NYSGA 2018
NEW ENGLAND INTERCOLLEGIATE GEOLOGICAL CONFERENCE
110TH ANNUAL MEETING
NEW YORK STATE GEOLOGICAL ASSOCIATION
90TH ANNUAL MEETING
Guidebook for Fieldtrips in New York and Vermont
A meeting in honor of Bruce Selleck and William Kidd
Edited by
Tim Grover
and
Helen Mango
Hosted by Castleton University and Colgate University
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Geology Program Coordinator
Natural Sciences Department
Castleton University
Castleton, VT 05735
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A1: MIGMATITES OF THE EASTERN ADIRONDACK MOUNTAINS: NEW CONSTRAINTS ON THE TIMING, PETROLOGY, AND TECTONIC SETTING OF PARTIAL MELTING
Michael L. Williams
Claire R. Pless
Kaitlyn A. Suarez
Department of Geosciences
University of Massachusetts
Amherst, MA 01003
Timothy W. Grover
Department of Earth and Atmospheric Sciences
University of Northern Colorado
Greeley, CO 80639
Sean P. Regan
University of Alaska Fairbanks
Fairbanks, AK 99775
Graham B. Baird
INTRODUCTION
The Adirondack Mountains of New York are a classic example of a high-grade, polydeformational terrane that has been used as an analog for mid/deep crustal continental collisional and extensional tectonism (Mezger, 1992; Selleck et al., 2005; Rivers, 2008). Numerous studies have characterized the nature and grade of metamorphism and deformation (Bohlen et al., 1985; Spear and Markussen, 1997; Storm and Spear, 2005). Many rocks show evidence for significant degrees of partial melting, and it seems likely that, at least in some rocks, significant amounts of melt have been lost from the local system. Previous workers have interpreted the melting to have occurred during one (or several) orogenic or thermal/magmatic events, but in many regions, questions remain about the timing, melt producing reactions, tectonic setting, and even the number of partial melting events. In order to interpret the tectonic history of the region and use the region as an analog for modern deep crust, it is critical to constrain the timing, setting, and rheologic implications of melting. Migmatitic rocks in the eastern Adirondack Mountains can be very similar to one another in appearance. They tend to be gray-colored gneisses with layers of garnet-rich Qtz-Fsp-Sil-Bt gneiss interlayered with pink or white Qtz-Fsp leucosome with or without garnet. Leucosomes range from centimeters to a meter or more in width. The proportion of K-feldspar and plagioclase in both leucosome and gray gneiss varies widely. Many of the leucosomes are interpreted to be the product of in-situ melting although some have been interpreted to have been injected from outside of the local system (Bickford et al., 2008). Distinctive garnet- and sillimanite-rich gneisses with little biotite, “khondalite”, are relatively common, and are interpreted to be restitic rocks, where a significant amount of melt has been lost from the local rock.
The timing of melting is one of the critical outstanding questions in the Adirondack Mountains. Different rocks in different regions have been interpreted to have melted during one (or several) of four possible events: the 1190-1160 Shawinigan orogeny, the 1160-1140 AMCG intrusive event, the 1090-1050 Ottawan orogeny, or 1050- 1030 extensional tectonism. In some regions, cross-cutting or contact relationships constrain the timing of melting although the interpretation is rarely unambiguous. In general, the gray gneisses and leucosomes form sub-parallel layers, and timing constraints tend to come from geochronology, particularly of zircon and monazite. Even though the migmatitic paragneisses can contain abundant zircon and monazite, interpretations of melting relationships are commonly constrained by relatively subtle textures and compositions of chronometer phases.
Multiscale compositional mapping combined with high-spatial-resolution (micron-scale), in-situ geochronology and geochemistry of monazite can provide significant insight into the nature of melting reaction(s), the timing of melting, the relationship to deformational events, and the ultimate significance of companion zircon geochronology (Williams et al., 2017). This field trip will visit localities in the eastern Adirondack Mountains where we have applied this technique to samples of migmatite. The results suggest that melting did indeed occur at different times, producing rocks of very similar appearance. Although the relative intensity of the different events may have varied regionally, the degree of melting and melt segregation also played a role in controlling how fertile a rock was for melting in a subsequent event. The results presented here have a number of implications for the tectonic history of the region and for the rheology of the deep crust in general. They also provide a template for future studies in this region and in other regions in order to compare and interpret the tectonic setting of migmatitic rocks.

A1: MIGMATITES OF THE EASTERN ADIRONDACK MOUNTAINS: NEW CONSTRAINTS ON THE TIMING, PETROLOGY, AND TECTONIC SETTING OF PARTIAL MELTING
M.L.Williams, C.R.Pless, K.A.Suarez, T.W.Grover, S.P.Regan, G.B.Baird

STOP 1. North Pond/Swede Pond (Rt. 8, Hague, NY)
Stop 1a. Dixon schist-marble-quartzite-khondalite
(Description partly taken from McLelland et al., 2002). Rocks along Rt. 8 are dominated by metamorphosed sedimentary rocks; partially melted metapelitic rocks are particularly common. Bedding and a strong early foliation are shallowly dipping (east or west). Recumbent isoclinal folds have been described in a number of outcrops, including those at Stop-1. McLelland et al. (2002) described a large recumbent isoclinal fold at the southwestern end of the stop-1 outcrop. The fold axis plunges gently to the east or west and has been folded about upright E-W axial planes.
Rocks directly across from the pull-off are dominated by Dixon schist and marble. Dixon schist is a rusty slabby sulfidic rock that varies from garnet-bearing impure quartzite to garnet-rich schist. The more schistose varieties are similar to sillimanite-garnet-quartz-feldspar (khondalite) gneiss that we will see throughout the trip. Of particular note here is the strong mineral lineation on shallow, east-dipping foliation surfaces. The lineation plunges 20° = 090° on a foliation oriented 012°,21°.
Both Dixon schist and the khondalite were mined for flake graphite during the early part of the 20th century. The region around Swede Pond is known as the Dixon National Forest and the former mining hamlet of Graphite is located just to the East.
Stop-1b – Khondalite
Walk westward (carefully – traffic moves very fast!) along the road to the outcrops near the sign: “Parking Area 800 Ft.”). Most of the rocks along the way consist of shallowly-dipping Adirondack marble. This is a highly recrystallized, primarily calcite marble. Although the rocks are almost completely recrystallized, there are a number of very heterogeneous folds varying from isoclinal to open.
Just west of the “Parking Area 800 Ft.” is an excellent exposure of khondalite gneiss. The name “khondalite” originated in India, and refers to sillimanite-garnet-quartz-feldspar gneiss with notably little biotite. Sample 16TG-154 (Williams et al., this guidebook; Williams et al., in review) came from this outcrop. Here, the rock contains biotite, garnet, quartz, sillimanite, apatite, rutile, zircon, and monazite. Much of the biotite is interpreted to be a retrograde phase. K-feldspar is extremely abundant making up much of the matrix of the sample. It occurs in continuous layers wrapping around garnet, forming strain shadows near garnet, and locally occurs as sigma-style porphyroclasts. Importantly, plagioclase is absent in the sample (Fig. 4).
The lack of peak biotite and plagioclase and the abundance of garnet and K-feldspar suggests that the following were important melting reactions.
Bt + Pl + Als + Qtz = Grt + Kfs + melt (1)
Bt + Als + Qtz = Grt + Kfs + melt (2)
Some initial melting may have been associated with muscovite dehydration, but modeling suggests that the amount of melting was probably limited (Storm and Spear, 2005; Yakumchuk and Brown, 2014). The complete lack of plagioclase suggests that Reaction-1 was exceeded. Peak conditions are interpreted to be in the range of 0.6-0.8 GPa, 700-800 °C based on phase relationships Storm and Spear, 2005; White et al., 2007; Yakumchuk and Brown, 2014) and also on calculated P-T conditions (Bohlen et al., 1985; Essene, Storm and Spear, 2005). We suggest that P-T conditions were within the region shown on Fig. 3.
The lack of plagioclase in sample 16TG-154 indicates that Reaction-1 was not significantly reversed during melt crystallization and retrograde metamorphism. The abundance of K-feldspar and garnet, similarly, suggests that Reaction-2 was not significantly reversed. Sample 16TG-154 does contain biotite, but at least some of this biotite may reflect late biotite growth associated with fluid influx long after melting. Because of the lack of reversal of the melting reactions and the lack of plagioclase-bearing leucosome, it seems likely that a significant portion of the melt component was removed from the local rock.
Monazite results. Figure 9 shows calculated monazite dates for sample 16TG-154 (see also Fig. 5). Figure 9 b,c shows Y-content in monazite vs. calculated date for the same sample. Arrows (Fig. 9c) connect core and rim analyses from single monazite grains. Figure 9d shows the sum of heavy rare earth elements (HREEs) in monazite vs. calculated date. HREEs and Y are strongly partitioned into garnet. The characteristic ‘U-shaped’ profiles in Figure 9 b,c,d are interpreted to result from significant garnet growth at ca. 1150 Ma and garnet breakdown after ca. 1000 Ma.
Figure 9e shows U content in monazite vs. calculated date. Like Y and HREEs, U decreases dramatically prior to 1150 Ma. However, unlike Y and HREEs, there is no late-stage increase in U. Partitioning data from Stepanov et al. (2012) indicate that the actinides (U, Th) have positive monazite/melt fractionation, but U has a significantly lower ratio than Th and most REEs. During partial melting U and other trace and REEs will be partitioned from the whole rock into melt. The decrease in U in monazite in sample 16TG-154 is thus, interpreted to result from partial melting of the sample. The fact that this reduction occurred at the same time as the Y and HREE reduction is consistent with melting by reactions 1 and 2. The fact that the U content of monazite remains low during cooling and during subsequent events is taken as evidence that a large component of the melt was removed from the local system.
Importantly, sample 16TG-154 shows little evidence for garnet growth or melting at ca. 1050 Ma, suggesting that little or no melting occurred in this sample during the Ottawan orogeny.
M.L.Williams, C.R.Pless, K.A.Suarez, T.W.Grover, S.P.Regan, G.B.Baird

STOP 2. Treadway Mountain
A large outcrop of biotite-quartz-feldspar ± sillimanite metapelite occurs along the north side of Rt. 8. Leucosome layers are abundant, ranging from centimeters to one meter in thickness. Similar to the North Pond outcrops, the foliation is shallowly dipping. Mineral lineations are poorly developed (preserved), but where present, trend to the east.
Several samples have been collected from this outcrop, including 16Tg151-a, b. In addition, this is Locality-9 from Bickford et al., 2008 (see below). The samples contain the common assemblage K-feldspar, garnet, quartz, sillimanite, with apatite, rutile, zircon, and monazite. Biotite is relatively abundant compared to samples from the previous stop. Most of the biotite is interpreted to be a retrograde phase, but some biotite may have been present at peak conditions. Samples contain both plagioclase and K-feldspar. The feldspars are fine to medium grained and are dispersed through the rock. Most samples have plagioclase richer and poorer layers, but distinct leucosome segregations with sharp boundaries are not present in the analyzed samples.
Monazite Results: Monazite date-composition relationships in sample 16TG-151 (Fig. 10) show some distinct differences from the North-Pond samples. Both Y and HREEs decrease significantly at ca. 1150 Ma, consistent with a significant period of garnet growth at this time (Fig.10 b, c, d). Yttrium and HREEs increase after 1000 Ma, consistent with garnet breakdown. However, there is considerably more variability in monazite composition at ca. 1050 Ma than in samples from Swede Pond. Importantly, several monazite grains show a distinct decrease in Y and HREEs at ca, 1050 Ma. This is interpreted to indicate a second period of garnet growth.
Uranium in sample 16TG-151 shows fairly little change from 1150 through 1050 Ma (Fig. 10e). There is some evidence for an averaging effect (that is, grain-to-grain variability decreases in younger monazite) and possibly a slight decrease in U is apparent after 1000 Ma. This is distinctly different from the trend in 16TG-154 where U was dramatically depleted at ca. 1150 Ma.
Two grains (M-3 and M-19) illustrate the behavior in this sample particularly well (Fig. 11, 12)). Monazite grain #3 has eight distinct domains (Fig. 11). HREEs drop at ca. 1150 from the innermost core to the next domain. HREES are relatively constant to the outermost core domain (1060 Ma) Then, there is a second decrease in Y and HREEs at approximately 1060-1050 Ma. This decrease is taken to be a second period of garnet growth at ca. 1050 Ma. The relatively constant Y and HREEs in grain 3 from 1150 to 1060 suggests that there was relatively little garnet growth (or breakdown?) during this 100 m.y. period. Grain 19 has no domains in the 1150 Ma range, but has at least nine domains that are ca. 1060 and younger (Fig. 12). This grain documents the progressive decrease in Y and HREEs at ca. 1050. The fact that U also decreases at this time is taken to indicate further melting, probably by a reaction such as reaction 1 and/or 2.
M.L.Williams, C.R.Pless, K.A.Suarez, T.W.Grover, S.P.Regan, G.B.Baird

STOP 3. Roadcut on NY-74, west of Ticonderoga, NY
The easternmost rock at this outcrop is a poikiloblastic garnet amphibolite that contains garnet, clinopyroxene, plagioclase, and hornblende. Fe-Ti oxides are present as well, although they can only be seen in thin section. The garnet crystals can reach upwards of 3 cm in diameter. The best examples are concentrated in the westernmost part of the garnet amphibolite. Schistosity, mostly defined by aligned hornblende, appears to both traverse and wrap around the garnet poikiloblasts, suggesting garnet growth was syn-kinematic. Geochemistry of the amphibolite (Fig. 13) is consistent with other metamorphosed mafic members of the AMCG suite, constraining intrusion of the protolith to ca. 1160-1140 Ma.
Further west along the outcrop is a transition zone, marked by a vegetated slope, containing some layers of marble and a rusty weathering paragneiss. West of the transitions zone the vast majority of the outcrop is composed of a garnet-biotite-plagioclase-perthitic microcline-quartz ± sillimanite migmatitic paragneiss. Leucosome proportion is estimated to be 30-50%. Leucosomes preferentially contains garnet, while biotite is concentrated in the melanosome. Both leucosome and melanosome layers are deformed and define a foliation parallel to a schistosity defined by aligned biotite. Quartz and feldspar are granoblastic in both the leucosome and melanosome.
Within the migmatitic paragneiss are a number of garnet rich boudins. These are possibly deformed and metamorphosed dikes. The boudins indicate stretching in all directions parallel to gneissic layering suggesting a flattening finite strain. Also present within the migmatitic paragneiss are bodies of amphibolite similar to what is present at the east end of the outcrop. Three, 5 to 30 meters wide, unstrained, amoeboid-shaped, pegmatitic granitic plutons also occur in the migmatitic paragneiss.
The pegmatites contain feldspar, quartz, and biotite and are likely an alkali-feldspar granite, but some feldspars are striated. Locally associated with the pegmatites, the host migmatite contains cm-size books of graphite.
The fabric is essentially parallel among the garnet amphibolite, the migmatitic paragneiss, and the transitional contact zone (Fig. 14). The in the eastern half of the outcrop the foliation is ∼062, 43 SE but changes to ∼103, 45 SW by the westernmost end of the outcrop. A mineral lineation, defined by aligned hornblende, biotite, or sillimanite, depending on location, predominately plunges gently to the ESE (∼20 → 100). This fabric is identified as the second regional fabric, forming after AMCG intrusion. Asymmetric tails on garnet poikiloblasts in the garnet amphibolite suggest top to the west transport.
Samples of the migmatitic paragneiss and of pegmatite were processed for U-Pb zircon geochronology (see also Regan et al., 2015). Presented here are preliminary results from a new sample that is solely of a structurally late migmatite leucosome that connects to adjacent leucosome boudinaged by S2.
Two distinct generation of zircon can be identified in this sample by both CL and geochronology. Oscillatory zoned zircon with a bright CL response are found as cores within grains or forms entire grains. The most concordant, best cluster of ages for this zircon type provide a weighted 207Pb/206Pb age of ca. 1177 ± 10 Ma (Fig. 15), interpreted to reflect zircon growth during Shawinigan migmatization, consistent with previous results from this outcrop. Oscillatory zircon with a dark CL response is found as rims surrounding the bright zircon or as entire grains. The oldest, concordant, tightly grouped analyses of this zircon produce a Concordia age of ca. 1067 ± 6 Ma (Fig. 15), which is taken to be the age of a second, Ottawan-aged migmatization event and formation of the penetrative S2 fabric.
One pegmatite body produced uniform oscillatory or sector zoned zircon with a dark CL response. Nine analyses of this zircon are concordant or near concordant and produce a Concordia age of 1015 ± 10 Ma (Fig. 16), thought to best represent the age of pegmatite formation.
We have completed analyses of monazite from several samples from the migmatite. The overwhelming majority of monazite are ca. 1050 Ma and younger (Fig. 17), with only one small core domain yielding ca. 1150 date. Decreasing Y for some monazite grains (Fig. 18), the presence of ca. 1050 Ma monazite inclusions in garnet, and given the zircon data, demonstrate most monazite crystallization occurred late during or following Ottawan leucosome crystallization. Increasing Y in 1030 Ma and younger monazite supports garnet breakdown and retrograde metamorphism beginning at this time, as seen throughout the region. With the zircon data suggesting an earlier period of melting during the Shawinigan orogeny, it is likely that virtually all earlier monazite was consumed during this early or subsequent melting events. No age pattern is convincingly evident in the monazite U concentration, suggesting melt largely did not escape during either melting event (Fig. 18). Figure 15. Concordia diagram for zircons isolated from the late-leucosome analyzed by LA-ICP-MS. Green analyses best date Shawinigan migmatization (weighted average given), blue analyses best date Ottawan migmatization (Concordia age). Representative CL images and spot ages shown.

STOP 4. Dresden Station: Garnet-sillimanite gneisses and coronitic metagabbro
Many field trips have stopped at this location over the years. This description comes partly from Grover et al. (2015). McLelland et al. (1988a) used this outcrop to show that there were multiple phases of metamorphism recorded by the rocks in the Adirondacks. A sharp contact between garnet-sillimanite gneiss and a coronitic metagabbro is well-exposed in this outcrop (Fig. 19). The gneiss is a garnet-sillimanite-plagioclase-K-feldspar-quartz (khondalite) gneiss with a small amount of biotite. The khondalite is associated with quartzite and calcsilicate rocks as in many other parts of the Adirondacks.
The garnet-sillimanite gneiss is penetratively deformed with a well-developed foliation and a lineation that plunges gently to the east (21° → 091°). Much of the metagabbro is undeformed and a coarsely crystalline texture is preserved throughout much of the unit. There are places within the gabbro however that are deformed and foliated. Adjacent to the khondalite contact, the gabbro is fine grained (Fig 19c). It appears to coarsen toward the interior of the body, suggesting that a chilled margin formed at the contact. The contact between the gneiss and the gabbro is, in places, at a high angle to the foliation in the gneiss. These field relationships suggest that the gneiss was deformed and metamorphosed prior to the intrusion of the gabbro (McLelland et al., 1988).
Coronitic metagabbro bodies are found throughout the Adirondacks (Whitney and McLelland, 1973; Whitney and McLelland, 1983; Regan et al., 2011). Figure 20 illustrates the typical corona texture with an olivine core, surrounded by a rim of orthopyroxene, which in turn is rimmed by symplectic intergrowths of garnet and clinopyroxene. Ilmenite is typically rimmed by Ti-rich hornblende. Clouding in the plagioclase is due to microscopic spinel inclusions. This mineral assemblage and texture developed via subsolidus recrystallization from a rock that was originally an olivine-clinopyroxene-plagioclase gabbro. Preliminary P-T estimates suggest this mineral assemblage developed at approximately 9 kb and 750 °C. The growth of amphibole also requires an influx of an H2O-bearing fluid.
To the south and west of this outcrop of gabbroic rocks, ferrogabbroic rocks and anorthosites, all thought to be associated with the metagabbro, have a very strong penetrative fabric and in many locations are mylonites. These rocks contain garnet, clinopyroxene, plagioclase, and local orthopyroxene, similar to the mineral assemblage found in the coronitic metagabbro. Kinematic indicators in these rocks suggest a west-directed thrust sense of motion. This period of deformation and metamorphism must postdate the emplacement of the AMCG rocks, but it is not clear if it reflects the late Shawinigan orogeny or the Ottawan orogeny. McLelland et al. (1988b) reported a U-Pb zircon, multigrain age of 1144 ± 7 Ma for the metagabbro. This is interpreted as an igneous crystallization age, consistent with the gabbro belonging to the AMCG suite. However, Aleinikoff and Walsh (2015) report a SHRIMP age of ca. 1134 ± 3 Ma which suggests emplacement was during the waning stages of, or slight after the Shawinigan Orogeny.
We dated monazite from several samples of the garnet-sillimanite gneiss. The data fall into three broad populations (Fig. 6). The oldest population has a weighted mean of 1179 ± 9 Ma. We suggest that this represents the timing of prograde Shawinigan metamorphism. The next population yields ages that cluster around 1151 Ma. This is interpreted to represent the peak of metamorphism and melting driven by a thermal perturbation resulting from the intrusion of the gabbroic rocks. Most of the remaining monazite dates are 1020 Ma or less. These are too young to correlate with the proposed timing of the peak of the Ottawan Orogeny (∼1090-1050 Ma). Instead, this generation probably represents post-Ottawan decompression and uplift.
The following model is consistent with the field observations and data from this outcrop. The mineral assemblage and the fabric in the garnet-sillimanite gneisses formed during the Shawinigan Orogeny. The gabbro was emplaced at approximately 1150 Ma, after the prominent fabric developed in the gneisses. This time frame is consistent with both the monazite dates from the gneisses and the multigrain age from the gabbro. Following the emplacement of the gabbro there was another period of deformation and metamorphism. This is when the nearby AMCG rocks were deformed and metamorphosed. This is also when the coronitic texture developed in the metagabbro. This requires in influx of an H2O–bearing fluid. Unlike the nearby rocks, the coronitic metagabbro was not pervasively deformed at this time nor did any new monazite grow in the garnet-sillimanite gneiss. Perhaps strain was partitioned around this outcrop and the lack of strain resulted in little recrystallization in the largely anhydrous garnet-sillimanite gneisses. The 1020 Ma and younger monazite ages in the garnet-sillimanite gneisses record monazite growth during post Ottawan extensional collapse.

STOP 5. Lineated Sillimanite Gneiss and the East Adirondack Shear Zone
Walk south (left) along Rt.22 to the first set of low outcrops.
Stop 5a. Sillimanite Gneiss. This outcrop contains interlayered garnet-biotite-sillimanite gneiss with greenish calcsilicate layers. The well exposed foliation surfaces of the sillimanite-bearing gneiss commonly have well-lineated, coarsely crystalline sillimanite. The lineation plunges gently to the southeast (18° → 130°). This lineation is distinctly different than that at the previous stops. It may represent a completely different stage of the tectonic history. These rocks are interpreted to have been affected by the East Adirondack shear zone (Wong et al., 2012), although kinematic indicators are better developed in outcrops just to the south (Stop 5b).
In-situ monazite analyses have been carried out as part of this work (Suarez et al., 2017) and also as part of the Wong et al (2012) study. Six distinct compositional populations (generations) have been recognized in the migmatitic metasedimentary samples: 1178 Ma, 1139 Ma, 1064 Ma, 1049 Ma, 1030 Ma, and ca. 1000 Ma (Fig. 8); see also Suarez et al. (2017). Uncertainties are on the order of 10-20m.y. (2σ). There is a distinct drop in Y and HREE between the ca. 1140 and ca. 1060 populations. An additional decrease occurs between the 1064 and 1049 Ma populations. Uranium also shows decreases in the 1060-1050 Ma range. Although work is still underway, there is little evidence in these rock for significant melting at ca. 1150 Ma. The data are consistent with the dominant melting event to have occurred at ca. 1060 Ma. The lowest Y and REE and U contents are associated with the ca. 1050 Ma population, suggesting that garnet growth and additional melting occurred at this time. It is possible that there was one prolonged melting event or alternatively melting may have occurred during the culmination of Ottawan shortening (1090-1060 Ma) and a second event associated with post Ottawan extension and intrusion of the Lyon Mountain granite (ca. 1050 Ma).
The 1030 Ma monazite population is interpreted to represent shearing associated with the East Adirondack shear zone (Wong et al., 2012). Y and HREE contents increase dramatically after 1050 Ma, probably reflecting garnet break-down during exhumation. The high(er)-Y overgrowths on monazite grains are commonly located along the foliation and in extensional quadrants of the grains, further supporting the connection between this monazite generation, extensional shearing, and exhumation (see Wong et al., 2012). This period of extension is synchronous with the extension along the Carthage-Colton shear zone (McLelland et al., 2001; Streepey et al., 2001; Johnson et al., 2004). Based on the extremely low Th-content, the youngest population (ca. 1000 Ma) is interpreted to represent monazite associated with fluid infiltration and hydrothermal alteration events.
Monazite is present in thin leucosome layers within the gray migmatitic gneiss and also in thicker leucosome layers in the outcrop. Monazite from a thin (∼ 1 cm) leucosome layer yielded dates 1050 Ma and younger. Because monazite is expected to dissolve into partial melt, i.e. most partial melts are undersaturated in monazite components (Kelsey et al. 2008; Yakymchuk and Brown 2014; Harley and Nandakumar, 2014), we interpret the ca.1050 date to be the time of crystallization of this leucosome layer. Monazite from larger, more discrete leucosomes also yielded ca. 1050 Ma dates. However, these monazite grains have much greater U-content than the gray gneiss or the thin leucosome. We suspect that these leucosomes represent injections of external partial melts that have fractionated and evolved to higher U contents.
We have carried out in-situ and bulk zircon (U-Pb) analysis of some rocks from this outcrop. Bulk separates from gray gneiss yielded only ca 1050 dates but in-situ samples of the gray gneiss yielded both ca. 1150 and ca. 1050 dates. We suspect that this reflects a sampling bias where only the larger monazite grains were recovered during mineral separation and these are dominated by the younger populations. An analysis of a zircon separate from the thin leucosome layer from which monazite was also analyzed, yielded both ∼1150 and ∼1050 dates. The older dates are interpreted to represent inherited zircon incorporated into the partial melt from the gray gneiss (see Suarez et al., 2017).
It is interesting to compare these data with that from stop 3. The monazite data from the stop 3 rocks had a strong Shawinigan, AMCG, and post Ottawan signature but virtually no peak Ottawan monazite growth. The monazite data from this outcrop have Shawinigan and Ottawan monazite populations, but the data suggest that melting occurred mainly during the Ottawan event(s).
Follow the outcrop to the south remaining on the east side of Highway 22 for the time being.
Stop 5b. East Adirondack Shear Zone. Note the interlayer folds in the upper part of the road cut on the east (Fig. 21). This exposure illustrates an older foliation that was transposed into a new foliation. Perhaps this is an example of a Shawinigan S1 transposed into a younger Ottawan or post-Ottawan S2. Across the highway is an asymmetric mafic boudin with a long tail that continues to the south in the upper part of the road cut (Fig. 22). Although it is clear that the boudin was involved in a period of ductile deformation, the road cut is oriented almost at right angles to the lineation direction so it is difficult to use the boudin for kinematic analysis.
The granitic rocks to the left of the mafic boudin in the picture below are mylonitic, LS-tectonites with megacrystic K-feldspar porphyroclasts Fig. 23). The prominent lineation plunges gently to the southeast. Data from this outcrop was cited by Wong et al. (2012) as evidence for the East Adirondack Shear Zone. They report a U-Pb zircon age obtained using the SHRIMP-RG at Stanford, of the granite in Figure 23 of ca. 1140 Ma. This date suggests that it is part of the AMCG suite and was emplaced towards the end or after the Shawinigan Orogeny.
Following this reasoning they suggested that most of the strain in the rock is therefore related to Ottawan shortening or post Ottawan extension. They examined the asymmetric K-feldspar porphyroclasts as kinematic indicators to document shear sense motion. Although they found some porphyroclasts that suggested top to the west, thrust sense motion more suggest top to the east, normal-sense motion. They concluded that this rock experienced both early shortening (west-directed thrusting) and later extension.
Optional: cross the road and return to the vehicles along the west-side outcrops. This side gives an excellent view of the east-side outcrops, and to the north, there are excellent examples of coarse biotite-sillimanite-feldspar gneiss.
Be careful crossing the road. It is busy and fast.

STOP 6. Mylonitic, Migmatitic “Straight” Gneiss
This is a beautiful exposure of a garnet-biotite-plagioclase-K-feldspar-quartz sillimanite gneiss. Most of the white layers are interpreted as leucosomes that formed during partial melting via a reaction such as Reaction-1. The strongly attenuated nature of the leucosomes, along with also remnants of pegmatite dikes that are now sheared into the foliation plane attest to the significant strain recorded by these rocks. The rocks have the same strong southeast trending, gently plunging lineation that was seen at the last stop. We suggest that some of the strain in these rocks may be the result of post-Ottawan extensional collapse in the East Adirondack shear zone.
This outcrop was part of a study by Bickford et al. (2008). They concluded that these rocks underwent partial melting at approximately 1050 Ma at the tail end of the Ottawan Orogeny. They suggest anatexis was facilitated by an influx of H2O-bearing fluids and decompression as a result of extensional collapse.
Figure 24 shows our preliminary U-Th-Pb electron microprobe monazite data from this outcrop. The data indicate monazite growth from approximately 1150 Ma through 1000 Ma with most monazite growth after 1100 Ma.
Post-1100 Ma monazite have distinctly lower yttrium content than ca. 1150 monazite. The data are consistent with garnet growth and melting associated with the Ottawan orogeny and post Ottawan extension in agreement Bickford et al. (2008). However, there is also evidence for Shawinigan/AMCG monazite growth, and some suggestion that some melting may have occurred at this time. Combining these results with those of the precious stop, we see evidence for significant partial melting during Ottawan/post-Ottawan time. The extent of melting during Shawinigan/AMCG time is less certain, but the evidence from monazite or from zircon is certainly much less compelling. The may partly due to the intensity of the ca. 1050 Ottawan metamorphism, melting, and deformation, but we suspect that the earlier melting was much less extensive. Because the rocks were clearly fertile for melting during the later event, we suggest that Shawinigan temperatures were not sufficient for extensive biotite dehydration melting. This may reflect decreasing temperatures toward the outer parts of the Adirondack dome (as originally suggested by Bohlen et al., 1985). However, the observation that Shawinigan melting was synchronous with the AMCG plutonism suggests that the larger distance to AMGC gabbro and anorthosite may have contributed to higher temperatures during late Shawinigan time. We are in the process of collecting more data in order to further explore the effects of the Shawinigan and Ottawan Orogenies and post Ottawan orogenic collapse on these rocks.

