry Valley (Fig. 1). The Cherry Valley immediately overlies a centimeter-scale bed of shale, which overlies a 10-25 cm bed of limestone. These are the Lincoln Park and Chestnut Street submembers of the Union Springs' Hurley Member. Dacryoconarid Event III, which is not regarded to be correlative with the Lower Kacak Event of Walliser (1995), occupies the concretionary interval of the Bakoven and the Chestnut Street bed. This is the total known range zone of Costulatostyliolina (Fig. 1) in the northern Appalachian Basin. Similarly, Nowakia also first occurs in the Bakoven and is present in the Chestnut Street bed. Nowakia is absent from the section above this bed until it reappears in the Berne Member of the Oatka Creek Formation 23-34 m above the top of the Cherry Valley. This 11 m thick section of the Berne is interval of Dacryoconarid Event IV. It is now regarded to be correlative with the Upper Kacak Event of Walliser (1995). This part of the section can be accessed either by scrambling up through the very sooty lower Berne above the Cherry Valley or by following a path up hill beneath the power line that intersects the road just above where we turned off of Rt. 20. Intermittent outcrops are present along the south side to this path and the hill side immediately above it. The crest of the section along the powerline is in a fine-grained arenite that lies above the highest occurrence of Nowakia, which marks the top of the Dacryoconarid Event IV interval.

Stop A. The upper Onondaga Fm and lower Bakoven Shale
The upper Onondaga Formation and lower Bakoven Shale are exposed in a ditch/ephemeral stream on the west side of the road. If the weather has been dry, it is possible to walk up through the uppermost limestone beds and locate the Onondaga-Bakoven contact. The lowermost shale bed is Daryoconarid Event I, the first occurrence of iStriatostyliolina. Watch your trip leader carefully for the optimal place to pull on to the right shoulder of the road and park.

Stop B. Union Springs Formation
This stop is a road cut in the Union Springs Formation on Camp Pinnacle Road south of John Boyd Thatcher State Park and above the Helderberg escarpment. The top of the Onondaga Formation is exposed on the flats below the intersection of Beaverdam and Camp Pinnacle Roads. The Cherry Valley Limestone is present in the ditch uphill from the shale exposure. Only Styliolina and Viriatellina are present in the upper section of the Union Springs at this location. Although Costulatostylioina is not present, Nowakia is abundant in the lower, darker part of the section. The form of Nowakia that occurs in this interval is very similar to N. otomari, though it does not appear to be that species. Nonetheless, this part of the outcrop is referred to the Dacryoconarid Event III interval and is regarded to be correlative with the Lower Kacak Event of Walliser (1995).

B8: GLACIAL LAKE ALBANY IN THE CHAMPLAIN VALLEY
G. Gordon Connally
Research Associate
Albany, NY 12230
Donald H. Cadwell
Research and Collections
The purpose of this article is to document and discuss the expansion of Glacial Lake Albany from the northern Hudson Valley into the southern Champlain Valley. The area under consideration extends ±115 km from Crown Point to the Hoosic River. Crown Point is on Lake Champlain at 44° 02' N. The Hoosic River is an eastern tributary to the Hudson River at about 42° 55' N.

B8: GLACIAL LAKE ALBANY IN THE CHAMPLAIN VALLEY

Diamond Point Flats
Diamond Point Flats are on the Rebounded Lake Coveville "water plane"

Bolton Landing, built on a Lake Coveville delta.

Picture Stop; view of southern Lake George

Sabbath Day Point; pit on left in a Lake Coveville delta.

Silver Bay Lake Coveville delta below, on right.
Rebounded Lake Albany, Stable Lake Albany, and Coveville features from here to the Hague.

The Hague

Trout Brook Valley, at left, filled with lake clay

Lake Champlain on the horizon, ahead to the left

Rise up the foreset slope of the Streetroad delta

STOP 1. Heustis Gravel Pit
This stop is in a gravel pit developed in one of the deep kettle holes in the 520 ft Streetroad Delta (p. B.8-7; Crown Point 7.5' Quadrangle) of Rebounded Lake Albany. Obviously this delta was deposited upon or within a large block of stagnating ice. This ice was left behind following the Bridport Readvance. Note the lithologies present and the size, sorting, and degree of abrasion of the clasts. There were no lower erosional or depositional levels on the foreset slope. Thus, the ice margin must have receded before any lower lake levels developed.

STOP 2. Raffety Sand and Gravel pit
This stop is in a recently developed gravel pit developed in an inversion ridge. This ridge, originally mapped as a kame terrace, is made of sediment deposited by glacial Putnam Creek, diverted southward by the ice of the Bridport Readvance (p.B.8-5). The sediment was deposited in a low channel held in by high banks of ice. When the ice left, the sediment became the high ridge above the area where ice had been. Thus, the topography is inverted. The inversion ridge is evidence of deposition from higher parts of an ice-free drainage basin, onto, into, or next to the ice. This deflected sediment is part of the evidence for the Bridport readvance.
The 520 ft, Rebounded Lake Albany level, is shared with the Crown Point Delta (p.B8-7; Crown Point 7.5 Quadrangle). When the ice margin receded, Putnam Creek again flowed eastward into the lakes of the Champlain Valley. Thus, Rebounded Lake Albany continued to exist for some time after northward withdrawal of the ice margin. We will cross lower levels of the Crown Point delta as we proceed to Stop #3.

Cross 440 ft delta flat of Stable Lake Albany

Descend to a 360 ft Coveville delta

Descend a 260 ft Lake Fort Ann delta

Then to a 200 ft Fort Ann delta; exposures of sandy beds on left

Putnam Creek bridge

Lake Greens Corners delta
rise back up to the sandy,160 ft, Lake Greens Corners delta inset into lake clays

Sandy 340ft Coveville delta at left against Bulwagga Mt.

STOP 3. St. Pierre Farm for Core #3
This is a brief stop to view the St. Pierre Farm where Core #3 was retrieved (p. B.8-l4; Port Henry 7.5' Quadrangle). The upper 4-4.5 m is Champlain Sea sediment with marine shells. Below this is at least 11 m of varved lacustrine clay.

STOP 4. Crown Point State Park
Turn right (east) into Crown Point State Park (Port Henry 7.5' Quadrangle). Here we will discuss the results of the coring program (p. B.8-16) and examine the polished, striated limestone bedrock beneath the Crown Point bluffs. After our discussion, this will be our lunch stop.

STOP 5. Champlainside Farm for Core #2
This will be another brief viewing stop to point out the location for Core #2 (p. B.8-15; Bridport 7.5' Quadrangle) on Champlainside Farm. The West Bridport section (p. B.&-Figure 10) is immediately southwest on the shore of Lake Champlain.

Two (or three) tills over striated limestone at lake level.

STOP 6. Type Locality for Bridport Readvance
This will be another brief viewing stop (Bridport 7.5' Quadrangle) to see the location of the former gravel pit that was the type locality for the Bridport Readvance (p. B.8-4 and Figure 4)

Drive up alluvial fan surface graded to 420 ft Lake Coveville

Climb to 540 ft Rebounded Lake Albany delta

Ice-contact inversion ridge on left
Ice-contact inversion ridge on left; marks channel that carried Middlebury River waters south to the delta; ice must have blocked westward drainage

Abandoned gravel pit to left
Abandoned gravel pit to left; in 1965 showed ice-contact deposits at proximal end of Lake Albany delta as a mirror image of the Streetroad/Crown Point deltas on the west side of the Champlain Valley

Descend the foreset slope of the East Middlebury delta.
Abandoned pits on left and right displayed the foreset slope in 1965; mirror image of foreset slope of the Streetroad delta of New York

Left would take you to 460 ft Stable Lake Albany delta
A left turn would have taken you into Brandon and the 460 ft Stable Lake Albany delta of the Neshobe river.

340 ft Stable Lake Albany delta

STOP 7. Bishops Commercial Sand and Gravel
At this stop (Thorn Hill 7.5' Quadrangle) we will examine the bottomset sands of the 280 ft. Fort Ann delta of the Hubbardton River (p. B.8-6) for signs of varves. Then we will climb up to the pit and look for ice-contact structures (and the toe of a Taconian thrust fault).

380 ft Stable Lake Albany surface of the Fair Haven delta

440 ft Rebounded Lake Albany leve
Climb to 440ft Rebounded Lake Albany level of Fair Haven delta of the Castleton River

STOP 8. 440 and 400 ft levels of Fair Haven delta
This will be a brief drive-by stop (Poultney and Thorn Hill 7.5' Quadrangles) on the 440 and 400 ft levels of the Fair Haven delta (p. B.8-6), unless a recent exposure is available.

Slate Quarry on left

To the left is the Lake Coveville delta
To the left is the Lake Coveville delta, to the right is the Poultney River. Large logs in the alluvium date at 5,580 yrs BP.

Cross 450 ft delta of Rebounded Lake Albany.

Enter the Hudson River Gorge

STOP 9. Type locality for the Luzerne Readvance
This is the type locality for the Luzerne Readvance (p. B.8-2 and Figure 3; Glens Falls 7.5' Quadrangle). Compare the texture, color, and clast lithologies between the lower grayish-black till and the overlying readvance till. Many times in the past it has been possible to observe a stone line at the base of the readvance till. This has been interpreted as a thin layer of lodgement facies beneath the subglacial melt out till.

460 ft Stable Lake Albany delta
Ascend to the 460 ft Stable Lake Albany delta of the Hudson ruver (Glens Falls delta)

B9: INFLUENCE OF GEOLOGY ON THE ACIDIFICATION STATUS OF ADIRONDACK LAKES AND STREAMS
Richard April
Department of Geology
Colgate University
Hamilton, NY 13346
Robert Newton
Smith College
Northampton, MA 01063
The effects of acid deposition on surface water quality in the Adirondack Mountains have been studied intensively since the 1970's. Findings show that the acidification status of streams and lakes is strongly influenced by the geology and hydrology of watersheds. Deposition input quantity and quality, the mineralogy and depth of surficial materials, the hydrological properties of soils, groundwater flow paths, wetland processes, snowmelt, etc., all contribute to the final chemical composition of surface waters. This field trip will explore the processes and conditions that influence the acid-base status of surface waters as we visit lake watersheds in the eastern and central Adirondacks that display a wide range of geochemical, geologic and hydrologic characteristics.
As it turns out, for reasons to be described later, most acidic lakes and streams (pH 5) are found in the western and southwestern Adirondacks, and also at high elevation and remote areas of the high peaks region. With the 2002 NYSGA/NEIGC meeting and field trips based out of Lake George, it was not feasible to design a one-day trip to encompass all of these areas; they were simply too far away, or too inaccessible. However, the stops selected for this field trip will illustrate how both geology and hydrology play a pivotal role in determining the acid-base status of surface waters contained within watersheds. Lakes and ponds to be visited include Clear Pond, Harris Lake, Arbutus Lake, and Echo and Little Echo Ponds.

B9: INFLUENCE OF GEOLOGY ON THE ACIDIFICATION STATUS OF ADIRONDACK LAKES AND STREAMS

STOP 1. Clear Pond
Clear Pond is a 73 hectare drainage lake occupying a watershed located just a few miles north of the Blueridge Highway (county route 2). The area of the watershed is 12.06 km2 and the lake surface elevation is 582 m. A number of fish species are present, including brook trout, cutlips minnow, common shiner, bluntnose minnow, white sucker, pumpkinseed, and banded killfish. The watershed sits at the southern edge of the high peaks region and is underlain by bedrock consisting primarily of metanorthosite with some granitic gneiss. Thin section examination of two bedrock samples (CL 102 and CL-108) revealed the first to be a garnet-biotite-quartz-oligoclase gneiss (metapelitic) with 5% garnet, 5-10% red-brown biotite, 40-50% quartz and 40-50% oligoclase. The second, CL-108, was a gabbroic anorthosite with 90% plagioclase (∼An45) and 10% hypersthene.
Areas of thick (3 m) till and stratified drift cover approximately 33% of the watershed, with the remaining area covered by thin (<3 m) glacial deposits and bedrock outcrops. (Some nice till cuts can be seen on the right side of Elk Lake Road as we drive into Clear Pond.) Tills in the watershed contain on average 68% sand, 28% silt, and 4% clay. Quartz and plagioclase are the major minerals present in both the soil and glacial sediments. Vermiculite, mixed layer mica-vermiculite and kaolinite constitute the bulk of the clay mineral fraction. Heavy minerals make up 16.2% of the soil, with major constituents as follows: garnet 36.9%, pyroxene 34.4%, opaques 10.1 %, epidote 8.6%, and hornblende 3.8%.
Comparative values of the chemistry of Clear Pond in 1982-84 and 2002 are shown in Table 1. This is a circum-neutral lake with an ANC of about 100 µeq/L. Compared to other lakes we have studied (Driscoll and Newton, 1985; Table 2), Ca concentrations are relatively high but do not approach values found in some carbonate influenced lakes. More importantly, the high Ca concentrations are coupled with high SiO2 suggesting that the Ca is ultimately derived from silicate minerals rather than carbonates. Although there has been a significant decline in SO4-2 (20 µmol/L) over the past 20 years, ANC has remained relatively constant. The nearly one-third drop in SO4-2 suggests that we are seeing the effect of reduced SO2 emissions over this period of time. The lack of response in ANC is due to a corresponding decrease in base cations. Ca concentrations have declined by about 10% in twenty years and may reflect decreases in the flux of particulate matter to the watershed.
Questions to consider at this stop are:
1) Is Clear Pond a typical Adirondack drainage lake? Does the preponderance of metanorthosite over granitic gneiss influence the chemistry of the lake water?
2) Is it possible to tie the decline in SO4-2 in Clear Pond over the past 20 years to reductions in SO2 emissions? Has sulfate (and nitrate) accumulated in soils over time, and if so, how long will sulfate (and nitrate) continue to leach out of soils into stream and lake water?
3) Does the increase in NO3- reflect the fact that Knox emissions have remained relatively steady over the past 20 years, whereas SO4-2 has declined? Has the proportion of sulfuric and nitric acid in deposition changed over time?
4) Do lower Ca values reflect a decrease in the influx of Ca from particulate matter, or does it suggest lower weathering rates in the watershed? Ca depletion on exchange sites has been measured in the Adirondacks (and the Catskills, Lawrence et al., 1999) and it has been suggested that high acid loadings in the Adirondacks have resulted in significant depletion of base cations in the soils. Because ANC values for most Adirondack lakes have not risen over time, is it possible that base cation depletion in soils will delay recovery of Adirondack lakes?

Cross Hudson River

STOP 2. Harris Lake (and Woodruff Pond)
Harris Lake is a 116 hectare drainage lake occupying a watershed located just north of route 28N. The area of the watershed is 183 km2 and the lake surface elevation is 473 m. Over a dozen fish species inhabit the lake, including northern pike, common shiner, white sucker, brown bullhead, rock bass, small mouth and largemouth bass, and yellow perch. The watershed straddles a thin belt of metasedimentary rocks comprising calcitic and dolomitic marble, calcsilicate rock, quartzite and interlayered gneisses (Geologic Map of New York, NYS Museum). Harris Lake has not been included in any of our recent studies, so we cannot provide more detailed information concerning the geologic and mineralogic characteristics of the bedrock and glacial deposits within the watershed. However, in walking the southern shoreline one observes outcrops of calcsilicate rocks, containing calcium carbonate. In a class mapping exercise in 1990, Colgate students identified outcrops of calcsilicates on the northeastern shoreline, and also found outcrops of granitic gneiss in the northern areas of the watershed and on the easternmost shore of the lake.
Harris Lake is classified as a salt-impacted lake using the ALSC classification. This reflects the proximity of the highway and other anthropogenic influences within this watershed. Throughout the Adirondacks, developed lakes with cottages and their associated septic systems, roads, etc. have relatively high concentrations of Cl-. We stop at Harris Lake to consider the effect of the presence of carbonate-bearing rock on the chemistry of the lake water. As Table 3 shows, Ca values are high, but compared with other Adirondack lakes the values do not seem so unusually high to suggest that the chemistry of Hanis Lake is controlled solely by the dissolution of carbonate minerals. In fact the lake does not have a high enough Ca/Si ratio (2) to qualify as a carbonate controlled lake. Perhaps a better indicator of the impact of carbonate minerals in this watershed is the high ANC of about 200 µeq/L, which is one of the higher values observed in the Adirondack lakes we have studied.
Only Woodruff Pond, a drainage lake that sits a stone's throw to the south of Harris Lake and whose watershed also sits within the calcsilicate - marble belt, has a higher ANC (and Ca concentration; Table 2). Three bedrock samples from Woodruff Pond were thin sectioned and analyzed with the petrographic microscope. The first was an amphibolite with 40% hornblende, 10% hypersthene, 50% plagioclase and 1-2% biotite. The second was a quartz-diopside calcsilicate with 30-40% quartz, 20-30% diopside, 20% phlogopite, 5-10% calcite and a trace of tremolite. The third was a quartzite with 60% quartz, 20% hornblende, 5% garnet, 5% hypersthene, and 2% calcite. Areas of thin (<3 m) till and stratified drift cover most of Woodruff watershed; calcsilicate boulders are common. Quartz, plagioclase, K-spar, hornblende and pyroxene are the major minerals present in both the soil and glacial sediments. Vermiculite, mixed layer mica-vermiculite, kaolinite and talc constitute the bulk of the clay mineral fraction. Heavy minerals make up 21.1 % of the soil, with the major constituents as follows: pyroxene 23.5%, opaques 23.5%, hornblende 20.3%, garnet 11.9%, and epidote 5.8%.
On the basis of the thickness of the surficial deposits one might expect Woodruff Pond to have low ANC. In fact, Woodruff Pond has the highest ANC (410-480 µeq/L) observed of all lakes in the Regional Integrated Lake Watershed Acidification Study (RILWAS; Driscoll and Newton, 1985). The relationship between base flow discharge and sum of the base cations (SBC) for several of the RILWAS lakes is shown in Figure 10. All lakes except Woodruff Pond fall between end members Woods Lake, with a watershed dominated by thin till (average depth 2.3 m), and Panther Lake, with a watershed dominated by thick till (average depth 24.5 m) (April and Newton, 1985). Woodruff Pond has a base flow discharge that is almost as low as Woods Lake, yet it has an ANC of 410 µeq/L and SBC of 600 µeq/L. Woodruff Pond watershed provides a dramatic example of the influence that carbonate minerals have on surface water chemistry. Although the water moves rapidly along shallow flow paths, there is sufficient time for carbonate minerals to react. The dissolution of calcium carbonate is rapid and base cations, in this case Ca, are readily supplied to the drainage water. Hence, flow path is not the major determinant in regulating the acid-base status of Woodruff Pond. Rather, the presence of carbonate in the till, and also in the underlying bedrock, is the chief controlling factor.
Harris Lake watershed probably falls somewhere in between and can be considered a carbonate-influenced (but not controlled) lake. Although carbonate is present in the watershed materials, the entire watershed is not within the calcsilicate-marble belt. Indications are that at least half of the watershed area is underlain by granitic gneiss, and a1though some of the till in the watershed may contain carbonate minerals, much of the glacial cover is likely 'granitic' in composition and devoid of carbonate minerals. Water in contact with carbonate-bearing outcrops along the shore may contribute to increasing lake water Ca concentrations and ANC. However, the relative contribution of this water to the overall chemistry of Harris Lake, although significant, is probably small. We still have much more to do in Harris Lake watershed, but it serves as a good example of an area where surface water chemistry is influenced, somewhat, by the presence of carbonate minerals.
Questions to consider at this stop are:
1) How do we measure the influence of carbonate minerals on lake water chemistry? Is it the carbonate in the bedrock, or the carbonate in the till, or both, that is most important in raising ANC and pH?
2) Does the Ca to Si ratio in surface water give us an indication of the amount of carbonate mineral weathering taking place? That is, can carbonate influenced water be differentiated from non-carbonate influenced water by calculating the molar ratio of Ca2+ to H4SiO4?
3) With ANC values of 200 µeq/L or more, are lakes like Harris and Woodruff essentially insensitive to changes caused by acid deposition?
4) Can we artificially create carbonate-influenced lakes by liming techniques? And is it best to lime the lake, or the watershed?

STOP 3. Marble Outcrop

Marble outcrops along side of Rt 28N

Adirondack Visitor Center -- Lunch Stop

STOP 4. Arbutus Lake - Huntington Wildlife Forest
Arbutus Lake is a 49 hectare drainage lake occupying a watershed located in the Huntington Wildlife Forest. The area of the watershed is 3.75 km2 and the lake surface elevation is 513 m. The watershed is underlain by granitic gneiss, but small outcrops of amphibolite and biotite-rich gneiss have been identified. The watershed has been gauged at the lake outlet since October 1991 with a V-notch weir. Huntington Forest has also participated in the National Atmospheric Deposition Program (NADP) and the National Trends Network (NTN) since Oct. 31, 1978 (see Figures 12 and 13 for trends in sulfate and pH). In May, installation of an EPA CASTNET dry deposition monitoring system was completed. Time permitting, we will take a short hike to view these instrumentation set-ups.
Areas of thin (<3 m) till cover approximately 80-90% of Arbutus Lake watershed; bedrock outcrops are numerous. Tills contain on average 73% sand, 24% silt,and 2% clay. Quartz, K-spar and plagioclase are the major minerals present in both the soil and glacial sediments. Vermiculite, mixed layer mica-vermiculite and kaolinite constitute the bulk of the clay mineral fraction. Heavy minerals make up about 16% of the soil, with major constituents as follows: hornblende 33.6%, opaques (ilmenite and magnetite) 29.5%. pyroxene 20.1 %, garnet 7.2%, and apatite 5.4%.
Comparative values of the chemistry of Arbutus Lake in 1982-84 and 2002 are shown in Table 4. This is a moderately sensitive water body with an ANC that ranges between 60 - 80 µeq/L. Ca concentrations are relatively high, and may reflect the presence of amphibolitic bedrock, and the presence of carbonate-bearing boulders in the till, especially up in the area around Archer Creek. Unlike Clear Pond, the 20 µeq/L reduction in SO4-2 here has resulted in a concomitant rise of 20 µeq/L in ANC.
Arbutus Lake watershed is a good example of a system in which the combination of thin till, short flow paths, and short residence times, which usually gives rise to surface waters with low ANC and high sensitivity, is modified by the presence of minerals in the watershed that weather at moderately fast rates (e.g., hornblende, pyroxenes, and garnet) to provide alkalinity to drainage waters. In June of 1989-1991, Colgate students mapping and collecting water samples in Arbutus watershed measured the ANC of all running inlet streams in the watershed. Most streams had ANC values that were in the range of 20-80 µeq/L, but several streams had ANC's that were as high as 160-180 µeq/L. The highest values were measured following a particularly dry spring, and lowest values were measured following a rainy period of several days. During periods of low precipitation, or drought, base flow becomes a proportionately greater fraction of the drainage water reaching the lake and ANC values in Arbutus Lake can rise to their highest levels. Highest ANC values also seemed to be concentrated in the northern part of the watershed, near or within the Archer Creek subcatchment. We are planning to conduct future field studies in this part of the watershed to better characterize the composition of the bedrock, till, and soils.
Questions to consider at this stop are:
1) How does one measure the contribution of mineral weathering to the ANC of surface waters? Can the relative rates of mineral weathering in soils and till be measured, and if so, how do we measure or calculate the relative contribution of each mineral to ANC?
2) Why did a reduction in SO4-2 result in an increase in ANC here and not at Clear Pond?
3) Several calcsilicate belts run east-west in the vicinity of the Arbutus watershed. Just a few miles to the east we noted the influence these rocks have on the chemistry of Harris Lake and Woodruff Pond. How much calcsilicate float is incorporated in the till of Arbutus watershed, and to what degree does this float influence the composition of soil and till in Arbutus? One clue, of course, is to look at the heavy mineral content of the till and soil, which includes significant quantities of hornblende and pyroxene.
4) A second possibility is to look at the bulk chemistry of till samples from the watershed. For example, XRF analyses of the 2 mm (sand) fraction of 13 till samples from the Big Moose watershed, a watershed underlain by granitic gneiss and containing many acid sensitive lakes, has a mean CaO value of 1.64 wt. %. The mean CaO value for till collected from watersheds underlain by metanorthosite or calcsilicates (including samples from Clear Pond and Woodruff Pond) is 4.13 wt. %. Till from Arbutus Lake watershed has a CaO content of 2.54 wt. %, an intermediate value. Does this indicate that source rocks for the till were both granitic gneiss and calcsilicates, and that the bulk chemistry reflects a mixture of the two? MgO values also suggest the same, as shown below.
Till from watersheds in: | wt. % Cao | Wt.%MgO
Granitic gneiss | 1.64 | 0.72
Metanorthosite/calcsilicates | 4.13 | 2.02
Arbutus Lake | 2.54 | 1.49

STOP 5. Echo and Little Echo Ponds
Little Echo Pond is a 0.8 hectare brown water seepage lake surrounded by thick (∼8 m) peat (sphagnum bog) deposits. Its watershed is estimated to be about 0.11 km2. Echo Pond, a stone's throw to the NNE, is a slightly larger seepage lake but without peat deposits edging the shoreline. The lake surface elevation of both ponds is 481 m. No fish species are present in Little Echo Pond; no fish population data is available for Echo Pond. The two Ponds sit in an outwash area at the eastern edge of the anorthositic massif, in the Fish Creek Ponds recreational area. Thin section examination of one bedrock sample taken in the vicinity (LE-105) revealed a typical Adirondack gabbroic anorthosite with 85% plagioclase (∼An45, 7% magnetite, 3% green clinopyroxene, and 5% garnet. Outwash deposits of quartz, K-spar, and plagioclase are composed of 97% sand and 3% silt. Heavy minerals constitute 12.8% of the sand fraction and include 46.3% opaques, 20.1% hornblende, 10.6% pyroxene, 5.6 % garnet, and 2.3% zircon. What little clay is in the soil consists of vermiculite, mixed layer mica-vermiculite, kaolinite, and possibly a trace of talc. Values of the chemistry of Little Echo Pond in 1982-84 and 2002, and Echo Pond in 2002 are shown in Table 5. What is immediately striking is the difference in water chemistry between these two seepage lakes located not more than 100 m apart. Little Echo is an acidic pond with an ANC that averages about of -50 µeq/L (range from ALSC data is-79,6 to -7.4 µeq/L). The pH is 4.3, the SBC and dissolved Si concentrations are low, and the DOC is high. Concentrations of organic anions are high and are comparable to SO42- (on an equivalence basis) (Driscoll and Newton, 1985). The kettle hole that formed Little Echo Pond and the surrounding peatland has almost completely filled in. There is at least one ephemeral channel that drains the lake-peatland system during periods of high water. This lake is virtually isolated from the groundwater in the surrounding outwash aquifer system due to the extremely low hydraulic conductivity of the peat and bottom sediments, both of which form a nearly impermeable barrier around the lake-edge and bottom. Recharge of lake water must depend primarily on direct precipitation, and from runoff through the upper "active" zone of the peat. Although much of the acidity of the pond can be attributed to atmospheric deposition of H2SO4, it appears that the peat deposits and the sphagnum mat surrounding the pond release substantial quantities of organic acids that contribute to the overall acidity. Sulfate concentrations in both the seepage lakes are much lower than in the drainage lakes both in the 1982 and 2002 data (Table 5 and Table 2). This is due to sulfate reduction occurring in the bottom waters of these lakes. The high DOC and long residence time of waters in seepage lakes provide conditions conducive to sulfate reduction. Measured profiles of dissolved oxygen show anoxic conditions in the deeper water of Little Echo Pond. Sulfate reduction is an ANC producing reaction making these lakes less acidic than they would have been had the water column been fully oxygenated. Silica concentrations are very low in seepage lakes. This is likely due to a combination of a reduced input of groundwater and the removal of silica by the action of diatoms. The long residence time of water in seepage lakes makes it likely that diatoms will remove a substantial amount of silica from the water column. Echo Pond is marginally classified as a "flow-through" seepage lake. There is sufficient groundwater contribution in this class of lake to provide enough ANC to neutralize acidic deposition. However, with a pH of 6.4, an ANC of 40 µeq/L, and relatively low SBC compared with drainage lakes we have studied (Table 2), the chemistry of Echo Pond is still greatly influenced by direct precipitation. Despite the good hydraulic connection between the groundwater and the lake, groundwater inflow is limited. This is likely due to the low hydraulic gradient within the groundwater system. Therefore, the influence of direct precipitation is still relatively high especially during rainy periods and snowmelt. We don't have much time series chemical or hydrologic data from Echo Pond, so, at best, our comments should be taken as suggestions based on work completed at similar seepage lakes. Questions to consider at this stop are: 1) Why doesn't Echo Pond have a surrounding peat layer? 2) Would Little Echo Pond be acidic if there was no acid rain? 3) How important is sulfate reduction in generating alkalinity? 4) Can we determine the relative proportions of groundwater and precipitation entering Little Echo and Echo Ponds?

C1: TIMING AND DEPTH OF INTRUSION OF THE MARCY ANORTHOSITE MASSIF: IMPLICATIONS FROM FIELD RELATIONS, GEOCHRONOLOGY, AND GEOCHEMISTRY AT WOOLEN MILL, JAY COVERED BRIDGE, SPLIT ROCK FALLS, AND THE OAK HILL WOLLASTONITE MINE
Cory C. Clechenko
Department of Geology and Geophysics
University of Wisconsin
Madison, WI 53706
John W. Valley
James McLelland
Geology Department
Colgate University
Hamilton, NY 13346
The depth and timing of intrusion of the Marcy anorthosite massif has critical implications for models of regional Adirondack geological development. Classic localities in the northeastern Adirondacks afford the opportunity to evaluate questions of anorthosite intrusion depth, as well as the timing and relation to granulite facies metamorphism. The first stops, at Split Rock Falls, Woolen Mill, and the Jay covered bridge, allow an examination of anorthositic to gabbroic rocks and magmatic textures and relations. The Split Rock Falls locality exposes a variety of cross cutting and block structures of anorthositic series rocks. At the Woolen Mill locality, clear cross cutting relations of anorthosite and gabbroic lithologies exist. Zircons from these rocks have been recently dated (SHRIMP II). Geochronology indicates intrusion of anorthositic rocks at ca. 1155 Ma. New δO data from zircons provide evidence for the magmatic contamination of the Marcy massif, causing an elevation in δO when compared to "normal" anorthosites worldwide. The data indicate the massif intruded as high δO magmas, and thus fluids did not elevate the δO of the anorthosite during overprinting granulite facies metamorphism. At the Jay covered bridge locality, coarse pyroxene-rich dikes cross cut anorthosite. Recent geochronology yields an age of ca. 1155 Ma for anorthosite here. The final stop, at the Oak Hill wollastonite deposit, exposes field relations of low δO calc-silicate lithologies to presumed marble protoliths and anorthosite. Combined stable isotope, major element chemistry, and trace element chemistry provide significant evidence for formation of the calc silicate lithologies in a complex hydrothermal skarn-forming system adjacent to shallowly intruding anorthosite. Garnet with oscillatory zonation of both major elements and δO from the related, nearby, Willsboro deposit further constrain skarn forming processes.
We plan to show, through field relations, geochemistry, and geochronology that the Marcy anorthosite massif intruded as a high δO crystal mush to depths less than 10 km in the crust at ca. 1155 Ma causing localized contact metamorphism. The anorthosite and contact metamorphic rocks were then subsequently overprinted by the regional metamorphism of the Ottawan orogeny after 1090 Ma.
This field trip guide is written in two parts. The first part includes general background information, including a brief description of anorthositic rocks in the area as well as a more detailed description of the geochemistry and petrogenesis of the Oak Hill and related wollastonite deposits. The second part includes a road log. The first three stops, anorthositic rocks at Split Rock Falls, Woolen Mill, and Jay, include detailed information about the field relations seen there, plus petrologic, geochemical, geochronologic information with interpretations. The last stop, at Oak Hill, includes detailed field descriptions of the rocks analyzed and interpreted in the first part of the guide.

C1: TIMING AND DEPTH OF INTRUSION OF THE MARCY ANORTHOSITE MASSIF: IMPLICATIONS FROM FIELD RELATIONS, GEOCHRONOLOGY, AND GEOCHEMISTRY AT WOOLEN MILL, JAY COVERED BRIDGE, SPLIT ROCK FALLS, AND THE OAK HILL WOLLASTONITE MINE

STOP 1. SPLIT ROCK FALLS LOCALITY. (OPTIONAL)
The roadcut across from the parking area provides evidence for multiple intrusions of anorthositic and gabbroic rock. The dominant rock type is gabbroic anorthosite that encloses altered xenoliths. Subophitic textures are preserved in the more gabbroic xenoliths. Garnetiferous gabbro truncates foliation in some xenoliths and has, itself, a different foliation. Some small xenoliths of anorthosite are elongated and are deformed parallel to the foliation in the gabbroic facies suggesting coeval magmatism. All of the above facies, including the garnetiferous anorthositic gabbro, are disrupted by a more mafic facies similar to Woolen Mill gabbro. Late mafic dikes (phanerozoic?) with well developed slickensides cut all other lithologies.
The outcrop not only gives good evidence for the composite nature of the Marcy anorthosite massif, but it also demonstrates the manner in which these rocks can acquire foliation during magmatism and without the need for regional strain. Numerous other localities exist in which different members of the anorthosite suite locally develop foliations that are crosscut by other anorthosite facies. The fact that these rocks involved are clearly contemporaneous, and that the fabrics are strictly local, provide compelling evidence that the foliations developed during composite magmatism when semi-consolidated blocks and magmas underwent differential movement.

STOP 2. WOOLEN MILL LOCALITY.
The outcrops exposed in road cuts along Rt. 9N and in the stream at the site of the former Woolen Mill afford an excellent opportunity to examine field relations of cross-cutting anorthosite series rocks. A sketch map is included here as FIG 7. Recent SHRIMP II geochronology has been undertaken on zircons from the different rock types exposed here to constrain the age of emplacement of the Marcy anorthosite massif and related mafic rocks (Clechenko et at. 2002). New oxygen isotope data have provided further evidence for the magmatic enrichment of the anorthosite in 18O/16O compared to "normal" magmatic anorthosite, as opposed to enrichment by metamorphic fluids during subsequent metamorphism (Clechenko et a1. 2002).
The roadcut on the south side of Rt. 9N shows anorthosite (east end) intruded by a dark, fine-grained rock, first described by Kemp and Ruedemann (1910) as the "Woolen Mill Gabbro". The Woolen Mill Gabbro is a fine to medium grained clinopyroxene-garnet-oligoclase granulite with magnetite-ilmenite, apatite, and minor K-spar and quartz. The gabbro also contains a population of cm scale andesine xenocrysts that have their apparent source from the anorthosite. These may be easily seen as one walks west along the outcrop from the anorthosite contact. The rock is clearly metamorphosed, as indicated by the presence of significant garnet, but lacks significant deformation features. Zircons obtained from the Woolen Mill Gabbro along the road had cores that exhibited oscillatory zonation, interpreted to be magmatic in origin, and yielded an age of 1154 ±9 Ma. Thick, structureless, U poor overgrowths imprecisely date a metamorphic zircon growth episode at 1008 ±32 Ma.
Cross the road to examine the exposures in the streambed. Starting from the west, where the stream flows out from under the Rt. 9N bridge, on the north shore of the stram, more Woolen Mill Gabbro is exposed, and as one walks east (downstream) on the north side of the stream, it may be seen becoming apparently finer grained near the contact. A weak apparent foliation also exists, defined by plagioclase lathes, and is best seen on the weathered surfaces in the streambed. Plagioclase xenocrysts and small anorthosite xenoliths are found in the gabbro in a number of spots near the contact. At the contact, the gabbro is clearly seen cross cutting anorthosite. The contact is irregular, and cuts across a variety of different anorthositic lithologies, including a coarse, pyroxene rich facies. At some points along the contact, anorthosite has broken off and floats in the gabbro. Dikes and veins of the gabbro extend into the anorthosite at a number of areas near the contact, and may also be seen as one moves further downstream away from the contact. Also apparent are a number of light colored, cm scale thick dikes that originate at the contact and extend into the gabbro. These appear anorthositic, as they contain plagioclase that has the blue gray appearance of plagioclase in the anorthosite, but contain significant quartz. These may be mixed lithologies, of the transitional "Keene Gneiss" type of Miller (1919).
Crossing the stream to south bank immediately below the parking turnout and across the stream from the breached dam, the outcrops are of anorthosite. The anorthosite here has a clearly developed "block structure" where several different types of anorthosite have been undergone brittle fracture and been intruded by a variety of mafic (anorthositic to gabbro) and felsic material. The angular blocks vary in size from 10's of cm to meter scale. The blocks are generally of a coarse anorthosite with large plagioclase phenocrysts and megacrysts. One block at the top of the outcrops has "deformed" anorthosite transitional to "undeformed" anorthosite. The occurrence in a block with a preserved transition indicates a magmatic origin of the foliation. Sub-ophitic pyroxene textures of a magmatic origin are also developed in a number of places in the anorthosite.
Some intruding material is equivalent to the Woolen Mill Gabbro, as mentioned above. This material is found as veining throughout the outcrop. In one location at the western edge of the outcrop and near the streambed, intruding Woolen Mill gabbro cuts across a plagioclase megacryst. At another location near the top of the outcrop, just adjacent to the foliated anorthosite block described above, gabbro intrudes the anorthosite in the shape of a reversed "7" when viewed looking to the south. Blocks of the gabbro are broken, forming a series trailing away form the main vein. Much of the intruding material is of an anorthositic gabbro composition. More than one "generation" of such gabbroic anorthosite material is found, varying in both mafic content and crystal size. Some of the dikes or veins are granitic in nature.
Anorthosite throughout the outcrops contains the characteristic post-metamorphic alteration assemblages of calcite &plumn; chlorite &plumn; sericite that are commonly seen as late-stage, hairline vein fillings or as alteration products of FeMg silicates throughout the Adirondacks (Buddington 1939; Morrison and Valley 1988). Average values of δ18O and δ13C for calcite are 12.6 and -2.2 per mil, respectively, which suggests that the alteration fluids were deep seated in origin and exchanged with igneous as well as metasedimentary rocks. These veins are related to the formation of at least some high density, CO2-rich fluid inclusions and the temperatures of alteration are estimated at 300-500°C (Valley et al. 1990).
The retrograde fluids that have infiltrated the anorthosite to precipitate calcite have not significantly altered its oxygen isotopic composition. Values of Δ(Cc-Plg) range from 0.9 to 6.6, indicating that the isotopic composition of the alteration minerals was controlled primarily by the hydrothermal fluid and that the δ18O of the host rock remains largely unchanged due to low fluid/rock ratios.
Values of δ18O (plag) for the blocks and their host anorthosite at this outcrop range from 8.5 to 9.3‰. In general, the metanorthosites in the NE part of the Marcy massif are somewhat more isotopically heterogeneous than those in the northwestern part of the massif, but they show the same roughly 2.5 per mil enrichment in the δ18O relative to normal anorthosites worldwide (Morrison and Valley 1988).
Travelling further downstream, the blocking texture becomes less pronounced. Populations of zircons were obtained from coarse-grained anorthositic pegmatites between the dam and the foundations downstream, as well as from gabbroic anorthosite taken from the outcrop at the dam. Prismatic zircons with sector and oscillatory zoning (interpreted to be igneous) from one of the anorthositic pegmatites yielded an age of 1151 ± 6 Ma. A subordinate population of zircons from the pegmatitic anorthosite have structureless (in CL) rims that give an age of metamorphic growth at 1012 ± 5 Ma. Upon reaching the foundations of the old mill buildings above some falls, more metagabbro is exposed, and below the falls, the rocks in the streambed and on the banks are nearly all Woolen Mill gabbro.
When all relations exposed here are taken in total, they suggest the intrusion of a plagioclase rich cumulate that was fractured and intruded by more mafic differentiates, as well as a small amount of granitic material. This event took place at ca. 1155 Ma, as determined from igneous zircons from both anorthosite and the metagabbro at this locality. These ages are in agreement with new ages determined from zircons separated from Marcy anorthosite at a number of localities across the Adirondacks, and are also in general agreement with ages determined from zircons from granitic rocks that cross cut anorthosite at a number of localities (Chiarenzelli and McLeHand 1991). The ages lend no credence to assertions of anorthosite intrusion in this area at ca. 1040 Ma (Isachsen et al. 2001).

STOP 3. COVERED BRIDGE AT JAY
(note: covered bridge no longer crosses river, and is set on side of road).
Walk a few tens of feet farther south along the paved road from the parking turnouts. At the sharp bend in the road there is a low, polished outcrop that exposes two exceptional examples of rafts of coarse, blue-gray andesine anorthosite of the Marcy facies. The more southerly of these is enveloped in medium grained gabbro similar to a xenolith bearing, gabbroic facies found on Giant Mtn. The gabbro is, in turn, surrounded by a fine grained anorthosite. The northern raft of Marcy anorthosite lacks the rim of gabbro and is directly surrounded by the white, fine-grained anorthosite. The northern raft also contains giant pyroxene whose edges show subophitic relates with plagioclase.
Within the river there occurs a wide area of water-swept exposures of white, fine grained, and highly altered anorthosite containing a few remnant blue-gray andesines as well as a few blocks of subophitic gabbro. The exposures are disrupted by two types of dikes: 1) late unmetamorphosed (Phanerozoic?) diabase paralleling the river, and 2) irregular, pyroxene-rich dikes and veins that trend mainly N-S and E-W but show other orientations and right angle turns, as well. Some of the dikes are 15-20 cm wide but most fall into the 2-5 cm range. Both sharp and gradational contacts exist. Several of the dikes intrude along zones of mafic mylonite that may be of the same composition as the dikes themselves.
Mineralogically, the dikes consist of coarse, emerald green clinopyroxene and Fe, Ti-oxide. Some dikes also contain small quantities of plagioclase, garnet, apatite and titanite. The apatite-free dikes contain very Mg rich clinopyroxene (XMg ∼0.80) while the apatite bearing dikes have less magnesian clinopyroxene (XMg∼0.65). The key to understanding these dikes is to note that some exhibit comb structure defined by pyroxene and plagioclase crystals growing perpendicular to the dike walls. The texture is diagnostic of growth from a fluid and provides compelling evidence that the dikes were intruded as liquid rich magma. Preliminary experimental work by D. Lindsley (pers. comm. to J.M.) indicates that representative dike material reaches its liquidus at 1200°C (max).
It is suggested that the pyroxene-rich, and sometimes apatite bearing, dikes represent immiscible silicate fractions complementary to magnetite-ilmenite deposits. Emplacements may have occurred when jostling about of essentially congealed plagioclase cumulates resulted in fracturing and the development of large blocks whose shifting provided passageways along which pyroxene-oxide magmas could be filter pressed. Depending on the batch of magma tapped, the intruding material would vary from Mg rich to Fe rich with apatite. Some dikes may have been cumulate-rich.
Dikes of the sort exposed in the river at Jay are common throughout the Marcy anorthosite massif, and their presence indicates a substantial, but small, amount of mafic material- and possibly mafic cumulate in the massif. Invariably these dikes consist of green clinopyroxene and Fe, Ti-oxide, and it is common to find the silicate and oxide phases physically separated within the same vein. It is possible that, where this occurs, it reflects liquid immiscibility operating on late-stage interstitial fractions that are filter-pressed into veins.
Zircons from the clinopyroxene-plagioclase dike contain very low concentrations of uranium and therefor yielded a poorly constrained SHRIMP II age of 1139 &plumn; 89 Ma. Because the zircons show no inherited cores, it is planned to re-date them by single grain TIMS procedures. A pegmatitic gabbroic anorthosite occupying a ductile shear zone in the riverbed yielded abundant zircons that gave a well constrained SHRMP II age of 1155 &plumn; 13 Ma. This age is consistent with ca. 1155 Ma SHRIMP II ages obtained from six other anorthosite samples from across the Marcy massif, including those described at Woolen Mill. These results make it clear that attempts to argue for a young age of ca. 1040 Ma (Isachsen et al. 2001) for the Marcy anorthosite are without merit and are wrong.