A2: GEOLOGY OF THE CARTHAGE-COLTON SHEAR ZONE AND LYON MOUNTAIN GRANITE AN ADIRONDACK FIELD TRIP IN HONOR OF BRUCE SELLECK
William H. Peck wpeck@colgate.edu
Martin S. Wong mswong@colgate.edu
Department of Geology
Colgate University
Hamilton, New York 13346
Eric L. Johnson johnsone@hartwick.edu
Hartwick College
Oneonta, New York 13820
INTRODUCTION
This field trip is in honor of Bruce Selleck, Thomas A. Bartlett Chair and Professor of Geology at Colgate University, who passed away unexpectedly on Monday, July 31 2017. Bruce had deep ties to the upstate New York region, growing up on a farm on the western edge of the Adirondack Park. Bruce received his AB in Geology in 1971 from Colgate University, and returned to his alma mater as a faculty member in 1974, after taking MA and PhD degrees at the University of Rochester. His scholarship initially focused on the Paleozoic sedimentology of New York, but as his career progressed Bruce later expanded the scope of his research to include study of the tectonic development of the Adirondack Mountains, with longtime colleague and friend Jim McLelland.
This field trip revisits some of the key localities first described by Bruce and his collaborators in the Carthage- Colton shear zone, a major structure in the southern Grenville Province that separates the Adirondack Highlands from the Adirondack Lowlands and underlies the crossroads of Selleck’s Corners, and the farm where Bruce grew up. This guide draws heavily on the Friends of the Grenville field trip to look at these rocks led by Eric Johnson and Bruce in 2005, Bruce’s papers on the Carthage-Colton shear zone and the syntectonic Lyon Mountain granite, and subsequent research by Bruce, the field trip leaders, and our colleagues.

A2: GEOLOGY OF THE CARTHAGE-COLTON SHEAR ZONE AND LYON MOUNTAIN GRANITE AN ADIRONDACK FIELD TRIP IN HONOR OF BRUCE SELLECK

STOP 0. Deformed AMCG Granitoids South Of Croghan
Extensive outcrops on both sides of NY-812
This stop, 1.2 miles south of the Citgo gas station, is included for field trip participants traveling from the south to the meeting place in Croghan as a ‘guide your own’ stop. Extensive outcrops of penetratively deformed and folded rocks are exposed on both sides of NY-812, showing a variety of granites, syenites, and amphibolites (Fig. 2). Lithologically these exposures are similar to the Diana syenite and other AMCG-suite granitoids elsewhere in the Adirondacks. These rocks are mapped as hornblende and biotite granites on the state geologic map.
In 2017 Bruce Selleck dated several rocks thought to be part of the Lyon Mountain granite suite to test the relationship with mineralogically similar, but deformed AMCG suite granitoids, including two rocks from this site (Fig. 3). The sample CRO-16-H is a deformed granitoid from this outcrop, and 29/31 analyzed euhedral zircon with oscillatory zoning yields a 207Pb/206Pb weighted mean age of 1176±10 Ma (MSWD=2.1), similar to other members of the AMCG suite. The sample CRO-16-PEG is coarse pegmatite dike that cuts fabric in the host syenite. Fourteen analyses of zircon with broad zoning in BSE from this sample have a 207Pb/206Pb weighted mean age of 1047±10 Ma (MSWD=1.7), which is identical to typical igneous ages determined in the Lyon Mountain granite elsewhere. Three older analyses overlap those of the host gneiss, and are likely inherited. Bruce Selleck (personal communication 2017) thought that this sample (lower left, Figure 2) might represent in-situ melting of AMCG suite granitoids to produce Lyon Mountain granite pegmatite.

Meeting Place

STOP 1. Intrusive Relationships on Indian River Road
This roadcut exposes dikes of leucogranite intruding mafic syenite. The location is south of the mapped extent of the Carthage-Colton shear zone and thus within the Adirondack Highlands. This outcrop and nearby exposures represent a zone of extension within the lower plate (Adirondack Highlands) of the CCSZ that accommodated granite emplacement. The undeformed quartz mesoperthite leucogranite dikes crosscut foliation in the syenite and contain xenoliths of the mafic country rock. The outcrop is also crosscut by later hematite-stained quartz-feldspar-calcite veins that resemble the late mineralized veins associated with the Lyon Mountain granite elsewhere in the Adirondack Highlands (Johnson and Selleck, 2005).
Zircon from the leucogranite are subequant to elongate with faintly zoned cores and finely-zoned rims (Selleck, et al, 2005). Four reliable core analyses yield an upper intercept age of 1195 ± 11 Ma (Fig. 5, MSDW = 1.3). Six analyses of zoned rims on yielded 1039 ± 10 Ma with MSDW of 0.63. The core ages are older than zircons dating intrusion of the Diana Complex (1164 ± 11 Ma) but overlap the ages of some older cores (1180–1190 Ma) in the Diana reported by Hamilton et al. (2004). The rims represent igneous overgrowths on xenocrysts, as observed in other studies of the Lyon Mountain Granite suite in the Adirondack Highlands (Chiarenzelli et al. 2018), and fixes the minimum age of penetrative deformation in this part of the Adirondack Highlands.

STOP 2. The Valentine Wollastonite Deposit
The Gouverneur Talc Company No. 4 Quarry is at the northern contact between the 1164 Ma Diana Syenite and marble country rocks, which exposes the spectacular blue calcite marble and coarse wollastonite of the Valentine deposit. This relatively quick visit to the deposit will examine representative lithologies dumped by the mine operators next to the entry gate, which is a working site.
Mining of wollastonite in the Valentine deposit began in 1977, and the deposit is well-known as a mineral collecting locality (Chamberlain et al. 1999). Calcite marbles (containing blue or white calcite) contain high-grade mineral assemblages including diopside, spectacular euhedral graphite, and other accessory minerals such as phlogopite, titanite, and chondrodite (Gerdes and Valley 1994). Secondary minerals include talc, prehnite, quartz, secondary vein wollastonite, and a host of other alteration minerals (Chamberlain et al. 1999). Visible in the dump next to the entry gate are also cemented crushed ore from the mine. This phenomenon was first noticed by Bruce Selleck in 2017 at the Lewis deposit, and now that we know what to look for, cementation of wollastonite mine wastes seems to be a common process. Crushed wollastonite rock is cemented by calcite and silica, probably by the carbon sequestration reaction wollastonite + CO2 = calcite + quartz occurring passively in wollastonite mine waste rock and old stockpiles. Current research on wollastonite weathering and carbon sequestration is ongoing.
Wollastonite is relatively common in marbles and calc-silicates of the Adirondacks, and often forms during closed-system metamorphism of protolith sediments via the reaction calcite + quartz = wollastonite + CO2. Large wollastonite ore deposits are found associated with 1155 Ma AMCG-suite rocks: around the Westport anorthosite dome in the eastern Adirondack Highlands (Fig. 1) and adjacent to the Diana Syenite at the boundary of the Adirondack Highlands and Adirondack Lowlands. Large wollastonite deposits require the infiltration of large volumes of water-rich fluids into the contact aureole, producing wollastonite via reaction with calcite country rocks, where fluids bring in dissolved silica and remove evolved CO2 (Gerdes and Valley, 1994).
The Willsboro-Lewis wollastonite district around the Westport dome has low oxygen isotope ratios (as low as - 1.3‰ SMOW), which is indicative of interaction with large volumes of heated meteoric water, and places important constraints on the depth of emplacement of the Marcy anorthosite in the Adirondack High Peaks. (Valley and O’Neil 1982; Clechenko and Valley 2003). Large volumes of surface water during contact metamorphism is strong evidence of shallow (10 km) anorthosite emplacement (Valley and O’Neil 1982). Interestingly, the main wollastonite ore of the Valentine deposit does not share the low oxygen isotope ratios of the Willsboro-Lewis wollastonite district (Gerdes and Valley 1994), nor does wollastonite at the contact of the the contemporaneous Morin or Lac St. Jean anorthosites of Quebec (Peck 1996; Higgins et al. 2001; Peck et al. 2005). This may indicate deeper emplacement in the crust or differences in hydrothermal flow for these other AMCG plutons.
The Valentine deposit shows steep gradients in oxygen isotopes ratios across skarn/hostrock boundaries, revealing that fluid infiltration was channelized and not pervasive (Gerdes and Valley 1994). Carbon isotope fractionations between co-existing calcite and graphite are consistent with equilibration at 675°C during overprinting Ottawan metamorphism (Gerdes and Valley 1994), although for the most part the rocks are not very deformed by this event.

STOP 3. Dana Hill Metagabbro
Park the vans and climb the hill on the north side of the road. On the trail up the hill we will pass several submeter width EVENT 4 shear zones.
The Dana Hill Metagabbro preserves multiple deformation and veining events ranging from granulite facies ductile to sub-greenschist facies brittle events (see appendix). In many cases, cross-cutting relationships allow for the determination of a sequence of events. To date, six major deformational/veining events have been identified (Johnson et al. 2004; Streepey et al. 2001). At this stop we will examine the complex deformational events recorded in the body. This outcrop along with the outcrops at the top of the hill across the road, exhibit all 6 deformational events. We will start at the far southern end of the outcrop and examine the deformational sequence of events recorded. From the oldest to the youngest, this outcrop preserves EVENT 1 mega-shearing, EVENT 3 hornblende veining, EVENT 4 sub-meter shearing, and EVENT 6 folding and brecciation. EVENTS 2 and 4 can be observed at the top of the hill across the road. Events 3 through 5 take place in the presence of a fluid or fluids that drive scapolite replacement of original plagioclase feldspar in the host metagabbro. In the Diana Syenite, cm to dm-wide shear zones (dated to 1052-1034 Ma) also exhibit scapolite replacement of plagioclase feldspar. This scapolitization event is widespread in and around the CCSZ from just north of Harrisville to Colton and is present in both the Highlands and Lowlands terranes, therefore, marks a common event for both terranes. Continuing north to the far end of the outcrop lone finds a weakly to un-deformed pod of cumulate gabbro.
The goal of this stop is to demonstrate that the Dana Hill Metagabbro body acted as a rigid block during deformation. In some instances, cumulate igneous textures have been completely preserved while in other exposures the gabbro is ultramylonitic in texture. The resistance to deformation in the Dana Hill Metagabbro resulted in an episodic response to the applied stress leading to discrete pulses of deformation. This body preserves individual and distinct events that record the much of deformational history of the region. The earliest shear zones are massive (30m wide) and mylonitic to ultramylonitic. These shear zones record recrystallization temperatures in excess of 700°C. Subsequent shearing events are dramatically different forming sub-meter wide anastomosing shear zones at recrystallization temperatures at or below 700°C. The last deformation events to affect the Dana Hill Metagabbro transition to brittle failure at low to sub greenschist facies conditions. The deformational history is one of an exhuming footwall with deformation beginning in the granulite facies and eventually passing through the brittle-ductile transition at greenschist to sub greenschist facies conditions. We will examine these events and the available geochronologic data for this complex outcrop.
EVENT 1 shearing accounts for the mylonitic character of the outcrop as a whole. The foliation here dips steeply yet transport lineation orientations plunge shallowly to the north-northwest. Kinematic indicators yield dextral shear sense. These mylonites contain recrystallized clinopyroxene + amphibole + sphene+ plagioclase (An45-51) along with accessory minerals (apatite, zircon, ±quartz). Amphibole compositions for these samples range from Ferroan Pargasite to Magnesian Hastingsite. Chlorine contents are high for all amphiboles studied ranging from 2 to 18% hydroxyl site occupation. Amphibole and plagioclase chemistries are presented in Johnson et al. (2004). All samples exhibit a well-annealed polygonal fabric with perfect 120° triple junctions between grains. Grain sizes show a narrow range of variation for these samples averaging in the range of 100-300 μm for polygonal plagioclase. Re-crystallization temperatures for event 1 samples using the quartz-free geothermometer Holland and Blundy (1994) range from 744 to 770°C (for 6 kbar). Scapolite replacement of plagioclase is not present in EVENT 1 shears at this location.
Hornblende veins cut the foliation at high angles in zone a. Hornblende veining belongs to EVENT 3 (we do not see EVENT 2 shear zones in this outcrop.). The hornblende veins are surrounded by reaction halos were scapolite replaces plagioclase feldspar in the host metagabbro. These halos can extend several mm into the surrounding metagabbro. In zone b (see Fig. 7), the metagabbro is folded and the open to nearly chevron folds are marked by EVENT 3 hornblende veins. What looks like a rotated cleavage fanning across the folds are in fact the old EVENT 1 mylonitic foliation surfaces. This zone transitions into the chaotic breccia zone (EVENT 6; zone c). Brecciation (EVENT 6) was accompanied by the growth of actinolite, biotite and chlorite after hornblende, and the breakdown of scapolite to a mixture of albite, epidote, and calcite Breccia sample H-6A preserves rafts and clots of scapolite-rich mylonitic metagabbro with an invasive matrix of fibrous mats of chlorite, epidote, and actinolite. Hornblende (ferroan pargasite) that has not suffered alteration to actinolite is fluorine-rich (average F = .75 wt %; average Cl = .57 wt %).
At the eastern margin of the breccia zone, an EVENT 4 sub-meter scale shear zone is exposed. This shear (sample CR-3-87 Johnson et al. 2004; Streepey et al. 2001) preserves deformation textures (little annealing) and contains the assemblage hornblende + recrystallized clinopyroxene + plagioclase An 32 + Fe-Ti oxides + scapolite (minor). Plagioclase-amphibole equilibrium pairs where present yield re-crystallization temperatures for event 4 shearing in the range of 730°C to 680°C ±50°C for a pressure of 6 kbar.
GEOCHRONOLOGY OF STOP 3
Figure 8 shows the U/Pb isochron data for shear-zone grown sphene (titantite) for this outcrop. U/Pb data presented represent EVENTS 1, and 4 and yield a tightly constrained age of 1020.7 ± 3.1 Ma. Since recrystallization temperatures for events 1 through 4 occur at temperatures above the closure temperature for U/Pb in titanite, these dates represent cooling ages for the body. The consistency of U/Pb titanite ages for samples throughout the body indicate that all (Events 1-5) shearing and veining occurred prior to 1020 Ma.
The 39Ar/40Ar results for hornblende in these samples is presented in figure 9. These data mark the date at which these samples cooled through the 500-550°C closure temperature for hornblende. The data from this and outcrop A-4 (opposite side of the road) are quite interesting. The results yield two cooling ages: one at ∼985-1000 Ma and a second (recorded in two shear zones) of ∼935-940 Ma. The latter and younger ages were determined from two samples at this outcrop. The 945 Ma age was used by Streepey et al. (2001) to constrain the timing of the last stage of movement along the CCSZ. Only one of these samples (CR-2A2) yields a statistically clear plateau. Both shear zones are overprinted by the later brecciation event and, therefore, may have suffered some Ar loss during this event. Conversely, hornblende overgrowths on undeformed cumulate textured Dana Hill Metagabbro sample from the outcrop across the road also yield a 945 Ma age indicating hornblende growth at this time. Whether or not the 945 Ma age represents renewed deformation remains a point of controversy. The bulk of the shear zones and veins studied in the Dana Hill Metagabbro (and surrounding Diana Syenite body) record 39Ar/40Ar hornblende cooling ages of 985-1010 Ma. The generally flat spectra for39Ar/40Ar data indicate rapid cooling with little to no post closure disturbance.
The Dana Hill Metagabbro outcrops visited today provide an overview of the deformational and thermal history associated with movement along the CCSZ. Since this body is located at the CCSZ detachment, it records the entire deformational history. During exhumation of the footwall (Highlands), the width of the deformation zone narrows with falling temperature and confining pressure and eventually, deformation is confined to regions directly adjacent to the detachment surface. Due to its the resistance to strain, preexisting deformational fabrics in the Dana Hill body were not completely overprinted, leading to the complexly deformed body that we see today. Deformation in the Dana Hill can be broken down into three distinct regimes with falling temperature and pressure: 1. mega-shearing, 2. sub-meter width shearing, 3. brittle failure. Early veining episodes may have been driven by fluid infiltration into the body (driving scapolite and high Cl hornblende-forming reactions). The origin(s) of these fluids (CO2 and HCl/NaCl rich) is unknown, but a likely source is from exhalation/mobilization of evaporite deposits in the adjacent lowlands hanging wall block. This origin for the metasomatic fluids fits well with the observed and widespread scapolite veining and scapolite replacement of original plagioclase feldspar in the lowlands near Pierrepont (Tyler, 1980; Selleck , pers. comm.).

STOP 4. Lyon Mountain Leucogranite, Mylonite, and Ultramylonite at Brouses Corners
Dark mylonitic gneiss and coarse calcsilicate granulite are intruded by coarse, pink leucogranite pegmatite at this locality (Figure 10). Strong foliation in the mylonitic gneiss and local bands of ultramylonite dip gently northwest with a prominent NW-trending lineation. The pegmatite contains xenoliths of mylonitized wall rock and pegmatite veins invade and crosscut mylonitic foliation. However, the pegmatite shows mylonitic deformation along portions of its intrusive margins. One m-scale mass of pegmatite forms a large strain sigmoid indicating deformation consistent with a top-NW sense of shear. These field relationships suggest that the pegmatite was intruded during the waning phases of mylonitization and provides a key age constraint on deformation. Pseudotachylite veins crosscut mylonitic foliation but do not apparently crosscut granite pegmatite. Note the quartz veins within the tabular pegmatite masses; the veins appear to have formed as extension fractures within the pegmatite. Some quartz veins are strongly mylonitic.
Zircons from sample BC-PEG, a coarse, undeformed pink granite pegmatite, are equant, and faintly zoned in CL and BSE images (Figure 11). These analyses (based on 9 spot ages) yield a mean age of 1044 ± 7 Ma (Selleck, et al. 2005), which overlaps within error the ages of leucogranite at Indian River Road and Selleck’s Corners. This age fixes the timing of mylonitic strain, as the pegmatite intrudes mylonitic rock but is itself mylonitized along its margins. The pegmatite does not contain pseudotachylite veins, suggesting that these veins developed prior to pegmatite intrusion, although no unequivocal crosscutting relationships have been observed.
Petrographic and electron backscatter diffraction (EBSD) analyses provide further insight into the kinematics of this late Ottawan deformation. Kinematic indicators in the mylonites and ultramylonites include sigma-type porphyroclasts and a strong oblique grain shape fabric that is well developed in quartz-rich mylonites (Figures 12 and 13). These kinematic indicators consistently indicate a top-NW sense of shear. Petrographic observations suggest that quartz largely was dynamically recrystallized by sub-grain rotation (SBR) recrystallization, which typically occurs between 400-500°C (Stipp et al., 2002). EBSD analyses of these mylonites yield strong lattice preferred orientations (LPO), with c-axes showing patterns that are a combination of single girdle and y-strain axis maxima (Fig. 13). This pattern suggests quartz deformation via a combination of rhomb and prism <a> slip, which is also consistent with deformation temperatures of 400-500°C (e.g. Law, 2014), matching the petrographic data. Axes patterns also illustrate a sense of rotation consistent with a top-NW sense of shear consistent with other kinematic indicators. Taken together, these data are consistent with the interpretation of extensional motion on the CCSZ during late Ottawan time and that deformation continued through cooling to upper greenschist-facies temperatures following peak metamorphism.

STOP 5. Syntectonic Lyon Mountain Granite With Quartz-Sillimanite Segregations
Numerous natural outcrops and small road cuts along Buck Pond Road expose quartz-microcline and quartz-mesoperthite leucogranite with common centimeter- to meter-scale quartz-sillimanite segregations, veins, and nodules. The leucogranite is locally gneissic with northwest-dipping banding. Quartz-sillimanite segregations (Figure 14) are commonly drawn out into elongate, cm-scale rods surrounded by equigranular quartz- mesoperthite granite. Sillimanite crystals form a mineral lineation in deformed segregations, but these oriented segregations are also crosscut by slightly later granitic veins that contain sillimanite. Sillimanite is often partially replaced by muscovite. Networks of quartz-sillimanite veins, quartz-sillimanite ribbons, and smaller quartz-sillimanite segregations and nodules are interpreted as representing progressively dismembered magmatic-hydrothermal features, forming via leaching of granite and granitic magma by Cl-rich, acidic fluids during granite pluton emplacement (McLelland et al. 2002). Equigranular, coarse, pegmatitic quartz-microperthite granite surrounds some quartz-sillimanite in strain sigmoids that indicate a top-down to the northwest transport. Outcrops of leucogranite south of Buck Pond Road are more massive and contain diffuse pegmatitic zones and quartz veins typical of the Lyon Mountain granite of the Adirondack Highlands. This quartz-sillimanite nodular granite is mappable for ∼ 12 km along strike to the southwest as a 1-2 km wide belt immediately southeast of the mapped extent of the CCSZ.
This outcrop was sampled for U-Pb SHRIMP zircon geochronology (Selleck et al. 2005). Zircons were separated from coarse-textured quartz-microperthite granite that forms a 10-cm-thick band irregularly intrusive into quartz-sillimanite nodule–bearing quartz-microcline gneiss. Nine age determinations on well-developed, oscillatory zoned rims (Figure 15) give a weighted mean age of 1046 ± 7 Ma. We interpret this result as documenting intrusion of 1046 Ma granite during active extension on the Carthage Colton shear zone at this site.

A3: STRUCTURAL AND STRATIGRAPHIC FEATURES OF THE TACONIC FORELAND, NW VERMONT
Adam Schoonmaker
Department of Geosciences
Utica College
Utica, NY 13502
adschoon@utica.edu
William S.F. Kidd
formerly of Geological Sciences
University at Albany
Albany, NY 12222
wkidd@albany.edu
INTRODUCTION
This trip will focus on some of the spectacular structural and stratigraphic features exposed in NW Vermont at the Highgate Gorge, and at Lessor’s Quarry, and the outcrop known as “The Beam”, on South Hero Island. The last two are Rolfe Stanley Memorial outcrops. The trip will cover the nearly continuously exposed upper Cambrian to Ordovician section in the Highgate Gorge. This section exposes the shelf to basin transition including bedded limestone and massive dolomite, a spectacular shelf slope limestone breccia, as well as basinal slates containing likely turbidite-transported carbonate shelf beds. The out-of-sequence cataclastic Highgate Falls Thrust is well-exposed, as well as a number of other Taconic-aged structures including folds, axial plane cleavage, and rotated en echelon fractures. The “Beam” is a superb meso-scale illustration of duplex faulting in thin-skinned foreland belts, while Lessors Quarry exposes fault-bend fold structures as well as a number of fault zone features, and an excellent example of an out-of-sequence thrust. This trip is highly recommended for students to see a number of well-exposed and instructive stratigraphic and structural features. Part of this material has appeared in a previous NEIGC field trip (Schoonmaker and Kidd, 2009).
Walking in the gorge can be treacherous and difficult both on the boulder field in the channel and the outcrop; both may be very slippery if wet. Care should be observed, and adequate footwear is advised.

A3: STRUCTURAL AND STRATIGRAPHIC FEATURES OF THE TACONIC FORELAND, NW VERMONT

Assembly Point - Price Chopper in Warrensburg

STOP 1. Upper Highgate Falls Gorge.
Continue straight and immediately bear right onto Mill Hill Road. Park at bottom of hill at terminus of pedestrian bridge and hydroelectric dam. A nice view of the gorge can be seen from the pedestrian bridge including the Highgate Falls Thrust. Proceed by foot along canoe portage westward through the woods along the north side of the gorge for about 50 feet. Find trail on the left that leads down to the gorge floor below the dam.
CAUTION!
ACCESS TO GORGE FLOOR WILL REQUIRE SCRAMBLING DOWN STEEP ROCKS AND ONCE ON THE GORGE FLOOR, WALKING ON THE BOULDER FIELD CAN BE TREACHEROUS, ESPECIALLY WHEN WET.
PLEASE ALSO NOTE THAT IF HORN SOUNDS, EXIT GORGE AS QUICKLY AS POSSIBLE IN THE EVENT OF AN UNLIKELY, BUT POSSIBLE DAM RELEASE.
The Mississquoi River laps against the walls in part of the gorge, so it is cut into upper (Stop 1) and lower (Stop 2) sections with different access points.
The entire (both upper and lower) east-west oriented Highgate Falls Gorge on the Mississquoi River continuously exposes a cross-section (Figure 3) of upper Cambrian to lower Ordovician rocks that include the upper section of the Gorge Formation (dolomitic arenites and quartz sandstones), the entire Highgate Formation (a variable unit containing bedded micrite, limestone clast breccias, calcareous slate, and argillaceous micrite), and the lower part of the Morses Line Formation (argillaceous slates with dolomitized, boudinaged micrite beds, bedded micrite, and limestone breccias;). The Highgate Falls Thrust repeats part of this section (Gorge Fm.) in the part of the gorge nearest the dam. In the upper gorge we can see the upper part of the Highgate Formation, lower section of the Morses Line Formation, and part of the Gorge Formation above the Highgate Falls Thrust.
Once in the gorge, the Highgate Falls Thrust is clearly visible in the north wall, approximately 200’ downstream from the point where the path descends (Figure 6E). A dolomite clast breccia with dolomitic matrix of the Gorge Formation is thrust over dark slates and a bedded argillaceous micrite and carbonate breccia interval within the Morse Line slates along a shallowly east-dipping surface characterized by phyllonitic fault zone cleavage (cataclastic breccia- to gouge-sized material); beds and cleavage in the lower block are bent in the direction of transport of the upper block. In the lower block, directly beneath the thrust, regional cleavage in the slates and bedded micrite can be seen to be bent as it approaches the fault zone indicating a westward sense of motion of the upper block. Because the regional cleavage is deformed by the motion of the upper block, the Highgate Falls Thrust is inferred to be an out-of-sequence thrust fault.
To the west and slightly downstream from the highway bridge that can be seen high above the gorge and upstream from the powerhouse and tailrace visible on the south bank, the contact between calcareous slates of the uppermost Highgate Formation are overlain by black slates of the Morses Line Formation (Loc. A, Figure 5). The contact on the north bank is mesoscopically folded and marked by a minor thrust fault containing a slickenlined and stepped calcite vein on the north limb (Figure 6C). On the south bank, the south limb is an observable depositional contact. Landing (1983) interpreted the micrite clast breccia (unit 29, Figure 5) that occurs beneath the calcareous slate as a tectonic breccia and the base of the black slates as an unnamed thrust. In contrast, Haschke (1994) suggested the slickensided surface on the north limb was a detachment that cut out a major thrust between the Highgate and Morses Line Formations, similar to relationships described by Hayman and Kidd (2002) for the northern part of the Taconic Allochthon. The observable conformable contact relationships on the south bank in the Highgate Falls Gorge do not support either of these hypotheses. Further, minor thrust faults are observed in the Morse Line slates (Figure 6D) and cleavage can also be seen in the Highgate Formation limestones, suggesting that a significant structural boundary is not present within this section, which contrasts to the Quebec and Taconic allochthons where deformation and cleavage deformation occurred prior to final emplacement (Rowley and Kidd, 1981; St. Julien, 1977; Lebel and Kirkwood, 1998).
Within the Morses Line Formation, dismembered dolomitized micrite beds can be observed, wrapped by the regional cleavage. Some of these beds are cut by small pre-cleavage faults with a normal sense of displacement. These normal faults are interpreted to be from passive continental margin downslope slumping. Similar dolomite beds in slate are observed further north in Vermont and in Quebec, well east of the strike of these beds, where they form part of the upper block of the Highgate Falls thrust and are mapped as Stanbridge Slates (Figure 3).
Structurally just below the Highgate Falls Thrust, there is a small fold of bedding in slate and thin limestone beds whose axial plane is parallel to the regional cleavage in the slates; this can be seen in the outcrop close to the river channel.

STOP 2. Lower Highgate Falls Gorge.
The lower gorge displays the contact between the Gorge and Highgate Formations, slightly below the Cambrian-Ordovician boundary (Figures 3 and 4). Here, the depositional contact the Gorge Formation and the overlying Highgate Formation can be observed (Figure 3, Loc. D). In contrast, Landing (1983) placed the contact at Loc. B (Figure 3). Very similar argillaceous limestones and limestone breccias are present both below and above this level. We prefer to define the contact at the lower horizon (Figure 3, Loc. D; Figure 6B), which separates dolomitic sandstones below (Gorge Formation) from the argillaceous limestones and limestone breccias above (Highgate Formation), as a much more readily identifiable contact in the field and likely representing a significant change in depositional environment from sand-dominated carbonate deposition to shale-dominated carbonate deposition. Regardless of the placement of the Gorge- Highgate contact (Figure 3. Loc. D vs. Loc. B) within the Highgate Falls Gorge, no major structural boundary is observed in this section, up to the position of the out-of sequence Highgate Falls Thrust.
The contact is readily identifiable in the gorge by the abrupt change from thick-bedded dolomitic arenites of the Gorge Formation below, to thinner bedded argillaceous micrites and limestone breccias of the Highgate Formation above (Figure 6b). In Quebec, this contact is inferred to be a major fault boundary between the autochthonous shelf platform rocks (Milton Dolomite) and allochthonous, deep marine rocks (Stanbridge Complex).