STOP 4. OAK HILL WOLLASTONITE DEPOSIT.
(Private property, seek permission at NYCO offices in Willsboro if coming on your own)
The Oak Hill wollastonite deposit provides excellent exposures of interlayered calc silicate skarn and anorthositic lithologies, typical of the low δ18O wollastonite skarns found between Lewis and Willsboro in a belt of rocks that form the northern and northwestern border of the Westport anorthosite dome. The rocks are exposed on an east-facing hillside of Oak Hill, cleared of trees and overburden within the past few years. The rocks occur in nearly parallel layers, dipping into the hillside, as known from drill core and outcrop. The lowest units known are anorthositic rocks of varying composition. Overlying this are the calcsilicate skarns of the wollastonite ore horizon, and above that, interlayered igneous and marble lithologies. This portion of the trip will entail a short walking transect across lithologies exposed on the hillside and in test pits at Oak Hill. This transect is the surface exposure of the drill core cross section shown previously in this field guide as well as in Clechenko (2001). A sketch map is included (FIG 8) showing the position of features described in this road log. Future mining activity at the Oak Hill deposit may alter the appearance of the area, and the features described below and shown on the map may be altered or simply not exist in the future.
The footwall of the layered sequence is anorthosite of the Westport dome. These are not seen in outcrop, but are known from drill core. The anorthosite varies in appearance, generally based on the amount of mafic minerals and also quartz, but is typically of the "Whiteface facies" of foliated white anorthosite with crushed and recrystallized plagioclase. Some samples contain noticeable quartz. Minor garnet is found in these anorthosites.
The structurally lowest outcrop exposures in the section are massive garnetites consisting almost entirely of grandite garnet, and are exposed as knobs on either side of one of the southern access road. The garnetites overlie anorthositic rocks at this locality as known from drill core. Garnet from the massive garnetite has a generally restricted compositional range (when compared to garnet from wollastonite ore) that is grossular rich. The garnetite has a range of textures. Some areas are comprised of small, equant sized garnet grains, while other areas are comprised of large (cm scale) euhedral crystals surrounded by feldspar and quartz. This texture is also observed in garnetite from the Willsboro mine, where some of the euhedral garnets have oscillatory zonation. The texture is interpreted to be growth of euhedral garnets into open void spaces. The clusters of large euhedral garnet behaved as rigid boudins during Ottawan metamorphism and deformation, while surrounding garnets could have been recrystallized and chemically homogenized. The low δ18O nature of the zoned garnet and the demonstrated chemical link between anorthosite and the oscillatory zonation of the garnet indicate that anorthosite intrusion took place at shallow crustal levels. Additional accessory minerals include titanite and clinopyroxene. On the west-side of the western knob, the garnetite is intruded in places by gabbroic anorthosite. Presumably, this represents a late stage anorthositic series magma that intruded the garnetite. The intrusive relation has been armored by garnetite during the major metasomatic wollastonite ore formation event and subsequent deformation.
Structurally overlying the garnetite. and exposed in two pits (one south of the garnetite, one uphill and west), are the low δ18O (garnet as low as -2‰) wollastonite skarn rocks CWo + Gt + Cpx), along with related lithologies, including garnetite (Gt ± Qtz ± Fsp) and garnet pyroxene skarn (Gt ± Cpx ± Ttn ± Ap). As a general rule, the wollastonite skarn rocks (the wollastonite ore) do not contain primary quartz or calcite, indicating their high variance and formation in an open system. The wollastonite skarn runs approximately at 50-70% wollastonite, though that is variable on the order of 10's of centimeters. Garnet from wollastonite ores have a range of composition XAdr= 0.1 to 0.9. Color is a general indicator of garnet composition, with darker garnets containing more Fe, and redder garnets more Al. Garnet δ18O varies with composition, higher XAdr correlates to lower δ18O. The walls of the test pits and numerous blocks lying about afford an excellent opportunity to examine the nature of the wollastonite ores. The ores are a coarse granoblastic combination of wollastonite, grandite garnet, and clinopyroxene. The ore has a weak foliation defined by the tabular wollastonite grains, and a gneissic habit defined by compositional layering of varying modal percentages of wollastonite and the mafic minerals. In the lower test pit (south of the garnetite knobs), typical wollastonite ore may be found. as well as blocks of garnet pyroxene titanite apatite skarn. In the upper test pit (just above the garnetite knob to the west), additional wollastonite ore rocks may be found, as well as a meter thick layer of garnet pyroxene skarn that pinches and swells along the boundary between the wollastonite ores and the overlying anorthosite (Whiteface facies). The anorthosite is typical "Whiteface facies" and is further exposed as one proceeds west above the test pit.
A number of interlayered lithologies, including anorthosite and anorthosite series rocks such as gabbro, marble, and rocks of granitic composition overlie the wollastonite ores. Generally, the interlayering is on the order of 10's of centimeter to meter scale. The anorthositic rocks here are generally similar in appearance to those in the footwall of the deposit that are known from drill core. In places they are more typical Whiteface facies, with occasional plag megacrysts and garnet. Included in one marble layer exposed at the base of the hillside immediately above the upper wollastonite test pit are complexly folded granulite calc-silicate layers. In some places the calc silicates folds and layers are weathering out of the marble, resulting in pieces of rock lying about that give the appearance of ditty dishrags. The complex folding is the result of the contrast in physical properties between the calc silicate layers and the marble. Marbles contain clinopyroxene, titanite, and graphite, and have a measured calcite δ18O of -20‰ for three different samples.
The overall given by the exposures and field relations at Oak Hill are as follows. (1) Skarn formation predates the major deformation and recrystallization event. Low δ18O skarn lithologies are deformed, they exhibit foliation and gneissic layering, and the whole of the deposit is drawn into parallelism with the other lithologies and in accord with the regional grain. These relations are well preserved in the quarry faces at Willsboro. (2) Deformation has served to obscure the original geometry and spatial relations of the hydrothermal system that formed the skarn. (3) The skarns are intimately related to anorthosite. This is true at Oak Hill and throughout the belt surrounding the Westport dome. (4) The fluid flow system that created the skarns was channelized, as marbles near the skarns have not been affected by infiltration of fluids and turned to calc silicates.
When field relations and textures are combined with geochemical analysis, our understanding of the Willsboro-Lewis skarn belt formation is significantly advanced. All results point to formation at shallow crustal levels adjacent to anorthosite (at ca. 1155 Ma) in a complex hydrothermal system, followed by regional deformation and granulite facies metamorphism at temperatures of ∼750° C and pressures of ∼7 kb. The metamorphism was fluid-absent, most likely by partitioning of water into small amounts of anatectic melt. Garnet U-Pb ages of ∼1035 Ma from skarns indicate closure of U diffusion and end of recrystallization at that time.

C2: PRECAMBRIAN GEOLOGY OF THE WHITEHALL AREA, SOUTHEASTERN ADIRONDACKS
Philip R. Whitney
140 CEC
Albany NY 12230
Glenn B. Stracher
Dept. of Geology
East Georgia College
Swainsboro, GA 30401
Timothy Grover
Dept. Of Geology
Castleton State College
Castleton VT 05735
br / At the easternmost edge of the Adirondack Highlands, between Fort Ann and Whitehall (Fort Ann and Whitehall 7.S' USGS sheets) is a tilted (?) fault block of severely deformed, granulite facies metamorphic rocks of Middle Proterozoic age, called the Pinnacle Range by Hills (1965). The gently sloping eastern side of this block is nearly a dip slope in many locations, and the rocks throughout are not far below the Proterozoic - Paleozoic unconformity, which is exposed at Stop 2. The northern end of the block is a zone of intense ductile shear at least several hundred meters thick that we will refer to as the Whitehall Deformation Zone (WDZ).
Only parts of the area have been mapped in detail. The southern half was mapped by Hills (1965), and the northern third by Stracher (1986, 1989). The map of Fisher (1985), at a scale of 1:48,000 is based on Hills' map and reconnaissance mapping by the New York State Geological Survey in 1982 and 1983. A somewhat different picture is presented by the 1:250,000 map of Thompson et al. (1990). Both the Fisher and Thompson maps cover large areas of which the Pinnacle Range is only a small portion; clearly more work is needed to reconcile the differences. Figure 1 shows the general geology of the Pinnacle Range; Figure 2, from Stracher (1992), shows the geology of the northernmost portion, including the WDZ.
The trip will focus on the extraordinary deformational features of the WDZ, as well as on evidence for a retrograde metamorphic event of probable Taconic age. Stop 1 will be a walking traverse in the most strongly deformed part of the WDZ. Stops 2 through 5 will examine road cuts along NY Route 4, from south to north.

C2: PRECAMBRIAN GEOLOGY OF THE WHITEHALL AREA, SOUTHEASTERN ADIRONDACKS

McDonald's closed in August 2015

STOP 1.
Park well off the road wherever space permits. A well-worn trail leaves this dead end road on the R. We will follow this trail W, updip along a dip slope about 0.8 miles to a point just S of Death Rock (Fig. 2). Numerous outcrops in and near the trail are garnet-sillimanite-K feldspar-quartz metapelite's, locally containing calcsilicates, quartzites, and graphitic schists, the "Hague Gneiss" of Alling (1917). Elsewhere, these rocks are host to graphite deposits (Alling's "Dixon Schist") that were worked commercially in the nineteenth century. Foliation and lineation are exceptionally well developed and the rocks exhibit severe grain-size reduction with locally mylonitic textures. The ESE-plunging lineation seen here is typical of much of the southeastern Highlands. Be alert for kinematic indicators. Approaching the height of land S of Death rock, the trail passes outcrops of biotite gneiss with K-feldspar megacrysts, some of which appear to be rotated porphyroclasts. Well exposed within the biotite gneiss to the N of the trail is a large lens of olivine metagabbro. Foliation in the gneiss wraps around the gabbro lens. Later, at stop 4 we will see more such gabbro lenses that provide (given certain assumptions) clear evidence of shear sense. Further north, at Dresden Station in the Putnam 7.5' Quadrangle, similar metapelites with less extreme deformation are crosscut by 1144 ± 7 Ma olivine metagabbro, a key piece of evidence for early (Elzevirian?) deformation and metamorphism in the Adirondacks (McLelland et al. 1988, 1996 ). Our interpretation is that the Whitehall shear zone, of probable Ottawan age, overprints the earlier deformation with resulting loss of the crosscutting relations. After examining the metagabbro and its surroundings, backtrack along the trail to return to the vehicles.

STOP 2.
The outcrop on the east side of Route 4 just north of the intersection exposes the unconformity between the Proterozoic gneisses and the Middle to Upper Cambrian Potsdam Sandstone. Missing: roughly 500 million years of the geologic record and 25-30 km of rock. The Potsdam here consists of coarse arkosic sanstones and quartz-pebble conglomerates, locally with carbonate cement. Good exposures of Potsdam are present in railroad cuts a short distance east and downhill. The dip of the unconformity surface is 10-15° east, roughly parallel to the eastern slope of the Pinnacle Range fault block. The unconformity is exposed again at a location 18.6 miles N along Rte. 22 from the intersection of Rtes. 4 and 22 in Whitehall; there, foliation in the gneisses is perpendicular to layering in the potsdam, and pockets of radioactive conglomerate are present at the unconformity surface. If time permits at the end of the trip this can be included as an optional extra stop for those heading N.
A short distance eastward along the S face of the outcrop, the gneisses are complexly fractured and the fractures are filled with a dark, fine-grained clastic rock rich in ferroan dolomite (Fig. 3), In thin section, abundant shreds of fresh biotite and angular grains of feldspar "float" in a matrix of dolomite. What is the origin of this fracture filling? The clastic grains are too fresh to be the result of weathering, and too quartz-poor to be clastic dikes of Potsdam age. Are they simply the host gneiss pervasively shattered by hydrofracturing caused in turn by tectonic overpressures during overriding of the area by Taconic thrust slices? The western edge of the Giddings Brook slice is less than five miles east of here. Ferroan dolomite is ubiquitous as fracture fillings here and at the next three stops, easily visible due to rusty weathering. In the next cuts to the south on the E side of the highway, folded, biotite-rich mafic layers in gneiss contain thin (sub millimeter) dolomite veinlets parallel to foliation. In cuts directly across from this stop, crosscutting veins contain both dolomite and adularia (low-T K feldspar).
Except for recent weathering of the fracture fillings, there is little evidence for extensive weathering of the unconformity surface, indicating a relatively short interval between erosion of the Proterozoic rocks and deposition of the Potsdam.

Outcrops at the edge of the woods on R are fine-grained white Potsdam ss.

Outcrops on L are extensively fractured granitic gneisses close to a N-S fault.

STOP 3.
Walk N on the E side of Rte. 4 to the roadcuts. The rocks at the south end of the cut on the east are migmatitic, strongly foliated (sillimanite)-garnet-biotite-quartz-two feldspar paragneisses. Present in the outcrop are several thin quartz-microcline pegmatites in various stages of tectonic disintegration. Large microcline crystals survive the tearing-apart process better than quartz, and locally form large porphyroclasts. Rotation of such feldspar porphyroclasts can provide sense-of-shear indicators under favorable circumstances, but the orientation of the rock face is nearly perpendicular to lineation, limiting their usefulness here.
Observe the variable shape of the garnets: some are equigranular, others are elliptical or flattened. Some are inclusion rich while others are nearly devoid of inclusions. Careful study of these variations, combined with microprobe analysis, might yield information on the relative timing of deformation and metamorphism.
From here, continue north (uphill) along the east side of the road; stay away from traffic. Following a gap in the outcrop, the next rocks are strongly foliated and lineated garnet-rich paragneisses with local lenses of calcsilicate rock. Roughly 30 m. north, these are underlain by mafic gneisses that comprise most of the remainder of the cut. These include numerous lenses and pods of calcsilicates and garnet hornblendite. A large, orange inclusion of grossular-diopside-quartz calcsilicate granulite is visible in the mafic gneisses on the opposite (W) side of the road. Near road level on this side, two coarse-grained ultramafic lenses contain the unusual assemblage actinolite-phlogopite-serpentine- talc-chlorite-diaspore. The purplish-red "mineral" is not garnet but a fine-grained aggregate of diaspore and an opaque mineral (retrograded spinel?). The lenses are surrounded by a zone of biotite-rich "blackwall". Two analyses of the ultramafic rock are given in Table 1; note the exceptionally high Al2O3 and MgO. Suggestions as to what these unique rocks are (were?) are welcome.
Near the north end of the cut, still on the east side, two large pods or megaboudins of massive, fine-grained, garnet-rich metagabbro are enclosed in strongly foliated amphibolites. The transition between foliated and unfoliated rock is abrupt. Patches of tourmaline-bearing pegmatite are present at the broken (?) end of one of the pods.

Park as far off the road as possible at the S end of a large, 2-sided cut. Cautiously cross road and walk N along the W side. The first rocks are garnet-rich biotite-quartz-plagioclase paragneisses, overlain by pink, inequigranular, strongly foliated garnet-biotite granitic gneisses ranging in texture from augen gneiss to mylonite and displaying prominent quartz-ribbon lineation. The contact between these two rocks is sharp, but note the spherical, undeformed garnets directly on the contact. Going north. the pink gneiss is followed by a deformed lens of metagabbro, broken near rood level and injected (?) with heterogenous, locally pegmatitic granitic rock. Looking S across the road and roughly parallel to strike, a similarly shaped gabbro is probably a continuation of the same body. Structurally above the metagabbro migmatitic paragneiss reappears. Note the prominent discontinuity in the foliation. Toward the north end of the cut is another gabbro body, also with a continuation on the opposite side of the road. These biotite-ilmenite-hornblende-two pyroxene-garnet-plagioclase metagabbros are fine-grained and massive with relict igneous textures except in immediate proximity to contacts. Although relict olivine is lacking in the gabbros at this location, larger meta gabbros elsewhere in the Pinnacle Range block contain olivine and are compositionally indistinguishable from olivine metagabbros elsewhere in the eastern and central Adirondack Highlands.
Consider the implications of the roughly sigmoidal shape of the metagabbro lenses. Opposite sense of shear is indicated depending on whether emplacement of the gabbro was pre- or syn-tectonic. If the former, the SE-over-NW shear sense is consistent with most (but not all) other kinematic indicators in the area. Cross the road and walk back S to the vehicles.

Migmatitic paragneisses
Outcrops on L are pale gray, migmatitic, graphite-biotite-sillimanite-quartz-2 feldspar migmatitic paragneisses.

South end of the next set of roadcuts.
An operating stone quarry is just behind the cut on the L.

STOP 5. North end of the same cuts.
Park off road on R and cautiously cross. We will start at the north end and walk south on the west side. Beware of high-speed trucks; stay well off the road. The rocks here include many of the lithologies found throughout the Adirondack Highlands, interlayered or tectonically interleaved. Foliation and compositional layering strike consistently NE with a gentle to moderate SE dip; walking S along the cut we will be traversing the section from structurally higher to lower. The northernmost rocks are interlayered marbles and paragneisses. Marble layers and lenses contain rotated inclusions of amphibolite, paragneiss, serpentinite, and calcsilicate granulite as well as coarse fragments of quartz and microcline suggesting tectonically disrupted pegmatites. Some amphibolite inclusions are surrounded by pyroxene-rich reaction haloes. Interiors of marble layers are calcite marble containing diopside, phlogopite, scapolite, quartz, microcline, and graphite. Near contacts with paragneiss, calcite is replaced by dolomite (Fig. 5), diopside is extensively serpentinized, and there is evidence of cataclastic deformation. In thin section, dolomite is in unreacted contact with quartz and microcline, consistent with low-temperature dolomitization superimposed on granulite facies rocks. Small lumps of cherty silica appear in thin section to be void fillings. Contacts between dolomitic marble, rusty brown on weathered surfaces, and the paler calcite marble are relatively sharp and easily seen. Marble-paragneiss contacts are commonly marked by thin graphite films, possibly originating through concentration of graphite through dissolution of marble by solutions channeled along the contacts.
Paragneisses contain the assemblage biotite-garnet-quartz-plagioclase ± sillimanite and abundant fine-grained quartzo-feldspathic leucosomes that may have been anatectic melts or intrusive granite sills, subsequently grain-size reduced during deformation. Near contacts of paragneiss with marble, celadonite and dolomite are locally present, especially in thin, dark layers layers within the paragneiss. Somewhat surprisingly, biotite and other silicates are ordinarily fresh and unaltered.
Continuing south along the cut, garnetiferous quartzo-feldspathic gneiss underlies the paragneiss and marble. The contact is offset by a small normal fault adjacent to a zone of dolomite-cemented breccia. Note the unmetamorphosed diabasic dike just back from the face of the roadcut and roughly parallel to it.
Following a covered interval, outcrop resumes in strongly foliated charnockitic (garnet-two pyroxene-quartz-plagioclase-perthite) gneiss, underlain by a thin dolomitic marble layer containing numerous rotated fragments of various silicate rocks. The marble here is parallel to foliation in the gneisses both above and below it; across the road the it locally truncates foliation in the charnockite. The marble may have acted as tectonic grease along a small thrust fault. The marble in turn is underlain by mafic garnet-hornblende-pyroxene-plagioclase gneiss with minor quartz and K feldspar, similar in most respects to the jotunite associated with many anorthosite bodies.
Immediately beneath the mafic gneiss is a layer of calcite marble and calcsilicate rock containing scapolite, clinopyroxene (Di56), calcium garnet (Gr91), and, locally, wollastonite. The localized occurrence of wollastonite in the marble, and the iron-poor garnet, suggest that the wollastonite here is of contact metamorphic origin, in contrast to the metasomatic wollastonite ores in the northeastern Highlands.
The structurally lowest rocks are a complex sequence of interlayered paragneisses, garnet-sillimanite-quartz-K feldspar metapelite, marble, boudinaged amphibolite, and gabbroic anorthosite gneiss, the latter most clearly seen across the road. Note the aligned, matchstick-size sillimanite crystals present on some foliation surfaces in the metapelite. Prominent fracture surfaces parallel to foliation and compositional layering are coated with chlorite and calcite and display slickensides with an updip sense of movement. Prior to DOT renovation of the roadcut, slight updip offset was visible on a vertical diabase dike that strikes NNE parallel to the road. The orientation of the slickensides is subparallel to the well developed, ca. S50°E granulite facies mineral lineation. This suggests that some foliation surfaces formed during Grenvillian deformation served to localize later, probably Taconic, movement.

C3: METAGABBROS OF SOUTHEASTERN ADIRONDACKS - EVIDENCE FOR SEPARATE SYN-KINEMATIC, AND POST-KINEMATIC SUITES
Barbara A. Fletcher and W. S. F. Kidd
Department of Earth and Atmospheric Sciences
University at Albany
This trip will explore some of the metagabbroic intrusions that dot geological maps of the eastern and central Adirondacks and are generally assumed to have been emplaced contemporaneously with the anorthosite massifs of the highlands. However, not only do most metagabbro intrusions in the southeastern Adirondack region lack the penetrative deformation fabric seen in the anorthosites and the host rocks, but many display chilled margins with sharp contacts that cut the foliation of the host gneisses.
Granoblastic texture and the presence of garnet coronas confirm that these rocks experienced granulite facies metamorphism, but without the strong deformation seen in most other rocks of the Adirondacks. Evidence of strain accommodation is limited to small, localized ductile shear zones well within the intrusive bodies. Could these rocks have been emplaced at shallow crustal levels, as is thought for the anorthosites, then buried to the 20-25 km depth necessary for granulite facies metamorphism, and yet somehow have escaped significant deformation? Stops will include several outcrops of undeformed metagabbros with clearly visible host rock contacts and chilled margins, as well as an example of deformed metagabbro with large, stretched garnet porphyroblasts.

C3: METAGABBROS OF SOUTHEASTERN ADIRONDACKS - EVIDENCE FOR SEPARATE SYN-KINEMATIC, AND POST-KINEMATIC SUITES

STOP 1. INTERCHANGE 21 ROADCUT.
STOP 1. INTERCHANGE 21 ROADCUT. At the N end of the outcrop E of the exit ramp, is gray granitoid gneiss with small migmatitic veins and pods, some of which contain 0.5 cm hornblende crystals. 40/Ar39Ar dating of biotite from this outcrop yielded a cooling age in the range of 750-850 Ma (Heizler and Harrison, 1998). Note that the gneissic foliation dips E (280°/40°E), as does most of the regional fabric. Close examination shows small areas of low-grade alteration, a result of contact with fluids during exhumation.
Carefully, climb to the top of the outcrop and proceed S, watching for the contact between gneiss and metagabbro. Partially buried in sand, it is best exposed at the edge of the outcrop where very fine-grained metagabbro sharply truncates the foliation in the gneiss. Just to the left (S) of the contact, a xenolith of gneiss is enclosed in metagabbro. Another gneiss xenolith is seen on the face of the outcrop (Fig. 2).
Continue to walk S along the top of the outcrop observing the coarsening of grain size in the metagabbro. Over ∼2-3 m the grains become large enough to show clearly the relict igneous texture, i.e., ophitic or subophitic, in which the spaces between the more numerous plagioclase grains are filled by mafic mineral grains. With a hand lens, tiny pale, pink garnets can be seen to "decorate" the plagioclase, evidence that the gabbro was metamorphosed at granulite facies conditions.
Within the coarse-grained metagabbro are several areas that can be seen to have reacted ductilely to localized, intense strain. Determine the sense of shear for each. Are they consistent? Why are they located well within the metagabbro body, rather than nearer the contacts?
Walking S, metagabbro grain size decreases as the contact with a quartz-poor gneiss is approached. Again, the contact is sharp, truncates foliation in the gneiss (although this gneiss is less strained than that at the N end) and the margin was very obviously chilled against the cooler host rock. A meter or so into the gneiss, an aplite dike can be seen. Look carefully at the gneiss for feldspar porphyroclasts.
Returning N, stop in the middle of the metagabbro and look W at the outcrop across the exit ramp. One of the trip leaders always assumed that the outcrop was a continuation of the metagabbro, until the other checked it out and found that it is is the same gneiss as that at the N end of the E-side outcrop. When the orientation of the contacts are mapped and projected, it is seen that the contacts are not parallel and their projections intersect in the middle of the ramp. In other words, this intrusion is not a planar structure.
The field evidence from this outcrop can be used to set up a geological scenario for the intrusion of the host rocks by the gabbro. The truncated foliation shows that the host rocks were already deformed when the gabbro intruded. The sharp contacts indicate that the host rocks were cool enough to behave brittlely. The chilled margins in the metagabbro indicate that a temperature difference existed, which is not surprising given that the host rocks were brittle and the gabbro was a melt. So far, this is a typical scenario for magmatic intrusion. The puzzle arises when we attempt to account for the facts that the metagabbro shows granulite facies mineral assemblages, as do the gneisses, but unlike the gneisses, the metagabbro is not pervasively deformed.

Park behind motel

Toll Booth: $5.00 per car

Park at Overlook 3 and take a moment to enjoy the view
Park at Overlook 3 and take a moment to enjoy the view of Lake George, which lies in an Ordovician aged graben.

STOP 2. METAGABBRO "FINGERS".
Carefully cross the road to the outcrop which places you in the vicinity of the two "fingers" labeled 1 and 1 in Fig. 2. You will be able to trace out several faces and contacts of these intrusions, showing that these are more planar in shape than the intrusion at the Stop 1. Notice that they are very fine-grained. This is likely because they chilled uniformly due to their thinness. But like the Stop 1 intrusion, the contacts here are very sharp. Close inspection shows truncated foliation.
Two more fingers occur to the left (SE). The second of these, 4 on Fig. 2, is indicated by a rounded shape with dashed lines because its contacts are vague. Take a look at this one and see what you make of its contacts. Why do we call these intrusions "fingers"? Because they occur in proximity, the assumption is that they are finger-like protrusions of a larger (hand-like) intrusive body, that worked their way into an array of fractures in brittle-regime host rock.
Contacts and faces of the fingers can be seen on top of the outcrop as well.

Park in parking lot

STOP 3. SUMMIT OF PROSPECT MT.
Walk north and find footpath to summit. The outcrops here are gneiss, largely lichen-covered. A flight of steps near the summit brings the footpath to an end at the summit roadway. (Fig. 4). Cross the road and head for the top. Then backtrack NW to the outcrop next to the roadway (1 on Fig.4). At the S end of the outcrop, reddish-weathered metagabbro is cut by a pegmatite dike. To the right of this, there is a block of metagabbro that looks at first to be detached. A closer look shows that this block is still in place and is underlain by gneiss. The contact strikes 350° and dips 40° N. Like the fingers seen at the previous stop, this exposure of metagabbro appears to be mostly fine-grained. There is no contact visible at the N end of the outcrop.
Continue around the roadway to the NE. Where it begins to loop back toward the SE, there is a large roadcut on the right (W) side of the roadway (2 on sketch map). About 150 ft. (∼50 m) into a mass of quartzo-feldspathic gneiss, a contact zone with fine-grained metagabbro is seen. Here there is no sharp contact, but rather a zone in which small blocks of gneiss appear as inclusions in the metagabbro, and vice versa. This intrusion shows coarsening away from the contact. Pegmatite intrusions can be seen with retrograde biotite in the adjacent metagabbro. The S contact of metagabbro and gneiss is a rounded, "bullnose" structure striking 352° with a 50°E dip.
Outcrop 3, on the sketch map (Fig A), will not be visited on this trip. It consists of a metagabbro intrusion into gneiss with sharp contacts, chilled margins, and localized shear zones within the metagabbro. Although it has not been established in the field, the three outcrops of metagabbro could be connected in the subsurface. Referring to Fig. 4, it can be seen that the intrusions share a similar orientation within the host rock. Follow the roadway past monotonous walls of gneiss to the parking lot.

Pass large pegmatite that intrudes metadiabase. DO NOT STOP.

Contact of metadiabase and gneiss (covered).

Park here for Stop 4.

STOP 4. I-87N EXIT 23 ROADCUT.
Begin at the N end of this large (∼250 m) body of metagabbro. The N contact with the host rock is not exposed. The first 100 ft. (∼30 m) shows fine-grained metagabbros enclosed in more coarse-grained rock, suggesting the coarser rock was back-intruded by the finer-grained rock. At about 100 ft. (30 m), the metagabbro is well-foliated for about 30 ft. (10 m). Then a complex zone is seen, in which a 1.5 ft. (0.5 m) -wide pegmatite intruded a 10-ft. (3-m) shear zone in the metagabbro. The zone has also been disrupted by a fault. The shear zone strikes 66°and dips 70° W and has a left lateral shear sense.
Continuing along the outcrop, many pegmatite intrusions of varying sizes are seen. The metagabbro displays subophitic texture and is quite coarse grained. At about 350 ft. (105 m), another shear zone occurs, about 3 ft. (1 m) wide, showing left lateral shear sense. Several more smaller shear zones are seen.
The metagabbro begins fining as the contact with Grenville metasedimentary rocks is approached. The contact is sharp. Look closely at the metasediments. They are strongly foliated, having been isoclinally folded. Metamorphic minerals in these rocks include phlogopite, diopside and garnet.

STOP 5. TRAIL TO INMAN POND.
Bring your lunch along.
The trail enters the woods from the parking lot and follows a stream on the right (N). Bedrock in the trail is quartzo-feldspathic gneiss, with occasional float (loose) blocks of metagabbro and marble/calc-silicate rock. At a little over a half mile (0.8 km) there is a faint trail on the right (N) that leads to the outcrops. For now, pass that by and continue W toward the pond. At 0.7 mile (1.2 km) a clear trail to the right leads to the north shore of Inman Pond. At 1 mile (1.6 km), there is a stone fire ring in a grove near the water. Lunch!
Return to the first side trail, now on the left (N) and turn onto it. Proceed at a slight angle to the left of the trail and continue right up to the outcrops (1 on Fig.6). The metagabbro here is quite different than what was seen at the previous stops. It has a coarser texture, and is studded with garnet porphyroblasts in the 3-cm range, that have amphibole rims. Look for a block that displays perpendicular faces. The garnet porphyroblasts appear fairly equant on one face of outcrop, but on the perpendicular face they are strongly elongated, with felsic pressure shadows.
A half foot (15 cm) -wide felsic layer is seen within the garnet-bearing metagabbro. It is presumed to have come from the nearby gneiss, perhaps by partial melting of the gneiss during metamorphism. Proceed to the E, noting the change in outcrop to gneiss, and continuing until the faint side trail is encountered. Turn left (N) on the trail. Outcrop 1 (Fig.6) is on the right side of the trail about 100 feet (30 m) ahead, depending on where the trail is picked up.
This outcrop displays a contact zone between metagabbro and gneiss. While there are remnants of the sharp contacts formed when the metagabbro intruded, they are severely disrupted and in some places foliation is continuous across both lithologies. A lens-shaped body of metagabbro is seen, as well as small areas of interfingering between the two rock types. There can be no doubt that this metagabbro was deformed along with the gneiss, and must have been emplaced prior to the peak of dynamic metamorphism.
From Outcrop 2, proceed southeast to outcrop 3, crossing a small ravine and scrambling up a small hill. Continue to the SE edge of the hill, where outcrops will be more numerous. Climb down a bit to find a 10 ft (3 m) -high ledge of interlayered marble. gneiss and strongly foliated metagabbro. The bottom layer is marble with small pods of coarse-grained metagabbro. Above that is a layer of gneiss, with an area of mylonitic fabric near the face of the outcrop. Another layer of gneiss is above that, but this layer has a pronounced oblique fabric that is truncated by yet another layer of horizontally layered gneiss. This is capped by marble that contains inclusions of folded gneiss.
While a fascinating outcrop in its own right, the relevance to the metagabbro question is that here the metagabbro has been severely dismembered, included in marble and incorporated into a mylonite. Very different from what was seen at the first 4 Stops of this trip.
From here, proceed down the slope to the S and the trail will be encountered in about 300 ft (100 m). Turn left (E) onto the trail and return to the parking lot

STOP 6. METAGABBRO BOUDINS
This spectacular roadcut provides another example of gabbros that were fully involved in the deformation event that produced the Adirondack gneisses, i.e., that were intruded syn-kinematically. On the east side of the road, two large pods of metagabbro are seen. They are continuous with those seen on the left (W) side of the highway. The formation of boudins is the result of competency contrast during deformation, usually folding of a more competent layer surrounded by a less competent layer. The competent layer deforms more brittlely than the less competent layer. However the stretched out tails of the metagabbro layers show that they deformed ductilely as strain continued after boudinage.

STOP 7. DRESDEN STATION METAGABBRO
Walk S on Belden Rd. and turn left (E) onto Rt. 22. BE VERY CAUTIOUS. Traffic moves swiftly and the curves in both directions limit the sight distance. Proceed about 300 ft (100 m) and find a sharp contact between metagabbro and gneiss. Here the gneiss is a khondalite, a name given to aluminous gneisses of sedimentary origin. This one contains purplish garnets, some of which are truncated by the metagabbro. Once again, there are small areas of ductile shear in the metagabbro.
The metagabbro continues across the road, but is harder to trace because the contact between gneiss and metagabbro is subparallel to the face of the roadcut, and the metagabbro is thin or absent in some places. However, on top of the W end of the outcrop the intrusion and its contact are clearly seen.

C5: GEOLOGY OF THE MOUNT INDEPENDENCE STATE HISTORIC SITE, ORWELL, VERMONT
Helen N. Mango
Department of Natural Sciences
Castleton State College
Castleton, VT 05735
History
(The following history is condensed from the Mount Independence information brochure.) Mount Independence is a Revolutionary War site. A fort was built here to guard against British attack from Canada because Fort Ticonderoga, across Lake Champlain, offered a poor view to the north. Lake Champlain narrows considerably between Mount Independence in Vermont and Mount Defiance in New York, and it was assumed that whoever controlled the lake here would have a natural advantage over the opposition.
American troops began clearing land on Mount Independence in June of 1776. By autumn, three brigades (12,000 men) had established camps, and had built two batteries and a picket fort. The combined strength of Mount Independence and Fort Ticonderoga caused the British to retreat quickly to Canada in October, 1776. Many American soldiers went home that winter, reducing the force at Mount Independence to 2,500. Many of those remaining fell ill, and a number froze to death during the winter. Not many replacements reached the camp in the spring, so when British troops staged an attack in July, 1777, the Americans had to evacuate Mount Independence. British and German troops remained there until learning of the British surrender at Saratoga in October, 1777. They retreated back to Canada, burning the structures on Mount Independence as they left.
Mount Independence is a National Historic Landmark, containing the remains of batteries, blockhouses, barracks, a hospital, and other historic features. Archaeological work continues and the trails are being modified and improved to incorporate more of the historical significance and natural beauty of the site. (Note: this trail improvement may mean that, in the future, some stops in this field guide may no longer correspond to described trail markers.)
Geologic Setting
The rock formations found at Mount Independence are typical of Cambrian and Ordovician sedimentary deposition on the passive margin that developed on the east coast of ancient North America (Laurentia). The following description of the geologic history and geologic setting of Mount Independence is taken from various sources, including Fisher, 1984; Rankin, et al., 1989; Thompson, 1990; Selleck, 1997; and Isachsen et al., 2000.
Beginning in the Late Proterozoic, rifting of a supercontinent began, leading to the formation of the Iapetus Ocean (forerunner of the present-day Atlantic Ocean) with Laurentia as the landmass to the west. Approximately 600 m.y. ago, sediment, at first of continental origin, began to accumulate on the eastern margin of Laurentia. As rifting continued, sedimentation on the developing passive margin became of shoreline and shallow marine affinity. The earliest of these sedimentary units occurring in western Vermont (Dalton, Cheshire, Dunham, Monkton, and Winooski Formations) are not found at Mount Independence; here, the oldest unit is the medial Late Cambrian Potsdam Formation, a quartz sandstone which unconformably overlies the Proterozoic metamorphic rocks of the Adirondack massif just to the west. (An excellent exposure of this unconformity is described in Selleck, 1997). Overlying the Potsdam Formation is the Late Cambrian Ticonderoga Formation, a sandy dolomite which marks the transition from a shoreline to more of a shallow marine environment. Overlying the Ticonderoga Formation is the Late Cambrian/Early Ordovician Whitehall Formation, which is mostly a massive dolomite, and then the cross-bedded sandstone at the base of the Early Ordovician Great Meadows Formation. The Whitehall and Great Meadows Formations belong to the Beekmantown Group. All formations dip gently to the north, at angles of between 4 and 10°, averaging 6 or 7°.
To the east of Mount Independence are numerous exposures of dolomites, limestones and shales of the overlying Chazy, Black River and Trenton Groups. The thick sequence of Cambrian and Early Ordovician clastic and carbonate rocks indicates that water depth throughout this time was quite consistent. Water depth did not increase substantially until the Middle Ordovician, when this deepening water heralded the onset of the Taconic Orogeny.
Mount Independence is offset from rocks immediately to the east by a high-angle normal fault, which brings the Cambrian Ticonderoga Formation in contact with the Early and Middle Ordovician Providence Island and Middlebury Formations. While there is no outcrop for at least a mile to the east of this high-angle fault, there is no doubt that the rocks to the east are indeed younger than the rocks of Mount Independence. The age of this high-angle fault is debatable; pre-Silurian movement is likely, although there may have been re-activation of the faults in the Mesozoic related to rifting that produced the present day Atlantic Ocean.

STOP 1. POTSDAM FORMATION.
This "stop" begins here and continues down to the water. Typical upper Potsdam sandstone is displayed in the numerous outcrops along this trail. It is in general a pinkish, tan, and gray quartz-cemented quartz sandstone. Fisher (1984 ) terms this upper arenitic unit the Keeseville Member.
Figure 1 is a photomicrograph of a representative sample, showing the well-sorted, well-rounded nature of the quartz grains. Many of the grains are laced with fluid inclusions; the salinity and temperature of homogenization of the fluid inclusions may help suggest the source rock of the quartz sand.
As the trail descends to the Carillon dock at the water's edge, various features of the Potsdam Formation are visible both in outcrop and in the steps, including grain size differences, shaly layers, worm burrows and cross-bedding. Some of these features cause the rock to weather with a moth-eaten appearance. In other beds, pure quartz-cemented quartz sandstone has the appearance of a quartzite. These various characteristics of the Potsdam Formation suggest an established, mostly intertidal sedimentary environment of deposition, with current action producing the cross-bedding, mudflats forming the thin shaly layers, and bioturbation by near-shore/beach-dwelling organisms.

STOP 2. TICONDEROGA FORMATION.
This outcrop is approximately on strike with the Ticonderoga Formation seen at the base of the first staircase on the Southern Defense Trail. The outcrop displays most of the main features of this formation. It is a light- to medium-gray siliceous dolostone, weathering various shades of gray. It contains rounded quartz grains and scattered knots of white calcite and quartz. Some terminated quartz crystals are also present.
Figure 2 is a photomicrograph of a sample of Ticonderoga Formation taken from one of the lower layers. All features indicate that deposition occurred in a shallow, warm, marine environment, near a quartz sand beach. The rounded, frosted quartz sand grains are windblown from the beach into the water. The peloids (rounded dark masses 75 - 100 µm in diameter) are probably aggregates of carbonate mud (Blatt et al., 1980). The oolith (center of picture) indicates warm, carbonate-saturated water and strong, periodic bottom currents, such as those found in a tidal bar or tidal delta environment (Blatt et al., 1980).
Cross-bedding is also found elsewhere in the Ticonderoga Formation, supporting the notion of strong currents. Figure 3 shows cross-bedding in an upper unit of the Ticonderoga Formation, found on the east side of Mount Independence.
Microscopic examination of the quartz content of the Ticonderoga Formation shows a bimodal distribution, with larger, rounded quartz sand grains and smaller, more angular quartz silt. Figure 4 is a photomicrograph of another sample taken near the one shown in Figure 2. Note the subparallel layering of the coarser and finer particles. This bimodal size distribution may represent the interplay of wind and fluvial or storm deposition.
Cross road to north and go into the woods at the end of the split rail fence. Be careful of old strands of rusty wire. Go straight uphill and slightly left (west). The outcrop here is of Potsdam Formation with some excellent cross-bedding, structurally higher than the outcrop of Ticonderoga Formation just seen at Stop 2. Therefore, there is a fault between the two outcrops.
This fault continues to the west, just south of the road leading to the marina (see geologic map, Figure 9). All along the fault, Potsdam Formation is on top of Ticonderoga Formation, although the actual fault surface is not visible anywhere. Because the trace of the fault is approximately parallel to contacts, this appears to be a dip-slip fault, either in keeping with the Ordovician "cross-faults" identified by Hayman and Kidd (2002) or the early Mesozoic faults of Stanley (1980) for other sections of the central and northern Champlain Valley of Vermont. As the road approaches the marina to the west of the entrance to the State Historic Site, the land flattens out, forming a terrace-like surface. This flat area is probably underlain by Ticonderoga Formation, although there is no outcrop between an occurrence of this unit at the water's edge and outcrops at the base of the cliff. The coincidence of this rock type (less resistant than the Potsdam Formation below and above) with the approximate level of the surface of the lake suggests that this flat area is indeed a terrace, carved into the softer rock when the lake surface was 5 or 10 m higher than present. Higher lake levels have occurred for much of the last 13,000 years, beginning with the waning of the Wisconsin stage of the last Ice Age (Hunt, 1980).
Walk back up the road to the entrance to the State Historic Site. Go left (west) of the museum and walk up the gravel path, past a low exposure of Potsdam Formation (on strike with what we just saw in the woods).

STOP 3. TICONDEROGA FORMATION.
Somewhere in the last 50 m or so just traversed is the contact between the Potsdam and Ticonderoga Formations. This outcrop is therefore near the bottom of the Ticonderoga Formation. The flat surface of the outcrop is approximately equal to bedding, and contains a raised ridge of quartz sandstone. Three possibilities for the origin of this sandstone are:
1. clastic dike.
2. sand filling a paleokarst solution cavity
3. sand filling a crack (either a desiccation crack or one tectonically formed)
If either (2) or (3) is correct, this implies a proximal source of sand that could wash into the cavity or crack. While the Ticonderoga Formation does contain quartz sand (as the main rock type of this very outcrop shows upon close examination), it appears to be windblown into the sediment rather than washed in by currents. For scenario (1), the source of the quartz sand would be below the dolomitic sediment. The proximity of the underlying Postdam sands makes this a reasonable prospect, especially if the Ticonderoga sediment was being deposited rapidly on as-yet-unconsolidated Potsdam sand, applying pressure and causing some of the sand to be squeezed upward.

STOP 4. POTSDAM/TICONDEROGA CONTACT.
This outcrop contains the contact between the Potsdam and Ticonderoga Formations. The contact is taken to be the undulating surface where gray quartz-cemented quartz sandstone (containing truncated cross-bedding) is overlain by pinkish-tan coarse-grained sandy dolostone. This suggests a period of erosion of the Potsdam sands before deposition of the dolomitic sediment; however, if the interpretation of the clastic dike at Stop 3 is correct, the Postdam sediment was not yet lithified when it underwent erosion. Figure 5 is a view of this outcrop, with the contact highlighted.
Continue along the White Trail, past White Trail Stop 2 (Foundation) and uphill over ledges of Ticonderoga Formation. Follow the White Trail to the Information Outpost, and continue west along the Red and Blue Trails. Where these trails diverge, follow the Red Trail all the to the end (Red Trail Stop 3). Here there is a lovely view of Mount Defiance west across Lake Champlain, and of Fort Ticonderoga to the north. The large road cuts at the base of Mount Defiance are in Proterozoic metamorphic rocks of the Adirondack massif. Therefore, the high-angle normal fault that bounds the massif on its eastern side lies somewhere under the lake. The boulder of pink granite near the bench is likely from the igneous rock of the massif, transported here by Ice Age ice sheets.
The outcrop just below the overlook is of dolostone containing chert, sandy layers/lenses and laminations. This outcrop is considered to be Whitehall Formation on the basis of the laminations and slightly fetid odor on a freshly broken surface, although the many similarities with the Ticonderoga Formation (dolomitic matrix, sand layers, chert) illustrate the difficulty in placing the contact between the two formations, especially in relatively flat areas where outcrop is discontinuous.