STOP 3. “The Beam”, a Rolfe Stanley Memorial Outcrop.
NO HAMMERS, PLEASE. “The Beam” is one of two Rolfe Stanley Memorial Outcrops visited on this trip. Over the years, many thousands of introductory geology students, as well as undergraduate and graduate students in Rolfe’s structural geology courses have visited these sites. Both beautifully expose important features of foreland deformation, including thrust faulting, fault duplexes, fault bend folds, ramps and ramp faults, en echelon fractures and rotation, and pressure solution cleavage. Detailed descriptions and interpretation for both of these outcrops can be found in Stanley (1988).
“The Beam” occurs in the Cumberland Head Formation, and is a ∼1 foot thick micrite bed sandwiched between calcareous shale, and is bounded by bedding parallel thrust faults that display slickensided and stepped calcite veins, and fault zone cleavage (Figure 7). The micrite bed is relatively uncleaved while the surrounding shale displays a strong, steeply east-dipping spaced pressure solution cleavage, similar to that seen throughout the region, containing clay selvedges. Along the thin fault zones, a fault zone cleavage dips more shallowly to the east. Near the fault zone, the spaced cleavage in the shale is rotated towards parallelism with the fault zone cleavage indicating westward transport along the thrust faults. This is consistent with fiber orientation and slickenside fracture steps developed within veins in the fault zones.
Several ramps have been cut through the micrite bed, merging with the bounding bedding-parallel floor and roof faults, creating several horsts. Ramp faults show progressive development from initial slight buckle folding, en echelon fracture formation along the west-dipping limb, and subsequent rotation and failure through the trace of the fractures. Some fractures show diachronous development where older fractures are rotated and cut by younger, unrotated ones. Fracture arrays dominantly dip to the east (west-climbing), parallel and are cut by ramps, but a few dip to the west (east-climbing) and are not cut by through-going faults.
Deformation in “the Beam” contrasts with that in the surrounding calcareous shale. “The Beam” has shortened through rigid body displacement along thrust faults with little cleavage development within the bed, and minor folding over the ramps (∼13%), while shortening in the shale is accommodated by pressure solution development (∼11-16%), with some minor bedding parallel faults present in the shale. Stanley (1988) interpreted the development of ramp faults to be time transgressive, from east to west across the outcrop. This diachronous shortening in “the Beam” was accompanied by a concurrent progressive pressure solution cleavage development from east to west in the surrounding shale (Figure 8).

STOP 4. Lessor’s Quarry, a Rolfe Stanley Memorial Outcrop.
NO HAMMMERS PLEASE. There is plenty of loose material on the quarry floor to collect. Abundant fossils can be seen both in the walls of the quarry and in the loose debris. This is the second of two Rolfe Stanley Memorial Outcrops we will be visiting.
Lessor’s Quarry exposes several bedding parallel and bedding cutting faults in the Glens Falls Limestone. The north wall displays several interesting features (Figure 9) including an out-of-sequence thrust that is both bedding parallel and bedding-cutting indicating that is preserves parts of its initial ramp and flat sections (labeled “Thrust Fault” in figure 9). It cuts an earlier folded thrust with fault splay and tip (labeled “Synformal Thrust Fault” in figure 9). The folding of the earlier thrust is thought to be the result of an unseen ramp beneath the quarry (Stanley, 1988); e.g. a fault bend fold similar to those seen in the roof thrusts at the “Beam”. Since the younger through-going thrust cuts the earlier folded thrust, the younger thrust must be an out-of-sequence fault.
Also related to the younger out-of-sequence fault is the orientation of bedding in the block above the thrust. In the eastern part of the wall, bedding parallels the thrust, while in the middle parts of the north wall, bedding is truncated by the thrust. Then, in the western part of the wall, bedding returns to a parallel position (Figure 9). This suggests that the currently flat thrust in the middle of the wall, initially developed as a ramp, cutting bedding in the upper block. This part of the upper block was then transported off the ramp and onto a nearby flat section of fault, resulting in counterclockwise rotation of the upper block in this section. Note also, the spaced pressure solution cleavage in the upper block is rotated in a counterclockwise manner in the middle part of the wall, over the “ramp” section. Because the cleavage was rotated during movement, the fault must have evolved subsequent to cleavage development, consistent with its interpretation as an out-of-sequence fault (recall that the Highgate Falls Thrust also rotates cleavage).
Along the western wall, next to the entrance, is a steeply dipping fault surface decorated with smeared calcite where the upper block has been removed by quarry operations. It is abundantly lineated with slickenfibers; well-developed steps indicate an east-over-west sense of transport for the removed upper block.

A4: NEW VIEWS ON THE DEGLACIATION OF MT MANSFIELD
Paul Bierman
Lee Corbett
Stephen Wright
Chris Halstead
University of Vermont
Burlington, VT 05405
P. Thompson Davis
Bentley University
Waltham, MA 02452
Jeremy Shakun
Boston College
Chestnut Hill, MA 02467
Mt. Mansfield, the highest peak in Vermont, preserves evidence of Pleistocene glaciation including erratic boulders, striations, and ice-sculpted rock outcrops. Over the past several years, we have been collecting samples of bedrock and erratic boulders and using them for cosmogenic dating. Mapping of recessional sedimentary features including moraines and striations where bedrock crops out on the mountain, place these data in context. This part-day day trip will (weather permitting) travel to the top of the mountain by car and visit locations we have sampled while discussing the ages that resulted from these analyses. We will discuss evidence for ineffective erosion at the mountain peak (could there have been cold-based ice?). If time permits, we may examine ice marginal deposits in the valley resulting from ice channelized by the topography. For those who are interested, the trip will end with a tour of the Vermont cosmogenic clean labs now supported by as a community sample processing facility by the US NSF.

A4: NEW VIEWS ON THE DEGLACIATION OF MT MANSFIELD
P.Bierman, L.Corbett, S.Wright, C.Halstead, P.T.Davis, J.Shakun

Meeting point
P.Bierman, L.Corbett, S.Wright, C.Halstead, P.T.Davis, J.Shakun

Start of driving trip,
Toll Road, Mile 0 – Parking lot off of 108 at the base of the Toll Road. From here we will drive 4.8 miles up a steep, one lane, mostly dirt road.
P.Bierman, L.Corbett, S.Wright, C.Halstead, P.T.Davis, J.Shakun

End of driving trip,
Toll Road, Mile 4.8 – Arrive at the parking lot at top of the road, elevation 1170 m asl. From here, we will walk along the ridge, weather permitting to a variety of sampled outcrops on the summit ridge trail. The trail is rough and rocky and not always well marked. Our journey and time spent on the ridge will depend on the weather which can be quite variable and potentially dangerous any time of year. The summit of Mt. Mansfield is 1339 m and about a 2 km walk each direction from the upper parking lot. We hope to be able to walk to the summit and return if the weather allows.
Since our route will depend on weather, we include below a table of GPS coordinates for sample sites allowing others to relocate within a meter of where we collected each sample.

STOP 1. Sample sites MM-01 and MM-02.
This is a site where a boulder and a nearby bedrock outcrop were collected. Both return 10Be exposure ages that are similar within error of 14.5 and 14.8 ka, respectively. We will speak at this location about sampling strategy, the use of high precision GPS for sample location, and the utility of matched boulder/bedrock pairs to test for inherited nuclides.

STOP 2. Sample site MM-03.
This erratic boulder has been known to science for nearly two centuries. Featured in the illustration pictured first Edward Hitchcock’s Report on the Geology of Vermont (1861) and reproduced below along with Thom Davis’s best vintage interpretation, it has a 10Be exposure age of 13.65 ka and a 14C exposure age that is older but within one standard deviation (15.6 ka). We will speak at this boulder about the implications of the dual isotope measurements for ice cover and deglacial history. Note that a nearby boulder (MM-04, 15 m away) gives a nearly identical 10Be age of 13.75 ka.

STOP 3. MM-06 and -07.
These are two bedrock samples collected very close to one another as a test of reproducibility. They have 10Be exposure ages of 14.1 and 13.6 ka. We will examine these sites closely and compare the weathering to that of the sites near the peak (Stop 4) which have higher, in some cases much higher 10Be exposure ages.

STOP 4. Sample sites MM-08, -09, -10, and -11.
These are four bedrock samples sites near the summit of Mt. Mansfield ranging in elevation from 1270 to 1300 m asl. They have 4 of the 5 oldest exposure ages (15.2 – 22.9 ka) that we have measured on the mountain (the other being MM-12 on Sunset Ridge, about 50 m lower). These relatively high 10Be exposure ages imply ineffective glacial erosion and the inheritance of some nuclides from a prior period of surface or near-surface exposure. Two of the samples, MM-08 and MM-09 have also have in situ 14C exposure ages. These ages are much younger, 11.7 and 9.7 ka, respectively. We will talk at these sites about the need for some type of cover being present on the summit (till, ice, snow fields) after deglaciation of the lower flanks of Mt. Mansfield at 13.9 ± 0.6 ka. After this stop, we will return to the vehicles and descend the Toll Road.

A5: IAPETAN RIFT‒PASSIVE MARGIN TRANSITION IN NE LAURENTIA AND EUSTATIC CONTROL ON CONTINENTAL SLOPE OXYGENATION, TACONIC SLATE COLORS, AND EARLY PALEOZOIC CLIMATE
Ed Landing
New York State Museum
222 Madison Avenue
Albany, NY 12230
ed.landing@nysed.gov
Mark Webster
Department of the Geophysical Sciences
University of Chicago
5734 South Ellis Avenue
Chicago, IL 60637
mwebster@geosci,uchicago.edu
This excursion along the eastern margin of the New York Promontory was prepared for the combined meeting of the New England Intercollegiate Geological Conference and the New York State Geological Association (October 2018). Parts of this guide have been modified from reports by EL and MW and are so indicated. There are two purposes of the trip:
(1) The oldest rocks of the Taconic allochthon (originally deposited on the east Laurentian margin of the Iapetus Ocean and at least 75 km east of its present location) and coeval shelf margin rocks in the Green Mountain anticlinorium (not seen on this trip) indicate a surprisingly late persistence or rejuvenation of Iapetan rifting. Although Iapetan rifting began in the later Ediacaran, the oldest record of sedimentary rocks deposited on the middle Proterozoic basement of the Grenvillian orogen in NE Laurentia is late Early Cambrian, with immature sandstone (arkose) deposition continuing well into the Middle Cambrian in the eastern Ottawa-Bonnechere aulacogen (Landing et al., 2009, In press; Figure 1).
(2) The second purpose of this trip emphasizes the relationship of major eustatic changes to macroscale alternations in oxygenation on the east Laurentian continental slope as exhibited in the Lower Paleozoic of the Taconic allochthon. One take away from the trip is that formation-level, type 1 depositional sequences on the Cambrian‒Ordovician shelf of eastern Laurentia are reflected by macroscale color alternations on the continental slope. Major eustatic rises and marine onlap of the cratons led to an increase in global insolation as reflective subaerial continental regions areas were inundated by shallow marine seas. These epeiric seas served as a heat sink, warmed the World Ocean, increased a key greenhouse gas (water vapor), all of which led to global temperature rise. Additional effects of high global temperatures include a decrease in storminess with decreased latitudinal temperature gradient as well as major intervals of deposition of organic-rich muds on the continental slope. Increased organic-rich mudstone deposition followed on the slope from the greater thickness and intensity of the mid-water low oxygen zone and in shallow shelf water with decreased oxygen solubility. Thus, high eustatic levels and consequent high global temperatures correlate with carbon sequestration and are the backdrop of the fossil fuel industries. This temperature increase is unrelated to changes in pCO2 or other greenhouse gasses, and is termed “global hyperwarming” (Landing, 2002, 2007, 2012). “Global hyperwarming” is also pertinent to global climate in the near future with accelerated global warming with eustatic rise due to the melting of ice on land in the high mountains, Greenland and other North Atlantic islands, and Antarctica. (modified from Landing et al., 2007, p. 25)

A5: IAPETAN RIFT‒PASSIVE MARGIN TRANSITION IN NE LAURENTIA AND EUSTATIC CONTROL ON CONTINENTAL SLOPE OXYGENATION, TACONIC SLATE COLORS, AND EARLY PALEOZOIC CLIMATE

Small road cuts in Grenvillian Proterozoic gneiss

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" [abandoned by Landing, 2003, 2012] as mapped by Fisher, 1984).

Pass ca. 1.0 mile of roadcuts in Proterozoic Hague gneiss
Pass ca. 1.0 mile of roadcuts in Proterozoic Hague gneiss on block uplifted by Welch Hollow fault.

STOP 1. Middle Proterozoic‒Cambrian Nonconformity
and Relationship of Thin Lower Paleozoic Shelf Succession to Continental Slope Oxygenation and Paleoclimate.
The stop shows several meters of east-dipping, medium–coarse grained, slightly dolomitic, brownish-weathering quartz arenite of the Potsdam Formation (Keeseville Member) with a basal quartz pebble conglomerate nonconformably overlying middle Proterozoic Hague gneiss. We are about 7 m below an upper Middle Cambrian trilobite-bearing horizon higher in the Keeseville with Crepicephalus, Komaspidella, and Lonchocephalus (Flower, 1964, p. 156; Landing et al., 2007, Stop 1.4).
This photogenic locality records the absence of ca. 600 million years of Earth history at this planar nonconformity. Further north in the Lake Champlain lowlands as the south flank of the Ottawa-Bonnechere aulacogen (OBa) is approached, the non-fossiliferous, feldspathic, often fluviatile facies of the lower Potsdam Formation (Ausable Member) appears under the Keeseville (Figures 1, 3). The oldest Paleozoic unit on the Grenville orogen in the OBa of northeastern New York and Adjacent Quebec is the Altona Formation―a thin (to 84 m-thick) heterolithic unit under the Ausable that includes lower wave-deposited feldspathic sandstones and red mudstones and a higher subtidal red-mudstone and sandstone with hydrothermal dolostones (compare Landing et al., 2009, In press, with Lowe et al., 2017, 2018; Figure 3).
Stop 1 emphasizes that the Lower Paleozoic of the eastern New York shelf is very thin by comparison with Great Basin and southern Canadian Rockies sequences (i.e., 100s of meters vs. several thousand meters). The ridge crest to the east is capped by the upper Little Falls Formation and only ca. 200 m of upper Middle‒Upper Cambrian overlies the Grenvillian basement. Thin Lower Paleozoic successions occur in this region of New York and Vermont 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., 2003A). For example, the Cambrian–Ordovician boundary is a type 1 sequence boundary everywhere on the New York and western Vermont shelf and across most of the Laurentian platform except parts of the Great Basin and Southern Oklahoma aulacogen (Landing, 1988A; Landing et al., 1996, 2003, In press).
The Crepicephalus Chron (a “Chron” is the age of a biostratigraphic “Zone”) is a time in the late Middle Cambrian in which epeiric seas covered almost all of Laurentia except for the Trans-Continental Arch (e.g., Lochman-Balk, 1971), as well as major parts of other paleocontinents. Such an interval of high eustatic levels and high global insolation would have corresponded to a time of global hyperwarming, to d/a conditions on the Laurentian continental slope (i.e., part of the Hatch Hill Formation; Figure 3), and the probability of off-shelf transport of Potsdam sands and their deposition as part of the Hatch Hill lithofacies as seen at Stop 7. (modified from Landing et al., 2007, p. 30)

Rte 4 follows the Proterozoic–Cambrian nonconformity.
Rte 4 roughly follows a topographic break corresponding to the Proterozoic–Cambrian nonconformity.

Grenville gneiss and intersection with Rte 22 on right
Road cuts along Rte 22 east of Comstock State Prison are important sections through conodont-, cephalopod-, trilobite-bearing Lower Ordovician Tribes Hill (=“Great Meadows Formation”, abandoned), Rochdale (=“Fort Anne Formation”, abandoned), and Fort Cassin formations (Landing et al., 1988B, 2003, 2012, Appendix 2).

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.

STOP 2. Potsdam Formation Roadcut.
Ca. 30 m of upper Potsdam (Keeseville Member) siliceous and dolomitic quartz arenite are exposed in the Whitehall area. Whitehall east of the Champlain Canal is on the Proterozoic, and the canal is cut in Potsdam sandstone (Figure 5). Herring-bone cross sets, dolomitic quartz arenite pebbles at the base of small channels/dunes, and U-shaped Diplocraterion burrows point to tide influenced, high energy Potsdam facies at Stop 1. The upper Middle Cambrian Potsdam is overlain by lower Upper Cambrian dolomitic quartz arenites and quartzose dolostones of the Galway Formation higher on Skene Mountain (=“Ticonderoga Formation,” term abandoned, also mistakenly termed “Theresa Formation” in reports on eastern New York; Landing et al., 2003, In press). The top of Skene Mountain is upper, but not uppermost, Little Falls Formation limestone (Landing et al., 2010). (after Landing et al., 2007, p. 31)

Pass cuts in Potsdam Formation on the right

STOP 3. Crossman Quarry - Upper Cambrian Little Falls Formation.
NO HAMMERS! The eastern end of the quarry (Figure 5) 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, represents Fisher’s (1977) undefined “Warner Hill Limestone” member of the middle “Whitehall Formation” (abandoned designations). As commonly seen in Cambrian‒ Ordovician carbonate platform successions in eastern New York, hydrothermal dolomitization likely during the Late Ordovician Taconian orogeny has led to abrupt lateral and vertical changes from limestones with well preserved primary structures to dolostones without primary structures is seen in the Little Falls Formation at Stop 3 (see Collins-Wait and Lowenstein, 1994; Landing et al., 1996; Landing, 2007). The limestones are increasingly dolomitized with the loss of primary features in the western part of the quarry. In the western quarry, neomorphic dolostones predominate and have been termed the “Skene Dolostone” (Wheeler, 1941) sensu Fisher (1977, 1984). These sucrosic “Skene” dolostones comprise the road cut opposite the western end of the quarry. Long term weathering of these dolostones yields buff-colored blocks that are present in the talus at the west end of Crossman quarry.
Well preserved primary carbonate structures appear in the middle part of the quarry. A photogenic block with a large, gray thrombolite capped by laminated thin, greenish weathering limestones with chert lenses and small SH-V stromatolites lies at the base of the cliff in mid-quarry.
A biostratigraphically important Late Cambrian trilobite assemblage appears ca. 3.5 m above the base of the eastern quarry wall (Taylor and Halley, 1974). A meter above this, the major trilobite replacement (biomere) that marks the Laurentian Upper, but not uppermost, Cambrian Sunwaptan Stage‒Ibexian Series boundary is located at a lithofacies break that includes an intraclast pebble-ooid packstone bed (with the biomere boundary), an overlying quartz arenite, and a distinctive bed with SH-V and LLH-V stromatolites (Landing et al., 2010). Above this, a biofacies dominated by snails and other mollusks with rare trilobites and sparse euconodonts occurs in a thrombolite-dominated succession to the top of the quarry, and indicates a restricted, relatively near-shore environment (e.g., Westrop et al., 1995).
One purpose of Stop 3 is to show the lithology of non-dolomitized Upper Cambrian limestones of the Little Falls Limestone (“=Whitehall Formation,” abandoned term; Landing et al., 2003A) for comparison with lithofacies of the overlying Tribes Hill Formation at Stop 4.
The second purpose of Stop 3 is to show the lithology of an extensive carbonate platform facies that extended in the Late Cambrian from eastern Laurentia and lenses out in coeval siliciclastic-dominated deposits of the Upper Mississippi Valley in western Wisconsin and Minnesota and passes west through the Black Hills and into western Laurentia. As the Late Cambrian was a time of particularly extensive epeiric seas and consequent global warming (e.g., Landing, 2010, 2011, 2012), it was also a time in which black shale deposits as those of the Hatch Hill Formation developed on the east Laurentian slope (Stop 7, 8). (modified from Landing, 2002, p. B6-7; Landing et al., 2007, p. 33, 34).

Ridges to the East
The first ridge to the east is the local top of the Cambrian‒Ordovician passive margin carbonate platform (Middle Ordovician Providence Island Dolostone). 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 synorogenic, lower Upper Ordovician “Snake Hill Formation” flysch and Forbes Hill mélange (Fisher, 1984).

Cross east end of Skene Mountain.

STOP 4. Rte. 4 - Lower Ordovician (Upper Lower Tremadocian) Tribes Hill Fm.
Stop at east end of high, whitish weathering road cut in Tribes Hill Formation. Although dominated by dolostone, this cut is only 0.3 km south on strike of identical, but non-dolomitized strata with well preserved primary depositional structures in the Tristates quarry (see Landing et al., 2003A; Kröger and Landing, 2007; for conodonts, trilobites, and cephalopods). Key stratigraphic divisions (Sprakers, Van Wie, and Wolf Hollow members) of the Tribes Hill Formation will be pointed out from the south side of Rte 4; only the uppermost Tribes Hill Formation (Canyon Road Member, highstand facies) is missing in the Rte 4 road cut, as well as in the Tristates quarry.
This east-dipping section brackets most of the Tribes Hill Formation (=“Great Meadows Formation,” abandoned, of Fisher and Mazzulo, 1976, and “Spellman Formation,” abandoned, of Fisher, 1968). Well preserved, dark grey limestones with primary structures in the Tristates quarry just 0.2 mi. to the north (see Landing, 2002, Stop 5; Landing et al., 2003), are replaced directly along strike by hydrothermal, whitish dolostones at Stop 4. As detailed by Landing et al. (1996, 2003a, 2012) and Landing and Westrop (2006), the Tribes Hill Formation, with a uniform succession of four lithologically distinctive members, can be recognized in outcrops over about 12,000 square kilometers from Dutchess County, SE New York, through the lower Mohawk River valley of east-central New York, and north through the Lake Champlain lowlands. Just north of the New York-Quebec line near Ste-Clotilde, the Tribes Hill Formation has been termed the “Ste-Clotilde Member” (Landing et al., In press).
With exception of the Great Basin and the Southern Oklahoma aulacogen, the Cambrian‒Ordovician boundary is an unconformity across the Laurentian craton (Landing, 1988, 2012; Landing et al., In press). Early, but not earliest, Ordovician (i.e., late early Tremadocian) eustatic rise led to the Stonehenge transgression (Taylor et al., 1992) across Laurentia with carbonate platform deposits reaching the Upper Mississippi Valley (Oneota Dolostone) and the Rocky Mountains (e.g., Manitou Formation). These onlap units form a type 1 depositional sequence, are about the same age, and have Rossodus manitouensis Zone conodonts. This depositional sequence unconformity is the contact of the Upper Cambrian Little Falls Formation and Tribes Hill, with some localities showing channels up to 2.0 m deep that are filled with Upper Cambrian carbonate blocks (Landing et al., 2003A).
Rossodus manitouensis Zone conodonts have been recovered from most outcrops of the Schaghticoke d/a interval in continental slope sequences in NE Laurentia (e.g., Landing et al., 1986). This means that the extensive deposits of the Stonehenge transgression correspond to intense continental slope d/a conditions, as would be expected by the global hyperwarming model (Figures 3, 4).
The three lower members of the Tribes Hill Formation are exposed at Stop 4 (Figure 7). These include the thinner bedded, silty, locally intraclast-rich limestones (now dolomitized) of the Sprakers Member (which includes Fisher and Mazullo’s, 1976, “Winchell Creek Siltstone,” abandoned).
The most interesting unit of the Tribes Hill is the Van Wie Member. Although the Van Wie is never more than about 1.5 m in thickness, it occurs throughout the outcrop belt of the Tribes Hill. In all sections where it is exposed, it invariably has a basal intraclast conglomerate (to 25 cm in thickness) and a middle, thinner intraclast bed. These beds are perfect time-markers and seem to represent major storm beds in a time when dark shale appears in the Tribes Hill Formation. The Van Wie Member is composed only of black shale to the south in Dutchess County (Landing et al., 2003A). To the north and west, the shale lightens and has more limestone/dolomitized limestone beds (e.g., Landing et al., 1996), as at Stop 4. This presence of organic rich shale and the indication of d/a conditions on the NE Laurentian shelf also corresponds to the global hyperwarming model which predicts increased preservation (not increased production) of organic material in hot, oxygen-poor epeiric seas.
The Wolf Hollow Member (Fisher, 1954; = “Fort Edward Dolostone member” of Fisher and Mazullo, 1976; =“Kingsbury Limestone,” abandoned of Fisher, 1984, Appendix 2) occurs above the Van Wie at the Rte 4 cut and is characterized by massive thrombolite buildups in all outcrop areas of the Tribes Hill Formation. These huge thrombolites can be seen in the upper part of the Rte 4 cut.
The highstand facies of the Tribes Hill is the Canyon Road Member (Landing et al., 1996). The Canyon Road features lower ooid wackestones and mollusk-rich carbonates above the Wolf Hollow thrombolites. Higher strata include sparsely fossiliferous carbonates with planar stromatolites and molds of gypsum and anhydrite crystals and nodules of the upper Canyon Road Member. However, the Canyon Road is absent in the Rte 4 and nearby Tristates quarry sections (Figure 5).

Low ridge in front of the vehicles is the front of the Taconic overthrust.

Cemetery on right is on Taconic mélange
Pass cemetery on right which is on Taconic mélange under the Taconic master thrust.

STOP 5. Taconic Overthrust.
Note low brownish outcrops on left (north) side of road. Stop on right side of road at top of low hill.
NO HAMMERS! This is primarily a stop for geological perspective as well as one of the best localities to photograph the base of the Taconic master thrust.
The sheared brownish sandstone at the east end of the roadcut is the upper Lower Cambrian Bomoseen Formation, which here forms the westernmost edge of the Giddings Brook slice. The ca. 2.0 m of thin bedded, gray weathering, dark gray, fossiliferous limestone under the Bomoseen is a block of Upper Ordovician Glen Falls Limestone (Trenton Group, global Katian Stage) in what has been called the Forbes Hill mélange (Zen, 1961; Fisher, 1984).
A look to the west shows just how “compact” the geology of eastern New York is (Figure 1). As noted above, the Proterozoic basement of the Adirondack Mountains massif is analogous to the Black Hills in South Dakota; the block faulted, uplifted Proterozoic‒Ordovician shelf succession in the southern Lake Champlain lowlands to the front of the Rocky Mountains; and the Taconic overthrust succession to the Nevada thrust belt.

Upper Cambrian part of the Hatch Hill Formation
The cross road is in the Upper Cambrian part of the Hatch Hill Formation, with EL having found disarticulated olenid trilobites in a thin limestone bed on the east side of Rte. 21. This trilobite locality is ca. 50 m North of the intersection. After the intersection, drive South on strike along roadcuts in interbedded black shale and reddish (pyritic and dolomitic) weathering sandstone of anoxic/highly dysoxic Hatch Hill Formation (upper Lower Cambrian‒ lowest Ordovician)

Hatch Hill Formation in road cut on right.
After crossing creek at curve in road, pass more Hatch Hill Formation in road cut on right. After this point, the road swings east and crosses a large westerly overturned syncline in the Giddings Brook slice.

Lower Ordovician Deep Kill Formation
Greenish gray roadcut of Lower Ordovician Deep Kill Formation (“Poultney Formation” in many reports, abandoned designation, Landing, 2012).

OVERTURNED upper Lower Cambrian Browns Pond Formation.
Drive south past low road cuts in black slates and thin sandsones in east-dipping, OVERTURNED upper Lower Cambrian Browns Pond Formation.

Overlook of Browns Pond to south.

Browns Pond Formation
Pass low roadcuts on east and west sides of Holcombville Road. These overgrown cuts comprise the upper part of the type section of the Browns Pond Formation of Kidd et al. in Fisher (1984), The southerly cut on the east side of the road includes black shale with thin, laminated sandstones that are almost devoid of fossils with exception of a large looping epichnial trace fossil. The northerly cut on the west side of the road is a carbonate clast debris flow (Holcombville Member of Landing et al., Submitted) that forms the top of the Browns Pond Formation from Washington Co. to Dutchess County 160 km to the south. The pasture further west is underlain by slaty green mudstones of the Middle Granville Formation. These strata (Stop 3B and 3C of Rowley et al. (1979), Stop 8 of Landing, 2002, Stop 6.2 of Landing et al., 2007) will be better seen at Stop 6.