STOP 5. WHITEHALL Fm - WARNER HILL LIMESTONE Mbr
STOP 5. WHITEHALL FORMATION - WARNER HILL LIMESTONE MEMBER. This rock is a thick-bedded, light- to medium-gray limestone or calcic dolostone containing large, rounded, frosted quartz sand grains and having an almost conglomeratic appearance. Weathered surfaces are almost white, and contain wavy raised lines that resemble the stromatolites seen in the Warner Hill Limestone Member of the Whitehall Formation in Whitehall. New York. There are patches and pebbles of dark blue/gray chert, and occurrences of coarse calcite crystals stained yellow-orange by iron oxide. To the east, on the other side of the Blue Trail and in the woods, are numerous discontinuous and moss-covered outcrops of limestone/calcic dolostone with iron oxide staining, occurrences of earthy-looking matrix, patches and stringers of chert, white weathering and variably fetid odor on fresh breaks. Another feature that is taken to be indicative of the Whitehall Formation is the thick to massive bedding, which makes it very difficult to take structural measurements. (In contrast, the Ticonderoga Formation is mostly medium-bedded, with well-defined bedding planes).
Continue north and downhill along the Blue Trail. Note that the trail is stabilized by black slate that has been brought in from elsewhere. The trail flattens out as the land opens up a bit. About 50 - 100 m past the two-log bridge, there is a Blue Trail marker on a tree on the left (west) side of the trail. Directly across the trail and about 5 m into the woods is a very large boulder of almost pure black chert. Welby (1961) discusses the genesis of the chert in the dolostones, concluding that is was formed penecontemporaneously with the dolostone, and transported by currents as semi-solid masses.
The talus slope to the east at the base of the cliff contains boulders of stromatolitic limestone and "two-tone" mottled and laminated dolostone, all indicating that the cliff is made up of Whitehall Formation.

STOP 6. WHITEHALL Fm - SKENE DOLOSTONE Mbr.
The outcrop is of massive dolostone with layers of chert containing nodules of dolostone, iron oxide staining on groundwater seeps, and a dark shaly layer. The dolostone is mostly medium- to dark-gray and finely crystalline, with a slightly fetid odor on a freshly broken surface. The massive bedding is characteristic of this part of the Whitehall Formation.
Continue north along the Blue Trail. Blue Trail Stop 5 (Spring) has steps made of slabs of laminated Whitehall Formation. Continue along to Blue Trail Stop 4 (Quarry). The rock here is a calcareous dolostone (it reacts weakly without scratching to dilute hydrochloric acid) suggesting it is near the top of the Whitehall Formation, as the reappearance of more calcareous layers is characteristic of the upper portions of the formation.

STOP 7. LUNCH.
The "beach" here is a large outcrop of the fairly featureless, massive gray dolostone that typifies a large part of the Whitehall Formation. Coarse calcite is visible in some cracks. A floating bridge once connected Mount Independence and Fort Ticonderoga at this point. There is also a geological connection, because Fort Ticonderoga is in the direction of dip; the stratigraphy is continuous from this stop, under the water, and onto the other shore, as noted by Brainerd and Seely (1890).

STOP 8. GREAT MEADOWS Fm - WINCHELL CREEK Mbr.
The basal unit of the Great Meadows Formation, the Winchell Creek Member, occurs all around the Horseshoe Battery in scattered outcrops on the slopes. To the north. along a now-abandoned section of the Orange Trail, are a few small outcrops of the distinct cross-bedding that exemplifies this unit (Figure 7). Bedding orientation is difficult to ascertain, but it appears that this unit is lying conformably above the Whitehall Formation. Other outcrops contain a sedimentary breccia that is also characteristic of the lowermost Great Meadows Formation (it is described for the equivalent Cutting Formation by Welby, 1961). The fragments in the breccia are laminated and cross-laminated. Figure 8 shows an outcrop of this breccia, with two fragments outlined for clarity. The structure of both the cross-bedding and the breccia show up best on weathered surfaces, and are almost invisible on fresh breaks.
The Winchell Creek Member in its type locality is described by Fisher and Mazzullo (1976) as a cross-bedded quartzo-feldspathic siltstone with thin lenses of sandstone, containing ripple marks, desiccation cracks, and trace fossils. Prolonged weathering accentuates the sedimentary structures (giving the look of "fine grained wood" of Brainerd and Seely, 1890). Fisher and Mazzullo (1976) give a Gasconadian (Early Ordovician) age for this unit, and conclude that the sedimentary environment of deposition was low- to high-energy tidal flats. Therefore, it appears that the sedimentary environment did not vary greatly throughout the deposition of the Ticonderoga, Whitehall and basal Great Meadows Formations. Indeed, Welby (1961) and Fisher and Mazzullo (1976) conclude that this tidal-intertidal environment continued at least until deposition of the limestones of the overlying Fort Cassin Formation (Bascom equivalent).

C6: ACADIAN EXTENSION AROUND THE CHESTER DOME, VERMONT
Paul Karabinos
Department of Geosciences
Williams College
Williamstown, MA 01267
Low-angle normal faults are important both in regions undergoing crustal extension and in mountain belts dominated by crustal shortening. Extension and shortening may be coeval in mountain belts, as has been proposed for the high Himalayas (Hodges et al., 1992) and Tibetan plateau (Molnar and Tapponnier, 1978), or extension may postdate shortening and thickening of continental crust, as has been suggested for the Basin and Range province (Coney, 1979; Sonder et al., 1987). Extension was also an important process in the evolution of ancient mountain chains (e.g. the Alps, Selverstone et al., 1984; Selverstone, 1985).
Although extension appears to be a fundamental process during orogenesis, major Paleozoic normal faults in the Appalachians have not been widely recognized. A notable exception is the Lake Char-Honey Hill fault in southeastern New England, which has been interpreted as an Alleghenian low-angle normal fault (Goldstein, 1989; Getty and Gromet, 1992). Before Goldstein's (1989) breakthrough work showed that the sense of displacement was normal on this fault, it was universally regarded as a thrust. Astonishingly little evidence exists for normal faulting during the Acadian orogeny in the New England Appalachians. Castonguay et al. (2001) presented thermochronological evidence for Silurian to Early Devonian exhumation and extension along the St. Joseph fault in Quebec, and Spear (1992) suggested that the Monroe Line along the New Hampshire-Vermont border might be a normal fault based on contrasting styles of metamorphism of rocks across the contact.
Karabinos (1999, 2000) proposed that faults previously interpreted as Taconic thrusts around the Chester dome in southeastern Vermont (Thompson et al., 1990; Ratcliffe et al., 1997) are really Acadian normal faults. This field trip presents evidence for an extensional shear zone around the Chester dome and explores the tectonic implications of Early Devonian crustal extension. If normal faulting occurred in western New England, an entirely new aspect of the Acadian orogeny will have to be taken into account in tectonic reconstructions, the interpretation of geologic maps, and models of ore genesis within the belt.

C6: ACADIAN EXTENSION AROUND THE CHESTER DOME, VERMONT

STOP 1. BASEMENT-COVER CONTACT, NW MARGIN OF CHESTER DOME.
Gneiss on E side of road, opposite pullout, belongs to the Middle Proterozoic Mount Holly Complex (i.e. basement) and is composed of quartz, plagioclase, biotite, epidote, ± muscovite, ± microcline. It contains abundant highly-deformed pegmatites that distinguish it from plagioclase schist and gneiss of the Late Proterozoic Hoosac Formation. Walk N along road. The first outcrop on the W side of the road belongs to the cover sequence. Before looking at it, head into the woods on the E side of the road to see more of the Mount Holly complex. The contact appears to be buried in the small drainage on the E side of the road. The basement gneisses are highly sheared and the mylonitic fabric dips steeply to the W. Outcrops up the hill contain large-scale C-S fabric and rotated boudins that indicate a tops-to-the-west sense of shear.
The first outcrop on the W side of the road is quartz, muscovite, biotite, chlorite, ± plagioclase, ± garnet schist. Many layers contain graphite. Further N are exposures of quartz-rich gneiss with a distinctive pinstripe texture. I think these rocks belong to the Cram Hill and Moretown Members of the Missisquoi Formation of Doll et al. (1961). Reverse direction and drive N on Route 103.

STOP 2. CAVENDISH GORGE, AGE OF THE CAVENDISH Fms.
Walk from gate up the dirt road to the power station. One of the tonalite samples dated by Ratcliffe et al. (1997) was collected from the outcrop next to the power station. Ratcliffe et al. (1997) argued that the 1.4 Ga tonalite intruded the metasedimentary units of the Cavendish Formation here in the gorge, the type locality of the Cavendish Schist of Richardson (1929), making them older than the tonalite and part of the Mount Holly Complex. I think that the contact between the tonalite and the marble, calc-silicate, quartzite, and schist is tectonic in origin. I collected a sample of quartzite from the gorge
about 50 m upstream from the power station and separated detrital zircons. Single grain Pb-evaporation and SHRIMP ages of 25 grains are all younger than 1.4 Ga (Karabinos et al., 1999), consistent with a tectonic contact between the tonalite and the metasedimentary rocks, but inconsistent with the intrusive contact interpretation. Furthermore, the plagioclase schist here in the gorge does not contain abundant highly-deformed pegmatite like the gneiss we saw in the basement at Stop 1. I correlate the plagioclase-rich schist with the Hoosac Formation and the marble, calc-silicate, and quartzite with the Tyson Formation.

STOP 3. Basement-Cover Contact, NE Margin Of Chester Dome.
Outcrop immediately E of pullout is quartz-rich, biotite, chlorite, ± plagioclase schist with graphite in some layers. Farther E are more outcrops of this lithology and some layers contain highly fractured garnet crystals partially altered to chlorite along the cracks. These rocks are part of the Silurian to Devonian sequence, probably the Northfield or Waits River Formation. As we will see, the intense zone of shearing along the northeastern margin of the Chester dome extends all the way from Middle Proterozoic basement through the pre-Taconic cover sequence and into the Silurian units.
Walk back down the road to the W. Outcrop beneath roots of fallen tree on S side of road is similar to the first outcrop and contains a fine-grained, unmetamorphosed mafic dike, probably related to Mount Ascutney, a Cretaceous intrusion. Outcrop along road is spotty but it is much better in the woods on both sides of the road. Below the schist of the Northfield Formation are highly deformed felsic and mafic gneisses that contain a mylonitic fabric. It is difficult to assign these rocks to a specific unit with certainty, because they are highly deformed and out of context, but they look more like Barnard Gneiss to me than anything else. This is the only rock type belonging to the Late Proterozoic to Ordovician sequence that I have found along this traverse.
Only 0.3 mile W of the pullout is an outcrop of highly sheared augen gneiss. This rock was mapped as Bull Hill Gneiss by Doll et al. (1961) but not all layers contain microcline, which is typical of Bull Hill Gneiss. In thin section it is clear that the eyes in some layers are composed of plagioclase aggregates. Doll et al. (1961) showed Bull Hill Gneiss directly below Hoosac Formation around much of the Chester dome. It is common to find Bull Hill Gneiss at the basement-cover contact in the southern part of the dome, but other feldspar-rich gneisses belonging to the Mount Holly Complex are more typically found at the contact in the northern part of the dome.

Outcrop on corner contains unmetamorpbosed dike.

STOP 4. GNEISSES IN MOUNT HOLLY COMPLEX WITH ACADIAN DIKES.
The tonalitic gneiss seen here is a common lithology in the Mount Holly Complex of the Chester dome. Also, the deformation fabric in this outcrop is typical of that found in the Mount Holly Complex rocks within the core of the dome, away from the basement-cover contact. Late cross-cutting felsic dikes, like those seen here, are found in many outcrops in the dome below the mylonite zone. Ratcliffe et al. (2001) reported a zircon SHRIMP age from a granitic dike near Gassetts of 392 ± 6 Ma.

Road parallels the W flank of the dome
The narrow valley that the road follows S of Grafton parallels the W flank of the dome near the contact between the Hoosac Formation and the Bull Hill Augen Gneiss.

STOP 5. BASEMENT-COVER CONTACT, SE MARGIN OF DOME.
Walk back about 20 m to W to first outcrop; look for signs of blasting dating from construction of the old railway bed parallel to the West River. Rock is quartz, microcline, plagioclase, biotite, epidote gneiss belonging to the Mount Holly Complex. This does not look like typical Bull Hill Gneiss to me, although that is what is shown below the Hoosac Formation by Doll et al. (1961). Walk E along road past pullout. There are some outcrops of plagioclase schist that look like Hoosac Formation. Continue E along road to where the river is just S of the road. River outcrop of mafic schist with some pelitic schist layers. The pelitic schist contains some large garnet crystals that were incompletely altered to chlorite. The extent of retrogression increases significantly to the E along this traverse. This outcrop and others further E contain cleavage surfaces with sprays of biotite after amphibole. There are some outcrops along the road to the E of quartz-rich schist with distinctive pinstripe texture. I would assign the mafic and pelitic schist found in the river outcrops to the Moretown Formation. I see no rocks here that I would assign to the Pinney Hollow, Ottauquechee, or Stowe Formations.

Covered bridge on left, outcrop of quartz-rich gneisses on right.
These rocks could belong to the lower part of the cover sequence (Rosenfeld, 1988) or to the basement. Approximately 0.3 mi. back to the southeast, near the sharp bend in the road, is sheared Bull Hill augen gneiss on the north side of Route 30. The Bull Hill Gneiss intruded the Mount Holly Complex at approximately 960 Ma (Karabinos and Aleinikoff, 1990), after the Grenville orogeny. Therefore, the deformation fabric in this unit is from the Taconic and Acadian orogenies.

STOP 6. TOWNSHEND DAM SPILLWAY, MYLONITE ZONE
SW MARGIN OF CHESTERDOME.
Walk around gate and down dirt road past the equipment sheds for Townshend Dam. Just before the road crosses a small and unappealing stream, walk through the reeds to the large outcrop to your right. Here is highly sheared plagioclase schist and gneiss of the Hoosac Formation. Some dark biotite-rich layers are present. Note the extreme isoclinal folds and boudinage in this outcrop.
If it is not too wet, it is possible to walk up the spillway and see rocks on both sides of the cut. Another approach is to walk back to the road, continue across the bridge about 100 m, and head into the pine trees on the right. Look for the end of the chain link fence and then walk around and past it to the N, toward the edge of the spillway using extreme caution - there are some steep cliffs with a lot of vertical relief. From the top of the spillway it is possible to look to the other side and appreciate the extreme deformation of these rocks. There are abundant folds, boudinage, and truncated layers. The rocks on this side of the spillway are made up of interlayered mafic and pelitic schist, with sprays of biotite after amphibole, typical of the Moretown Formation. Between the rocks of the Hoosac Formation just seen and the Moretown Formation, there are some layers of garnet-rich pelitic schist that could be part of the Pinney Hollow Formation. They contain quartz, plagioclase, muscovite, biotite, chlorite, and garnet up to 3 cm in diameter. Some of the garnets have quartz inclusion trails that indicate a counter-clockwise sense of rotation when viewed to the N. This is consistent with other kinematic indicators that show tops to the W. An interesting feature of this section is that the degree of retrogressive alteration of garnet to chlorite is highly variable. Furthermore, some layers contain graphite, usually as inclusions in plagioclase and garnet. I believe that the graphite was precipitated from infiltrating fluids during metamorphism and that the variable retrogression reflects irregular fluid flow, probably during the normal-sense shearing that preceded the doming phase of deformation.
This is the location of garnet dated by Christensen et al. (1989) using 87Sr/86Sr variations; they found that the garnet grew at approximately 380 Ma over a 10 m.y. interval. Thompson et al. (1993) assigned rocks in the spillway to the Pinney Hollow and Ottauquechee Formations. If this assignment is correct, these units are very thin compared to their outcrop widths on the E flank of the Green Mountain massif.

STOP 7. TOWNSHEND DAM ROADCUT
MORE OF THE MYLONITE ZONE, SW MARGIN OF CHESTERDOME.
Please be very careful looking at this roadcut. It is a dangerous curve and cars drive by very fast. Stay off the pavement!
The most prominent feature of this long roadcut is the strong mylonitic fabric parallel to the compositional layering. There are abundant folds, boudinage, and rotated garnets. However, what impresses me is the extent of disrupted layering throughout the outcrop. The pelitic layers contain quartz. plagioclase, muscovite, biotite, chlorite, and garnet. The mafic layers contain quartz, plagioclase, hornblende, muscovite, biotite, garnet, and epidote.
Rosenfeld et al. (1988) called these rocks Ottauquechee Formation. I would prefer to assign them to the Moretown Formation. The rocks on the W side of the roadcut show the distinctive pinstripe texture of the Moretown Formation. Rocks in the middle and eastern end of the roadcut are admittedly more difficult to assign, but look like Moretown Formation to me based on their lithology. Whatever units these intermediate rocks belong to, there is not much distance between rocks belonging to the Hoosac Formation to the E and rocks that are clearly part of the Moretown Formation to the W. Compared with the typical sequence seen on the E flank of the Green Mountain massif, only several kilometers to the W, the section between the Hoosac and Moretown Formations here is highly attenuated. Outcrops W of here along Route 30 are mapped as Moretown Formation and show the distinctive pinstripe texture. They do not, however, have the strong mylonitic fabric seen in this roadcut.

C7: EURYPTERIDS - CENTRAL-EASTERN NEW YORK FIELDTRIP
S.J.Ciurca, Jr.
Joseph LaRussa
Rochester, New York

C7: EURYPTERIDS - CENTRAL-EASTERN NEW YORK FIELDTRIP

Outcrops on both sides of road.
Important sections revealing a number of Early Devonian stratigraphic units including the Kalkberg and included metabentonites.

STOP 1. Outcrops of Thacher Formation and overlying units.
This section exhibits a wonderful stratigraphic sequence of limestones of various characteristics, but is little different from the type Olney Limestone at Split Rock near Syracuse, New York. These units are within the Early Devonian Erieopterus Biozone. Howellella (occurring in enormous numbers), and Megambonia are characteristic here as they are in the type Olney Limestone at Syracuse. In this region (Cherry Valley), Erieopterus occurs in beds with these typical marine forms and in ostracodallimestone interbeds. Progressively westward, typical marine forms disappear, especially in the dolomitic units (Chrysler Formation of the Syracuse area, Honeoye Falls Dolostone south of Rochester, and the Clansbrassil Formation of Ontario, Canada.

Petrified Creatures "Museum of Natural History"

STOP 2. Litchfield Town Hall Site
While a relatively shallow roadcut, opposite the Litchfield Town Hall, the small exposure of the Phelps Waterlime at the top of the Fiddlers Green Formation here has yielded hundreds of eurypterid specimens. This is not a classic locality as Tollerton (1997) implied. Within Litchfield Township and nearby areas, Clarke & Ruedemann (1912) described that in the 1800s, eurypterids were collected from rock fences and barn foundations. No specimens were obtained from bedrock and the stratigraphy of the eurypterid-bearing beds was poorly known at the time. In the 1960s, the Litchfield Town Hall Site was a glacially polished and striated mass of waterlime that was difficult to examine. A eurypterid carapace appeared on the polished surface and I and some associates began to slowly excavate the upper portions of the Fiddlers Green here. Over the following several years, we extracted hundreds of specimens indicating a fauna identical to that found at the well-known Passage Gulf roadcut. One scorpion has been recovered from this site as well as cephalopods, ostracods, a phyllocarid and several species of eurypterids-all part of the Eurypterus remipes Biozone.
At this stop we will examine the transition from the underlying Victor Member to the overlying Phelps Member of the Fiddlers Green Formation. Various sedimentary structures, including stromatolites and mudcracks, are present and, of course, we will look for eurypterid remains.

STOP 3. Syracuse Formation (Salina Group) on right.
Small section exhibiting typical Syracuse Formation. The Waeringopterus Biozone is found within the lower portions of the Syracuse Formation in this part of the state.

STOP 4. Red shale of the Vernon Fm (Salina Group) on left and right.
Approaching the base of the Salina Group, we see here exposures of the bright red Vernon Formation. In this part of the state, the Hughmilleria Biozone occurs within basal beds of the formation. Pterygotids are also well-known, especially those of the Downing Brook Member.

STOP 5. Roadcut on east side of the highway.
This roadcut reveals one of the most important (eurypterid-bearing) sections in this part of the state. The top of the Fiddlers Green is exposed in a drainage ditch, and upsection, several other Siluro-Devonian units are exposed. Eurypterid horizons begin with the top of the Fiddlers Green Formation (Bertie Group)-the Phelps Waterlime, exposed in the ditch, has yielded fragmentary eurypterid remains and salt hoppers. The Cobleskill Formation is also well-exposed on the east side of the road. The formation consists of 3 members: a lower dolostone, middle limestone member, and an upper dolostone. The dolostones are massive, and the upper dolostone is important for the Eurypterus fauna it contains. The Silurian-Devonian boundary is placed above this unit. Early Devonian sedimentation is revealed by the waterlime beds (Chrysler Formation) in the sequence above the Cobleskill Formation. Erieopterus microphthalmus is characteristic of these waterlimes and also the lower beds of the overlying Thacher Limestone.

STOP 6. Roadcut on E side of the highway, Chittenango Falls Roadcut.
This cut begins with the Akron-Cobleskill. While an occasional stromatoporoid has been found in the Akron-Cobleskill here, note the lack of the fossiliferous limestone, observed at Clockville, in the section exposed here. Overlying Chrysler dolostones are known for the fine celestite and calcite specimens obtained here over the years. Large pieces of eurypterid integument have been observed on some massive chunks of fallen Akron-Cobleskill at this site and at Clockville. Above is a thick sequence of limestones, including the Erieopterus-bearing Thacher Limestone.