STOP 6. Late Early Cambrian Anoxic‒Oxic‒Anoxic Macroscale Alternations.
Stop at entrance to small newly reactivated quarry on west side of road. This stop consists of two parts. Do not descend into the quarry.
A sharp black‒green mudstone transition (Browns Pond‒Middle Granville formation contact; see Figure 3) is exposed at the top of the steep ramp that descends into the quarry. The Holcombville Member consists of thin limestones interbedded with black shale and capped by a massive calcareous sandstone debris flow at the top of the ramp. These strata are replaced ca. 50 m north along strike by fossiliferous, bedded turbiditic limestones that show fracturing and incipient transition into a debris flow. The second part of this stop is an examination of the abundant trace fossils in oxic facies of the Middle Granville Formation. Bedding surfaces of the Middle Granville Formation are present on large blocks of green and red slate at the south end of the Holcombville quarry and in the hillock of scrap slate on the east side of the road. These bedding surfaces occur because of the local bedding plane-parallel cleavage. These bedding surfaces are crowded with shallow trace fossils that range up to several centimeters in diameter.
Browns Pond Formation and Holcombville Member
The Holcombville Road quarry is an overturned section. It begins with 2.0 m of black siliciclastic mudstones with light gray, nodular lime mudstones. The black mudstones are capped by a 30–40 cm-thick, white weathering, light gray, conglomeratic pack- to grainstone composed of limestone sand–small intraclast pebbles and trilobite fragments with black phosphatic sand grains.
This ca. 2.4 m of section is the type locality of the Holcombville Member of the Browns Pond Formation (proposed in Landing et al., Submitted). This is the top of the Browns Pond Formation and the end of the late Early Cambrian Browns Pond d/a macroscale interval on the east Laurentian continental slope (Landing and Bartowski, 1996; Figure 3). The 30–40 cm-thick limestone thickens to 2.0 m about 50 m to the north where weathered cracks on the top surface of the unit show initial fracturing and formation of a debris flow. Micro- and macrofaunal elements of the lower Elliptocephala asaphoides assemblage (Lochman, 1956; Bird and Rasetti, 1968; Landing and Bartowski, 1996) occur in this top member of the Browns Pond Formation (E. Landing and M. Webster, unpub. data).
Middle Granville Formation
A very abrupt change into the gray-green mudstones of the Middle Granville Formation takes place right at the top of the 30–40 cm-thick limestone clast debris flow. The lower part of this greenish macroscale interval formed by the Middle Granville Formation has several beds of nodular lime mudstone. A thin limestone clast debris flow (less than 1.0 m-thick), which may be visible if the water level is low, is composed of slumped fragments of these nodular limestones. All of the thin Middle Granville Slate is exposed in the quarry. Several beds of purple slate in the Middle Granville Formation are present higher in the quarry.
Hatch Hill Formation.
A transitional interval from the Middle Granville Formation into the Hatch Hill Formation is seen in the west wall of the quarry, where the Middle Granville changes color into a grayish hue. Black, pyritiferous mudstones of the lower Hatch Hill Formation [i.e., the Hatch Hill interval, or terminal Lower Cambrian–lowest Ordovician dysoxic/anoxic macroscale interval (Landing et al., 2002)] appear ca. 5 m higher. Finally, beds of dolomitic quartz arenite characteristic of the Hatch Hill Formation are found not more than 10 m stratigraphically above the west wall of the quarry.
Middle Granville Slate and Middle Granville oxic interval (MGOI) (15 minutes)
Trace fossils Trace fossils are abundant in the red and green mudstones of the Middle Granville Formation. The crude retaining wall composed of large slate slabs at the south end of the Holcombville Road quarry and the large large pile of brownish red and minor green slate on the east side of the road came from the Holcombville Road quarry.
If the light is good, dense Planolites and large grazing traces up to 2 cm-wide can be seen on many of the bedding plane-parallel cleavage surfaces of the green and reddish slate. 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. Evidence for skeletalized metazoans is not present in the reddish, purple, and green mudstones of the Middle Granville Formation. Either these metazoans were not present on the bottom, or their calcareous remains were dissolved away during diagenesis.
MGOI Landing et al. (2002) noted that the uppermost Lower Cambrian (upper but not uppermost Olenellus Zone) on the New York and Québec portions of the eastern Laurentian slope is composed of red and green siliciclastic mudstones. Unfortunately, Landing (2007, also Landing et al., 2007) equated this interval of improved oxygenation of slope waters 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 dysoxic/anoxic interval) took place in the terminal Early Cambrian (Landing et al., 2002).
This interpretation has been re-evaluated with the obvious fact that the lower Elliptocephals asaphoides assemblage of the upper Browns Pond and Middle Granville formations is not terminal Lower Cambrian (upper Dyeran Stage in Laurentia) but is best regarded as middle Dyeran. Thus, it is older than the “Hawke Bay regression” and the improved slope oxygenation represented by the Middle Granville Formation is now called the Middle Granville Oxic Interval (MGOI, Landing et al., Submitted; Figure 3). In the course of this re-evaluation, Landing et al. (Submitted) questioned the existence of a global “Hawke Bay regressive interval,” restricted the “Hawke Bay” to its type area in NE Laurentia, and regarded it as epeirogenic in origin. They concluded that units such as the Hatch Hill Formation really reflect shore-line derived sand transport from Laurentia at a time of slowly rising sea levels that led to a global hyperwarming interval. In short, Early–Middle Cambrian boundary interval eustatic rise and climate amelioration led the development of a poorly oxygenated (hot) continental shelf in such widely separated areas as northern Vermont (Parker Slate) and eastern California (Mule Springs Formation) (Landing and Bartowski, 1996). (text modified from Landing et al., 2007, p. 64)

STOP 7. Tanner Hill Road
Hatch Hill D/ Interval (Late Early Cambrian‒Earliest Ordovician) Through Early Late Ordovician Paleooceanographic Changes on the East Laurentian Slope(30 MINUTES).
OHM 2020: Locations are very confused and should certainly not be trusted along this road. Why such vague location data is used in 2018, when everyone has GPS in their phone, boggles the mind!
Park just before (east of) lowest exposed rocks on right (north) side of road. Walk uphill and examine east dipping (overturned) section in core of large syncline. Walk through section in Hatch Hill, Deep Kill, Indian River, and Mount Merino formations (Figure 3). Stop 7 ends just west of hill crest with reappearance of Indian River Formation red mudstones on west limb of syncline.
Hatch Hill Formation deposition.
Dolomitic quartz arenites; conglomeratic sandstones; and interbedded, minor dark gray and black siltstones and shales of the Hatch Hill Formation form the lowest part of the section. (Remember that the section is overturned.) The medium- to massively bedded, coarse, lensing, conglomeratic sandstones and conglomerates of the lowest part of this interval have typically been compared to submarine channel-fill deposits in reports on Taconic geology (e.g. Keith and Friedman, 1977, 1978; Friedman, 1979; Rowley et al., 1979). However, significant erosion at the base of purported channels and vertical stratigraphic successions characteristic of channels have never been described in the Taconic regions. These coarse-grained sandstones and conglomerates may simply be sand- and debris-sheets that originated at many places along the shelf–slope break and upper slope, and not from a persistent point source (submarine canyon head).
The sandstones of the Hatch Hill Formation become thinner-bedded and finer-grained higher in the section, and black mudstones become dominant. This entire interval with black mudstones up to an abrupt transition into the 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 [lowest Tremadocian] Hatch Hill dysoxic/anoxic 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 d/a mudstone deposition that has been termed “Poultney A” (abandoned designation, Landing, 1988b) by Theokritoff (1959; Zen 1964).
Deep Kill Formation.
A sharp transition from the Hatch Hill Formation into the overlying green-gray mudstones of the Deep Kill Formation is present in the drainage ditch on the north side of Tanner Hill Road. Limited outcrop of the Deep Kill Formation likely explains the apparent absence of the black mudstone-limestone mesoscale intervals characteristic elsewhere of the formation (Figure 3).
Indian River Formation
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 during early burial to a grayish sediment color, and an alternative explanation for the color of the Indian River must be found.
Landing (1988b) noted three lines of evidence 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. These lines of evidence include: 1) occurrence of radiolarian cherts and thin volcanic ashes undiluted by background argillaceous sediment; 2) thorough burrow-homogenization of much of the unit; and 3) presence of large, up to 3 cm-wide burrowers that were active on a well-oxygenated bottom).
Indian River Formation: oldest synorogenic deposit of Taconic orogeny and its age
The cherty, red slaty mudstones of the Indian River Slate mark an important stage early in the Taconic orogeny. Landing et al. (1992) discussed comparable red mudstones in a number of orogens, where they always underlie green mudstones and higher flysch (e.g., Taconian allochthons from New York to western Newfoundland, southern Uplands of Scotland, Hercynian Rheinisches Schiefergebirge and Harz Mountains, and Jurassic of Japan).
These data suggest that red, cherty, oxygenated shales in collisional orogens reflect the following history: 1) passage of a peripheral bulge through passive margin successions; 2) consequent flexural uplift and restriction of sedimentation on the peripheral bulge to slowly deposited pelagic muds, radiolarian cherts, and thin volcanic ashes of the Indian River-type; and 4) final flexural down-warping and increased rates of deposition as sediment provenance changes to the emergent accretionary prism (initial cherty green mudstones of Mount Merino-type and overlying Austin Glen flysch). 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 north of the road.
The interpretation of the Indian River Formation as reflecting the first indication of the Taconian orogeny in continental slope facies seems also to explain why its color (red and minor green) is so distinct from other continental slope mudstones in the Taconics (Figure 3). Rather than being part of the “standard” black or green macroscale alternations that reflect eustatic and climate change by the hyperwarming model, the red color of this apparently condensed unit and presence of volcanic ashes suggest depositional controls associated with advance of the Taconic arc toward the Laurentian margin.
The age of the Indian River Formation can only be approximated because biostratigraphically useful fossils (e.g., graptolites, conodonts) have not been recovered from it―although it is possible that the radiolarians in the cherts (Ruedemann and Wilson, 1936) might allow correlation of the unit. Geochronologic dating of the ashes is certainly possible, although it has not been done.
The Indian River is likely upper Middle Ordovician (Darriwillian Stage). It is underlain by lower Middle Ordovician rocks (upper Dapingian Stage) of the Deep Kill Formation (Landing, 1976) and overlain by the lower Upper Ordovician (lower Sandbian Stage) Mount Merino Formation with Nemagraptus gracilis Zone graptolites (e.g., Berry, 1962).

STOP 8. Cambrian‒Ordovician Boundary Interval in Dysoxic-Anoxic Hatch Hill Formation
OHM 2020:Locations on this road are very confused. Landing 2007 Stop 6.4 is 1.4 miles from Holcombville road, whereas in this guidebook it is only 1.0 miles.
(Stop 6.4 of Landing et al., 2007)
Graptolite and conodont biostratigraphy
As discussed above, the Cambrian–Ordovician boundary is a type 1 sequence boundary/unconformity across most of the Laurentian craton. However, stratigraphically unbroken Cambrian‒Ordovician boundary intervals occur in successions marginal to NE Laurentia. This is the case in the upper Hatch Hill Formation in northern and southern localities in the Taconic allochthon (Landing, 1993). An understanding of the Cambrian–Ordovician boundary interval in the Taconics came from a re-examination of three supposed “Upper Cambrian” dendroid graptolite genera that had been repeatedly cited from Stop 8 (Berry, 1959, 1961; Bird and Rasetti, 1968; Fisher, 1984).
Examination of these specimens in the NYSM Paleontology Collection showed the assemblage was not Upper Cambrian because it is comprised exclusively of rhabdosomes of the earliest Tremadocian form Rhabdinopora flabelliformis praeparabola Erdtmann, 1982 (Landing, 1993). This monospecific, lowest Tremadocian dendroid assemblage was relocated in a shaly interval (42.9–43.4 m) near the top of this cut on Tanner Hill Road (Figure 22). Twelve samples were processed for conodonts, but only a sample at 22.8 m yielded conodont elements—an upper Sunwaptan Stage fauna [Eoconodontus notchpeakensis Zone with E. (E.) alisoniae Landing, 1983; E. (E.) notchpeakensis (Miller, 1969); Proconodontus muelleri Miller, 1969]. These limited biostratigraphic data indicate that the base of the Ordovician lies in the interval 22.8 to 42.9 m, an interval without any evident stratigraphic break or lithofacies change (Figure 8).
Lithostratigraphy and regional comparisons
The Stop 8 succession in the upper Hatch Hill Formation is dominated by thin- to medium-bedded dolomitic quartz arenites and dark gray silt shales with debris flows with pebble-sized dolomitic sandstone clasts. By comparison, bedded limestones and carbonate clast debris flows dominate the Cambrian–Ordovician boundary interval at a locality approximately 160 km south in the southern Taconics near Hudson, New York (Landing, 1993). As in the late Early Cambrian (Landing and Bartowski, 1996), southern localities in the Giddings Brook thrust seem to have occupied a somewhat higher position on the continental slope and have more prominent bedded limestones and carbonate clast debris flows (Landing, 1993).
The presence of black mudstones of the Hatch Hill d/a interval through the boundary interval obviously does not seem to follow the global hyperwarming model. Perhaps the explanation lies with erosion of boundary interval strata across most of the Laurentian craton with eustatic fall that was not latest Cambrian but earliest Ordovician.

STOP 9. Raceville: Early Middle Ordovician Dysoxic/Anoxic Interval
(Stop 11 of Landing, 2002)
This vertically dipping section shows a black mudstone-limestone, d/a macroscale alternation in the lower (northern) part of the section and overlying greenish-gray, siliciclastic mudstone without limestone intercalations. The ca. 2 m-thick interval of interbedded limestone and black mudstone is repeated three times in the northern half of the roadcut with the northernmost and central repetitions comprising a faulted antiform and the southern repetition fault bounded.
Most of the limestone in the lower part of the section consists of fine-grained, argillaceous lime mudstones 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; Landing, 2012; Figure 3). The climate maximum, eustatic high, and resultant interval of intensified dysaerobic slope water recorded by the Raceville cut is equated with earliest Middle Ordovician (Whiterockian) eustatic rise across Laurentia and with deposition of such units as the Providence Island Dolostone (Figure 3). (modified from Landing, 2002, p. B6-14)

STOP 10. Best Exposed Taconic Section: Browns Pond and Bomoseen Formations on Mettawee River
OHM 2020: This is Lower Turnpike on Google Earth, not Upper
Cross single lane bridge on Mettawee River; park immediately north of the bridge. Walk east for ca. 0.6 km on dirt track that follows the Mettawee River. Descend to river bank at east side of ruins of a small mill(?) foundation. Stop 10 features an overturned (east-dipping), almost completely exposed section from the base of the Middle Granville through the Browns Pond Formation and down into the upper Bomoseen Formation. (Stop 5A of Rowley et al., 1979)
Overview
An exception to the limited outcrops typical of the Taconics is the almost completely exposed succession that extends along the Mettawee River. This is an overturned (ca. 30° E) Cambrian‒Ordovician succession in the Giddings Brook slice (Jacobi, 1977; Rowley et al., 1979). It was measured stratigraphically downward (i.e., upstream) from a horizon low in the Middle Granville Formation. Sample horizons are recorded as meters below the top of the section (e.g., Mett-neg 29.0; Figure 9).
The Lower Cambrian (Bomoseen‒base of Middle Granville formations; Figure 9) extends upstream from the falls over the Bomoseen Formation to 0.6 km east of the Upper Turnpike bridge. The Mettawee River section is on the SE margin of Theokritoff’s (1964) map area, but he mapped the interval between the Bomoseen and Hatch Hill formations as green and purple green slates, although an intervening black slate unit is prominent in this interval (Figure 3). Subsequently, Jacobi (1977; also Rowley et al., 1979) mapped this unit of black siliciclastic mudstone and limestone within the Bomoseen‒Hatch Hill interval.
This black unit is the Browns Pond Formation (Jacobi, 1977; Kidd et al. in Fisher, 1984). Recognition of the Browns Pond Formation as distinct from the lithologically similar, higher Hatch Hill is important for mapping purposes (Rowley et al, 1979; Fisher, 1984) and demonstrates that the two formations represent temporally distinct intervals of strong d/a on the east Laurentian continental slope. This has implications for regional and global geological and climatic history (Landing, 2000, 2007, 2012; 2013A, B).
The Mettawee River section received little early paleontological attention. Walcott (1912, p. 188, USNM locality 21a) reported a linguloid brachiopod and trilobites he identified as “Olenellus?”, “Ptychoparia sp.”, and “Microdiscus connexus (Walcott)” from a “limestone below the first fall of Mettawee River, above the North Granville bridge…” but did not illustrate the specimens and gave limited provenance information. Without modern study, Walcott’s (1912) identifications only indicate a Laurentian late Early Cambrian (Waucoban) age. This fauna could have come from any Lower Cambrian unit in the section as high as the lower Hatch Hill Formation (Figure 3).
The turbiditic limestones and thin debris flows at the top of the Browns Pond Formation (Holcombville Member of Landing et al., Submitted) earlier yielded the oldest acid-resistant microfossils from the southern Taconic allochthon (Landing and Bartowski, 1996). This suggested that older limestones and calcareous clast debris flows exposed through the Browns Pond and extending down into the upper Truthville Formation (Figures 3, 9) on the Mettawee River might have biostratigraphically significant faunas that could provide an upper age bracket on the origin of the east Laurentian passive margin in offshore facies, and also a bracket on the earliest d/a interval on the east Laurentian continental slope.
Lithofacies and sediment provenance
Most of the Browns Pond and Truthville formations on Mettawee River are a low energy facies of planar laminated mudstone and scattered mm-thick sandstones. The succession demonstrates for the first time that the Browns Pond is dominated by thin (cm- to several mm-thick), argillaceous lime mudstone-rich distal turbidites that may have sand- to rare granule-sized clasts of black lime mudstone. A near lack of trace fossils in the Browns Pond (noted below; also Landing, 2012) suggests the substrate was low in oxygen and metazoans were probably rare. The most abundant body fossils are echinoderm sclerites that comprise a few thin lenses and small ripples (to 1 cm-high) and were likely washed off the east Laurentian shelf or upper slope (horizons Mett-neg 112.5; -neg 62.5, -neg 28.5; Figure 9).
Decollement surfaces in the middle Browns Pond (Figure 9) suggest rapid deposition and easterly transport of weakly compacted mudstone. Weak compaction, soft mud at the sediment‒water interface, and low oxygen bottom waters that led to black, organic-rich mud deposition (Landing, 2012) with locally abundant pyrite (Figure 9) suggest that Browns Pond and Truthville shelly fossils are allochthonous (e.g., Thayer, 1983).
Carbonate clast debris flows occur in the Truthville and Browns Pond. The clasts are fine grained, dark colored, and laminated and were locally derived, and not from an Iapetus Ocean-facing shelf or upper slope. Thus, debris flows clasts in the Browns Pond and Truthville yield limited skeletal fossils.
As proposed earlier (Landing, 2007, 2012, 2013a, b), the succession of the Mudd Pond Member quartzite and late Early Cambrian age of the uppermost Browns Pond (Landing and Bartowski, 1996) suggest these units are the slope equivalent of the shelf Cheshire Formation (quartz sandstone tidalite) and Dunham Formation (now dolostone-dominated) (Figures 3, 9). This lithologic correlation and the soft, strongly d/a substrate of the Browns Pond would mean that the abundant, fine-grained carbonate component of bedded carbonates and secondarily transported debris flows came off the Dunham shelf (Figure 3) and that much of the fossil debris was similarly transported.
Trace fossils and bottom water oxygenation
Planolites burrows at the base of thin sandstones of the upper Bomoseen persist into the Truthville Formation where they appear at the base of a few thin sandstone laminae (Figure 9). The Browns Pond has rare trace fossils with a few Planolites at Mett-neg 71.6 to Mett-neg 82 m on bedding surfaces and in slabbed sections.
Most of the Browns Pond Formation lacks any traces, which is consistent with it’s interpretation as a strongly d/a facies (Landing, 2007, 2012, 2013A, B). The only large traces found in our study are black anastomosing forms limited to an interval of green weathering, medium grey mudstone (Mett-neg 55.0 to Mett-neg 56.5) that records a more oxic time on the sea floor (e.g., Landing, 2012). The large traces suggest Megagrapton Książkiewicz, 1968, a distinctive graphoglyptid known from large polygonal nets with common right angle branches (e.g., Książkiewicz, 1970). However, recent analysis of graphoglyptids (Fan et al., 2018) shows that Megagrapton actually has burrow segments intersecting at a variety of angles and thus allows a tentative assignment of the Browns Pond traces to “Megagrapton?” These large traces are not restricted to a bedding surface; they are endichnial and the product of a burrowing organism.
Ichnofossils of the Taconic allochthon have not received modern study, and the Browns Pond specimens can only be compared with the spartan original description and illustration of Fucoides flexuosus Emmons, 1844. They were compared with specimens in the NYSM Paleontology Collection that were collected from the “Cambrian roofing slates” in the vicinity of Middle Granville village that have been assigned to Emmon’s ichnospecies. Both the Browns Pond and Middle Granville specimens have been tentatively assigned to Megagrapton? flexuosus (Landing et al., Submitted). Walcott (1890, p. 603) complicated an understanding of Taconic Cambrian ichnofossils by synonymizing all previously named, elongate taxa with the thin (i.e., thread-like) looping trace Helminthoidichnites Fitch, 1850. This synonymy meant that the much wider looping traces of the genotype Gordia marinus Emmons, 1844, became H. marinus (Emmons, 1844). Walcott (1890, pl. 52) illustrated large and anastomosing traces comparable to those from Mettawee River as H. marinus?
Late Early Cambrian age of syn-rift and passive margin (Truthville‒lower Hatch Hill Formation)
A significant “younging” of the oldest units deposited along the margin of the New York Promontory (Figure 3) is indicated by re-evaluation (Landing et al., Submitted) of the biostratigraphic information on mineralogically immature, likely syn-rift deposits in the Taconic allochthon (i.e., Rensselaer and Bomoseen formations). In addition, small skeletalized faunas appear below and persist through the Browns Pond d/a interval to provide an upper age bracket (late Early Cambrian; middle Dyeran Age in Laurentia) on the rift‒ passive margin transition along the New York Promontory.
Reactivation of rifting or persistence of rifting into the late Early Cambrian (late Dyeran) and into the middle Middle Cambrian is suggested by the first marine transgression of the Ottawa-Bonnechere aulacogen and the coeval, on-strike subsidence of the Franklin Basin on the northwest Vermont shelf (Landing, 2007, 2012, 2013A, B; Landing et al., 2009, In press; Webster and Landing, 2016; Figure 1). The occurrence of Oldhamia-bearing trace fossil assemblages from the Rensselaer and Bomoseen formations comports with Dyeran and not Ediacaran or earliest Cambrian rifting and production and accumulation of immature sediments on the Iapetus-facing margin of the New York Promontory, as well as the Quebec Reentrant and Newfoundland Promontory. Recent study (Herbosch and Vaniers, 2011) shows that the trace fossil Oldhamia is biostratigraphically useful and not a form that could range even down into the Ediacaran. Indeed, Oldhamia is known worldwide from strata that are no older than the late Early Cambrian, and its occurrence in the Rensselaer and Truthville formations thus brackets the transition in NE Laurentia from a rifted margin in the late Early Cambrian in the Taconics of New York and Quebec (see Swett and Narbonne, 1993; Landing et al., Submitted) and western Newfoundland (Lindholm and Casey, 1990).
The trilobite Olenellus from the Cheshire Formation has been reported as the oldest body fossil on the northeast Laurentian passive margin (Osberg, 1969; Allen et al., 2010). However, Olenellus was actually reported to occur much below the Cheshire in chemically immature (syn-rift) arkoses at Clarksburg Mountain, northwest Massachusetts (Walcott, 1888). The Clarksburg Mountain rocks with purported Olenellus were mapped as “Cheshire” by Ratcliffe et al. (1993) with the “Cheshire” claimed to nonconformably overlie gneisses in southern Vermont and adjacent Massachusetts (e.g., Rankin et al., 1989; Hatcher, 2010; Allen et al., 2010). However, Walcott’s (1888) Olenellus report is from rocks that occur only ca. 30 m) above the nonconformity with middle Proterozoic gneisses. This occurrence in conglomeratic arkoses brought to the “Mendon Formation” (Brace, 1953, p. 33; Skehan, 1961), and now best assigned to the Pinnacle Formation (Landing, 2007, 2012, 2013a, b; Webster and Landing, 2016; Figure 3). Walcott’s (1888) supposed Olenellus specimen was never illustrated and a search of the USNM collections by MW did not lead to its recovery. Thus, a presently unidentified trilobite has been found in the oldest syn-rift sedimentary rocks on what became the NE Laurentian shelf, with the trilobite indicating a late Early Cambrian (Epoch 2) correlation comparable to that of the oldest known rocks (i.e., Rensselaer or Bomoseen formations) of the Taconic allochthon. This generalized correlation supports the conclusion that Iapetan rifting persisted or was re-established on the NE Laurentian margin in the late Early Cambrian.
Higher Mettawee River strata provide information on post-syn-rift deposits on the continental slope of NE Laurentia. The age of these deposits are provided by fossils from the Mettawee River section and other localities. The middle Dyeran Elliptocephala asaphoides assemblage spans the upper Browns Pond‒lower Hatch Hill formations. This assemblage brackets the d/a intervals known from these two formations and the more oxic strata of the Middle Granville Formation (Landing and Bartowski, 1996; Landing 2012; Figure 3). The Mettawee River section shows that small skeletalised taxa known from the E. asaphoides assemblage persist lower through the Browns Pond and into the relatively oxic facies of the upper Truthville Formation (Landing et al., Submitted).
The likely sediment source of the carbonate-rich Browns Pond Formation and its fossils was the Dunham Formation on the eastern shelf of the New York Promontory (Figure 3). This slope‒shelf correlation accords with correlation of the Mudd Pond Member quartzite with Cheshire Formation tidalites on the shelf (Landing, 2007, 2012). This correlation of shelf and continental slope units would make the lowest calcareous mudstones of the Browns Pond Formation equivalent to the shelf-margin Moosalamoo Phyllite that underlies the oldest NE Laurentian carbonate platform carbonates (Forestdale Marble; Figure 3).
These correlations mean that a significant amount of rift‒passive margin rocks on the NE Laurentian shelf (Pinnacle‒Dunham) and slope (Rensselaer‒Browns Pond) are lower and middle Dyeran. This remarkable history of events that took place in the earlier Dyeran is actually incomplete as the Middle Granville Formation (and Middle Granville Oxic Interval, or MGOI), its shelf-equivalent HST, and the lowermost Hatch Hill Formation, are all middle Dyeran (upper Elliptocephala asaphoides assemblage interval). Thus, reestablishment of reduced slope water circulation and deposition of dysoxic/anoxic mudstones of the lower Hatch Hill Formation beginning in the middle Dyeran in the Taconics of eastern New York and adjacent Vermont are significantly older than the upper Forteau and lower Hawke Bay formations (upper Dyeran) of western Newfoundland.
The MGOI is proposed as replacement for the earlier named “Hawke Bay Oxic Interval” in theTaconics by Landing (2012; “HBOI” abandoned by Landing et al., Submitted). Correlation of the lower Hawke Bay Formation in western Newfoundland as upper Dyeran means that it is the lower, but not lowermost, part of the lengthy Hatch Hill d/a interval (middle Dyeran to lowermost Ordovician; Landing 2012) that really correlates with the Hawke Bay Formation and “events” (Landing et al., Submitted). As concluded by the latter authors, the designation “Hawke Bay Events” (“HBE”) should probably be limited to NE Laurentia where a comparable rifting history that persisted into or was reinitiated in the late Early Cambrian along the Newfoundland and New York promontories and included similar, likely coeval, lower passive margin successions (Bradore‒Hawke Bay and Cheshire‒Monkton formations).
The paradigm that the Hawke Bay Events necessarily featured major eustatic regression and an interregionally extensive unconformity should be questioned as noted by Nielsen and Schovsbo (2015) who regarded the “Hawke Bay regression” in Baltica as an epeirogenetic event. Similarly, Knight et al. (2017) also suggested that the “type” Hawke Bay Event unit, the Hawke Bay Formation in western Newfoundland may also be the record epeirogenic activity, not eustatic changes. Indeed, available evidence can be used to interpret the Hawke Bay Formation not as a regressive sedimentary unit but as an HST lithosome. By this proposal, the Hawke Bay, and possibly Monkton Formation in NE Vermont, were deposited either during an interval of reduced rate of sea-level rise and development of a shallow-water siliciclastic by-pass shelf with significant offshore transport of shoreline-derived quartz sand (Landing et al., Submitted). The qualification “local epeirogenic uplift” is appropriate in NE Laurentia as approximately coeval (upper Dyeran‒Middle Cambrian) deposits of the Parker Formation in NW Vermont that are correlated with the Hawke Bay Formation record an abrupt and continuing epeirogenic foundering of the Franklin Basin and at least the eastern part of the Ottawa-Bonnechere aulocogen (Landing et al., 2009; Webster & Landing, 2016). (modified from Landing et al., Submitted)

West edge of Taconic Allochthon
The ridge to the left (east) of Rte 40 is formed of more resistant 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 allochthon and edge of the Giddings Brook slice. Rte 40 meanders back and forth from the edge of the allochthon and onto erosionally less resistant, structurally underlying Upper Ordovician “Snake Hill” synorogenic flysch and wildflysch (see Landing et al., 2003b; English et al., 2006) and, in the area of Middle Falls, onto Lower and Middle Ordovician shelf carbonates.

High road cut on left (east) is in overturned Deep Kill Formation.
The road cut shows repetitions of meter-scale Logan cycles that are relatively tectonized (faulted). Better preserved Logan cycles will be seen at Stop 11. This road cut is Stop 12 of Landing (2002).