Stop 7. Split Rock Quarry
This old abandoned quarry, which we can only hope will be preserved, contains a wonderful section of the Olney Limestone. The Early Devonian Erieopterus microphthalmus occurs abundantly in strata just beneath the Olney Limestone, in the lower beds of the Olney, and in several feet of strata near the top. The occurrence is analogous to the famous Bertie Group eurypterid horizons in that the eurypterid remains occur in windrows (current induced accumulations of animal and plant remains). While pterygotid remains are relatively rare here, they been found within the Erieopterus-bearing horizons in this quarry.

C8: LATE CAMBRIAN (SAUK) CARBONATE FACIES IN HUDSON-MOHAWK VALLEY OF EASTERN NEW YORK STATE
G.M.Friedman
Brooklyn College and Graduate Center of the City University of New York
Brooklyn, NY
Northeastern Science Foundation, Inc.
affiliated with Brooklyn College
Rensselaer Center of Applied Geology
Troy, NY
For correspondence - Northeastern Science Foundation
15 Third Street
P.O. Box 746, Troy,
NY 12181, USA
gmfriedman@iuno.com
GEOLOGIC SETTING
During the Late Cambrian, eastern New York State lay within the tropical zone in the southern hemisphere, in which an enormous tropical carbonate platform was bounded on both north and south (respectively west and east, today) by deep ocean basins. During Cambrian-Ordovician time, most of the North American continent was a shallow epeiric sea, known as the Great American Bank. At its edge terrigenous and carbonate sediment moved by slides, slumps, turbidity currents, mud flows, and sandfalls down a relatively steep slope to oceanic depths coming to rest at the deep-water basin margin (rise) (Sanders and Friedman 1967; Friedman 1972, 1979; Keith and Friedman 1977, 1978; Friedman and Sanders 1978; Friedman et al. 1982). The paleoslope was probably an active hinge line between the continent to the west and the deep ocean to the east, resembling the Jurassic hinge line of the eastern Mediterranean (Friedman et al. 1971). Such hinge lines in the early history of mountain belts are localized by contemporaneous down-to-basin normal faulting (Rodgers 1968, quoting Truempy, 1960). Later thrusting has lifted the deep-water facies into juxtaposition with the shelf facies along hinge-line faults. This displacement was so great that the Cambrian and Early Ordovician deep-water sediments were shifted far west of their original basin margin. Figure 1 reconstructs the paleo geomorphic profile across the Late Cambrian passive margin of North America in New York State. Here, the Sauk Sequence is composed predominantly of carbonate rocks and it outcrops in two discrete provinces separated by the "Rickard's line" (R in Fig. 1), across which the Sauk Sequence abruptly thickens toward the east. The absence of Lower to Middle Cambrian rocks in west-central New York in part produces the abrupt change in thickness across Rickard's line. "Rickard's line" corresponds to the "edge of Middle/Lower Cambrian rocks", (plate 3 of Rickard 1973) which has been named by John E. Sanders in Friedman et al. (1993).
The Sauk Sequence interfingers with the Taconic Sequence. On the shelf the Beekmantown Group includes the Galway Formation and the Wappinger Group includes the Pine Plains Formation (Fig. 1). The Taconic Sequence includes the Hatch Hill Formation (Fig. 1).

C8: LATE CAMBRIAN (SAUK) CARBONATE FACIES IN HUDSON-MOHAWK VALLEY OF EASTERN NEW YORK STATE

STOP 1. Modern Nonmarine Limestones (Travertine)
STOP 1. Saratoga Spa State Park - Modern Nonmarine Limestones (Travertine) This stop can also be reached through the Main Gate of the State Park on Route 9.
Route of Walk
Walk downhill into a small wood with picnic tables towards a stone building. Bicarbonate-charged, saline waters issue from a faucet at the side of the building and pass through a pipe below dirt road, re-emerging on the bank of Geyser Brook. Calcite precipitates on the steep slope of this bank, forming a terrace of travertine. The water is known as Orenda Spring and the terrace as Orenda Terrace; we will get a better view of this terrace later from below. A walk of a few hundred feet along road leads to the Hayes Well Spring at bridge across Geyser Brook. A taste of this water is "rewarding" for its initial effect. Prolonged drinking is not recommended. From Hayes Well Spring you see the Island Spouter "Geyser" a few hundred feet upstream. The water of this fascinating "geyser" is from a well; the water spouts from a small orifice to a height of about 30 feet. This well was drilled about 80 years ago and the large cone of travertine has formed since.
Follow the path along Geyser Brook upstream. In the bank on the left is an exposure of Middle Ordovician Canajoharie Shale, an outer shelf to slope facies. On occasion, graptolites can be found in this shale. Continue to Orenda terrace to study travertine. Note the rippled surface of the travertine and the brown iron-oxide coloration that occurs in streaks. Up to approximately 4 cm of calcite may precipitate annually at the foot of the terrace. Note "caves" and dripstone at far end of terrace and search for calcite-coated twigs and leaves or impressions of leaves in the travertine (Fig. 2). Pisolites occur in abundance on the walkway between travertine cone and bank of brook, spheroids and oolites are likewise present (Fig. 3) (Friedman 1997a,b; Schreiber et al. 1981).
Discussion
Although in the classical studies on limestones, including the authoritative work of Pia (1933), nonmarine limestones are accorded some measure of importance, the literature on modern carbonates neglects nonmarine carbonates. At this stop you'll see an excellent example of travertine, a rock which, like reef-rock, crystallizes in an initially stony condition.
The term travertine is derived from the Italian word, travertino. a corruption of tiburtino, "the stone of Tibur", which is a former name of the locality now called Tivoli (see Sanders and Friedman, 1967, p. 176). The type locality at Tivoli has been classically described by Lyell (1830, p. 207-210) and Cohn (1864). Some authors make a sharp distinction between the terms travertine, tufa and sinter: others use these terms synonymously (Pia, 1933; Gwinner, 1959, Sanders and Friedman, 1967, p. 176).
Travertine in Saratoga Spa Park gathers around the orifice of wells, on terraces from which water descends, or as a cone around a "geyser". Waters enriched in calcium bicarbonate issue from the subsurface, lose their carbon dioxide and insoluble calcite precipitates:
2HCO3- + Ca++ ⇌ 2CaCO3 + CO2
Twigs and leaves of beech, maple, and oak are preserved as they become coated with calcite or leaves form impressions in the travertine. The calcium bicarbonate-enriched waters originate nearly 1,000 feet below the surface in the underlying Cambrian-Ordovician limestones and dolostones, (Friedman 1997a,b). The waters are confined as in an artesian well beneath a thick cover of impervious Canajoharie Shale from which drilling recovers them. In the early years, the springs issued from natural crevices in the rocks, particularly from the prominent MacGregor fault. Later, pits were dug; the present wells flow through pipes set in bore holes.
The composition of the Saratoga mineral waters is unique among waters that precipitate travertine (Back et al. 1995; New York State Department of Health 1959; Young and Putnam 1979). Most waters that make travertine, especially those of the classical areas in Europe, drain areas of karst, and are of low salinity. By contrast, analyses of Saratoga waters give salinities that geologists classify as "brackish" (approximately 11 %). Inspection of tables of analyses (e.g. Kemp, 1912) indicates the closeness of the composition of these waters to that of formation waters. This is especially true of the high concentration of NaCl. As the waters in the subsurface apparently dissolve limestones and dolostones, the concentration of the calcium, magnesium, and bicarbonate is higher than that of many formation waters. As in most formation waters, the sulfate content is low. Although the origin of the mineral waters is controversial (Hewitt, McClennan, and Nilsson, 1965.), this controversy parallels that of the origin of formation waters. The mineral waters are probably formation waters whose salinity has been lowered as a result of mixing with meteoric water.
Newly precipitated calcite from the Orenda Springs gives a strontium isotopic value of 0.716429 (10) 87Sr/86Sr (±2 S.D), a continental crust signature (Fig. 4). If the waters were derived Upper Cambrian formation water, an Upper Cambrian seawater signature would have been obtained (see Fig. 4). The Sr analyses are normalized to sup86/supSr/sup88/supSr= 0.11940. Analyses of NBS 987 averaged 0.710241 (09) (n = 39) during the period of these analyses. Errors on sup87/supSr/sup86/supSr are given as 2 sigma (95%) in the last two digits.
According to Siegel (1996) carbon isotopic analyses likewise hint that the carbon may be expelled from the mantle many kilometers below the fault zone.
The aquifer containing the mineralized water is under carbon-dioxide pore pressures which are estimated to be as high as five atmospheres that spew mineral water and gas several meters high (Siegel 1996). The acidity of the mineral waters depends on how much carbonic acid they contain; pH typically ranges from 5.5 to 6.5. The mineral springs discharge cold water (Sneeringer and Dunn, 1981). The distribution of the wells, a total of about 200, is controlled by the MacGregor Fault and its subsidiary faults. The mineral waters always occur on the eastern (downthrown side) of the fault.
Puzzling is the observation that the precipitate is calcite rather than aragonite. Travertine in other places is usually calcite, but the composition of the responsible mineral waters shows depletion in magnesium. By contrast the mineral waters of Saratoga are enriched in magnesium and because the magnesium ion inhibits the formation of calcite, aragonite should result, but apparently does not. This question deserves study.
Hollocher (2002) notes that the spring waters are saline and carbonated, and derived from the Beekmantown Group carbonate rock aquifer at depths of 100 m in the Spa State Park. All springs and drilled wells occur just to the east of the Saratoga Springs-McGregor fault, a probable Taconian normal fault.
Comparing the strontium isotopic composition of the modern travertine of the Saratoga Spa State Park with the strontium isotopic composition of nearby exposures of Cambrian Beekmantown (Galway) carbonate rock (Fig. 5) shows that the travertine composition overlaps that of continental crust, whereas the Beekmantown Group has the composition of Late Cambrian seawater. The strontium isotopic composition of this travertine reflects extant meteoric freshwater that modified the marine Upper Cambrian Beekmantown carbonate and hence yields a continental crust isotopic ratio (Fig. 5).

STOP 2. Lester Park
This locality is the site of one of the finest domed microbial mats to be seen anywhere preserved in ancient rocks.
It is significant in the history of geology as the area where stromatolites were first described and interpreted (Figs. 5-16). These structures are part of the Hoyt Formation of Late Cambrian (Late Franconian to Early Trempeleauan) age (490-505 Ma).
The stromatolites ("layered stones") were constructed by microscopic cyanobacteria, formerly known as "blue-green algae", on a shallow epeiric sea floor, when New York and the eastern United States lay south of the equator. They represent the activities of organisms that generated the earth's first abundant oxygen. The Hoyt Formation contains trilobites assigned to the Saukia zone of the Late Cambrian (Trempeleauan Stage); it unconformably overlies the Galway or Theresa Formation and underlies the Ordovician Gailor Formation (Fig. 9) (Cushing and Ruedemann 1914; Fisher 1968, 1977, 1980; Flagler 1966; Mazzullo et a1. 1978; Landing 1979; Ludvigsen and Westrop 1983; Sternbach and Friedman 1984; Friedman 1985).
The earliest reference to stromatolites was that of Steele (1825) whose first description of North American oolitic limestone also called attention to the presence of what we recognize today as domal stromatolites. "Great quantities of calcareous concretions of a most singular structure; they are mostly hemispherical but many of them are globular and they vary in size from half an inch to that of two feet in diameter"(Steele 1825, p. 17). Hall (1847, 1883) considered these "singular structures" to be the remains of sea plants. In later descriptions Hall (1883) assigned them genus status and proposed the Linnaean binomial, Cryptozoön proliferum. Hall recognized several kinds of species to designate "a peculiar form and mode of growth". Early descriptions of these stromatolites include those of Cushing and Ruedemann (1914) and Goldring (1937,1938). Burne and Moore (1993) termed these structures "the first formally-named stromatolite".
Kalkowsky (1908) coined and defined the term "stromatolith" 83 years after Steele's discovery and 25 years after Hall assigned the term Cryptozoön. As Walter (1976, p. 1) notes "the original definition is now of historical interest only". He prefers the definition of stromatolites as "organosedimentary structures produced by sediment trapping, binding, and/or precipitation as a result of the growth and metabolic activity of micro-organisms, principally cyanophytes", an unpublished definition proposed by S.M. Awramik and L. Margulis in 1974. Burne and Moore (1987) proposed the term microbialites for these organosedimentary structures and for deposits of laminated internal structure, such as the cabbage-head microbialites, they employ the term stromatolitic microbialites. Jackson (1997) uses the term microbial laminites.
Interestingly, Grabau and Shimer (1909) placed these domed stromatolites with the stromatoporoids; Rothpletz (1916) concurred with this interpretation. In the interim stromatoporoids which formed large reefs in the Silurian and Devonian were ascribed to coelenterates and sponges (particularly sclerosponges). Hall (1883) noted that these domed structures which we now recognize as stromatolites "have long been known under the name of stromatopora, from their general resemblance in form and structure to that fossil", but objected because "a careful examination of the nature of these bodies proves that while having a concentric structure common to stromatopora they have not the regular succession of tubuli characteristic of species of that genus". Only recently have we returned to Grabau and Shimer and Rothpletz in the idea that Paleozoic stromatoporoids represent fossilized stromatolites (Kazmierczak and Kempe 1990).
Cushing and Ruedemann (1914) assigned the name Hoyt Formation to the stratigraphic interval containing the domed stromatolites referred to by them as Cryptozoöns, and Goldring (1938) published an extensive paper on these stromatolites which she considered to be reef builders. The title of her paper speaks for itself: "Algal barrier reefs in the lower Ozarkian of New York with a chapter on the importance of coralline algae as reef builders through the ages". Apparently she considered these microbial structures to be the skeletons of coralline algae. Fisher and Hanson (1951) revised the geology of Saratoga Springs and vicinity, especially the section in which the spectacular stromatolites occur. More recent studies of these domed stromatolites are those by Owen and Friedman (1984) and Friedman (1985, 1987a,b, 1988a,b, 1995, 1997a,b).
Walcott discovered an abundant trilobite fauna and through his publications the Hoyt Limestone became the only fully illustrated and completely described Upper Cambrian trilobite fauna from epeiric carbonates of the eastern United States (Walcott 1879, 1890, 1912). The trilobites are disarticulated but unabraided and transportation appears to have been minimal (Taylor and Halley, 1974). Ludvigsen and Westrop (1983) revised all of the trilobites known to occur in the Hoyt Limestone.
In addition to inarticulate trilobites the fossils present include simple conical conodonts (Landing 1979), and echinoderm debris.
LOCATION AND DISTRIBUTION
Lester Park includes a quarry, named Hoyt Quarry, which serves as geologic reference for the Hoyt Limestone, a rock unit that extends around the southeast side of the adjacent Precambrian Adirondack Mountains. Several LateCambrian-age fossils are found only here and farther south along the Hudson River.
Goldring (1938) recognized the presence of three species of stromatolites (which she referred to as calcareous algae): Cryptozoön proliferum Hall, C. ruedemanni Rothpletz, and C. undulatum Bassler. A glaciated surface at Lester Park exposes the spectacular internal geometry of Cryptozoön proliferum Hall, noted by her as domal stromatolites (Figs. 5-10).
Across the road (west) from Lester Park Rothplatz (1916) described this same "species", which occurs as a bed of stromatolites resting upon oolitic limestone and oolitic limestone overlies it again. This section of stromatolites is composed of dolomite.
GEOLOGIC AND SEDIMENTOLOGICAL SETTING
In places glaciated surfaces expose horizontal sections of domed heads, known as cabbage-head structures (Fig. 5a-i). These cabbage-heads, many of them compound, occur as vertically stacked, hemispherical stromatolites. They are composed of discrete bulbous and domal structures built of hemispheroidal stromatolites expanding upward from a base, although continued expansion may result in the fusion of neighboring colonies (Friedman et al. 1982). The heads, many of them compound, are circular in horizontal sections, and ranging in diameter from a few centimeters to a meter. These structures apparently possessed relief of up to 0.75m above the surrounding substrate. The size of the larger heads suggests that they formed in highly turbulent waters.
The Hoyt Limestone was deposited in a peritidal setting along a prograding shoreline (Owen and Friedman 1984). Oolite shoals and stromatolitic buildups restricted oceanic circulation causing locally increased salinity which contributed to dolomitization.
The lower part of one of the sections (Lester Park) provides the most complete sequence observed in the Hoyt Formation. Five lithofacies have been recognized. A vertical sequence from Lithofacies 1 to Lithofacies 4 is represented in Figure 13. Lithofacies I consists of ooids, some crossbedding, and sparry calcite cement. Lithofacies 2 is composed of partially dolomitized, domal stromatolites containing abundant carbonate particles (skeletal fragments, especially of trilobites, ooids and oncolites) and horizontal burrows. Lithofacies 3 consists of dolomitic micrite to siltstone containing composite particles. These carbonates are dark gray in color and are interpreted as deposits that formed under somewhat reducing conditions. Bioturbation features and lenses of skeletal trilobite debris are present. Channeled indistinct laminae between domal stromatolites have been designated lithofacies 4. Lithofacies 5 consists of stromatolitic laminae with local mudcracks and birdseye textures (not shown in Fig. 13).
Figure 14 shows the depositional environments ranging from the lee of an oolite shoal upward to the lower intertidal zone. The sequence seen here may have resulted from the lowering of sea level (regression) or from the depositional buildup of carbonates (progradation).
Figures 14 and 15 present a cross-section of the interpreted Hoyt depositional model. The following observations support the model: 1) the presence of well-developed ooids in the Hoyt lithologies suggests that the offshore energy barrier was an oolite shoal perhaps similar to that of the western Bahama Bank (Ball 1967) or to that of Abu Dhabi on the Trucial coast (Kendall and Skipwith 1969; Friedman 1995); 2) the presence in the Hoyt of high-relief domal stromatolites immediately overlying cross-bedded oolite, and the dependence of stromatolite morphology on energy (Logan et a1. 1964), suggest that the large heads were restricted to high-energy areas; and 3) the presence of coarse calcarenite between the heads. Storm surges may account for mixing of carbonate grains in the different lithofacies.
Figure 15 is a block diagram showing the facies relationships in which a peritidal setting includes oolite shoals, lagoons, and peritidal flats. The evidence for deposition under peritidal conditions for the Hoyt Limestone at Lester Park includes: (1) mud cracks, (2) flat-pebble conglomerate, (3) ripple marks, (4) small channels, (5) cross-beds, (6) birdseye structures, (7) syngenetic dolomite, and (8) the presence of stromatolites (Friedman 1969; Owen and Friedman 1984).
MACRO- AND MICRO-STRUCTURES OF STROMATOLITES
The high-relief domal and bulbous stromatolites contain distinct, generally uniform laminae of dolomite and calcite (Figs. 5-10). The structure of the stromatolites is sharply defined on weathered outcrop surfaces by the differential weathering of the calcite and dolomite. The laminae reflect spurts of the cyanobacterial mats that may correspond to tidal cycles and hence have been designated tidal rhythmites. Successive laminae of these structures drape over the ends of previous laminae, many of them curling underneath the heads to form "overturned" laminae.
The alternating calcite and dolomite laminae permit an analogy with modern microbial mats in hypersaline pools of the Red Sea coast (Friedman et al. 1973). In these modern stromatolites aragonite and high-magnesian calcite laminites alternate. In these modern analogs the high-magnesium calcite laminites contain abundant organic matter in which magnesium has been concentrated to form a magnesium-organic complex. Between the magnesium concentration of the high-magnesian calcite and that of the organic matter sufficient magnesium exists in modern microbial laminates to form dolomite (Friedman et al. 1973, 1982). Hence the observation in ancient stromatolites, such as observed in the Hoyt Limestone, that calcite and dolomite in the stromatolites are interlaminated, with calcite probably forming at the expense of aragonite and dolomite forming from high-magnesian calcite. However Burne (1995) pointed out that in modern settings, while some cyanobacteria may provide a substitute for mineralization, no microbiological evidence exists to support the view that cyanobacteria are capable of secreting carbonate minerals. Yet sediment-binding by cementation, i.e. precipitating or secreting carbonate minerals, and/or mat trapping is necessary to preserve stromatolite structure. Whether the morphology of stromatolites is biologically or environmentally controlled is controversial (Duane and Al-Zamel 1999).
Dark centers of the dolomite crystals suggest the concentration of organic material during the early stages of formation of the dolomite. The microcrystalline calcite laminae are commonly compound, i.e. they may consist of two or more micrite laminae without intervening dolomite. The micrite laminae display a wide range of textures, ranging from dense micrite to grumous and pelletoidal textures. Abundant detrital quartz particles are found in the micrite laminae; in contrast, the dolomite laminae only rarely contain detritus. The detrital grains consist almost exclusively of rounded quartz sand or angular quartz silt; carbonate grains are common in the domed stromatolites and are generally completely dolomitized. Quartz-sand grains obtained from insoluble residues of the microbial heads show strong pitting of surfaces. Between cabbage-head structures, skeletal fragments, especially those of trilobites, brachiopods, and pelecypods are abundant and in places ooids (Fig. 16) and quartz-sand particles are common. Fragments of trilobites are abundant. They belong especially to an assemblage known as Plethopeltis (Ludvigsen and Westrop 1983).
MODERN ANALOGUES OF CAMBRIAN SARATOGA SPRINGS STROMATOLITES
Friedman (1987a,b, 1988a,b) compared the petrography of the microbial stromatolites with that of modern cyanobacterial mats in hypersaline pools of the Red Sea coast (Friedman 1985; Friedman et a1. 1973). Modern cyanobacteria secrete radial ooids, oncolites, and grapestones which occur in these Cambrian rocks; interlaminated calcite and dolomite which in part compose the stromatolites of the New York Cambrian correspond to alternating aragonite and high-magnesian calcite laminites which modern cyanobacteria secrete. In modern cyanobacteria of sea-marginal ponds of the Red Sea the high-magnesian calcite laminites contain abundant organic matter in which magnesium has been concentrated to form a magnesium-organic complex. Between the magnesian concentration of the high-magnesian calcite and that of the organic matter sufficient magnesium exists in modern microbial mats to form dolomite. Hence the observation in ancient microbialites, such as observed in this classical New York Cambrian locality, that calcite and dolomite are interlaminated, with calcite probably forming at the expense of aragonite and dolomite forming from high-magnesian calcite. The structure of the stromatolites is sharply defined on weathered outcrop surfaces by differential weathering of the calcite and dolomite.
Discovery of the stromatolites at Hamelin Pool in western Australia (Fig. 17) (Logan 1961; Logan et a1. 1964) revived interest in the spectacular domed microbial stromatolites of Saratoga Springs, New York, which as Cryptozoön proliferum Hall, were the first formally named stromatolites.