STOP 11. Logan Cycles in the Deep Kill Formation: Mesoscale Cycles in the Milankovitch Band.
Stop on edge of Rte 40 close to ca. 10 m-high cut with prominent banded slaty mudstones (Logan cycles) on left (east) side of road. CAUTION! Traffic can be heavy so cross road after discussion only with care. No hammers on this superb locality, take pictures only. When examining road cut, stay close to the steel barrier or walk inside the barrier and hug the rocks.
This is one the most accessible and well preserved outcrops that shows the color and lithologic alternations characteristic of Logan cycles (Landing and Benus, 1985; Landing et al., 1992) in the Lower‒ lower Middle Ordovician Deep Kill Formation. NO HAMMERS AT THIS STOP!
The road cut lies close to the leading (west) edge of the Taconic allochthon and Giddings Brook slice. The section is overturned and dips gently to the east. The utility of Logan cycles (Figure 10) in geological field work is that the upward transition of muddy intervals provides tops-sense in successions that otherwise lack much in the way of primary structures. Thus, successions with ABCABC, BCBC, or ABABABCABC, etc. alternations can be invariably used to determine tops-sense in a “monotonous” mudstone or slate sequence because A-units form the base of tripartite Logan cycles and C-units form the top.
About six Logan cycles are present in this ca. 8 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 10). The buff-weathering, dolomitic C-intervals are often the thickest part of the Logan cycles at Stop 11 and grade imperceptibly down into greenish B-intervals. Several of the Logan cycles show the abrupt upward transition from the C-interval into the black mudstone (A-unit) of the overlying Logan cycle.
Well preserved Logan cycles that extend through a completely exposed section through the Floian Stage at the leading edge of the Taconian allochthon at Levís, Québec, appear to have an average duration of less than 100 k.y. (Landing et al., 1992). EL’s experience is that only the more proximal parts of Laurentian slope successions (i.e., the Levis area in Québec and the central and southern Giddings Brook slice in eastern New York) show Logan cycles, while more distal successions are comprised of completely oxic mudstones with green, purple, and red color.
Logan cycles should be seen as potentially economically significant as they comprise 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). However, the rocks of the Giddings Brook slice record a much higher burial temperature, with euconodont elements typically opaque black in color and showing a C.A.I. (Color Alteration Index) of 4.0, which means that any organic materials have been transformed into graphite (e.g., Landing, 1976, 1993).

Pass by intersection with Bald Mountain Road on right.
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 (Floian) 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.

Note cut in vertical beds of lower Trenton Group
Note cut in vertical beds of lower Trenton Group (Upper Ordovician, Katian) limestones and thin interbedded black mudstones along east side of road and opposite wide dirt pull-off 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 southwest, the Proterozoic Grenvillian 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, possibly early Middle Ordovician and referable to the Providence Island Dolostone).

STOP 12. Schaghticoke Gorge: Late Early Tremadocian Dysaerobic Interval.
Turn right on curve immediately after red brick church and park in Hoosick River overlook parking area. Until recently, the rocks in Schaghticoke gorge below the dam could be accessed by walking downslope under the Rte 40 bridge. Unfortunately, the response to drinking parties and a few drownings has meant that an iron fence has been put along the gorge, and this makes access difficult. What will be done for Stop 12 will be a photographic overlook downstream from the eastern part of the bridge. Two areas on the north bank and on the upstream side of the island in the middle of the Hoosick where black mudstones of the Schaghticoke dysoxic/anoxic interval crop out can be seen. After this, CAREFULLY cross the bridge roadway, and the same Schaghticoke d/a interval can be seen tectonically repeated several times from a viewpoint directly down from the NE end of the bridge.
(Stop 13 of Landing, 2002).
The section appears to comprise a tectonically isolated slice at the leading edge of the Taconic allochthon and Giddings Brook slice. This is the type section of the upper lower Tremadocian (Lower but not lowermost Ordovician) “Schaghticoke Shale” (term abandoned by Landing, 1988b, and grouped as the lowermost part of the Deep Kill Formation).
Schaghticoke gorge features the structural duplication of a thin (ca. 4 m) black mudstone-limestone macroscale alternation in greenish-gray slaty mudstones. 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 Québec. Unfortunately conodonts from limestones in the Schaghticoke d/a interval in the Hoosick River gorge 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 (based on graptolites) black shale-limestone macroscale alternation exposed for 400 km along the south shore of the St. Lawrence River in Québec and which yields Rossodus manitouensis Zone conodonts (Landing et al., 1986). The Schaghticoke d/a interval of climate maximum and intensified dysaerobia on the Taconic continental slope and elsewhere along the Laurentian margin (Landing, 2012) is equated with the eustatic rise and the Stonehenge transgression (Taylor et al., 1992) that led to deposition of the Tribes Hill Formation as a type 1 depositional sequence (Stop 4).

Most of route is on low-lying terrane developed on Late Ordovician Utica Fm
Most of route is on low-lying terrane developed on Late Ordovician Utica Formation (black mudstone with abundant graptolites). A few miles south of Lake George village, note road cuts in middle Proterozoic Grenvillian orogen (Adirondack basement).

B1: GEOLOGY OF THE COPPER-KILN LANDSLIDE: A GLIMPSE INTO THE MARCY MASSIF DETACHMENT ZONE
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S.P.Regan
Dept of Geosciences
University of Alaska, Fairbanks
Fairbanks AK 99775
V.E.Guevara
T.Drauschak
Skidmore College
Saratoga Springs NY 12866
J.R.Chiarenzelli
Dept of Geology
St. Lawrence University
Canton NY 12230
R.M.Gaschnig
Env, Earth and Atmospheric Sciences
University of Massachusetts, Lowell
Lowell MA 01854
The Mesoproterozoic Grenville Province of eastern North America is a classic Proterozoic orogenic belt that formed over a protracted period resulting in the assembly of Rodinia (Hoffman, 1988). Despite the routine interpretation that the Grenville Province exposes the orogenic infrastructure of modern collisional orogens, the Grenville is unique from Phanerozoic orogenic belts, as anorthosite and related mafic rocks comprise 20% of the exposed region (Fig. 1a; Corriveau et al., 2007). This is in stark contrast to all Phanerozoic orogens, which lack massif-type anorthosite, let alone in such quantities or volumes (Ashwal, 1994; 2010). Owing to a high rheologic strength resulting from exceptionally coarse grain sizes, high temperatures of crystallization, and relatively low density, it follows that anorthosite complexes may have had a profound impact on the thermal and structural architecture of the Grenville Province as a whole.
The Adirondack Mountains in northern New York are a Mesozoic domical uplift of Grenvillian basement and form the southern extension of the contiguous Grenville Province (Fig. 1b; Roden-Tice et al., 2000). The region has served as a testing ground for petrologic inquiry and new analytical techniques for decades (Kemp, 1898; Buddington, 1939; Postel, 1954; Valley and O’Neil, 1982; Bohlen, 1985; Spear and Markussen, 1997; Bonamici et al., 2015; Quinn et al., 2017; among many others). The Adirondack Mountains contain several anorthosite complexes including the Marcy massif, the Oregon dome, and Snowy Mountain dome. The Marcy anorthosite massif (Mm) is a 3000 km2 classic Proterozoic-type anorthosite complex underlying the majority of the High Peaks region of the Adirondack Park in northern NY, and is thus an excellent place to study the impact of volumetrically abundant anorthosite on the orogenic infrastructure of the Grenville Province.
In late August, 2011, Tropical Storm Irene catalyzed over 40 landslides in the High Peaks region (Mackenzie, 2017). The highest concentration of landslides occurred within the eastern High Peaks area, and provide an unprecedented window into the internal compositional heterogeneity and structural complexity within the relatively undeformed Marcy massif (Chiarenzelli et al., 2015). As a companion to a 2015 NYSGA field trip (Chiarenzelli et al., 2015), which focused on the Bennies Brook slide on the northwestern slope of Lower Wolfjaw Mountain, we looked at a landslide in proximity to the northern margin of the Marcy massif. The Copper-Kiln (or Cooper-Kill) landslide occurred north of Wilmington near Whiteface Mountain, and combined with detailed 1:24,000-scale mapping along the southeastern margin of the Marcy massif, provide insight into how the anorthosite behaved during the polyphase metamorphic and structural evolution of the Adirondack Mountains.

B1: GEOLOGY OF THE COPPER-KILN LANDSLIDE: A GLIMPSE INTO THE MARCY MASSIF DETACHMENT ZONE

Meeting Point - Wilmington Town Beach

Trailhead

STOP 1. Deformed gabbro and gabbroic anorthosite
The first stop is not connected by continuous exposure to the rest of the slide. It is located in the outwash of the landslide. The pavement outcrop exposes deformed gabbroic anorthosite and gabbroic gneiss. There is very minor obliquity between the contact between the two lithologies and penetrative gneissosity indicating that a higher angle relationship may have been progressively rotated into parallelism. There is abundant syn-kinematic hornblende ± biotite at this outcrop consistent with decompression from peak granulite-facies metamorphism during shear.

STOP 2. Bottom of the Copper-Kiln landslide
The base of the landslide composed of gabbroic anorthosite to fine-grained anorthosite with diffuse contacts. Coarse orthopyroxene porphryoclasts are aligned in the predominant foliation and rimmed by hornblende. Several garnet porphyroblasts are present, but almost entirely replaced with randomly oriented hornblende. The predominant foliation dips moderately to the west northwest. There are also two Neoproterozoic mafic dikes approximately 10 cm wide.

STOP 3. Sub-concordant leucogranite to syenite intrusions in anorthosite
Along the landslide there are abundant meter-scale sub-concordant syenite to leucogranitic sills, preserving varying degrees of obliquity with host rock gneissosity. They are difficult to spot at first as they weather to a pale gray/pink and are initially difficult to distinguish from recrystallized anorthosite. The dikes contain minimal evidence for high degrees of strain, but the extent of post kinematic annealing still remains problematic. However, the preservation of perthitic textures, and subhedral and isolated mafic minerals, we interpret minimal subsolidus strain to have occurred subsequent to intrusion. They lack the very recognizable mesoscopic fabrics contained in surrounding anorthosite to gabbroic anorthosite gneisses. This dike was analyzed for U-Pb zircon geochronology and Ti in zircon thermometry. Based on results discussed above, we interpret the leucogranitoids to be members of the Lyon Mountain ferroan leucogranite suite (Chiarenzelli et al., 2017), and to have been emplaced synchronous with extensional collapse of the Adirondack Highlands (Regan et al., in review). The Mm detachment zone may have aided in their initial emplacement as seen here.
As we continue walking up the slide, there will be countless opportunities to discuss the relationships of the leucogranitoid sheets and host rocks. A Google Earth file of all data available on NYSGA website with all measurements, locations, photographs, and descriptions.

STOP 4. Skarn near top of landslide
As the landslide steepens and we approach the headwall, you will notice an increase in the number of pods, screens, and partially digested and subsequently deformed calc-silicate xenoliths. Xenoliths range in scale from cm-scale to several meters in thickness and display a varied mineralogy of garnet, diopside, wollastonite, dolomite, titanite, among others. Titanite can be seen in hand sample and is the root beer colored mineral in the calc-silicate lithologies. Thick quartz-veins several cm in thickness are also common in this package of rocks.

B2:THE PISECO LAKE SHEAR ZONE: A SHAWINIGAN CRYPTIC SUTURE IN THE SOUTHERN ADIRONDACK MOUNTAINS, NEW YORK
David W. Valentino
Department of Atmospheric and Geological Sciences
SUNY Oswego,
Oswego, NY 13126
david.valentino@oswego.edu
Jeffrey R. Chiarenzelli
Department of Geology
St. Lawrence University
Canton, NY 13617
jchiaren@stlawu.edu
ABSTRACT
Highly deformed Piseco granitic gneisses occur in an arching east-west transpressional ductile shear zone (Piseco Lake shear zone) that spans the width of the exposed southern Adirondacks. The highly deformed granitic gneisses have restricted silica content, are metaluminous, alkali-calcic to calc-alkalic, continental arc trace element signatures. These granitic rocks intruded supracrustal gneisses resulting in extensive Shawinigan partial melting. The Piseco Lake shear zone (20-30 km wide) formed in this belt of granitic rocks and correlate with a pronounced arcuate-shaped high magnetic anomaly. The magnetic anomaly extends well beyond the exposed Adirondack basement window.
The shear zone is 20-30 km wide and is believed to be the location of a cryptic suture because it occurs between the Adirondack Highlands (underlain primarily by anorthosite and related granitic rocks, AMCG suite) and the Southern Adirondack Terrane (underlain by calc-alkaline tonalitic arc rocks) (Valentino et al., in press). Within the shear zone, the original megacrystic granite contains lineated quartz and rodded feldspar aggregates up to a meter long in places. Along the axis of the shear zone there are thick (1-2 km), subvertical zones of granitic L-S and L-tectonites.
The northern domain of the zone is defined by large foliation domes that are cored by L-tectonite. The southern limbs of the domes steepen toward the south and merge with a wide zone (up to 15 km) of steeply dipping granitic mylonite. Overall, the shear system (domes and steep mylonite zone) forms the core of a region of intense ductile deformation with left-lateral kinematic indicators and subhorizontal E-W ribbon lineations.
The Piseco granitic suite are highly deformed suture-stitching arc plutons that intruded within a sinistral, oblique-convergent, shear system in the deep crust during the Shawinigan orogeny. This is ductile shear zone is the most continuous and largest in the entire Adirondack massif. The shear zone, associated granitic rocks, and the magnetic anomaly abruptly trends toward the south in the eastern Adirondacks. Just beyond this location, the magnetic anomaly appears to be truncated by a branch of the NY-AL magnetic lineament. Following the trace of the magnetic anomaly toward the west, suggests that the shear zone continues for a considerable distance beyond the Adirondack window. It’s magnitude, in addition to the magnitude and extend of the associated magnetic anomaly, suggests that the Piseco shear zone penetrates the Moho.
The current field trip is an update on our very long research project, and it’s geared toward an undergraduate student audience. All field locations were picked to accommodate large student groups. Sampling is generally prohibited by NYS law in the Adirondack Park, and we encourage future instructors to help preserve the field locations presented herein by showing and discussing, and not removing the spectacular bedrock features. Note that the field guide included here in was pirated and modified from the Friends of the Grenville field conference run by D. Valentino, J. Chairenzelli, D. Piaschyk, L. Williams and R. Peterson in 2008.

B2:THE PISECO LAKE SHEAR ZONE: A SHAWINIGAN CRYPTIC SUTURE IN THE SOUTHERN ADIRONDACK MOUNTAINS, NEW YORK

Meeting Point - Charlie Johns Store

STOP 1. Supracrustal rocks of the Adirondack Highlands, just north of the Piseco Lake shear zone
There are several rock types that can be observed at this outcrop, along with the contacts and primary compositional layering. The northern ¼ of the outcrop consists of garnet amphibolite, the next ∼ ¼ of the outcrop is calc-silicate gneiss and the southern ½ is fine-grained quartzo-feldspathic gneiss locally interlayered with highly folded marble (Figure 7). Complex sheath folds with sub-horizontal tight, isoclinal and sheath folds of foliation and compositional layers dominate the rock, particularly in the gneiss (trend to 110o). Excellent views of the folds are seen on top of the southern part of the outcrop. The contacts between the rock types are also folded in similar fashion. Matrix minerals define lineations that are E-W and sub-horizontal, but are variable in orientation due to folding. Overall, the diverse lithologies and style of deformation are typical for the supracrustal rocks located north of the Piseco Lake granitoids. At this point, note that primary compositional layering is preserved at this outcrop.

STOP 2A
The Piseco Lake shear zone includes the northern foliation domes that merge with the steeply dipping shear zones to the south. This series of field stops shows variations in the attitude and type of fabrics that occur in the core of the dome at Piseco Lake. Stops 2A to 2D are a driving traverse around Piseco Lake, the type location of the Piseco antiform. At all of these localities, dynamically recrystallized feldspars and quartz form spectacular ribbon- and rod-shaped mineral lineations (McLelland, 1984), in addition to accessory biotite and magnetite. In many places, the alignment of ribbons forms the foliation in this outcrop. Individual quartz-ribbons have aspect ratios upward of 60:1. At Stop 2A, the foliation is weakly to moderately developed, and dips shallowly southward at the eastern part of the outcrop, but is steeply dipping at the western end of the outcrop. The transition between these different foliation attitudes is difficult to determine because the intensity of the foliation is variable. Lineations are penetrative on all scales, and consistent in attitude (Trend: 110°; Plunge: 05°). Stop 6A occurs on the southern flank of the map-scale dome portion of the Piseco antiform. At the western end of the outcrop there are rods of amphibolite (10-30 cm diameter) within the granitic gneiss.
There is an apparent Mesozoic fault that traces down the western side of Piseco Lake and has locally displaced parts of the Piseco dome (Figure 3). At Stop 2B, again the foliation is weakly developed with a penetrative shallowly plunging lineation. However, the foliation, where it can be observed, dips northerly. As you drive southwest along the western side of Piseco Lake, note that the foliation gradually shallows and then dips southerly. There are several outcrops that can be observed where this transition in foliation attitude occurs. The rock fabrics at Stop 2C are similar to those at Stops 2A and 2B, but again the variably developed foliation dips toward the south. Stop 2D is located at the intersection of Rts. 8 and 10, and the foliation and lineation is penetrative.
Common to all of these field stops is that biotite and sometimes chlorite blades form microscopic lineations and foliation parallel to the macroscopic structure. Rare grains of hypersthene have been found, but they always have well developed overgrowth textures that include biotite and chlorite. The biotite and chlorite are the most abundant index minerals in the granitic gneiss, and suggest the deformation was last active under low- to moderate- metamorphic conditions, although probably began at much higher conditions to account for the relict grains of hypersthene.

STOP 2B:
The Piseco Lake shear zone includes the northern foliation domes that merge with the steeply dipping shear zones to the south. This series of field stops shows variations in the attitude and type of fabrics that occur in the core of the dome at Piseco Lake. Stops 2A to 2D are a driving traverse around Piseco Lake, the type location of the Piseco antiform. At all of these localities, dynamically recrystallized feldspars and quartz form spectacular ribbon- and rod-shaped mineral lineations (McLelland, 1984), in addition to accessory biotite and magnetite. In many places, the alignment of ribbons forms the foliation in this outcrop. Individual quartz-ribbons have aspect ratios upward of 60:1. At Stop 2A, the foliation is weakly to moderately developed, and dips shallowly southward at the eastern part of the outcrop, but is steeply dipping at the western end of the outcrop. The transition between these different foliation attitudes is difficult to determine because the intensity of the foliation is variable. Lineations are penetrative on all scales, and consistent in attitude (Trend: 110°; Plunge: 05°). Stop 6A occurs on the southern flank of the map-scale dome portion of the Piseco antiform. At the western end of the outcrop there are rods of amphibolite (10-30 cm diameter) within the granitic gneiss.
There is an apparent Mesozoic fault that traces down the western side of Piseco Lake and has locally displaced parts of the Piseco dome (Figure 3). At Stop 2B, again the foliation is weakly developed with a penetrative shallowly plunging lineation. However, the foliation, where it can be observed, dips northerly. As you drive southwest along the western side of Piseco Lake, note that the foliation gradually shallows and then dips southerly. There are several outcrops that can be observed where this transition in foliation attitude occurs. The rock fabrics at Stop 2C are similar to those at Stops 2A and 2B, but again the variably developed foliation dips toward the south. Stop 2D is located at the intersection of Rts. 8 and 10, and the foliation and lineation is penetrative.
Common to all of these field stops is that biotite and sometimes chlorite blades form microscopic lineations and foliation parallel to the macroscopic structure. Rare grains of hypersthene have been found, but they always have well developed overgrowth textures that include biotite and chlorite. The biotite and chlorite are the most abundant index minerals in the granitic gneiss, and suggest the deformation was last active under low- to moderate- metamorphic conditions, although probably began at much higher conditions to account for the relict grains of hypersthene.

STOP 2C
The Piseco Lake shear zone includes the northern foliation domes that merge with the steeply dipping shear zones to the south. This series of field stops shows variations in the attitude and type of fabrics that occur in the core of the dome at Piseco Lake. Stops 2A to 2D are a driving traverse around Piseco Lake, the type location of the Piseco antiform. At all of these localities, dynamically recrystallized feldspars and quartz form spectacular ribbon- and rod-shaped mineral lineations (McLelland, 1984), in addition to accessory biotite and magnetite. In many places, the alignment of ribbons forms the foliation in this outcrop. Individual quartz-ribbons have aspect ratios upward of 60:1. At Stop 2A, the foliation is weakly to moderately developed, and dips shallowly southward at the eastern part of the outcrop, but is steeply dipping at the western end of the outcrop. The transition between these different foliation attitudes is difficult to determine because the intensity of the foliation is variable. Lineations are penetrative on all scales, and consistent in attitude (Trend: 110°; Plunge: 05°). Stop 6A occurs on the southern flank of the map-scale dome portion of the Piseco antiform. At the western end of the outcrop there are rods of amphibolite (10-30 cm diameter) within the granitic gneiss.
There is an apparent Mesozoic fault that traces down the western side of Piseco Lake and has locally displaced parts of the Piseco dome (Figure 3). At Stop 2B, again the foliation is weakly developed with a penetrative shallowly plunging lineation. However, the foliation, where it can be observed, dips northerly. As you drive southwest along the western side of Piseco Lake, note that the foliation gradually shallows and then dips southerly. There are several outcrops that can be observed where this transition in foliation attitude occurs. The rock fabrics at Stop 2C are similar to those at Stops 2A and 2B, but again the variably developed foliation dips toward the south. Stop 2D is located at the intersection of Rts. 8 and 10, and the foliation and lineation is penetrative.
Common to all of these field stops is that biotite and sometimes chlorite blades form microscopic lineations and foliation parallel to the macroscopic structure. Rare grains of hypersthene have been found, but they always have well developed overgrowth textures that include biotite and chlorite. The biotite and chlorite are the most abundant index minerals in the granitic gneiss, and suggest the deformation was last active under low- to moderate- metamorphic conditions, although probably began at much higher conditions to account for the relict grains of hypersthene.

STOP 2D

STOP 3. Steeply dipping mylonite zone of the southern PLsz
Along Powley Road, depending on the time of year and the amount of road maintenance, there are a sub-continuous series of pavement exposures located in the road bed, as well as in the gutter on the east (northeast-bound) side of the road. Due to the nature of this location, the extent of the exposed rock at this stop changes yearly, so some or all of the rock described here may be viewable depending upon the time of the season in which the stop is visited (best later in the summer). The best exposures occur along the road between West and East Notch Mountains.
All rocks in this region are very similar in mineral content, and vary only in detail with regard to mineral percentage and fabric type and intensity. The rock is dominantly granitic gneiss with intense subhorizontal to shallowly plunging mineral elongation lineation that trends on average about 095 o, with steeply dipping generally east-west striking foliation. Both fabric elements are defined by ribbons of quartz, and ribbons of aggregate feldspar and quartz (generally 1-5 cm long depending upon grain size). Intensity of the fabric varies across strike at the 50 cm scale, with local layers of significantly coarser-grained fabrics (grains up to 1 cm in diameter). There are also places where the foliation intensity varies as seen at the field stops around Piseco Lake. Rare amphibolite bodies that are 10’s of cm thick occur within the granitic gneiss. Shear sense indicators are abundant in the granitic rocks and consistently show sinistral shear sense.

STOP 4. Southern extent of steeply dipping mylonite
Here the granitic gneiss fabrics contain both a penetrative foliation and lineation. The foliation is steeply dipping and strikes about east-west. Mineral elongation lineation defined by linear aggregates of quartz and feldspar are subhorizontal. The extent of readily available bedrock exposure diminishes south of this location, so this may be the southern-most exposure of the Piseco Lake shear zone. Note that this location is about 21 kilometers across strike from the northern side of the Piseco dome where the pronounced lineation occurs.

Parking for Stop 5

STOP 5: L>>S and L-tectonite in the core of the Piseco antiform
Excellent outcrops on the northwestern side of a small hill just west of the parking area. Follow the dirt road to a path through the woods, and then head up hill to the south to the outcrops. This outcrop of granitic gneiss contains domains of LS and LS. The LS domains contain large and numerous σ-type shear sense indicators, some δ-type are present but are much less frequent. The porphyroclasts are large about 1-3 cm and the recrystallized porphyroclastic material is often wrapped with a quartz ribbons. The interpreted shear sense is low-angle and left lateral. The foliation strikes east-west and dips gently to the south. Return to the vehicles by following the path. At this point the trip is over. Retrace your trip to Route 8. If you are headed back to Lake George, then follow Route 8 back to Speculator, NY, then turn south on Route 30.

B3: FAULT SYSTEMS OF THE TACONIC FORELAND; WHITEHALL, NY TO WEST HAVEN, VERMONT [ALL KINDS OF FAULTS!]
William S.F. Kidd
Chul Lim
formerly of Geological Sciences
University at Albany,
Albany, NY 12222
wkidd@albany.edu
Stephen S. Howe,
Atmos. & Environmental Sciences
University at Albany
INTRODUCTION AND SUMMARY
Map-influencing faults of all displacement kinds are observed in the anomalously narrow Taconic orogenic foreland belt in the region where it crosses the New York - Vermont border between Whitehall, NY and West Haven, VT. Outcrop and map evidence for the relative age and displacement of these faults can be seen on this trip. In general, transported Cambro-Ordovician rocks of the continental shelf/platform and of the off-shelf strata in the Taconic Allochthon are affected by thrust faults of a range of ages relative to development of folding and cleavage in particular rock packages. The thrusts all predate large-scale strike-parallel east-side-down normal faulting, in most places in the area to be visited apparently localized, in a large displacement (1000m or more), on one such fault, the Mettawee River Fault. The youngest of the fault sets consist of cross-strike, strike-slip faults, mostly ENE- to NE-trending, some of which have mappable apparent left-lateral displacement, cutting thrusts and the normal fault. Movement on the strike-slip fault system at West Haven was accompanied in this area by fluids shown by our measurements of fluid inclusion homogenization to have been at temperatures sufficiently high to be most likely of Ordovician (Taconic) age. These oblique faults probably had a prior history as components of a trench forebulge normal fault system; they may have influenced at that time and perhaps also earlier the sedimentation on the Laurentian continental shelf, and therefore they may have been localized by an old inherited fracture/fault system in the underlying Grenville basement of this area. More generally, forebulge normal fault scarp topography, whose development mostly predated thrust arrival at any position, are perhaps also responsible for the repeated detachment and stacking of thin slices of the medial Ordovician limestones of the uppermost carbonate shelf which are extensively developed in the area of this field trip. The lower contacts of these slices show evidence of thrust fault transport, and do not support a model of large-scale olistostromal origin. Because the trip is confined to the area affected by thrusts, we will not see unmodified forebulge-type normal faults in autochthonous basement and platform strata, but they occur just to the west along southernmost Lake Champlain. And there are other kinds of fault affecting geological maps of the area which may be mentioned during the trip.

B3: FAULT SYSTEMS OF THE TACONIC FORELAND; WHITEHALL, NY TO WEST HAVEN, VERMONT [ALL KINDS OF FAULTS!]

STOP 1A. Mettawee River Fault at the Type Locality
[NO HAMMERS OR SAMPLING PLEASE - NY DEC REGULATIONS]
The carbonates here along the west bank of the river are mapped by Fisher (1985) as early Ordovician Providence Island Formation, and dip about 10-20 degrees east; they are the upper part of a thrust transported section above the Comstock (= Pinnacle) Thrust. The lowest part of the section here consists of dolostones; the upper 4 meters contains limestone beds; we think it is possible that the part of the section containing limestones is either basal Chazy or Black River. These show some bedding surfaces with well developed ripples, and small mudcracks, and selectively dolomitized burrows. Local meter-scale folding of periclinal geometry affects these beds near this end of the section. Across the river at this point, you can see the dark mid-Ordovician non-calcareous shales which are there partly disrupted with phacoidal melangetype fabric. In the low cliff, a gently east-dipping planar vein of fibrous quartz cuts the deformed shales; the vein has top to west, thrust sense of shear. Walk north along the west bank of the Mettawee River on the mostly dolostone outcrop. About 170 meters north [43.473241; -73.368196], 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 this and other lower units 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 could infer either that the normal fault mapped to the north passes east of this outcrop within the shale belt, or west within the transported shelf carbonate section (of course badly exposed in that particular place), or that the motion is confined to a surprisingly narrow zone along the contact; the full exposure at Stop 6B suggests the latter possibility is more likely than one might be inclined to think based on this outcrop at Stop 11 alone.

STOP 1B. Deformed Flysch and Melange
Deformed Flysch and Melange Below Contact With Transported Chazyan Shelf Carbonates, in the Banks of the Mettawee River
Walk back to the south end of the carbonate outcrop on the west bank of the river [43.471240, - 73.368354]. If the water is high or it is raining, Stop 1B will be omitted. If conditions permit, scramble along the west bank to the south, and through the scrub along the west side of the old mill race cut at the waterfall, then walk south along the west bank of the river for about 1 km.
The Mettawee River here provides a superb series of exposures (location in Fig 3) that cross from parautochthonous shelf carbonates (mostly dolostones) at the access point, south past the waterfall (43.470990; -73.367826) through alternating zones of flysch-type dark shales, and shale melange, to a large sliver of Chazyan carbonate (mostly limestones, starting at 43.462450; -73.362534) (Fig. 3). The eastern part of the carbonate sliver is poorly exposed in the river, but mapping demonstrates that this fault contact (the Taconic Frontal Fault) cuts obliquely across Taconic stratigraphy and large-scale fold axial traces, and is therefore a post-slaty cleavage generation structure (Fig. 3).
The allochthonous carbonate at this locality and along the western edge of the Allochthon in general has been interpreted and shown on a published map by Fisher (1985) as blocks-in-shale (i.e., olistoliths) rather than as coherent fault slivers. This locality (and stop 5) provide excellent examples to test these two alternative hypotheses. The carbonate/melange contact exposed in the woods across the river here can be walked for nearly 2 kilometers as a continuous, unbroken structure along strike, with internal fold axes in the carbonates approximately parallel to the general contact trend (Selleck and Bosworth, 1985, Plate 1A). The carbonate must be in the form of a large, probably composite sheet. Minor secondary disruption near its margins is undoubtedly present (of structural and perhaps sedimentary causes), but we maintain that the origin of this and other km-(length)-scale carbonate sheets adjacent to the Allochthon was by detachment of thin slices from the top part of the underthrusting shelf rocks and their incorporation to and imbrication at the base of the advancing thrust pile. We think based on our mapping of all this belt north to West Haven that it is highly misleading to describe the geometry of the large carbonate bodies along the western edge of the Allochthon as "blocks-in-shale", and that it is most unlikely that they arose as primary sedimentary slump features (further discussed in Rowley and Kidd, 1982).