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

STOP 3. Fort Hunter Quarry
Alight at slight bend in road and walk to Fort Hunter Quarry which is across a former railroad track (now a bicycle pass) close to Mohawk River. (Fort Hunter Quarry cannot be seen from road; another small quarry visible from road is approximately 0.1 mile farther east, but will not be visited on this trip).
Products of Tidal Environment: Stromatolites
Stromatolites in the Fort Hunter quarry consist almost entirely of dolomite in the form of irregularly bedded, finely-laminated, undulating structures. The rocks in this quarry are part of the Tribes Hill Formation of earliest Ordovician (Fisher 1954). The lithofacies of the Tribes Hill Formation have been studied in detail by Braun and Friedman (1969) within the stratigraphic framework established by Fisher (1954). Figure 18 is a columnar section showing the relationship of ten lithofacies to four members of the Tribes Hill Formation. At Fort Hunter we will study the lowermost two lithofacies of the Fort Johnson Member (see column at right (east) end of the section, in Fig. 18.
Two lithofacies are observed: (1) lithofacies 1, mottled feldspathic dolomite, and (2) lithofacies 2, laminated feldspathic dolomite. Lithofacies 1 is at the bottom of the quarry, and lithofacies 2 is approximately halfway up.
Lithofacies 1
This facies occurs as thin dolostone beds, 2 cm to 25 cm, but locally more than 50 cm thick, with a few thin interbeds of black argillaceous dolostone which are up to 5 cm thick. In the field, the dolomite shows gray-black mottling and in places bird's eye structures. In one sample, the infilling of the bird's eyes shows a black bituminous rim which may be anthraxolite. In the field, trace fossils are abundant, but fossils were not noted. Authigenic alkali feldspar (microcline) is ubiquitous throughout this lithofacies. The insoluble residue makes up 22 to 54% by weight of the sediment in samples studied with most of the residue composed of authigenic feldspar.
Lithofacies 2
This lithofacies is mineralogically identical to the previous facies but differs from it texturally and structurally in being irregularly bedded and in containing abundant undulating stromatolitic structures ("pseudo-ripples") (Fig. 19) as well as disturbed and discontinuous laminae. In places there are a few thin interbeds of black argillaceous dolostone. The thickness of the laminates of this facies ranges from 0.5 mm to 2 or 3 mm; on freshly broken surfaces the color of the thinner laminae is black and that of the thicker ones is gray. The insoluble residue, for the most part composed of authigenic feldspar, constitutes between 35 and 67% by weight in samples studied.
These two lithofacies which form the basal unit of the Ordovician, were formed on a broad shallow shelf.
40Ar/39Ar spectrum analysis of the K-feldspar of Sauk-Sequence carbonates of Lithofacies 2, laminated feldspathic dolostone of the Tribes Hill Formation, yielded an age range reflecting the original detrital cores of Sauk age (470 Ma) and an uplift of Pennsylvanian age (320 Ma) (Matt Heizler, analyst) (Friedman 1990). Hearn et al. (1987) using 40Ar/39Ar analyses for the authigenic feldspar in Cambro-Ordovician carbonate rocks of the central and southern Appalachians obtained a comparable age for uplift (278-322 Ma). For carbonate rocks in Nova Scotia 40Ar/39Ar clocks in feldspar were reset at 300-320 Ma.
Stromatolites, bird's eye structures, scarcity of fossils, bituminous material, syngenetic dolomite, and mottling suggest that these rocks were deposited in a peritidal environment (Friedman et al. 1992). Based on analogy with the carbonate sediments in the modern Bahamas, Braun and Friedman (1969) concluded that these two lithofacies formed under supratidal conditions. However in the Persian Gulf flat microbial mats prefer the uppermost intertidal environment, and along the Red Sea coast they flourish where entirely immersed in seawater, provided hypersaline conditions keep away burrowers and grazers (Friedman et a1. 1973). Hence on this field trip we may conclude that the stromatolites indicate peritidal conditions without distinguishing between intertidal and supratidal. For more details on these lithofacies refer to Braun and Friedman (1969).
Data obtained from the analysis of fluid inclusions in calcite-healed fractures of these Lower Ordovician carbonate strata, in which microcline crystals are so prominent, indicate higher paleotemperatures and greater depth of burial than have previously been inferred for the rocks of this region (Urschel and Friedman 1984; Friedman 1987a,b). Average fluid-homogenization temperatures range from 96°C to 159°C. These high paleotemperatures are supported by oxygen-isotope and conodont-alteration data (Harris et al. 1978). A former depth of burial 7 km is implied when a geothermal gradient of 26°C/km (Friedman and Sanders 1982,1983) is used.
Thus, following subsidence to great depth, Pennsylvanian to Permian epeirogeny uplifted the strata, resulting in deep erosion. This leads to the surprising conclusion that isostatic unroofing following uplift has stripped off thick sections of strata whose presence was previously unsuspected.

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

Fulton-Montgomery Community College

STOP 4. North Tribes Hill Quarry
Route of Walk
Take the trail towards old abandoned crusher, but instead of heading towards the quarry move uphill to the first rock exposures. The rocks to be examined are near the edge of steep cliff.
Description and Discussion
At this stop microbial reef mounds are exposed (Friedman, 1996a,b). Ordovician domal thrombolites, termed here microbial reef mounds, occupied the basal part of meter-scale shallowing-upward cycles (Fig. 20). They are part of a high-energy facies that a sharp transgressive surface separates from an underlying low-energy peritidal setting. This erosional surface served as the surface on which one of the reef mounds established itself during initial transgression before further deepening. The others overlie a floor of skeletal grainstone reflecting a high-stand sea-level facies tract. Skeletal grainstone composes the fill between the mounds. A channel and several aggrading hummocks occupy inter-reef mound areas resulting from storm events in a subtidal setting.
Reef mounds formed at or near the base of upward-shallowing parasequences. They are part of a high-energy facies which a parasequence surface of emergence or near emergence separates from an underlying parasequence which terminated a low-energy peritidal setting. This erosional surface between the two parasequences is interpreted as a partly lithified hard ground. As elsewhere in the Cambro-Ordovician of North America, microbial reef mounds commonly occur near the bases of upward-shallowing cycles and most rest directly on underlying cycle caps (Osleger and Montafiez 1996).
Mound-Foundation Facies
Braun and Friedman (l969) designated the facies underlying the reef mounds as Lithofacies 7: Mottled Dolomitic Micrite and Biomicrite of the Wolf Hollow member (Fig. 18). This lithofacies is made up of a well-bedded, mottled limestone in which the mottles are composed of irregular patches of dolomite. On weathered surfaces, the limestone is whitish and the dolomite buff, but on freshly broken surfaces both limestone and dolomite are very light gray with the limestone somewhat darker and the dolomite of granular appearance. The outlines of large gastropods and cephalopods stand out, in places, on weathered bedding planes. A list of fossils found in this lithofacies was given by Fisher (l954, p. 88-89). Bird's eye structures are present in some beds as are pyrite patches. The limestone contains abundant dolomite-filled burrows, most of which are horizontal (or sub-horizontal) to the bedding plane, but some burrows have oblique to perpendicular orientations with respect to bedding. Many gastropods, especially Ophileta and Ecculiomphalous are found with the dolomite-filled burrows suggesting that these burrows may have been made by gastropods rather than by worms. However, the morphology of the shells suggests that these gastropods were not burrowers. Hence, worms or other soft-bodied organisms must have been abundant and produced the burrows.
This facies which is part of the Wolf Hollow Member (see Fig. 18), represents a low-energy deposit of micrite of a shallowing-upward cycle terminating in a sharp, planar erosion surface. This erosional parasequence surface is a marine flooding surface and represents a transgressive event for the next high-energy cycle in which the reef mounds formed. Below one of the reef mounds the underlying micrite of mound-foundation facies compacted (Fig. 21).
Differential compaction of this former lime mud, as the solid reef grew, suggests that the lime mud had not yet fully lithified (Friedman 1996a,b). This observation differs somewhat from that of reef mounds in Virginia and Argentina where a solid hard ground served as foundation for the reef mounds (Read and Grover 1977; Canas and Carrera 1993). However, the original lime mud of the mound-foundation facies of the Tribes Hill mounds was sufficiently lithified to support the growing mounds (Friedman 1996a,b).
Microbial Reef-Mound Facies
Reef mounds are prominent in the Wolf Hollow Member of the Tribes Hill Formation (Figs. 20, 21). They occur as isolated mounds (Friedman 1996a,b). These mounds are approximately one meter in length (measurements vary from 95 cm to 135 cm) and 60 to 70 cm in thickness. These measurements are approximate dimensions because the mounds do not stand out freely, and disappear in the enveloping facies. Moreover in the exposure, it is difficult to differentiate between short and long axes of the mounds.
The reef mounds are composed of clotted and peloidal microcrystalline matrix which was microbially precipitated, comparable to that in modern reefs (Friedman et at 1974). In modern reef settings, peloids display the pattern of calcified coccoid cells which cyanobacteria or chemoorganotrophic bacteria degrading the cyanobacterial organic matter precipitate (Friedman et al. 1985; Krumbein 1983). These textural features are products precipitated by the micro- environment of cyanobacteria (Nadson 1903; Kalkowsky 1908; Pia 1927; Johnson 1954; Endo 1961; Friedman et al. 1973). Peloids have been described as calcified algal filaments (Friedman et al. 1974); such calcification may be the result of precipitation of calcium carbonate on cyano-bacterial filaments in the presence of live bacteria (Chafetz and Buczynski 1992).
Because they are composed of microcrystalline carbonate matrix, the microbial reef mounds have a texture similar to that of the mound-foundation facies lithofacies 7 (see Fig. 18), a micrite. The term matrix has been commonly misapplied as a synonym of micrite, but the mounds are not composed of micrite. Mechanically deposited lime mud, following lithification, is known as micrite (Folk 1959). Identifying even modern reef rock is an experience in frustration: submarine microcrystalline or cryptocrystalline matrix that is biologically precipitated within millimeters to centimeters of the surfaces of reef rock is identical in appearance to micrite of mechanical origin. Case histories abound in which unwary geologists have misidentified the reef rock for low-energy facies composed of micrite (Friedman 1985, 1994a,b). Since micrite of low-energy origin and microcrystalline or cryptocrystalline matrix of reefs are indistinguishable it is easy to confuse high-energy reef facies for low-energy lime-mud facies (Friedman 1985, 1994a,b). This textural similarity led initially to an interpretation that mounds may be blocks of lithofacies 7 (micrite and biomicrite) that foundered and became lodged in channels (Braun and Friedman 1969). These reef mounds resemble blocks of micrite in tidal channels of the Bahamas that are derived by undercutting of the banks of the channels (Braun and Friedman 1969).
The reef mounds were included with lithofacies 8 of Braun and Friedman (1969) designated intrasparite and biosparite; the lithology is for the most part a skeletal grainstone. This lithofacies was referred to as channel fill, comparable to the Lower Ordovician reef mounds of western Argentina, of which Cañas and Carrera (1993, p. 169) noted "the reef mounds are dissected by conspicuous channels filled with coarse crinoidal grainstone and lithoclastic rudstone" (Fig. 21). This same observation applies in part to the setting of the Tribes Hill Formation. As in western Argentina, the reef mounds of the Tribes Hill Formation formed in part on a previously lithified or partly lithified sediment surface (Cañas and Carrera 1993, p. 169); and as in other places in North America, they occur near the base of an upward-shallowing cycle. The surface on which one of the mounds developed is a sharp transgressive parasequence surface separating the underlying peritidal lithofacies 7 (micrite and biomicrite) from the overlying subtidal reef-mound facies. The other reef mounds nucleated near the transgressive parasequence surface, but on top of underlying skeletal grainstone, during the initial transgression before rapid deepening occurred, a setting which is similar to that of comparable facies in the Great Basin, U.S.A. (Osleger and Montanez 1996).
Inter-Reef Mound Facies
Skeletal grainstone composes the fill between the reef mounds. One channel and several hummocks occupy the inter-reef mound areas (Fig. 22). The top of one of the hummocks rolls into the channel fill. The grainstones form lenses that build on top of one another. The channel displays the typical asymmetric profile of a tidal channel with a steep cut-bank and a low-angle slip-off slope (Fig. 22), However, the channel is entirely within grainstone, hence it is not a normal tidal channel which would be fine-grained on the steep side and coarse-grained on the opposing side. In normal tidal channels, as the channel shifts it leaves behind a layer of coarse debris at the bottom of the channel (Friedman et al. 1992). No such channel-floor lag layer is present in the inter-reef mound facies. Hence the channel must be related to storm deposition of the grainstone since filling is not the result of the shifting of the channel. Truncation by the channel and aggradation of the hummocks occurred at the same time. The channel and hummocks formed as a result of storm events in a subtidal setting. Following transgression, storm tides and currents generated this channel between which reef mounds and inter-reef mound facies accumulated. This channel, which was incised down to 30 cm into the underlying grainstone, is conspicuous and displays sharp margins (Fig. 22).
Of the various reef mounds one rests directly on the mound-foundation surface; grainstone of inter-reef mound facies makes up the floor of all the others.
The reef mounds formed in shallow-subtidal to possibly low intertidal settings in an agitated environment devoid of lime mud.

STOP 5. Exposures of Gailor Formation
STOP 5. Exposures of Gailor Formation On the north and south sides of the road are exposed a massive dolostone unit overlain by a bedded dolostone unit. Roughly 10' of section are exposed here. The massive unit is dark gray in color while the upper bedded unit is lighter gray and coarser as well. Large clasts in a variety of shapes and sizes, composed of micrite and medium-textured dolostones are scattered all over the outcrop and concentrated in the basal massive unit. Pods and lenses of chert colored black and white are profuse. Calcite mineralization is also a common feature, occurring in patches and veins in orange, white and black. Stromatolites are observed in the section on the north side of the road.
The breccia observed in the section perhaps represents dissolution collapse. The massive unit in the lower part of the section may be a microbial build up. The stromatolites suggest a peritidal origin for these dolostones.

STOP 6. Farther east on the north and south side of 67 are exposed roughly 27' of section. The following features are observed:
-wavy beds,
-massive units alternating with bedded dolostone units,
-the massive units appear to have a lenticular mound-like form, perhaps representing former microbial build-ups,
-the beds overlying the mounds display dips on flanks of the mounds perhaps indication fore-reef slopes,
-pervasive dolomitization seems to have obliterated original depositional features,
-alternatively the wavy bedding may represent tidal channels or hummocky cross stratification,
-other features observed in these units are fine laminae in the dolostones, burrow mottling, stylolites, presence of pyrite and white calcite mineralization in vugs and fractures.

Farther east are exposed roughly 15' of bedded dolostones displaying planar stromatolites, intraclasts, and bird's eye structures indicating peritidal environments of deposition.

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

Farther east on Route 67 near Manny Corners is a section exposing roughly 26' of bedded dolostones. The following features are observed here:
-planar and domal stromatolites,
-intraclasts ranging in size from 1 to 2 inches,
-vugs and fractures partially or completely filled with black and white calcite,
-bird's eye structures,
-bedded and nodular chert,
-breccia representing dissolution collapse?
The above mentioned features displayed in these fine- to medium-textured dolostones point to a peritidal environment of deposition.

C9: GEOMORPHOLOGIC FACTORS lN THE FAILURE OF GENERAL BURGOYNE'S NORTHERN CAMPAIGN OF 1777
Kenneth G. Johnson and Krista Reichert
Department of Geosciences
Skidmore College
Saratoga Springs, New York 12866
Throughout human history the natural landscape has imposed limitations and offered opportunities to those who would use it. Certainly, in warfare this has always been a fundamental truth - a truth that Lieutenant-General John Burgoyne understood fully when in 1775 he proposed his plan for the British Northern Campaign of 1777. This strategic plan, "Thoughts for Conducting War on the Side of Canada", which was in fact not particularly original, called for a three-pronged invasion (see figure 1) which would involve using waterways in the upper St. Lawrence and Mohawk Valleys and in the Hudson-Champlain lowland to transport armies which would converge on Albany, New York and thereby isolate the fervently rebellious New England colonies from those colonies south of New York. A possible follow-up element of the plan was to then turn east and advance on Boston (Ketchum, 1997). Albany was important also because it served as the supply base for the Northern Department of the American Army.
This field trip will focus on the path of advance of Burgoyne's army from Fort Anne south through Fort Edward and Saratoga (now Schuylerville) to Bemis Heights, where 225 years ago the American forces under Major General Horatio Gates defeated the British army and turned the tide of the American Revolution. The final stop will be at the Saratoga National Historical Park, where from 2:00 pm to 3:15 pm there will be period music, interpretive activities and a mass tactical demonstration at the Barber Wheatfield, ground of the second of the Battles of Saratoga.

C9: GEOMORPHOLOGIC FACTORS lN THE FAILURE OF GENERAL BURGOYNE'S NORTHERN CAMPAIGN OF 1777

Bloody Pond, a kettle lake and scene of two French and Indian War skirmishes.

On the right (west), an outcrop of Precambrian (Grenville) gneiss.
French Mountain, on the left, is a horst on the east side of the Lake George graben. The military road passed below the present road and this was the site of the French ambuscade of the Provincials and Iroquois, known as "The Bloody Morning Scout". Colonel Ephraim Williams, whose estate funded the founding of Williams College, was killed standing on a glacial erratic boulder. The colonials and Mohawks managed to withdraw from "Rocky Gulch" to Lake George, leaving some 100 casualties behind.

Kame terrace capped by Lake Albany clays.

Low area on left - an extension of Glen Lake, an ice block lake.
Immediately north of the Great Escape Amusement Park, on the left, is a low area which is an extension of Glen Lake, an ice block lake. Rt. 9 crosses Five Mile Run, the site of several French and Indian War ambuscades.

Round Pond (and Paradise Beach)
Round Pond (and Paradise Beach), a kettle lake with an esker on each side at the west end.

Blind Rock Rd. leads down from the kame and esker complex to Bay Rd.

Adirondack Community College
On left, Adirondack Community College, which is situated on glacio-lacustrine clays overlying Ordovician Beekmantown carbonates.

Drainage divide
Drainage divide between the Lake Champlain/Lake George basin and the Hudson River basin

Dix Ave follows part of the route of the old military road
Dix Ave follows part of the route of the old military road from Fort Edward to Lake George, which skirted a low area and then swung west and north to avoid the swamp through which Quaker Rd passes.

Wood Creek Valley, which Burgoyne used to advance
River St. hugs the northeast bank of the Hudson River and climbs to the Hudson Falls Village Park. Drive to the right around the park and continue south on Rt 4 (Main St.). Rt 4, which enters Hudson Falls from the northeast follows the northwest side of the Champlain Canal. This is the Wood Creek Valley, which Burgoyne used to advance from Skenesborough (now Whitehall) through Fort Anne to Fort Edward (see p. C.9-7).

STOP 1. UNION CEMETARY
This location is on the gently sloping top of the Glens Falls glacio-lacustrine delta. Connally (1973) maps these delta facies as foreset and bottomset sands, underlain by shale bedrock. It is easy to understand why the settlement that preceded present-day Hudson Falls was called "Sandy Hill". Jane McCrae, murdered by Burgoyne's Indians, is buried in the cemetery. It is ironic that she was engaged to a junior officer in Burgoyne's forces and members of her family were British sympathizers. The Americans made maximum use of her murder for propaganda purposes.

Site where Jane McCrea was murdered
Approximate location of site where Jane McCrea was murdered. A blockhouse in the defensive network of Fort Edward was located where the school athletic field is now situated.

Site of Fort Edward
On Right, site of Fort Edward (see p. C.9-7). Only a small portion of the moat of this fort remains, even though prior to 1800 it was the largest British military installation in North America. The fort was never besieged or defended, except against small raiding parties.

Bridge over Fort Edward Creek.
ON LEFT, an old canal aqueduct over the creek.

On right, Old Fort House Museum
On right, Old Fort House Museum, built prior to the Revolution of materials said to have been scavenged from buildings at Fort Edward

Cross Champlain Canal just north of Lock 7.
This water barrier did not exist prior to the 19th Century and consequently the military road extended down the east side of the Hudson River.

First burial site of Jane McCrea and the grave of a militia Lieutenant
On the right, the first burial site of Jane McCrea and the grave of a militia Lieutenant who was killed the same day, during the Indian attacks. near Fort Edward.

Cross Moses Kill

On the right, across the canal, FORT MILLER.

Old lock foundation on the left.

Bridge over Hudson River into Saratoga County.
Bridge is just upstream from the rapids, now a dam, that were exploited by the British to build a pontoon bridge to the west bank of the Hudson. General Fraser's advance corps completed the bridge the same day as the Battle of Bennington, but Burgoyne's army did not cross it until almost a month later.

STOP 2. STARK'S KNOB
At the junction of Rt. 4 and Stark's Knob Road. Immediately west of this point an abrupt hill was used by the Americans to set up an artillery position that commanded the road and the river crossing that Burgoyne planned to use in his retreat north after the Battles of Saratoga. The knob is a mass of pillow basalt, that probably was pushed westward in one of the Taconic thrusts. It is generally agreed that it is allochthonous. The knob has been quarried for road metal and is much smaller than it originally was.

General Philip Schuyler Commemorative Highway.

Breastworks of Burgoyne's fortified camp
At the entrance to the Schuylerville Central School, Rt. 29 crosses the position of the breastworks of Burgoyne's fortified camp and west of this at a higher elevation is an area of stabilized sand dunes composed of reworked material deposited in Glacial lake Albany.

Colonel Morgan held this position to prevent a British retreat to the west.

STOP 3. The SARATOGA BATTLE MONUMENT
This is a location that overlooks the Hudson Valley about 10 miles north of Bemis Heights, where the Battles of Saratoga were fought.