STOP 1C. Mettawee River Fault Along Upper Turnpike
**Please stay on the road and its margins; we do NOT have permission to enter the field or woods here**
Two outcrops of shale melange occur a short distance into the field to the east, where outcrop on the west side of the road and into the woods adjacent exposes dolostones of the Cambro-Ordovician shelf (here Fort Ann Fm according to Fisher, 1985; see Fig. 3). These outcrops demonstrate the existence of the Mettawee River Fault as an east side down normal fault and closely constrain its position.

STOP 2. Potsdam/Ticonderoga Arenites and the Comstock (Pinnacle) Thrust.
In low roadside exposure on the west side of the road gently-dipping Upper Cambrian sandstones can be seen, mapped by Fisher (1985) as Ticonderoga and upper Potsdam Formations. The map shows they can be systematically traced for 9 kilometers to the south and, for a shorter distance, to the northeast of this place. Along with the overlying carbonate shelf sequence they demonstrate a major thrust duplication of almost the whole thickness of the local Cambro-Ordovician continental shallow marine platform/shelf sequence. The thrust carrying these strata was termed the Comstock Thrust by Fisher (1985); it climbs rapidly at a lateral ramp 9 km south of this locality (see Fig. 4, and Hayman & Kidd, 2002a; Stop 12) and can be traced a further 11 km carrying limestones of the uppermost part of the Ordovician shelf sequence before it climbs again, into the shale-melange belt, near Smiths Basin.
This small outcrop is not photogenic, and the thrust itself is not exposed here (exposures of it have been found in the woods to the west and southwest), but the map demonstrates its reality, and also that it and the sheet of shelf strata it carries are cut and truncated by the Mettawee River Fault (normal, east side down). Limestones of medial Ordovician age form a low outcrop at the sharp corner in the road to the north, and a larger outcrop on the other side beyond, and are shown by Fisher as Isle La Motte Formation (= Middlebury limestones) and form the foot wall of the Comstock Thrust here. It is probable that the original thrust relationship has been modified at this location by a later ENE-trending fault, but this does not significantly change the basic overthrust relationship.
Hayman & Kidd (2002a, b) provide reasons why this thrust should be specifically identified structurally as the southern extension of the Pinnacle Thrust exposed near Shoreham VT, but it is also helpful to understand that this thrust is the lowest (westernmost) thrust of the overall Champlain Thrust system in this area from Whitehall south to Fort Ann; lower components seen in Vermont (Shoreham Thrust; Main Champlain Thrust) die out southward with their displacements transferred upward along lateral ramps.

STOP 3: Frontal Fault of the Taconic Allochthon as a Post-Cleavage Thrust
**Warning - poison ivy occurs on this outcrop, and most abundantly in the road verge in front of it**
This outcrop (location shown on Fig. 4) is one of a very few places where the western fault contact of rocks of the main Taconic Allochthon is fully exposed; you will see three of them on this field trip. The roadcut exposes cleaved arenites and siltstones belonging to the early Cambrian/latest preCambrian Bomoseen unit of the Taconic Allochthon stratigraphy (a member of the Bull Fm. of Zen, 1967). A gently east-dipping fault contact on a meter or less thickness of dark phacoidal melange-fabric shale is present, although it may require a small amount of digging to complete the exposure. Below the shale, fine-grained medial Ordovician limestones (Glens Falls/Orwell) dipping gently eastward form the western part of the outcrop. The limestones map as a thin (c.10m.) slice above a much larger thickness of medial Ordovician shale and shale melange exposed to the west. This is the same relationship inferred in the Mettawee River section at Stop 1B, although there a different and younger Taconic rock unit locally maps adjacent to the fault. Note that the cleavage in the Bomoseen dips moderately to steeply east and must be truncated by the more gently east-dipping fault. This relationship is consistent with the oblique truncation of map-scale tight-isoclinal folds and their axial traces at the fault in this region, folds to which this cleavage is an axial surface fabric. Bomoseen arenites and siltstones are typically greenish chloritic colors, but in this exposure they are grey to dark grey, particularly near the fault contact with dark medial Ordovician shales affected by phacoidal melange-fabric cleavage. We think that fluid hydrocarbons migrating in the fault channel while it was active penetrated into the lowest Taconic rocks here and (now dehydrogenated carbon) are the cause of the colour difference. The low-angle attitude of the fault exposed here (and shown by local detailed mapping into the stream just to the north) suggests it is a thrust.
Near Troy, NY, in the gorge of the Poestenkill, sense of shear evidence suggests that the more steeply dipping fault exposed there is an east-side-downthrown normal fault which has cut and displaced the original Taconic thrust contact (Lim & Kidd, 2008). Similarly, a large roadcut on NY route 9G west of Hudson, NY exposes a sharp vertical fault contact that is also probably a normal fault. Such post-thrust normal faulting is also seen in calcite veins developed within the Ordovician melange terrain of the NY Capital District (Lim et al, 2005), and in the Mettawee River Fault seen at stops 1, 6B, and 7C on this trip (the trace of which projects well to the west of this locality; see Fig. 4). In this outcrop and more generally in the northern part, however, we think that an older thrust relationship is preserved at the western margin of the Taconic Allochthon, with the transport at this place being post-folding and cleavage development in the Taconic Allochthon strata.

STOP 4. Folded Thrusts of Taconic Allochthon Slates (Plude's Quarry).
Locality marked on Fig. 6. Plude's Quarry (Fig. 5) provides another of the few known exposures of Taconic Allochthon continental rise rocks lying upon the underlying melange/flysch sequence. Along Carlton Road the shales and minor thin silty wackes of the Taconic flysch are broken into a phacoidal melange fabric, seen in the lowest part of the western face of the outcrop. Going up the western outcrop face, a contact of this material overlain by slates with a planar slaty cleavage can be found. This is the Taconic Frontal Fault, again; but mapping by us (as shown in Fig. 6) clearly demonstrates that here, this fault, and the immediately underlying carbonate slice (not exposed in this outcrop), are folded on a map scale. Higher up the outcrop face, harder grayish-green slates contain thin fine-grained ribbon arenites and are consequently identified as belonging to the Poultney Formation of the Taconic Allochthon sequence. In the northern outcrop face of the old quarry, the regional slaty cleavage in those rocks passes down into the grey slates without arenites, that are above grey "shales" with phacoidal melange fabric at the western base of this face. The Poultney rocks above are clearly folded with the slaty cleavage as an axial surface fabric. A previous description of this outcrop (Bosworth and Kidd, 1985), suggested that there is an older, folded (thrust) fault, crossed by the slaty cleavage, that occurs between the Poultney slates above, and the grey slates below, and that this structure may be of the same generation as the early, folded, "Taconic Basal Thrust" of Rowley and Kidd (1981). The gray slates in this interpretation would be a transported constituent of the Middle Ordovician flysch ("Hortonville" lithology). Alternatively, but perhaps less likely, they might be an arenite-poor section of the Poultney unit and that no significant fault occurs at this contact internal to the rocks showing slaty cleavage in this outcrop. What do you think?
In either case, the local geological map shows that the fault expressed by melange fabric at the lower western side of the quarry is also folded, but this fault structure is not required to pre-date or overlap in time the cleavage affecting Taconic Allochthon rocks. The underlying limestone slice that is also affected by folding shows a solution cleavage, variably developed; while this is plausibly the product of the local strain that transported and folded the limestones, this does not mean that it is the same age as the slaty cleavage expressed in the Poultney slates in the quarry. This solution cleavage is more likely to have developed later in the overall Taconic imbrication/transport sequence, just before or during the transport of the shelf limestone, and after the thrust emplacement of the Taconic Allochthon above the limestone.

STOP 5. Thrust of Medial Ordovician Limestones Over Shaly Melange
**Warning - be cautious about where you step up to the rock face - in places, large blocks are visibly detached, and may be very unstable**
Locality indicated on Fig. 6. Carbonate exposures such as this one form lenticular belts bounded by medial Ordovician shales and shale melange (flysch) in a zone near this vicinity up to a few kilometers wide between the western edge of the Taconic Allochthon and the eastern edge of gently east-dipping, unfolded strata that rest with intact unconformable relationship on crystalline Grenville basement (such as those exposed along Rt 4 just east of Whitehall, and on Warner Hill visible a few kilometers due west of this locality). Where contacts are exposed, such as here, evidence for faulted lower contacts of limestone over shale are seen. In particular, the underside of the limestone is coated with fibrous vein-type slickensides, the lineation plunging ∼120º close to down-dip. Truncation of stratification in the limestone is seen locally. An abundance of veins in the limestone within about 50cm of the thrust surface suggests hydrofracturing and high fluid pressures during thrusting. A crude, lenticular cleavage in the underlying shales is deflected adjacent to the fault surface. A small horse of shale, isolated above the main thrust surface by a duplex mechanism, can be seen about half way along the exposure of the thrust (Fig. 7).
Towards the western limit of the overlying carbonate sheet a detached sliver of carbonate about a meter long lies in the cleaved shale just below the prominent fault surface. This may be a structurally detached piece or (less likely because of its shape) an olistolith. Similar lozenge-shaped pieces of medium-grained graywacke up to about ½ m long occur sparingly in the phacoidally-cleaved shale in the western part of the cut. Slickensided surfaces suggest that they too are products of structural disruption ("structural slicing"), although an olistolithic origin for these small objects cannot be discounted.
The shale and siltstone is locally bedded, with minor folds in part of the outcrop. Most is pervasively disrupted by the phacoidal cleavage whose microstructural character, with abundant evidence for shear offsets, is clearly related to faulting (Bosworth and Vollmer, 1981; Bosworth, 1982, 1984). This carbonate and other exposures like it are shown on the New York State geological map and by Fisher (1985) as giant olistolithic blocks. We find this interpretation of the outcrops to be unconvincing and prefer an interpretation, as shown on Fig. 6, where the carbonates and stratigraphically overlying shales form thin thrust sheets accreted beneath the Taconic Allochthon. Demonstrable tight folds and internal duplex faults within these carbonate sheets account for the many places where stratigraphic continuity is disrupted within them.

STOP 6A. Shoreham Thrust Exposure on Book Road
In the center portion of this roadcut exposure, grey-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 (Fig. 8) shows unequivocally that it is here the westernmost thrust in the Champlain System - in other words that the "Main Champlain", or Orwell Thrust farther north has disappeared (Hayman and Kidd, 2002a, b). To the south across the Vermont-New York border, near the Poultney River, the Shoreham Thrust also disappears, first because it is truncated by the Mettawee River normal Fault, although minor folding and incipient ramps can be detected in shelf rocks at its expected position in a outlier which occurs east of Whitehall. A major WSW-ENE cross-fault (the Book End Fault) passing north of this outcrop at stop 6A 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 Stops 6B and 7 below.

STOP 6B. Mettawee River Fault Near West Haven
ASK PERMISSION AT BOOK FARM BEFORE ENTERING - find Book Farm along Book Road 1.1 miles south of specified parking place
From parking place, enter field on east by walking between the two fences, opening the gate if necessary (and closing it behind you) - please do not climb over it or the fence. 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. 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 60 meters 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 [43.647590, -73.341250] . This contact is clearly a steep (60º) 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" shelf 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 (stop 1A) north through this locality at least as far north as the latitude of Shoreham, VT, and all along this length it truncates regionally the thrust stack of the Champlain Thrust System.
We suspect that it extends significantly farther north and south than this, but 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 cross-sections, and which demand at least 1000 meters of throw, and perhaps more than 1500 meters. This fault cuts out large, significant parts of the Champlain Thrust stack, and is one of the main reasons why previous attempts failed to trace the major thrust faults of the System southward. 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 Fig. 9 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).
Evidence to be shown in the outcrop of the next stop (Stop 7A) demonstrates that these apparent strike-slip offsets on the WSW-ENE faults are from actual strike-slip displacement; discussion of the age and origin of this youngest faulting type and orientation is included with Stop 7A.

Parking for stop 7

STOP 7A. Veins and Left-Lateral Strike-Slip Faulting in Limestone at West Haven
ASK PERMISSION AT THE FARMHOUSE TO VISIT THIS OUTCROP AND THE OUTCROP OF STOP 7B. DO NOT GO TO THE OUTCROP UNLESS YOU HAVE PERMISSION. NO HAMMERS ALLOWED.
At the barns, follow the track west 50 meters to the low lying limestone outcrop surfaces.
On the main part of the outcrop, abundant calcite veins of two generations can be seen. The older set strike about N-S and mostly dip moderately to very steeply east. Their displacement sense is not particularly significant for this field trip, and is undetermined in most cases, but east-over-west thrust sense has been seen, as well as some strike-slip component in places, of dextral sense. The significant vein set for our purpose is the one striking between E-W and WSW-ENE, which in all cases cut N-S veins, and can be shown in a number of examples to displace subvertical instances in a sinistral sense. From this outcrop, the only one around West Haven where we have found this evidence, we infer that the apparent displacements on the WSW-ENE faults mapped here resulted from real left-lateral strike-slip displacement, not entirely from steeply directed north-side-down slip cutting gently-moderately east-dipping strata and thrust faults.
We sampled by drilling short cores of some of the veins in this outcrop, and looked for suitable fluid inclusions in doubly polished sections to measure fluid homogenization temperatures, and freezing temperatures, to determine the salinity for the fluids forming the veins. Refer to the vein fluid inclusion discussion in the introduction for overview and details. To summarise the essential point: low salinity fluid inclusions having homogenization temperatures up to 205-215C were found in the younger WSW-ENE veins having left-lateral strike-slip offset. Fluids in the fractures at burial depth when the veins formed would have been above this temperature, the amount depending on the depth, which we do not know. However, by comparison with results from Ordovician (Taconic) and Devonian (Acadian) vein fluid inclusions in the Albany area, we think this temperature is high enough to be indicative of an Ordovician age. Devonian veins (thrust sense only) from the Acadian foreland fold and thrust belt there are distinguishably lower, most around 120-130C. We are not aware of any reason to suppose that Acadian fluids are likely to have been significantly hotter in western Vermont compared with the Albany area, and we note an absence of both strike-parallel normal faults and WSW-ENE cross-faults cutting the Acadian Helderberg fold-thrust belt. So we think the WSW-ENE faults crossing the Taconic foreland thrust belt in this area are strike-slip faults (at least in their later/last slip episodes), and Ordovician (Taconic) in age, and are the youngest of the sets of structures of the Taconic event.
Whether they are Ordovician, or younger, what is their tectonic origin? Thrust faults are easy to understand in the context of the Taconic arc-continent convergence (Bird and Dewey, 1970), and strike-parallel normal faults in that general context similarly have modern exemplars, both of pre-thrusting origin from slab-pull on the down-going lithosphere (Schoonmaker et al, 2005), and extension late in and/or after convergence from slab detachment (Davies and Blanckenburg, 1995) or regional stress removal by large- scale plate boundary reorganisation. Strike-parallel strike-slip faults also are no problem to accommodate by strain partitioning of plate motion (McCaffrey, 1996). Cross-strike strike-slip faults of small displacement, on the other hand, are usually secondary linking structures (tear faults) for thrust (or perhaps with oblique slip for normal fault systems), and these commonly also serve as fault segment delimiters for earthquake slip zones. We think this may be the explanation for these WSW-ENE faults in the Taconic foreland but there is a problem; there are no known contemporary thrust or normal faults of the right relative age, nor any of a compatible slip direction! The WSW-ENE faults cut the Mettawee River normal fault, which itself cuts the youngest thrust in the Champlain Thrust stack. Even if we are wrong about the age constraint inferred from the fluid inclusion minimum temperature, the Acadian thrust slip direction for this sector of the northern Acadian Appalachians is also not towards WSW. Perhaps there could be small displacement blind thrust segments in the flat-lying platform sediments of the Taconic foreland, with WSW-directed slip, that would be the very last activity of the Taconic orogen, and to which these strike-slip faults form tear fault linking structures?
There are other features of the geology apparently associated with the West Haven cross-fault zone that we have noted: a) the abrupt change in the medial Ordovician strata across it from non-calcareous shales occurring to the south, to significant thicknesses of calcareous shales (Cumberland Head and Stony Point Fms) to the north. In addition, b) changes in thickness of the early Ordovician Ward Siltstone appear to be associated with this zone, and perhaps also with other faults of this orientation near and south of Whitehall. The implication is that these faults have developed using older fractures in the basement, and that these older structures also may have influenced subsidence patterns and deposition of sediments of the shelf both before and during the approach of the Taconic thrust stack.
Also, c) the map complexity near West Haven implies that the Shoreham Thrust interacted with an already fault-displaced irregular floor, from normal faulting on these oblique structures before the thrust stack interacted with it; it is admitted that inadequate outcrop and want of 3-D information prevent any clear demonstration of this. However, the same structural history inferred here can be observed if you visit outcrop adjacent to the Split Rock Point Fault [44.267592; -73.337305] near Westport NY where normal displacement on the NE-striking fault is expressed by the juxtaposition of Grenville gneisses and medial Ordovician calcareous shales; the dipping shales bent up by the normal-sense shear strain are cut by calcite veins showing clear evidence of strike-slip displacements, of both dextral and sinistral sense.

STOP 7B. Flat-Lying Thrust Fault of Taconic Slates Over Calcareous Shales and Limestones South of Coggman Creek
The small waterfall cliff, and the stream bed above, expose grey slates with a well-developed, moderately east-dipping cleavage. At the base of the small cliff, these slates rest on a practically flat thrust fault carrying them over dark calcareous shales. These calcareous shales pass down into limestones which form the rest of the outcrop in the stream bed and on the valley walls adjacent. The limestones are medial Ordovician Glens Falls, Orwell, and perhaps Middlebury equivalents, similar to those seen at stop 3, 5, and 7A (but are, we think - see map of Fig. 8 - not part of the same thrust slice as at stop 7A). The transported slates possess a well-developed planar slaty cleavage and so in a structural sense belong with the Taconic Allochthon. Their lithology, the same seen at stop 6B, monotonous, slightly silty, pale to medium grey slate, with thin silty laminations seen in some places, has been given the name Hortonville Slate by Zen (1967) in the area of the main Taconic Allochthon to the east. It is probably a medial Ordovician-age deposit, a facies of the orogenic flysch clastics accompanying the Taconic convergence, but significantly farther-transported than the melange, shales and turbidites of the Hudson Valley. Along structural strike, to the NNE near Benson, these rocks adjoin the Sunset Lake slice of the Taconic Allochthon, and we regard them as an enlarged addition/extension of the Sunset Lake slice. If this interpretation is valid, then the thrust exposed here is essentially another exposure of the western marginal fault of the Taconic Allochthon.
If you are using this field guide independently, you need to return to the Main Road at West Haven, go west along the road 0.3 mile, and ask permission at the house on the north side of the road to visit the outcrop of Stop 7C. If permission is obtained, the outcrop is about 150 meters north of the house.
If permission was previously obtained, return from Stop 7B to the edge of the woods and walk west along the edge for about 450 meters, passing limestone outcrops in the woods extending from near stop 7B until reaching the point where the ridge crest starts descending towards the west and enters the woods. Down the slope to the north at this point, the last limestone outcrop occurs above an exposure of dark melange-fabric shale just below it, showing that the thrust at the base of this limestone sheet runs here. From the point on the ridge, go approximately southwest contouring around the ridge end, and descend slightly into another (usually dry) streambed running roughly west.

STOP 7C. Metawee River Fault and Quartzites of the Pinnacle Thrust West of West Haven
See Fig 8 for location and geological surroundings. Exposures of dark partly melange-fabric shale occur intermittently in the bed and low bank of the stream to just before a sharp increase in the gradient over a rock ledge exposing well-lithified quartz arenites/quartzites. Below these, down stream, dolostones also outcrop. These arenites are part of the Potsdam/Ticonderoga unit of the Cambro-Ordovician shelf sequence, lying above the (here unexposed) Pinnacle Thrust, and must be separated from the melangy shales seen a few feet upstream by the principal strand of the Mettawee River Fault.
These quartzites also outcrop on the crest of the hill across Coggman Creek to the north, where they form an isolated erosional outlier of the Pinnacle Thrust, which overlies mainly dolostones of the Shoreham (locally termed Shaw Mountain) Thrust. We think the quartzite exposures near Root Pond a few kilometers to the northeast also form isolated remnants of the Pinnacle Thrust sheet, and that isolation by truncation of the Mettawee River normal Fault is a key to understanding their relationships, as it is for the similar occurrences of such quartzites in the same structural setting all the way from here at Stop 7C to Whitehall NY, where they can be shown to continue in the hanging wall of the Comstock thrust, as seen at Stop 2.

B4: NEW INSIGHTS INTO GLACIAL LAKES VERMONT AND ALBANY
John A. Rayburn
Dept. of Geology
SUNY New Paltz
New Paltz, NY 12443
rayburnj@newpaltz.edu
David J. DeSimone
DeSimone Geoscience Investigations
Petersburg, NY 12138
hawkeye272david@yahoo.com
Amy B. Frappier
Department of Geosciences
Skidmore College
Saratoga Springs, NY 12866
afrappie@skidmore.edu
The first comprehensive study of the proglacial lakes in the Hudson and Champlain Valleys was published by J.B. Woodworth in 1905 (NYSM Bulletin 84). He noted that “Lake Albany doubtless began on the south in the waters standing in front of the retreating ice sheet prior to the opening of the Mohawk outlet of the great glacial lakes to the west. As soon as the ice retreated in the valley to a position north of Albany and the drainage of Lake Iroquois came into the Hudson Valley Lake Albany properly came into existence”. While glacial lacustrine sediments in the Champlain Valley were originally attributed by Upham (1889), Baldwin (1894), Peet (1904), and others to be a separate Glacial Lake Champlain, Woodworth notes that its confines were somewhat larger and proposed the name of Glacial Lake Vermont, and that to the south it was certainly confluent with Glacial Lake Albany as he defined. “The outlet of theses ice-dammed waters at this early stage of confluence across the present divide of the Hudson and Champlain basins is a matter which concerns the interpretation of Lake Albany on the south and is considered in that connection. Lake Vermont may be said properly to have come into existence when in consequence of a local lowering of waters south of Fort Edward a discharge began across a barrier into the Hudson Valley to the south.”
While Woodworth (1905) only described one Lake Albany level, he noted under Lake Vermont that there were the Quaker Springs, Coveville, and Fort Edward levels. Of these, Woodworth stated that “The question of the outlets of Lake Vermont … has not been completely exploited as yet by field work”. The Fort Edward level outlet, better defined by Chapman (1932) as the Fort Ann outlet (see also DeSimone et al., 2008 NEFOP Field Guide Stop “Chapman’s Potholes”), is the only one that truly meets his definition of Lake Vermont. Although Chapman recognized several separate Fort Ann levels, only the lowest of these may actually have been confined north of the Champlain-Hudson divide. Woodworth and later Fairchild (1917), Stoller (1922), and Chapman (1937), suggested that the Coveville level threshold is represented by a channel overhanging the Hudson River near Coveville, NY. There are two significant problems with this hypothesis however: 1) the bedrock that this channel is formed in is generally too soft to have withstood the estimated steady-state discharge of the pro-glacial lake and 2) more recent mapping demonstrates that Coveville level strandlines exist south of this location.

B4: NEW INSIGHTS INTO GLACIAL LAKES VERMONT AND ALBAN

STOP 1. Glacial Lake Warrensburg
Along the upper Hudson between Luzerne and Warrensburg, the bluffs expose fine sand terraces, till, and bedrock. A few locations preserve elevated deltas and varved clays, in horizontal orientation and tilted large-scale slump blocks. Glacial Lake Warrensburg was named and mapped by Miller (1911, 1914), who inferred a single long, narrow ice-dammed glacial lake that integrated lacustrine features from Corinth in the south (Stoller’s 1916 moraine-dammed Lake Corinth) through Warrensburg and into the upper Hudson and Schroon valleys north of North Hudson (Lake Pottersville, Miller 1914), including the Paradox Lake valley east of Ticonderoga. Figure 5 shows Miller’s 1925 general map of the paleo-lake extent (dashed line indicates a Hudson-Champlain ice margin). Lake Warrrensburg must be older than the terrace at the Luzerne ice margin (Hanson, 1977). We infer that Lake Warrensburg flooded the Hudson-Schroon Valley north of Luzerne for a brief ∼few-century period, beginning with the ice blockage around ∼15.3-15.1ka.
Miller (1925) reports a graduated level of lacustrine deltas and sand plains, from 208 masl at Corinth to 300 m at Deadwater Pond about 110km to the north. He gives a slope estimate for the entire lake of 4.5 feet per mile, or about about 1 m/1.16 km. Miller locates the lake outlet between the ice and the Palmertown range highland pass “cut by a stream of large volume” near South Corinth. When the ice front retreated past Corinth into the Lake George valley, the upper Hudson outlet eroded below 194m, shifting flow to the east at Corinth through the Hudson valley into Lake Albany, or down the Kayaderosserass to the Milton delta. Lake Warrensburg presumably drained by ∼14.7ka to 14.8ka, during retreat of the ice blockade in the Bølling-Alleröd.
In 2011-12, 2015-16, and 2017-18, undergraduate teams in Amy Frappier’s Honors Paleoclimatology Practicum course have investigated GL Warrensburg varves using a series of short Ridge-style cores from a horizontal varve deposit overlain by a flat massive fine sandy unit several meters thick. The overlying unit is very silty, coarsening-upwards, and contains a kettle bog. Recent field observations of the varve deposits in this area by R.H. Lindemann, Frappier and students, and John Rayburn, indicate that varves can be found in protected locations in the upper terraces, along the banks of the Hudson, and into the river channel. The varve deposits form characteristically steep slopes, and are prone to slumping where undercut by stream incision. Our discontinuous varve cores from our upper terrace site reveal large ∼10cm-scale and small ∼1-2cm-scale varve sequences (Brill et al., 2012; Nolan et al., 2018). At least one low-sedimentation rate period a few decades in length occurring between ∼1-decade-long intervals when annual deposition was higher (Khan et al, 2016). The data are insufficient to determine whether this general sequence relates to glacial melting directly, or to hydro-climate fluctuations at regional or watershed scale that may integrate glacial melting and postglacial sedimentary redistribution processes. Longer, more continuous varve records from this and other sites in the lake basin may help to resolve this and other questions related to deglaciation of the upper Hudson Valley.
Rock flour, silt, and fine sand deposited in Lake Warrensburg was most likely sourced, at least in part, from valley glacier outwash in the adjacent Adirondack highlands, rather than solely from the Laurentide ice sheet. This suggests that GL Warrensburg varve sequences are likely to reflect high-altitude local conditions in the upper Hudson/eastern Adirondack region, particularly in the southern and western Hudson parts of the paleo-lake basin. Whether these varve sequences can be successfully cross-dated to the NAVC remains to be seen, but recovering longer sequences are essential (Nolan et al, 2018). Varves in the tilted slump block (Fig. 6 above) provide a promising target for future work, and additional field scouting further north may reveal additional sites where varves are preserved, including sites with the potential for local pro-glacial conditions that may be sensitive to larger-scale
Laurentide ice dynamics. The differing sensitivities of the southern/western and northern/eastern reaches of Glacial Lake Warrensburg present opportunities for future investigations of this interesting interval of deglaciation in the eastern Adirondacks.
En route to Stop 1, we pass by a typical elevated sand terrace in Warrensburg at ∼720 ft (Fig. 7). Stop 1 offers roadside access to a Lake Warrensburg delta perched on the Western flank of the Hudson Valley. Topset beds with rounded river cobbles evident in the upper section are mostly obscured by grain flows down the exposed section. The delta clearly exposes tilted foreset beds (Fig. 8). Their southerly dip indicates flow from up-valley, rather than from a lateral tributary. Known varve exposures found below the sand unit outcrop just 1km downstream on the eastern side of the Hudson in the lee of a river bend, but are not accessible by car. Split varve cores will be available for inspection along with the exposed delta, as we discuss the emerging picture of how Lake Warrensburg fits into the regional deglaciation history.

STOP 2. The Cove at Coveville
Woodworth suggested that the fully prograded Batten Kill delta into the Hudson Valley may have dammed lower lake levels, forcing discharge through the uplands onto the western uplands (Fig. 2). The flow would therefore have been forced between Bacon Hill and Northumberland and across a divide into the Fish Creek channel at Grangerville (Fig 9). From there it would follow the channel to Victory Mills, but then continue southeastwards and re-enter the Hudson Valley at the rock ledge at The Cove. His justification for this morphology is that “The cove at Coveville in its relation to this hanging valley shows clearly that a large stream at one time flowed southward over the wall of the gorge at this place into the main gorge of the Hudson river, and was arrested after a slight amount of cutting had been accomplished”. Chapman (1932) sites Woodworth (1905), and further states that “At Coveville, where the southern end of this channel overhangs the Hudson River by more than a hundred feet, is a rock ledge which acted as the controlling threshold for the waters during the Coveville stage of Lake Vermont”.
There are several circumstances which strongly argue against the Coveville threshold being located in any part of this channel, however. The primary observation is that no part of this channel is at the Coveville elevation level. The highest modern divide elevation just off Grange Hall Road correlates to the highest Fort Ann level (Fig 3), and The Cove incision is only a little lower than the lowest Fort Ann level. Secondly, the bedrock type along the channel path is not sufficiently strong to have held up to steady state Coveville phase discharges. While some rock at the northern end of the channel is allochthonous (Starks Knob, for example), the primary rock type through the rest of the area is a weak shale that demonstrates significant incision in equilibrium with modern base level and discharges. Since the Coveville stage remained constant and steady for hundreds of years until the catastrophic discharge associated with the breakout of Lake Iroquois at Covey Hill, the threshold for this lake must be elsewhere. The formation of Woodworth’s channel was more likely caused quickly by lake levels falling below the Coveville level with abandonment occurring as the discharge more easily cut through the Batten Kill delta and lacustrine deposits in the main Hudson channel. The Cove itself may just be an abandoned meander bend formed by Fort Ann level discharge through the valley.