CROSS FISH CREEK
CROSS FISH CREEK, which drains Saratoga Lake about 15 miles to the west. On the left, the Schuyler house, manorial home of General Schuyler, which for tactical reasons was burned by Burgoyne during his retreat after the Battles of Saratoga. The house was rebuilt and is now a National Park historical site.

On the right, an outcrop of Ordovician Normanskill shale.
Other outcrops are present along the road farther south.

On left, site of Dovegat House, which served as Burgoyne's headquarters.

Coveville - The Cove, a Hudson River backwater
Coveville - The Cove, a Hudson River backwater, is part of a plunge basin formed when the waters of present-day Fish Creek entered glacial Lake Fort Ann at this point. The British used this backwater to anchor some of their transport bateaux just before the Battles at Bemis Heights.

Wilbur Road, on right, leads to the Saratoga National Veterans Cemetery.

Entering National Park Lands.

ENTRANCE to Saratoga National Historical Park.
Our manner of entrance and activities there will be determined by the events connected with the 225th Anniversary of the Battles of Saratoga that are taking place on September 29th.

C10: SELECTED TILL AND STRATIFIED DRIFT DEPOSITS BETWEEN GLEN'S FALLS AND AMSTERDAM, NEW YORK
Donald T. Rodbell
Geology Department
Union College
Schenectady, NY 12308-2311
rodbelld@union.edu
This trip is intended to provide participants with a review of the glacial stratigraphy of the portion of eastern New York State between Glenn's Falls on the northeast and Amsterdam on the southwest, and to correlate the regional stratigraphy with the better dated sequences of central New York. All of the data presented here were gathered by students in four different classes of Glacial and Quaternary Geology (Geo 52) at Union College between 1995 and 2002.
The region covered by this field trip contains a wide variety of glacial deposits-from subglacial till to subaqueous fan deposits to varves to eolian sands. Perhaps what is most interesting about this region is the evidence present for a dynamic history of ice marginal advances and retreats, and the numerous glacial lakes that occupied much of the region during the latter third of marine isotope stage 2 (Martinson et al., 1987). While there has been a plethora of prior work conducted in the area, a near absence of numerical age control coupled with stratigraphic uncertainties generated by a laterally discontinuous and variable lithostratigraphy provides us with considerable room for debate over the details of the glacial history of the region.

C10: SELECTED TILL AND STRATIFIED DRIFT DEPOSITS BETWEEN GLEN'S FALLS AND AMSTERDAM, NEW YORK

STOP 1. Basal Till of Luzerene Gorge and Multiple Diamicts
This stop is on the north side of the Luzerne Gorge, in sight of the Hudson River. Exposed along the north side of the Gorge are a thick sequence of diamicts, sand, gravel, and varves (Stop 2). At this stop we will examine the basal till of the Gorge. In addition to a strong NNE-SSW fabric (Fig. 2), this till yields composition indicators that reflect a provenance in the Hudson-Champlain lowlands and from regions underlain by crystalline rocks.
Magnetic susceptibility, carbonate content, texture, and percentage crystalline clasts provide a means to distinguish till deposited from Hudson-Mohawk ice from that deposited by Adirondack ice. Tills deposited by Hudson-Mohawk ice contain >3% carbonate, whereas Adirondack tills are devoid of carbonate (Fig. 1). Similarly, the matrix of tills deposited by Hudson-Mohawk ice contain >10% clay and yield relatively low magnetic susceptibility values (</200 x 10sup-3/sup SI units). These trends are most obvious when comparing tills in the eastern Mohawk Valley (e.g., Stops 1-1, 2- 1, 2-2 in Fig. l) with tills in the Adirondacks (e.g., Stops 5-1, 5-3 in Fig. l). The tills of the Luzerne Gorge have compositions that are intermediate between these two end members; nonetheless, these tills are clearly distinguished from Adirondack till.

STOP 2. Varves Between Upper and Lowere Tills of Luzerene Gorge
Here along this portion of the roadcut are exposed ∼ 10 meters of varves. These are stratigraphically above the basal till of Stop 1 and below the uppermost till, which is exposed along the roadcut several hundred meters up the hill to the southwest of this outcrop. The site has been slumped and colluviated for years, but we have excavated through the colluvium and found varves at nearly all levels of this outcrop. These lake sediments may have been deposited in the southern-most segment of Glacial Lake Warrensburg (Connally and Sirkin, 1971), or in a smaller lake restricted to within the Gorge.
The presence of lacustrine sediment between diamicts reflects an interstadial period. In order to estimate the duration of this interstadial in the absence of a clean and continuous outcrop of varves, we have over the years determined sedimentation rate (couplets/cm) for varves in several positions of the outcrop. Typical values are 3.5 couplets/cm. If this rate is representative of the average sedimentation rate for the entire outcrop, then the duration of this interstadial period was ∼3500 years. Owing to the presence of several thick sand layers within this varve sequence, the calculated duration is best considered to be a maximum limiting duration.
Correlation of this interstadial period with the regional record of glaciation is tenuous, but it seems likely that the interstadial recorded here correlates with the Shed Brook Discontinuity in the western Mohawk Valley (Ridge, 1997) and thus with the Erie Interstade (14-16 ka). As noted above, the entire sequence here must be older than 13,150 14C yr BP, and this is consistent with correlation to the Erie Interstade.
The uppermost till is not very accessible in the Gorge, but a colluviated section of it is exposed about 200 m to the southwest of Stop 2. The till there is compositionally very similar to the basal till, again implying an origin in the Hudson-Champlain lowlands.

STOP 3. Delta on North End of Sacandaga Basin
This site, which is located on the north side of the Great Sacandaga Lake, exposes more than 10 m of well-sorted sand and some gravel, in discrete foreset beds that dip southward. Superimposed on these beds are well-preserved Type A climbing-ripple cross-laminations (Harms et al., 1982) that are exposed in planes parallel to flow. Both the foreset beds and the climbing ripples indicate a strongly southward paleocurrent direction (Fig. 4); or into the Sacandaga Basin. This deposit represents a delta or kame delta sufficiently distant from the ice margin to exclude the coarse fraction of glacial outwash.
A delta with southward paleocurrent directions on the north side of the Sacandaga Basin requires a lake in the basin, much as exists today. The elevation of this lake surface at its highest was 300 masl (∼985') based on the elevation of the topset-foreset transition of a large kame delta at the mouth of the Sacandaga River exposed in a gravel pit ∼3 km northwest of Northville.
To generate a lake of this elevation in the Sacandaga Basin, the outlet of the Basin near Conklingville or Lake Luzerne must have been dammed by the Hudson-Mohawk lobe. In addition, to generate a lake level of ∼300 m, the southern margin of the Sacandaga Basin must also have been dammed by the Hudson-Mohawk lobe. The top of the Broadalbin-Interlobate moraine (Stops 4 and 5; Dineen and Hansen, 1995), which is ∼20 km south of the aforementioned highest foreset beds has an elevation of ∼260 m. Even a high gradient of isostatic tilting of ∼1 m/km is insufficient to make the BIM a viable dam for Glacial Lake Sacandaga; the surface of the Hudson-Mohawk lobe must have been at least 20 m higher than the surface of the BIM. The presence of Glacial Lake Sacandaga requires substantial retreat of the Adirondack lobe while the northern side of the Hudson-Mohawk lobe at the BIM remained relatively stationary.

STOP 4. Herba Sand and Gravel Pit
This site is located on the northern edge of the BIM, ∼5 km west of Broadalbin. It exposes more than 25 m of stratified drift, which ranges in grain size from sand-cobble. Dineen and Hansen (1995) reported the presence of lodgement and flow till at this site. The dominant lithology here is metamorphic-mostly granitic gneisses.
Much of the BIM is underlain by stratified drift, and it seems likely that the BIM was deposited as a kame into standing water between the Adirondack lobe to the north and the Hudson-Mohawk lobe to the south. The distance between these lobes must have increased as Adirondack ice retreated northward and the northern edge of the Hudson-Mohawk lobe remained more-or-less stationary to the south of the BIM. This scenario is supported by paleocurrents which vary widely but which are dominantly southwestward (Fig. 4), along the axis of the BIM - thus meltwater was apparently transporting sediment off the sides of both lobes, but these currents were eventually shunted southwestward toward what must have been an outlet to Glacial Lake Sacandaga.

STOP 5. Broadalbin Town Gravel Pit
Located near the center of the BIM, this site exposes a complex stratigraphy of outwash sands and gravel, horizontally laminated lacustrine sands, all of which is overlain by a diamict. This latter has many of the hallmarks of Adirondack till; namely, a high magnetic susceptibility, an absence of carbonate, low % clay, and high % crystalline rocks (Fig. 1). In addition, this till yields NNE-SSW till fabric (Fig. 5), consistent with Adirondack ice.
The stratigraphy appears to record an advance of the Adirondack lobe from N-S over the BIM. However, Dineen and Hansen (1995) have interpreted the stratigraphy here to reflect an advance of the Hudson-Mohawk lobe (Yosts Readvance). The presence of rotten shale clasts supports their interpretation, and it is possible that the apparent Adirondack composition reflects Hudson-Mohawk flow lines that overtopped the Kayaderorseras Range ENE of this site. In this case the, it is possible that this till is correlative with the upper till in the Luzerne Gorge (Stop 1), and reflects a Luzerne Readvance that was considerably larger than envisaged by Connally and Sirkin (1971). If so, then the Mt. McGregor moraine of Connally and Sirkin (1971) either represents a younger readvance or a recessional ice position during deglaciation from the Luzerne Readvance.

STOP 6. West Glenville Gravel Pit
This site exposes a kame delta on the northern edge of the Glenville Hills. The deposit is about 20 m thick and contains sand and gravel beds that are inclined at about 20° to the west. This west-southwest flow direction (Fig. 6) likely reflects glacial meltwater off the western side of the Hudson-Mohawk lobe after it had shrunk to open a basin to the west. The elevation of this lake was ∼215 m (700') based on the elevation of the top of this kame delta, and on the elevation of numerous sand and gravel pits that demarcate the southern extent of the lake along West Glenville Road to the west of this locality.
Although LaFleur (1965) maps this kame delta at the eastern edge of his "Early Lake Amsterdam", it seems more plausible that this kame was deposited into a much smaller, restricted lake. This smaller lake would have been dammed to the east by the Hudson Mohawk lobe, to the south by the Glenville Hills and to the west and north by bedrock hills within 2 km of this locality. The most likely outlet for this lake was southward down Wolf Creek Hollow and Hoffman's Delta (Stop 7) to the Mohawk River. This is consistent with the elevation of the outlet to Wolf Creek Hollow (∼215 m), and explains: 1- why Wolf Creek Hollow is devoid of till, and 2- why Hoffman's Delta is so much larger than other deltas in the eastern Mohawk Valley. Proglacial lake drainage down Wolf Creek Hollow would have stripped the valley of drift and induced considerable down cutting. Much of this drift would have been deposited in Hoffman's delta.

STOP 7. Hoffman's Delta
Hoffman's delta is the largest delta in the Mohawk Valley east of Amsterdam. It was deposited at a lake level of ∼130 m (∼420'), and according to LaFleur (1965) it was deposited into a late phase of Lake Amsterdam, when the Mohawk Valley was plugged to the east, at approximately Scotia, NY, by the eastward retreating front of the Mohawk lobe.

STOP 8. Auiresville Exposure
Here are exposed 40 m of diamicts, gravel and sand. The base of the exposure is marked by a typical lodgement till from the Mohawk lobe that is at least 12 m thick. It contains several lenses of gravel, which may reflect intervals of meltout till. This till is overlain by at least 12 m of sandy colluvium, which appears to be derived from a sand and gravel unit that is at least 4 meters thick. This gravel unit possesses strata that dip northward at ∼20°. The top of the exposure is marked by a second till, at least 6 meters thick, which is oxidized at its surface, and which is, in turn, overlain by well-sorted fine-medium sand. The sequence appears to reflect multiple ice advances over the site. The two tills are nearly identical in composition, and clearly reflect a Hudson-Mohawk source (Fig. 1). What few fabric data we have from the lower till indicate that it was deposited along an E-W (or W-E) flow path (Fig. 7). The sand and gravel unit separating the tills appears to be deltaic and indicates the presence of a glacial lake in the Mohawk Valley at an elevation of ∼85 m (280').

C11: GEOCHEMISTRY AND SOURCE OF THE SARATOGA SPRINGS
K.Hollocher
Geology Department
Union College
Schenectady, NY 12308
Lacie Quintin
Department of Earth Sciences
Boston University
Boston, MA, 02215
Dan Ruscitto,
The City of Saratoga Springs is situated in the lowlands southeast of the flanks of the southeastern Adirondack foothills. The crystalline rocks of the Adirondacks are in contact with the overlying Paleozoic sediments along an early Cambrian unconformity and along faults. The faults are mostly high angle, typically trending north-northeast, and normal, mostly down-dropped to the east. Faulting was probably associated with downwarping of the Laurentian continental margin during Late Ordovician collision with the Taconian volcanic arc, though this may have involved reactivation of earlier faults. The McGregor fault, and its southerly branch of interest that runs through downtown Saratoga Springs, hosts a remarkable collection of springs (actually now all drilled wells) that tap a carbonate aquifer in the Beekmantown Group at depths ranging from 75 to 300 m. All of the original springs and producing drilled wells are on the east side of the Saratoga Springs-McGregor fault where shale cap rock exists. To the west of the fault the Beekmantown Group carbonates are exposed at the surface, and no waters of the Saratoga type are found. Saratoga Springs, New York is 300 km north ofNew York City and is situated on the western margin of the Hudson Valley. Streams within the drainage basin around Saratoga Springs generally flow towards the east and southeast, more or less parallel to the regional dip towards the Hudson River (Fish, 1881; Kemp, 1912). Subsurface water flow follows the dip of the strata, apparently paralleling overall surface water flow (Fish, 1881). Not surprisingly, Saratoga Springs experiences four distinct seasons throughout a year. Summer temperatures can reach 40°C in the summer and can drop below -30°C in the winter. Despite this large temperature range, early analyses of the springs indicated that the temperature and flow of the springs is practically constant over time (Fish, 1881), an observation that we can partially confirm. The rising of the spring waters through natural fractures and drilled wells is driven principally by rise of CO2 bubbles degassing from the rising water, though in some cases there is also a significant artesian head. Some wells, if turned off, apparently have to be primed to get them to start flowing again (e.g., the "water seal" of Kemp, 1912). CO2 degassing causes the waters to become supersaturated with respect to calcite, and precipitation gradually seals off the water conduits. The fact that natural springs were present at !east in the 1700's hints at recent fault movement opening new fractures.

C11: GEOCHEMISTRY AND SOURCE OF THE SARATOGA SPRINGS

A small pavilion for the Old Red Spring is at this intersection.

STOP 1. OLD RED SPRING. (10 minutes)
The Old Red Spring (0.24% TDS, intermediate type, A in Figure 7) is the north-easternmost and one of the least saline of the existing Saratoga-type mineral springs. Some older literature describes this spring as being highly iron-rich, though chemical analysis shows this water to be unremarkable. It has the second lowest salinity of all of the flowing springs and so is more popular than some of the other saltier and smellier waters.

STOP 2. CONGRESS PARK AND ENVIRONS. (1 hour 15 minutes)
This is a walking stop at which we will see all of the flowing springs in the Saratoga Springs downtown.
High Rock Spring, originally known as the Big Medicine spring, was the first mineral spring in the area to be introduced to European settlers in 1767 (Siegel, 1996). Though the spring conduit has been cleaned with hand tools repeatedly, the original spring has long since sealed its own conduit by calcite precipitation. Any flowing water in the High Rock Pavilion actually flows by gravity through a pipe from the nearby Governor spring. The cliff to the west of the pavilions is of the Beekmantown Group (Amsterdam Limestone), faulted above the Utica Shale on which the pavilions are built. The fault displacement, calculated at the now vanished Star spring that was located ∼100 m northeast of the High Rock Spring, is ∼49 m, much less than the total carbonate section here of ∼140 m (Cushing and Reudemann, 1914).
Peerless Spring (0.15% TDS, intermediate type, B in Figure 7) is the least saline of the flowing springs. Its salinity has been decreasing over the past decade. Based on the salinity curve over time, we predicted that it would stop flowing in 2001, which in fact it did. However, we found in spring 2002 that it was flowing again if only a trickle. It is called Peerless probably because it is the least saline and so least bad-tasting of the Saratoga Springs.
Governor Spring (0.35% TDS, intermediate type, C in Figure 7) is in the same pavilion as Peerless, though both are piped from the well head across the street.
Hathorn 1 Spring (0.71%TDS, saline type, D in Figure 7) is the first and northernmost of the saline-type springs as defined in Figures 3B and 4. The drilled well is just over 300 m deep and penetrates ∼80 m into Cambrian clastic rocks. This spring outlet and pavilion were moved in 1999 from across the street to make room for the parking garage currently at the original spring location.
Congress Spring (0.28% TDS, intermediate type, E in Figure 7) is perhaps the most frequently visited of the springs, its handsome pavilion being a prominent landmark in Congress Park.
Columbian Spring (0.026% TDS, fresh water, F in Figure 7) is actually from the municipal water supply for the City of Saratoga Springs, derived from the Loughberry Lake reservoir. This was once a natural spring or drilled well, converted to municipal water in order to give justification for the pavilion.
Deer Park Spring (0.30% TDS, intermediate type, G in Figure 7) has a miniature pavilion that is unique among the Saratoga Springs. Curious ducks sometimes congregate to watch people collecting samples and making field measurements. Ducks are not, however, big fans of the mineral waters.

Museum of Dance on the right.

Lincoln Baths on the right
Lincoln Baths on the right, the only operating public bath house in Saratoga Springs. This is the location of the Lincoln Spring, for which there is a small monument. The spring has been turned off for many years, presumably to supply mineral water to the baths.

This road travels through a remarkable tunnel made of pine trees.

The "Bottling Plant" is on the right (historic; now a car museum).

STOP 3. SPA PARK SPRINGS AND OTHER WATERS. (2 hours)
This is also a walking stop at which we will see all of the flowing springs in the Saratoga Spa Park and much of the park itself including the Saratoga Performing Arts Center (SPAC), the new antique car museum ("Bottling Plant"), and Geyser Brook, the valley for which hosts several of the springs.

State Seal Spring
(0.030% TDS, fresh water, A in Figure 8) is by far the most popular spring in the area. There are usually droves of people filling up jugs, canteens, and carboys to bring home for drinking. Little do they realize that this fresh water is completely unrelated to the mineral springs, but in fact is pumped from a nearby shallow well in unconsolidated glacial outwash. The proximity of the well to the surrounding golf course causes this water to have the highest NO3- concentration (24 ppm) of any water sampled in this project by a factor of ∼50. Doubtless there are organic residues as well, though these were not analyzed by us. The water is tested periodically and is safe for drinking. This water is also piped to a State Seal Spring outlet to the south, that we will see later today.

Geyser Spring
(0.51 % TDS, intermediate type, B in Figure 8). The original location of this spring was just west of the "Bottling Plant" across the street. As of this writing it is unclear if this water will be piped to the new pavilion at the new location of State Seal Spring. If so, you will see that people are not lining up to fill jugs with this real mineral water.

Charlie Spring
(0.54% TDS, intermediate type, C in Figure 8) is the only spring water accessible from within the gates of SPAC. When it is flowing, it comes out on both sides of an ornamental wall. Note trace fossils on the stone slabs surrounding this spring.

Orenda Spring
(1.18% TDS, saline type, D in Figure 8) Orenda is the most boisterous of the mineral springs, enthusiastically belching forth CO2 and traces of H2S. This is also the most energetic precipitator of travertine, the amount of travertine deposition being estimated at ∼2000 kg/year. The travertine mound is mostly along the slope down to Geyser Brook, on the banks of which it forms a fine travertine platform along which a path goes.

Island Spouter Spring
(1.05% TDS, saline type, E in Figure 8) is the most enthusiastic geyser among the flowing springs, spouting these days up to ∼7 m high. Originally drilled in the early 1900's for CO2 production, it was turned into an ornamental island in Geyser Brook in 1930. The ornamentation has either been eroded away or has been buried beneath the nearly 2 m of travertine deposited by the spring in the past 70 years. This spring usually runs through a cycle of varying geyser height that lasts perhaps 30 seconds, terminating in a short period of gas venting without water. Sampling this spring is a wet experience. Figures 9 and 10 show how the appearance of island Spouter has changed as erosion has taken its course and as nearly 2 m of travertine has been deposited.

Hayes Spring
(1.14% TDS, saline type, F in Figure 8) is the third most saline spring after Hathorn 3 and Orenda.

Polaris Spring
(0.42% TDS, intermediate type, G in Figure 8) is a gentle continuous geyser typically spouting ∼1 m high.

Ferndell Spring
(0.44% TDS, intermediate type, H in Figure 8, BEWARE OF POISON IVY) flows to the surface without a notable geyser. It has deposited a broad, low travertine mound that is the third most notable after those of Orenda and Island Spouter.

Coesa Spring
(not sampled, presumably saline type, I in Figure 8) bubbles to the surface beneath a small pond. Though it is occasionally seen to be connected to a hose, the pond and fence have so far discouraged sampling.

Hathorn 3 Spring
(1.81% IDS, saline type, J in Figure 8) is the most saline spring currently available. It is an enthusiastic belcher that has deposited a modest travertine sheet on the floor of its pavilion and downstream. Crystalline calcite crusts can commonly be found floating on the surface of quiet pools of the spring water, precipitated as CO2 degassing raises the pH and supersaturates the water with respect to calcite. This spring, which has a total dissolved solids content ∼½ that of sea water, is the worst tasting and smelliest of the currently flowing springs. This spring is used in most of our modeling as being the most saline water end member (D), though modeling suggests that more saline waters can exist at depth. In our data set of old spring analyses, the Hathorn 2 spring was probably equivalent to Hathorn 3 in salinity, and the Glacier Spring was slightly more saline. Neither of these two springs are currently flowing and there are no complete analyses of them.