STOP 3. Batten Kill Delta
One of the most well preserved deltaic landforms can be seen where the Batten Kill joins the Hudson River across the valley from Schuylerville. An excellent road traverse is to follow Rte 29 east from Schuylerville. Once across the Hudson River, the road rises moderately and climbs from bottomset beds through foreset beds to a large, delta topset flat. Minor drainage ditch exposures reveal the sediment of the delta that generally fines upward and reflects the pro-gradational nature of the landform. In 2008, the FOP trip (DeSimone et al., 2008) visited an active sand and gravel pit in the topsets and upper foresets but that pit is now long closed and well vegetated. A walking traverse along Windy Hill Road would give students an opportunity to follow the delta fore-slope and examine the sediment in drainage ditches.
An altimeter determined elevation of the topsets places them at 96m (315ft) correlating to the Coveville lake level (Fig. 3). Woodworth first recognized this beautiful delta does not fall on a Lake Albany water plane. Subsequent detailed surficial mapping by DeSimone confirmed Woodworth’s placement of the delta on a lower lake level. The important factor here is that Woodworth first had the foresight to recognize a major delta was correlative to a lake level lower than Albany. A short distance farther east in Greenwich there is a remnant of a higher deltaic surface that records deposition in a 115m (380ft) lake, a level that would place it on an Albany II water plane. This is the highest deltaic surface recognized from mapping along the Batten Kill and it does not fall on the highest level of Lake Albany as correctly noted by Woodworth. Either a higher delta has been removed by erosion or Lake Albany lowered before ice retreat opened the Batten Kill valley.

STOP 4. The Saratoga National Historic Park (Lunch Stop)
The short drive from Route 4, the old Colonial river road, up along the Kroma Kill to the park visitors center basically followed a transect perpendicular to the river. We went from flood plain alluvium up through a thick deep water glaciolacustrine clay section deposited in Glacial Lake Albany at its different levels. The fine grained lake silt and clay is capped along the flat river bluffs by a thin layer of lake sand deposited in the shallower waters of the last true lake in this section of the valley, Lake Coveville. We ascended to the gently rolling terrain at the top of the road where we parked and are in an area of generally thin till with rock outcrop.
Your vista from the terrace and lawn of the visitors center embraces a wonderful perspective across the Hudson Valley that is the consequence of both pre-glacial bedrock erosion that established the general geomorphology and Pleistocene glacial erosion and deposition that left the valley infilled with sediment (Fig. 10). Subsequent Late Pleistocene and Holocene erosion carved into the sediment and bedrock and left the landscape we see today.
The underlying bedrock topography and the more obvious glacial geomorphology were responsible for the terrain over which the 2 Battles of Saratoga were fought in September and October 1777. A single image is provided here to summarize the impacts of the geologic history on the battles (DeSimone 2016) and both bedrock and surficial geologic maps of the 4 quadrangles encompassing the SNHP will be available to view and discuss during lunch (DeSimone 2015a, 2015b). The maps are digitally published and those interested in the data layers can find them from the National Park Service and/or see the notations at the bottom of both maps.
The generation of the two maps is the result of a long, eight-year desire that finally came about in 2014- 2015. DeSimone represented the Vermont Geological Survey at a two-day meeting held at UMass in 2007 to discuss the status of mapping in regional Park Service properties. Bruce Heise and Tim Connors represented the National Park Service (NPS). When the SNHP came up for discussion, DeSimone volunteered that most of the park had already been mapped by students of Bob LaFleur during the late 1970s. Eric Hanson (1977, 1980) mapped in the Mechanicville, Schaghticoke and Quaker Springs quadrangles, Jack Dahl (1978) mapped in the Mechanicville and Schaghticoke quadrangles, and DeSimone mapped the Schuylerville quadrangle (1977, 1985). DeSimone suggested it should be an easy task to finish a map of the park. Fast forward to late 2014; Bruce and Tim contacted DeSimone to see if he was interested in doing the project. However, the surprise was they wanted a map of the 4 quadrangles, not just the park and there were little funds to do much, if any field work (Fig. 11). In early 2015, they indicated they also wanted a bedrock map of the 4 quadrangles. The result is the 2 maps you have to view. Both maps largely represent compilations with editing on DeSimone’s part. Meshing the surficial maps together and field checking the contacts consumed nearly all of the allotted field time. Map units were re-designated in an updated fashion with some compromises and simplifications for clarity. Both printed maps were intended to be wall hangers for the SNHP staff and the text on the maps was intended to be educational to the non-geologist.
The bedrock map is work by Kidd, Plesch, and Vollmer (1995), Rickard (unknown date), Landing et al. (2003), Fisher et al. (1970), and a beautiful hand colored map of the Schuylerville 15-minute quadrangle attributed by Rickard to Reudemann (unknown date). My role was to compile and edit the data, merge and extrapolate contacts and convert map units in older terminology to those of Kidd et al (1995). DeSimone chose to add pre-glacial bedrock channels to this map using data from Dineen and Hanson (1983) and Bruehl (1969) with extensions from both data sources.
The lunch view has the gently rolling to sloping till terrain of the visitors center giving way to the clay flats that were farm fields in 1777. The park has tried to maintain the vegetative cover as close to what it was during 1777. The Hudson Valley Lake Albany shoreline is both in the near distance and in the middle distance across the valley where the land slopes up abruptly into the Taconic highlands with the highest summit in view being Willard Mountain, a local ski area. That break in slope approximates the Lake Albany shoreline and follows Route 40 where we’ll drive after lunch. The track of Route 40 also largely coincides with the trace of the Taconic Frontal Thrust or TFT as seen on the bedrock map. On a clear day, the view extends into the far distance where the summit ridge of the Green Mountains is visible.

STOP 5. Hoosic Delta
The Hoosic River and the Batten Kill are two large Hudson River tributaries that drain the Taconic Mountains along the east side of the valley. From the Batten Kill, we proceeded south along Rte 40 following the shoreline of Glacial Lake Albany and also largely following the trace of the Taconic Frontal Thrust or Emmons’s Line. A very distinct lowland to upland break in slope is present.
We continued south to the junction of Rte 40 and Rte 67W and turned west to head back toward the Hudson River. We drove across several of the four readily distinguishable deltas of the Hoosic River deposited into different glacial lake levels. Each succeeding delta breached the older, higher delta. The expression of the deltas as landforms is harder to discern. However, this trip stop lets you see some deltaic sediment stratigraphy. There are two pits across the road from each other and at the time of this writing, we cannot know for certain which pit we can get access to on the day of the trip. Both pits are fairly small and if they are working in them, there will be no permission granted.
The sediment exposed is largely medium to fine sand with predominantly horizontal thin beds, some ripple laminations and no evident ice contact deformation. The sand beds are overlain by variably thick pebble gravel beds with medium to coarse sand. The pebble gravels are cross bedded with some cut and fill structures. The pebble gravels truncate the underlying sand beds. At the smaller pit on the north side of Rte 67, the gravel beds are thin and some of the topsoil has been removed from the exposed face. The original thickness of the gravel facies cannot be determined. At the larger pit on the south side of Rte 67, the gravel facies appears to be intact in the high wall of the pit beneath the more mature vegetation. Here, the pebble gravel facies is approximately 2m thick.
The elevation of the flat top of the landform in the south side pit is 107m (350ft) using the topographic map spot elevation. The pit has been excavated into an isolated terrace within the 107m (350ft) Hoosic delta (Fig. 3). The north side pit was excavated into a lower 98-99m (325ft) Hoosic delta. The largest preserved fragment of the Hoosic deltas has an extensive topset plain between the 113-118m (370-390ft) elevations. A lower terrace of the Hoosic River that may be deltaic lies at 82m (270ft).

STOP 6. Speigletown Kame Moraine and Strandline
This large pit with multiple open excavations exists in the Speigletown kame moraine-kame terrace (Fig. 12). The entrance to the operation is from NY Rte 40. Route 40 follows the old Mahican Indian trail later used as a Colonial coach road. The road generally follows the eastern shoreline of Glacial Lake Albany. Along this route from the Batten Kill south to the Hoosic River and continuing south to the pit, your drive follows the glacial lake shoreline. The shore face is sometimes veneered with beach sand overlying finer grained sediments. The Fane pit began decades ago at the road level. The landform is a kame terrace here that caps bedrock sometimes visible in deeper excavations. Across the road from the pit entrance, shale bedrock is exposed. The depositional environment was that of a classic kame terrace deposited between the Hudson-Champlain ice and the valley wall. The deeper parts of the pit reveal sedimentation from subglacial meltwaters that represent an esker distributing sediment into a subaqueous fan. There are remnants of ice contact sand and gravel deposits that can be seen north of Rte 142, west of Rte 40 and continuing down to toward the Hudson River in Pleasantdale. However, these are not shown on the 1963 surficial map of Schock. Mr. Fane reports hunting along trails leading down to the river and finding a large area of gravel buried by the silt-clay. In contrast, the kame terrace likely received sediment from a meltwater and meteoric water stream following the ice margin. Indeed, some of the sediment may be from the Hoosic River to the north as ice still blocked the river outlet.
Beach sand and even pebble gravel caps the sequence and will be our first vantage point in the pit along the power line. There is a pocket of silt-clay below the beach berm that may represent a small pond behind a sand spit. This is one of the finest sections of Lake Albany beach preserved in a still rural setting. Look across to the road bordering the south face of the pit and you can see the beach profile. The berm here is exceptionally high and this appears to be a natural feature, not built up at all for the power line construction.
Wander to the deepest excavations in the pit. Excavations have reached bedrock at least locally capped with a veneer of till. Ground water seeps out along the base of the gravel and sand as further downward infiltration of ground water is inhibited by the comparatively impermeable till and bedrock. The deeper portions of the pit reveal subaqueous fan sand and gravel. Previously, we could observe arched anti-form bedding from a classic esker. Finer grained lacustrine sediments have an on-lap relationship to the ice contact sediments indicating the quieter water facies were deposited after retreat of the ice removed the source of the proximal gravel and sand facies. One can think of the subglacial meltwater source as a fire hose discharging water and sediment from beneath the glacier into the bottom of Lake Albany. The point source of sediment deposits coarse grained sediment in one locale while adjacent areas have finer grained sediment. The fire hose analogy works if you think of the hose spraying sediment first in one direction and then in another.
We’ll work our way up section toward the east where the sediment is interpreted to have been cascading from the kame terrace environment into the deeper lake waters. Slope instabilities and debris flows must have been common. At any given moment, the pit face may show soft sediment deformation, load deformation and evidence of subaqueous sediment flows. Nearshore sediments in the higher parts of the pit truncate and overlie an ice contact facies composed of interbedded gravel and sand with typical ice contact deformation.

STOP 7. Boulder Moraine and Lag Terrace in Ballston Creek Channel at Shenantaha Creek Park
This last stop takes us to the Shenantaha Park on the Round Lake quadrangle that was most recently mapped by Hanson (1977b). A possible alternative to the Fish Creek/Coveville channel as a threshold and discharge channel for the Coveville level is a drainage divide at East Line between the Morning Kill, Drummond Creek, and Ballston Creek. Besides being at an elevation consistent with the Coveville level it would provide a threshold and drainage network that would effectively control the Coveville level if the main Hudson Valley were dammed between Speigletown and Halfmoon (Fig. 2). This network together with the Ballston Lake channel is all part of the Mohawk distributary channel network and may have been earlier formed in equilibrium with both northward discharge from the IroMohawk as well as southward discharge from glacial Lake Warrensburg. Early investigation of this system was published by Stoller (1911) who inferred that: “A flood of waters once swept northward through Ballston channel, dividing in the vicinity of East Line into three currents which pursued the several courses described above. The time in glacial history when this took place was subsequent to the general disappearance of the ice and also subsequent to the stage of maximum development of Lake Albany” (Fig 4). Stoller (1916) recognized the southern end of a lake that he named Lake Corinth that the Hudson River drained via Kayderosseras Creek southward into Lake Albany forming the Milton delta (Fig. 13). However at 125m (410ft) it is significantly above the Albany water plane (Fig 3) and may actually represent a separate higher-elevation ice marginal lake.
The much larger extent of Lake Corinth was recognized by Miller (1925) and renamed Lake Warrensburg (Stop 1), and that its existence must have been caused by ice damming the modern Hudson channel east of Corinth. It is possible that when the ice was farther south than depicted in Figure 13 the western region of the Hudson Valley including Milton and the distributary channels was confluent with a higher elevation glacial lake in the Mohawk Valley. Northward recession of the ice margin would subsequently drain this lake to Albany level allowing Mohawk Valley discharge directly into Lake Albany through the distributary channels. Eventual ice margin retreat to Glens Falls would open the modern upper Hudson channel to Lake Albany completely draining Lake Warrensburg.
For the distributary channels to form the threshold and drainage route for the Coveville level around the Speigletown moraine (Fig. 2) they would require flow southward through the Morning Kill and/or Drummond Creek channels, and/or westward drainage through the Ballston Creek channel to Ballston Lake to the Mohawk Valley. This would in turn necessitate that the IroMohawk flow directly eastward to the Hudson Valley through its modern channel at Cohoes rather than the distributary channels. It would also necessitate a bedrock exposure of significant strength to withhold steady state discharge from Lake Coveville. These conditions lead to a series of problems. First, like Woodworth’s proposed Coveville outlet, the local exposed bedrock is a weak shale/greywacke mélange that demonstrates easy incision to base level even under modern discharge conditions. Second, like Stoller, we observe no southward flow indicators in the channels and only very strong eastward flow indications in the Ballston Creek channel. This includes an imbricated boulder terrace (also recognized by Stoller, 1911) near the head of the Ballston Creek channel that could only have resulted from strong eastward flow towards Round Lake (Fig. 14).
On the positive side, however we have evidence that the IroMohawk must have been flowing directly eastward through the modern channel at least at the very end of the Coveville Phase. As stated in the introduction, the base level for Cohoes Falls which clearly migrated back through the shale-rich mélange from the Hudson Valley is at Hudson River level, and that discharges only on the scale of the IroMohawk could have caused this knick point migration. This now under-fit falls is apparently frozen in time with the reversal of drainage of Lake Iroquois at Lake Iroquois into Lake Vermont causing it to drop from Coveville level to Cobblestone Hill and Fort Ann levels (Rayburn et al., 2005; Franzi et al., 2007; DeSimone et al., 2008; Franzi et al., 2016). Interestingly there are a number or terraces carved into the shale of the Ballston Creek channel ranging in elevation from Coveville level at the head of the channel Shenantaha Creek Park near East Line down to through the Fort Ann levels and below to the east, suggesting that this channel was at baseline equilibrium with Lake Albany and Lake Vermont (Fig 15). Also the imbricated boulder moraine surface along the creek at the edge of the park correlates to Cobblestone Hill level, suggesting that it may have been the Lake Iroquois breakout flood that caused the lagged surface in the boulder moraine and that flow was eastward into the Hudson Valley at this location.

C1: THE CHEEVER AND MINEVILLE IRON OXIDE-APATITE (IOA) DEPOSITS
Marian V. Lupulescu
Research and Collections
New York State Museum
Albany NY
marian.lupulescu@nysed.gov
Jeffrey R. Chiarenzelli
Dept. of Geology
St. Lawrence University
Canton NY
jchiarenzelli@stlawu.edu
David G. Bailey
Geosciences Dept.
Hamilton College
Clinton NY
dbailey@hamilton.edu
Sean P. Regan
University of Alaska
Fairbanks, AK
sregan5@alaska.edu
Jared W. Singer
Earth and Environmental Sciences
RPI
Troy, NY
singej2@rpi.edu
The iron deposits of New York State were an important source for the U.S. iron industry during the nineteenth and twentieth centuries. The main component of the ore was magnetite; it was accompanied in some deposits by hematite, with gangue minerals such as feldspar, quartz, amphibole, and fluorapatite. Most of the medium-to highgrade magnetite ore was mined underground.
Iron mining in New York State has a long history. There are two main regions where iron mining was developed, the Hudson Highlands in the south, and the Adirondack Mountains in the north. After more than a century of iron production, Smock (1889) completed the first report on the iron ores of New York State and the first classification based on a “geologico-geographical arrangement.” His classification included almost all the iron occurrences known at that time and all the major ore types (magnetic iron, hematite, limonite, and carbonates).
The iron deposits from the Adirondack Mountains (Figure 1) can be divided into two major types: (a) iron oxide – apatite (IOA) deposits, and b) titanium iron oxide deposits. The most significant IOA deposits are those associated with the A-type Lyon Mountain Granite (gneiss), the Benson Mines magnetite-sillimanite-garnet ores, and the Jayville and Clifton magnetite ± vonsenite deposits, all of which are in St. Lawrence County. The most significant Ti-Fe oxide deposits are those found at Tahawus in association with rocks of the Anorthosite-Mangerite- Charnockite-Granite (AMCG) suite (nelsonites associated with gabbroic anorthosite), and smaller occurrences at Split Rock (“cumberlandite”), Craig Harbor (hornblendite), Tunnel Mountain (gabbro), all in Essex County, and the Port Leyden nelsonites in Lewis County.
The IOA deposits of the eastern Adirondacks display many similarities with Kiruna-type iron deposits. The IOA deposits found in Scandinavia, the Americas, Russia, Iran, Australia, and China, vary in age from Archean to Pleistocene (Williams et al. 2005), are all associated with volcanic to subvolcanic rocks, and were intensely metasomatized (Harlov 2016 and references therein). Their origin is strongly debated by modern researchers, but three hypotheses prevail: (a) primary igneous origin altered by secondary hydrothermal replacement (Sillitoe and Borrows 2002); (b) igneous derivation by liquid immiscibility (Nyström and Henriquez 1994; Naslund et al. 2002; Chen et al. 2010; Tornos et al. 2016); (c) flotation of silica melt-derived magnetite microlites (Knipping et al. 2015).
During this trip we will visit two IOA deposits, the Cheever Mine located north of Port Henry, and the # 21 mine at Mineville (Essex County) where we will examine the mineral composition of the ore, the contact with the host rock, the petrographic features of the host, and discuss the geochemistry, geochronology, and origin of the deposits. Both were significant iron producers but are also widely known for the association of REE-bearing apatite, and the exquisiteness of the magnetite crystals that can be seen today in many museums around the world. The mining history of the Mineville-Port Henry district is well presented by Farrell (1996) and summarized by Lupulescu and Pyle (2008) in their report for the NYSGA Annual Meeting at Lake George.
New data and comments for this field trip improve and bring new knowledge to the prior field trips offered by NYSGA in 2008 (Mineville) and 2015 (Cheever). Some parts of the text are comparable with those from Lupulescu and Pyle (2008) and Lupulescu et al. (2015) field guides.

Meeting Place - Port Henry Boat Launch Site

STOP 1. Cheever Mine.
The Cheever Mine is located above Lake Champlain north of the village of Port Henry. The ore strikes approximately north-south, dipping toward the west. The old mine workings are preserved as several trenches that follow the ore body for more than 0.5 km. The local bedrock geology consists of rocks of the Grenville Supergroup (marble and pelitic gneiss), coronitic gabbro, amphibolite, and pink to green variably altered rocks of the Lyon Mountain granite.
The metasedimentary rocks are represented by marbles containing folded and disrupted lenses of calcsilicates and gneisses. The coronitic metagabbro is spatially associated with the leucocratic rocks containing the magnetite-apatite ore. A similar situation can be seen at the Barton Hill mines at Mineville. The nature of the contact (tectonic vs. intrusive) is not obvious, being obscured by recent alluvium and cover in both cases. The coronitic metagabbro from Cheever mine is sheared toward the contact with the leucocratic rocks which contain quartz, albite, pyroxene, ± microcline. Lenses of mafic rocks containing pyroxene, amphibole, fluorapatite, and magnetite can be found in the tailings (the underground works are not accessible).
The coronitic metagabbro occurs in massive or layered form. The massive form contains plagioclase “clouded” with spinels, pyroxenes, garnet and annite. Both ortho- and clinopyroxene grains are present, and most display tiny elongated exsolution lamellae of other pyroxenes and ilmenite as the result of the sub-solidus re-equilibration during slow cooling. Rare grains of ilmenite and pyrrhotite pepper the rock. The coronitic metagabbro contains some layered facies with the same composition as the massive variety.
The footwall rock which hosts the ore is a granitic rock composed predominantly of quartz and albite. Minor phases are grains of relict clinopyroxene (largely replaced by chlorite), zircon (some grains are partially metamict), and magnetite with ilmenite lamellae. The hanging wall rock contains more relict pyroxenes than the footwall, and the magnetite/ilmenite grains are rounded and, in places, associated with pyroxene replaced by chlorite. In other areas, the rock displays a gneissic texture with alternating “bands” of quartz and albite, and pyroxene, amphibole, annite, magnetite, ilmenite, and rare grains of apatite and pyrrhotite (altered to goethite). Exsolution textures within the pyroxenes are common. This rock grades into a microcline, albite, quartz, pyroxene, amphibole, and ilmenite-rich gneiss. The ore-hosting rocks are undeformed (this observation is supported by field and microscopic study) and contain pyroxene, and in places rounded, droplet-looking, grains of magnetite and or ilmenite.
The rocks have high concentrations of Na2O (3.48 wt.% in the metagabbro, to 6.31 wt.% in the footwall leucogneiss); K2O varies from 0.36 % in the footwall rock to 3.63 % in the highest stratigraphic unit of the hanging wall leucogneiss. All rocks display normative “hypersthene”. The rocks are enriched in REEs and have steep profiles, like those from Mineville.
The magnetite-fluorapatite ore at the Cheever Mine exists as dikes / sills that have sharp contacts with the host rock. The ore contains abundant magnetite, fluorapatite, and augitic pyroxene. Other mineral phases present include ilmenite, titanite rimming magnetite, zircon, monazite-Ce, stillwellite-Ce, allanite-Ce, and thorite. The amphibole tremolite is very rare, and mostly is the result of the low temperature interaction of clinopyroxene, quartz, and later fluids. A relatively F-rich (1. 19 to 1.58 wt.%) amphibole with blue pleochroism under plane polarized light seems to be the last igneous mineral in the succession. Rare spinel phases were exsolved along the {111} crystallographic planes of the large magnetite grains. Lupulescu et al. (2017) interpreted the fluorapatite textures as result of the intense metasomatism; the products of the metasomatic reactions are monazite-(Ce) and allanite-(Ce) (Figure 2). The fluorapatite grains display high REE concentrations, especially heavy REEs and Y (Lupulescu et al. 2017). A detailed overview of the geology of the Cheever mine is in the report for the NYSGA Annual Meeting 2016 (Lupulescu et al. 2015).
The field relations and results of samples of the Lyon Mountain Granite collected for U-Pb zircon geochronology at the Cheever mine are shown in Figure 4. At these localities zircon xenocrysts were absent in several of the samples and ranged in age up to 1242 Ma where present. Lyon Mountain Granite samples range in age from 1040-1066 Ma in concert with ages obtained from numerous samples of the unit throughout the Highlands, suggesting multiple intrusive events along a magmatic conduit or fault. Note that the Lyon Mountain Granite is intruded by a gabbroic rock. The ore sample here yielded a sparse population of zircons, interpreted as igneous in origin, whose age was 1033.6±2.9 Ma.

STOP 2. Mineville
The host rocks at the Mineville mines have both mafic and felsic compositions and display igneous features. The metasedimentary rocks from the area include marbles, calc-silicates, and gneisses that structurally overly the igneous sequence. One of the first researchers at the Mineville mines, Kemp (1908), described the host rocks as “augite syenites and related types” with granite and diorite compositions being the “related types.” Alling (1925) suggested an igneous and sedimentary origin for the various protoliths that were later metamorphosed during the Grenville orogenic cycle. Buddington (1939) considered the composition of the host rocks as a petrographic hybrid between granite and metasedimentary or metagabbroic rocks. McKeown and Klemic (1956), using the maps of the Republic Steel geologists, described a structural sequence of metamorphic rocks, starting with a basal metagabbro (Kemp’s mafic syenite) followed by the magnetite ore and granite gneiss from the “Old Bed”, passing into a diorite, then to gabbroic rocks and the associated magnetite ore of the “Harmony Bed”.
Our observations show that the first unit above the magnetite-apatite ore in the hanging wall is a pinkish rock that compositionally grades from granite to syenite. Above this unit, there is a leucocratic granite, which is in turn overlain by amphibolite (Figure 3A). The foot wall is a gabbro intruded by two sheets of leucocratic granite and a pegmatite (Figure 3B).
The main minerals that were identified at Mineville in association with the ore are: magnetite, hematite (martite), fluorapatite, stillwellite-(Ce), allanite-(Ce), monazite-(Ce), edenite, actinolite, ferro-actinolite, scapolite, titanite, and zircon and dolomite, smoky quartz, calcite in late veins, and ilmenite and titanian hematite as tiny disseminations in the host rock. Micron-size mineral phases of secondary thorite, allanite-(Ce), parisite, and monazite-(Ce) in some apatite crystals were recognized under the polarizing microscope and by SEM – EDS in thin/polished sections. Bastnaesite-(Ce) and lanthanite-(Ce) previously reported by McKeown and Klemic (1956), and Blake (1858), respectively, were found as very rare and tiny crystals based on the electron microprobe data (Figure 4A). A detailed mineralogy of the Mineville deposit was presented in the report for the NYSGA Annual Meeting 2008 (Lupulescu and Pyle 2008).
Fluorapatite from the Mineville deposit contains high concentrations of LREEs and especially HREEs. According to Molycorp estimation in 1980, the 5 million cubic meter tailings (Figure 5) contain “8-9 million kilograms of Y2O3 with an average grade of 0.12 wt.% Y2O3 and 0.6 wt.% REO” (Mariano & Mariano 2012).

Mineville tailings - Figure 5
Location of Figure 5 in the text.

C2: GEOLOGY OF THE NORTHERN TACONIC ALLOCHTHON - STRAIN VARIATION IN THRUST SHEETS, BRITTLE FAULTS, AND POSTRIFT DIKE EMPLACEMENT
Jean Crespi
Jennifer Cooper Boemmels
Jessica Robinson
Geosciences
University of Connecticut
Storrs, CT 06269
This field trip has three purposes: (1) to present a synthesis of structural and strain data for the Taconic slate belt, which lies in the northern part of the Giddings Brook thrust sheet in the Taconic allochthon, (2) to present an analysis of fault-slip data for post-cleavage faults in the region, and (3) to present the results of work on mafic dikes
in the Taconic lobe of the New England–Québec igneous province. The stops have been selected to illustrate the along-strike variation in structure and strain, to show representative post-cleavage faults, and to include dikes in a variety of orientations. The data presented at the stops cover a region that extends for about 60 km along strike in the Giddings Brook thrust sheet, from Hubbardton, Vermont, in the north to Salem, New York, in the south (Figs. 1 and 2).
Emplacement of the Taconic allochthon took place during the Early to Late Ordovician Taconic orogeny. In the northeastern United States, the Taconic orogeny has traditionally been interpreted as resulting from the collision of a west-facing volcanic arc with the eastern margin of Laurentia (Bird and Dewey, 1970; Rowley and Kidd, 1981; Stanley and Ratcliffe, 1985). In this interpretation, emplacement of the Taconic allochthon occurred in a proforeland setting. New geochronological data, however, indicate that the orogeny involved several phases (Karabinos et al., 1998, 2017; Macdonald et al., 2014, 2017). During the first phase, a west-facing volcanic arc built on peri-Gondwanan basement collided with a fragment of a hyperextended eastern Laurentian margin. During the second phase, this composite ribbon terrane collided with Laurentia proper. Slab break-off and a reversal of subduction polarity resulted in the establishment of an east-facing volcanic arc. During the third and final phase, the Taconic allochthon was emplaced. In this revised interpretation, emplacement of the Taconic allochthon occurred in a retroarc setting.
The eastern North American margin developed after Paleozoic contractional tectonics culminated in the formation of Pangea. Early Mesozoic rifting of Pangea was accompanied by the development of a system of rift basins of Late Triassic and Late Triassic–Early Jurassic age and by massive outpourings of basalt of the ∼201 Ma central Atlantic magmatic province (CAMP). The extension direction during rifting was generally northwest–southeast (Withjack et al., 2012). In the northeastern United States, CAMP dikes trend northeast–southwest, consistent with the extension direction for rifting. Postrift magmatic bodies include the plutons and sheet intrusions of the Early Cretaceous New England–Québec (NEQ) igneous province (McHone, 1984; McHone and Butler, 1984). The origin of these intrusives has long been debated. They may be related to passage of the Great Meteor hotspot (Morgan, 1972; Sleep, 1990), edge-driven convection (Matton and Jébrak, 2009), opening of the North Atlantic Ocean (McHone and Butler, 1984), delamination of over-thickened crust, or some combination of these mechanisms.

C2: GEOLOGY OF THE NORTHERN TACONIC ALLOCHTHON - STRAIN VARIATION IN THRUST SHEETS, BRITTLE FAULTS, AND POSTRIFT DIKE EMPLACEMENT

Parking for Stop 1

STOP 1. High Strain Subdomain (Poultney/Deep Kill Formation)
This exposure of the Lower to Middle Ordovician Poultney/Deep Kill Formation was created in the mid 1990s when County Road 18A was constructed to improve road traffic safety. The strata lie on the overturned limb of a regional-scale syncline and in the immediate footwall of the out-of-sequence Middle Granville thrust. Our analysis of the orientations of slaty cleavage and the stretching lineation and of syntectonic fibers places the exposure in the high strain subdomain on the eastern edge of the central domain.
Slaty cleavage strikes north-northeast and dips moderately east-southeast, and the stretching lineation plunges approximately down the dip of slaty cleavage. Syntectonic fibers associated with slaty cleavage development are curved in XZ sections consistent with top-to-west-northwest non-coaxial flow. The fibers yield stretch values for the X direction of about 2.2–2.3. They also record a small amount of extension in the Y direction. The central domain, in contrast, is characterized by north-northwest strikes for slaty cleavage, a moderately raking stretching lineation, and lower stretch values for the X direction.
Many mesoscale folds are present in the exposure (Fig. 3). They are unusual because the hinge lines plunge moderately to steeply southeast. Hinge lines of folds in the Taconic slate belt are typically subhorizontal to gently plunging, and, in the central domain, the plunge direction is south-southeast. The bedding-cleavage intersection lineation also plunges to the southeast although not as steeply as the fold hinge lines. This difference is likely because the mesoscale folds are mostly in the middle of the exposure and the bedding-cleavage intersection lineation was measured at the ends of the exposure where bedding is planar. The folds in the exposure are also asymmetrical such that they have an S shape when viewed down plunge.
The fold geometry can be explained by initial development of parasitic folds and subsequent rotation of the hinge lines of the parasitic folds toward the stretching lineation. The S-shaped fold asymmetry is consistent with an origin as parasitic folds on the overturned limb of a west-vergent, regional-scale fold. In addition, the orientation of the hinge lines is consistent with an initial gentle south-southeast plunge and counterclockwise rotation toward the downdip stretching lineation during top-to-west-northwest non-coaxial flow. This interpretation implies that this area was a long-lived region of high strain, the folds starting to rotate during folding and continuing to rotate during slaty cleavage development. The Middle Granville thrust may be a late-stage expression of this high strain zone.
Mesoscale, post-cleavage faults in the strata typically contain laminated veins with well-developed steps. Some of the faults have formed along folded surfaces, a good example lying on the northeast side of County Road 18A near the middle of the exposure (Fig. 4). These faults are normal faults on the long limbs of the folds and reverse faults on the short limbs of the folds. Slip was subparallel to the hinge lines. In the example in Figure 4, the outer part of the fold has moved up relative to the inner part of the fold, and the striae show a systematic variation in orientation around the fold such that they are closer to exactly parallel to the hinge line in the hinge.

STOP 2a. Northern Domain (Mettawee Slate/Middle Granville Slate)
Cedar Point quarry exposes the hinge zone of a north-northeast-trending, regional-scale syncline. The folded strata belong to the Lower Cambrian Mettawee Slate/Middle Granville Slate, a purple and green slate that contains reduction spots useful for strain analysis. The quarry lies in the northern domain where slaty cleavage strikes north-northeast and dips gently east-southeast and where the stretching lineation plunges approximately down the dip of slaty cleavage. Here, slaty cleavage is overprinted by a crenulation cleavage, which becomes the dominant cleavage to the east of the approximate longitude of Lake Bomoseen and Lake St. Catherine. Although this crenulation cleavage is locally present in the Taconic slate belt, it does not appear to affect the orientation of slaty cleavage. In the northern domain, where the crenulation cleavage is more common than in the central and southern domains, the mean orientation of slaty cleavage for sites without a crenulation cleavage is essentially the same as the mean orientation of slaty cleavage for sites with a weakly to moderately developed crenulation cleavage.
At Stop 2a, the concave side of a folded bedding surface is visible through the graffiti. The intersection of the axial planar slaty cleavage with bedding indicates the hinge line of the syncline is gently plunging, which is typical of folds in the Taconic slate belt. At Stop 2b, the convex side of a folded bedding surface and the profile plane of the syncline are visible. The tight to isoclinal nature of the syncline is typical of folds in the Taconic slate belt.

STOP 2b. Northern Domain (Mettawee Slate/Middle Granville Slate)
Reduction spots are common in the slabs in the slate pile at Stop 2b. Two types of reduction spots are present: those with irregular shapes and those that are elliptical. The latter have been measured and reported on by Wood (1974), Hoak (1992), and Goldstein et al. (1995). The reduction spots here and elsewhere in the Taconic slate belt plot in the flattening field of a Flinn plot, and Goldstein et al. (1995) interpreted their horizontal distribution on a
Flinn plot as indicating the slates underwent a minimum of 55% volume loss during slaty cleavage development. Our strain data for the northern domain from strain fringes around pyrite framboids in black slates are not consistent with the reduction spot data. The strain fringes do not show fiber growth in the Y direction, and so we infer plane strain deformation for slaty cleavage development. In addition, the strain fringe data yield internally consistent data when combined with the orientation of slaty cleavage if constant volume is assumed. The 2.28 mean stretch value for the X direction obtained from strain fringes for the northern domain gives a strain ratio for the XZ plane of 5.2 for constant volume and a strain ratio for the XZ plane of 10.4 for 50% volume loss. These values can be plotted on an RXZ versus θ’ graph for simultaneous simple shear and volume change (Fossen and Tikoff, 1993) using the 28° mean dip of slaty cleavage in the northern domain and an appropriate dip for the Giddings Brook thrust sheet. We use a thrust sheet dip of 5° east, giving a θ’ value of 23°. A value of 5.2 for RXZ results in the northern domain plotting close to the curve for simple shear, which is by definition a constant volume deformation. A value of 10.4 for RXZ results in the northern domain plotting in the volume gain field, which is at odds with the 50% volume loss assumption used to obtain RXZ.

STOP 3a. Bomoseen Dike at Glen Lake
Glen Lake Trail (visit vtstateparks.com/assets/pdf/bomoseentrailmap.pdf for a copy of the trail map) is well marked with blue trail blazes, and the walk to Stop 3a is relatively easy with little elevation change, the occasional fallen tree, and some irregular terrain where the trail passes over exposed tree roots. About 7 minutes (∼0.25 miles) from the trailhead, the trail splits shortly after passing a foundation. Take the right fork, which passes on the east side of a cellar hole. The left fork follows the water’s edge and is not the main trail. About 10 minutes (∼0.35 miles) from the trailhead, the trail climbs up a small ridge. A hemlock tree with a girth of about 2 m and a blue trail blaze stands at the top of the ridge. The south-southwest side of the Bomoseen dike is exposed in the vicinity of this tree. Continue to follow the trail as it descends into a small valley to the north of the ridge. The valley is before the trail bends downhill to the left. Walk off trail just upslope from the valley floor to where a recently fallen tree has exposed the north-northeast side of the Bomoseen dike. Return to the trail and follow it downhill to the right. Rubble next to the trail of the dike and wall rock marks the dike-wall rock contact on the north-northeast side of the dike. Follow the trail up- hill back to the valley floor and walk off trail to the east-southeast toward Moscow Rd. Several exposures of the dike and of the wall rock to the south-southwest of the dike are present in this area. Return to the trail and follow it back to the trailhead.
The exposure of the Bomoseen dike at Stop 3a consists of several small and discontinuous outcrops of the dike between the eastern shore of Glen Lake and Moscow Road. Outcrops of the wall rock (West Castleton Formation on the Bedrock Geologic Map of Vermont (Ratcliffe et al., 2011) and Poultney and Hatch Hill formations on Rowley’s map of the Lake Bomoseen area (Rowley, 1983, Plate 1)) are also present.
The Bomoseen dike was named by McHone and McHone (1993) for an exposure of the dike on VT-30 along the eastern shore of Lake Bomoseen. The dike appears on Dale’s map of the Taconic slate belt (Fig. 5) (Dale, 1899) and is reproduced on Fowler’s map of the Castleton area (Fowler, 1950, Plate II) and on the Bedrock Geologic Map of Vermont (Ratcliffe et al., 2011). It trends west-northwest–east-southeast and is relatively wide and long compared to other NEQ dikes in the Taconic lobe.
McHone and McHone (1993) reported a width of 12 m for the Bomoseen dike at the VT-30 exposure (the entire width of the dike is no longer preserved). Our observations of the dike at the Glen Lake exposure give a comparable width. The location of the dike-wall rock contact on the north-northeast side of the dike can be located fairly closely on the trail by the change from dike rubble to slate rubble. A sample we collected from the southern side of the exposure of the dike closest to Moscow Road contains the dike-wall rock contact on the south-southwest side of the dike.
The Bomoseen dike is shown as extending continuously for nearly 5 km on Dale’s map of the Taconic slate belt (Dale, 1899). In addition to the Glen Lake and VT-30 locations, the Bomoseen dike is well exposed along North Road.
Of interest at Stop 3a is the well-developed pencil cleavage in the wall rock adjacent to the dike-wall rock contact on the north-northeast side of the dike. This pencil cleavage has been formed by the intersection of the well-developed slaty cleavage with closely spaced dike-parallel fractures.

STOP 3b. Scotch Hill Syncline
The exposure of the Scotch Hill syncline at Stop 3b (Fig. 6) is one of the best-known exposures in the Taconic slate belt. It lies in the northern domain. Slaty cleavage forms a divergent cleavage fan with the overall orientation of slaty cleavage similar to the mean orientation of slaty cleavage in the northern domain. Because sampling is prohibited at the exposure, the orientation of the stretching lineation was not determined, and the characteristics of strain fringes around subspherical core objects were not assessed.
The folded strata, which belong to the Poultney/Deep Kill Formation, contain two sets of veins, one that predates folding and one that formed at the same time as folding. A subhorizontal, north-northeast-trending fold axis for the Scotch Hill syncline is shown by both the pole to the best-fit plane to the poles to bedding and by the bedding-cleavage intersection lineation at the exposure (Figs. 7A and 8A). The poles to slaty cleavage lie on the best-fit plane to the poles to bedding (Fig. 7A). The prefolding veins are similarly oriented throughout the fold, and their poles do not plot on the best-fit plane to the poles to bedding (Fig. 7B). The synfolding veins lie in a range of orientations, and their poles plot on the best-fit plane to the poles to bedding (Fig. 7B).
The prefolding veins are long, thin, and planar (Fig. 8B–D). They are commonly boudined (Fig. 8B), and they are crosscut by the synfolding veins (Fig. 8C, D). The synfolding veins have a less uniform morphology but are commonly short, thick, and sigmoidal (Fig. 8C–F). They are either boudined or folded depending on the angle between the vein and slaty cleavage. Synfolding veins that are especially thick, however, do not appear deformed. Although the prefolding veins appear approximately parallel to slaty cleavage on the profile plane of the fold, their non-coaxiality with the fold can be seen on west-northwest-facing bedding surfaces on the subvertical limb of the fold. The trace of cleavage on the bedding surfaces is subhorizontal whereas the trace of the prefolding veins plunges to the south. The coaxiality of the synfolding veins with the fold can also be seen on these surfaces. Both the trace of cleavage and the trace of the synfolding veins are subhorizontal (Fig. 8A).
The asymmetry of the synfolding veins is consistent with flexural folding. On the gently dipping limb of the fold, the synfolding veins are commonly irregularly shaped, but a large vein has a classic sigmoidal shape, the S shape when viewed to the north-northeast indicating thrust sense of shear (Fig. 8E). On the subvertical limb of the fold, the synfolding veins are commonly sigmoidal, the Z shape when viewed to the north-northeast indicating westside- up sense of shear (Fig. 8F). These opposite senses of shear are as predicted for flexural folding with a pinned fold hinge. The mystery of the prefolding veins having similar orientations on the two limbs of the fold despite being prefolding is solved when shearing related to flexural folding is applied to their orientation. The change in orientation of the prefolding veins from limb rotation was essentially offset by the change in orientation from flexural folding- related shearing within the bedding layers.

STOP 4. Central Domain (Indian River–Pawlet/Austin Glen Formations).
The bedrock geology along Tanner Hill Road was mapped by Jacobi (1977) and is described in Rowley et al. (1979) and Landing (2002). The road transects the Tanner Hill syncline, and, although many of the exposures are small, the stratigraphic succession is readily observed and contacts between units are easily located. Exposures on the west limb of the syncline are somewhat inaccessible because of a new maple sugaring operation.
The small quarry immediately east of the parking area for this stop is in the Middle Ordovician Indian River Formation, a red slate interpreted as the oldest synorogenic unit in the Giddings Brook thrust sheet (Landing, 2002, 2012). Macdonald et al. (2017) dated a sample from a half-meter thick silicified layer of green slate in this quarry using LA-ICPMS and CA-IDTIMS. Zircon grains interpreted to be of primary volcanic origin and in the youngest population obtained by LA-ICPMS analysis yielded a CA-IDTIMS age of 466.06 ± 0.21 Ma.
Stop 4 lies in the central domain between the Taconic Frontal thrust and the Lee Road fault. The syncline extends for the length of the Taconic slate belt, the axial trace trending north-northeast–south-southwest in the northern and southern domains and north–south in the central domain. Folds between the Lee Road fault and Middle Granville thrust in the central domain have axial traces that trend north-northwest–south-southeast. The anomalous orientation of the fold axial traces is interpreted to be a result of the central domain having undergone inclined transpression (Crespi et al., 2010). Consistent with such a three-dimensional deformation, the stretching lineation is not downdip, and the strain fringes record extension in the Y direction. Moreover, in addition to the top-to-west-northwest sense of shear given by fibers in strain fringes in the XZ plane throughout the Taconic slate belt, fibers in strain fringes in the central domain are locally curved in the XY plane, the sense of curvature indicating sinistral shear.
Mesoscale, post cleavage normal faults are present in the exposures along Tanner Hill Road, most commonly in the Indian River Formation. Some of the faults contain laminated quartz and chlorite veins, but most lack fault-zone vein material, most likely because the faults are weathered. Faults with thick panels of laminated quartz and chlorite veins and well-developed steps can be found in out-of-place blocks in the quarries on either side of Tanner Hill Road. Although some faults are parallel to slaty cleavage, most strike parallel to but dip more steeply than cleavage (Fig. 9). The angle between the faults and slaty cleavage is on average about 10°–15°.

STOP 5. Southern Domain (Hatch Hill Formation) and Bardin Dike.
This exposure of the Lower Cambrian to Lower Ordovician Hatch Hill Formation consists of interbedded black slate and quartz arenites. Subvertical bedding is visible in the small roadcut on the north side of the large bend in Bardin Road, and subhorizontal bedding is visible in the small roadcut on the south side of the large bend in Bardin Road. Subvertical to steeply east-dipping and subhorizontal bedding is visible in the outcrops in the woods immediately west of Bardin Road. The exposure lies on the upright limb of a regional-scale syncline on Shaw’s map of portions of the Granville, Hartford, Hebron, and Pawlet townships (Shaw, 1984) and on the overturned limb of a regional- scale syncline on the Bedrock Geologic Map of Vermont (Ratcliffe et al., 2011). The bedding-cleavage relations at the exposure indicate a north-northeast-trending anticlinal hinge between the area with subvertical to steeply east-dipping bedding and the area with subhorizontal bedding. This anticline is parasitic to the regional-scale fold.
The characteristics of Stop 5 are typical of the southern domain. Slaty cleavage strikes north-northeast and dips about 30° east. The stretching lineation is about down the dip of slaty cleavage. Syntectonic fibers associated with slaty cleavage development are generally straight in XZ sections. Where they exhibit subtle curvature, the sense of curvature is consistent with top-to-west-northwest non-coaxial flow. The fibers yield a mean stretch value for the X direction of 1.6 and a mean stretch value for the Y direction of 1.1. The main differences between the southern and northern domains are that the fibers in the southern domain are not as curved as they are in the northern domain and that the stretch values for the X direction in the southern domain are not as high as they are in the northern domain. These differences are consistent with the southern domain lying at a higher structural level, i.e., farther from the base of the Giddings Brook thrust sheet, than the northern domain.
An igneous dike, which we have named the Bardin dike, lies in the hinge zone of the anticline indicated by the bedding-cleavage relations at the exposure (Fig. 10). This dike together with many of the other dikes/dike segments in the southern part of the Taconic lobe was mapped in 1895 by Florence Bascom, a pioneer in the advancement of women in geology. Bascom’s work contributed to Dale’s efforts to map the Taconic slate belt, and the dikes that she mapped appear on Dale’s map of the Taconic slate belt (Dale, 1899). The Bardin dike is about 2 m wide. Its trend is subparallel to the axial traces of the regional-scale folds in the southern domain and to the axial trace of the local anticline that it has intruded. Like the Bomoseen dike, a well-developed pencil cleavage is present in the wall rock adjacent to the dike. It can be observed on both sides of the dike and has been formed by the intersection of the well-developed slaty cleavage with closely spaced dike-parallel fractures. Bascom’s mapping shows the Bardin dike as one of a series of discontinuous dike segments that extends for about 12 km from Belcher, New York, to South Granville, New York (Fig. 10).

STOP 6. Central Domain/Southern Domain Boundary (Mount Merino Formation) and Granville Dike.
This exposure of the Middle to Upper Ordovician Mount Merino Formation changes from interbedded black slate and chert in the northwestern part of the exposure to mainly chert in the southeastern part of the exposure. Slaty cleavage, which is well developed in the northwestern part of the exposure, dips moderately east, and bedding is subvertical. The exposure lies on the upright limb of a regional-scale anticline on Kidd’s map of the Lee Road area (Kidd, 1976)/Fisher’s map of the Glens Falls–Whitehall region (Fisher, 1985) and on the Bedrock Geologic Map of Vermont (Ratcliffe et al., 2011). The bedding-cleavage relations here, however, indicate the exposure is on an overturned limb. This implies the strata are on the limb of a fold that is parasitic to the regional-scale fold.Stop 6 lies near the boundary between the central and southern domains where the trend of the fold axial traces changes from north-northwest–south-southeast in the central domain to north-northeast–south-southwest in the southern domain. Slaty cleavage strikes slightly east of north and dips moderately east. The stretching lineation is about down the dip of slaty cleavage. We have not analyzed syntectonic fibers at this location.
A narrow igneous dike lies near the southeastern end of the exposure. This dike does not appear on Dale’s map of the Taconic slate belt (Dale, 1899). It is shown, however, on Kidd’s map of the Lee Road area (Kidd, 1976) and is reproduced on Fisher’s map of the Glens Falls–Whitehall region (Fisher, 1985). We have named this dike the Granville dike. It is about 25 cm wide and has been mapped for a distance of about 0.4 km.
Like the Bardin dike, the Granville dike is one of a number of northeast-trending dikes/dike segments in the southern part of the Taconic lobe. These dikes have the same general trend in the southern domain, where the structural grain trends north-northeast–south-southwest, and in the central domain, where the structural grain trends north-northwest–south-southeast. Thus, while the dikes generally follow the structural grain in the southern domain, they cut across it in the central domain. The northeast-trending Granville dike cuts across the structural grain, but this is not immediately apparent at Stop 6 because both the dike and bedding are subvertical and because the orientation of bedding is difficult to see in three dimensions in the chert layers next to the dike. About 2 m to the northwest of the dike, bedding strikes about 035–040 (Fig. 11). Because the dike is slightly magnetic and somewhat curved, we used the orientation of dike-parallel fractures in the vicinity of the dike as a proxy for the orientation of the dike. These fractures strike about 060–075 (Fig. 11), which is oblique to bedding and similar to the 055–060 trend of the dike shown on Kidd’s map of the Lee Road area (Kidd, 1976). The oblique relation between the Granville dike and the structural grain is readily seen on Kidd’s map of the Lee Road area (Kidd, 1976)/Fisher’s map of the Glens Falls–Whitehall region (Fisher, 1985) where the dike is shown cutting across the contact between the Indian River and Mount Merino formations (Fig. 11).
Several brittle faults are present in this exposure. They are either oblique to both bedding and slaty cleavage or parallel to bedding.

C3: A TRAVERSE THROUGH THE SUTURE ZONE BETWEEN LAURENTIA AND THE
MORETOWN TERRANE IN NORTHWESTERN MASSACHUSETTS
Paul Karabinos
Williams College
Williamstown
MA 01267,
pkarabin@williams.edu
Francis A. Macdonald
Department of Earth Science
Webb Hall
UC Santa Barbara
Santa Barbara, CA 93106,
francism@geol.ucsb.edu
James L. Crowley
Boise State University
Boise, ID 83725,
jimcrowley@boisestate.edu
ABSTRACT
U-Pb dates on magmatic and detrital zircon from samples in the hinterland of the Taconic orogen place new constraints on the timing and plate tectonic geometry of terrane accretion and magmatic arc activity. The Moretown terrane, a Gondwanan-derived exotic block, extends from the Rowe Schist-Moretown Formation contact in the west to the Bronson Hill arc in the east. Arc-related plutonic and volcanic rocks formed above an east-dipping subduction zone under the western leading edge of the Moretown terrane from approximately 500 to 470 Ma, until it collided with hyperextended distal fragments of Laurentia, represented by the Rowe Schist. Magmatic arc rocks formed during this interval are primarily located in the Shelburne Falls arc, although some rocks that formed above the early east-dipping subduction zone are located in the northern part of the Bronson Hill arc to the east. Metasedimentary rocks in the Shelburne Falls arc contain detrital zircon derived from mixing of Gondwanan, Laurentian, and arc sources, suggesting that the Moretown terrane was proximal to Laurentia by 475 Ma. Explosive eruptions at 466 to 464 Ma preserved in the Barnard Volcanic Member of the Missisquoi Formation in Vermont and as ash beds in the Indian River Formation in the Taconic allochthons may record slab-break-off of subducted lithosphere following collision of the Moretown terrane with distal Laurentian crustal fragments. Between 466 and 455 Ma a reversal in subduction polarity lead to a west-dipping subduction zone under Laurentia and the newly accreted Moretown terrane. Magmatic arc rocks in the Bronson Hill arc formed above this west-dipping subduction zone along the eastern trailing edge of the Moretown terrane at approximately 455 to 440 Ma. The western boundary of Ganderia in New England is east of the Bronson Hill arc, buried beneath Silurian and Devonian rocks deformed during the Acadian orogeny.
Before plate tectonic theory was applied to the Appalachians (e.g. Bird and Dewey, 1970; St. Julien and Hubert, 1975), and even before absolute radiometric ages were available, three major orogenies were recognized based on paleontological constraints on deformed rocks below angular unconformities: the Ordovician Taconic, the Devonian Acadian, and the Pennsylvanian to Permian Alleghenian orogenies. During the 1970s and 1980s, it was common for geologists working in western New England to ascribe deformation to either the Taconic or Acadian orogenies. This simplified view of early Paleozoic tectonism held that the Taconic orogeny resulted from the collision of Laurentia with a ‘Taconic arc’ (for example, Rowley and Kidd, 1981; Stanley and Ratcliffe, 1985) and that the Acadian orogeny marked the collision of a Gondwanan-derived microcontinent called Avalonia (for example, Rast and Skehan, 1993). The Alleghenian orogeny occurred when the main Gondwanan continent arrived in the Late Paleozoic, and it was believed to have only affected rocks in southeastern New England (for example, Quinn and Moore Jr, 1968).
Detailed work in New England and in the Canadian Appalachians, where oceanic tracts, arcs, and Gondwanan-derived microcontinents are better preserved, has resulted in a much more intricate and complex history of arc and microcontinent accretion, reversals in subduction polarity, and intermittent back-arc rifting (for example, Van Staal et al., 1998; Zagorevski et al., 2008). More recent tectonic syntheses have expanded the number of orogenies to the Taconic, Salinic, Acadian, Neoacadian, and Alleghenian, and some of these involve multiple phases (for example, Van Staal and Barr, 2012). Gondwanan-derived microcontinents now include the Moretown terrane (Macdonald et al., 2014), Ganderia (Van Staal et al., 1998), Avalonia (Rast and Skehan, 1993), and Meguma (Schenk, 1997, White and Barr, 2010). Furthermore, we now know that Alleghenian deformation affected rocks as far west as the Bronson Hill arc and Connecticut Valley Trough in western Massachusetts (for example, Robinson et al., 1992). The expanding geochronological database in the northern Appalachians suggests that the Laurentian margin was continuously active from approximately 475 to 270 Ma.
Hibbard et al. (2006) mapped the suture between peri-Laurentian and peri-Gondwanan terranes in the New England Appalachians along the western margin of the Bronson Hill arc (Fig. 1). This interpretation required that the Shelburne Falls arc (Fig. 1) formed on a Laurentian-derived microcontinent as proposed for the Dashwoods block in Newfoundland (Waldron and van Staal, 2001), and that the Bronson Hill arc formed on the western leading edge of Ganderia (Hibbard et al., 2006). Based on detrital zircon data from the Moretown Formation in Massachusetts and Vermont, Macdonald et al. (2014) and Karabinos et al. (2017) argued that the suture between Laurentia and Gondwanan-derived terranes is located approximately 50 km further west along the contact between the Rowe Schist and the Moretown Formation, which is characterized by a high concentration of mafic and ultramafic lenses, interpreted as remnants of subducted oceanic lithosphere. Because the evidence for a Gondwanan source for the metasedimentary rocks came from the Moretown Formation, we called this newly identified Gondwanan-derived block the Moretown terrane.
The relocation of the suture has some important tectonic implications for the deep crustal structure of the New England Appalachians. 1) It requires that the Shelburne Falls arc formed on Gondwanan-derived crust rather than Laurentian crust. 2) Instead of the Bronson Hill arc having formed above an east-dipping subduction zone on the western leading edge of Ganderia, we suggested that it formed above a west-dipping subduction zone along the eastern trailing edge of the Moretown terrane after a reversal in subduction polarity. 3) If the Moretown terrane is distinct from Ganderia, the suture between these two Gondwanan-derived terranes must be located somewhere under the Silurian-Devonian rocks of the Central Maine basin (Fig. 1).
The focus of this trip is the evidence for a plate suture between Laurentia and a Gondwanan-derived microcontinent, the Moretown terrane, proposed by Macdonald et al. (2014) and Karabinos et al. (2017). In Massachusetts, this suture coincides with the boundary between the Rowe Schist and the Moretown Formation. There is nothing cryptic about this suture; it is characterized by a dramatic difference in sediment source as revealed by detrital zircon populations, and is characterized by a high concentration of mafic and ultramafic lenses, which we interpret as remnants of oceanic crust and mantle from an east-dipping subduction zone. It also coincides with a sharp 10 to 15 km decrease in depth to MOHO from the Laurentian margin to the Moretown terrane (Li et al., 2018). We will also discuss evidence suggesting that much of the 475 Ma arc magmatism in the Shelburne Falls arc occurred in close proximity to Laurentia even though the arc was built on the Gondwanan-derived Moretown terrane.

C3: A TRAVERSE THROUGH THE SUTURE ZONE BETWEEN LAURENTIA AND THE
MORETOWN TERRANE IN NORTHWESTERN MASSACHUSETTS

STOP 1. View to west and Hoosac Schist.
The gift shop has been closed for a while, but the platform next to it still provides a nice view of Mount Greylock (south), the Taconic Crest (distant ridge), and the southern termination of the Green Mountain massif (north). Across the road from the gift shop is an outcrop of Neoproterozoic to Cambrian Hoosac Schist. It is very similar lithologically to the Greylock Schist on Greylock and the Nassau Formation on the Taconic Crest. They were all deposited as silty mudstones on the continental slope and rise of the Laurentian margin following rifting of Rodinia. Metamorphic grade decreases from our location on Hoosac Summit to the west. Here the Hoosac Schist is a quartz, albite-porphyroblast, muscovite, biotite, garnet schist with lots of quartz veins. The Nassau Formation on the Taconic crest in the distance is a fine-grained phyllite; the Greylock Schist is at an intermediate grade of metamorphism. Small relict garnet crystals can be found in albite porphyroblasts on the eastern most ridges of Mount Greylock. The southern end of the Green Mountain massif is a south-plunging anticline cored by Mesoproterozoic basement unconformably overlain by Neoproterozoic to Cambrian Dalton Formation and Cambrian Cheshire Quartzite. The east-west valley that connects Williamstown and North Adams between the Green Mountains and Mount Greylock must contain a lateral thrust ramp with many kilometers of structural relief because the Taconic sequence rocks of Mount Greylock must have formerly covered the Green Mountains before erosion stripped them away.

Outcrop of Rowe Schist on right.

STOP 2. Rowe Schist.
There is abundant outcrop of Rowe Schist all along Whitcomb Hill Road. This is the best one and it contains a prominent contact between graphitic and non-graphitic schist near the bend in the road. Both contain the same mineral assemblage except for the greater graphite and sulfide content in the rusty-weathering rocks. The Rowe Schist is characterized by alternating graphitic and non-graphitic layers. The graphitic variety tends to be more concentrated in the structurally upper part of the unit close to the contact with the Moretown Formation and near the abundant lens of mafic and ultramafic rocks. The Rowe Schist is a typical pelitic schist with quartz, plagioclase, muscovite, biotite, and garnet. Locally it is more aluminous and contains paragonite (Na-white mica) instead of albit and chloritoid instead of biotite. It also contains a surprisingly large amount of quartz for a sediment that was presumably deposited on the distal Laurentian margin; the quartz-rich layers are best observed on faces perpendicular to the foliation. The two samples of Rowe Schist used for our detrital zircon study came from along this road (Figs. 1C and 3). Both zircon populations are typical of Laurentian-derived sediment with prominent peaks between 1.0 and 1.2 Ga corresponding to the Grenville of North America, and minor Neoproterozoic peaks at approximately 550 Ma, reflecting rift-related magmatism from the breakup of Rodinia.
Note that the boundary between the Rowe Schist and the Moretown Formation is a first-order geophysical feature- the depth to MOHO, as determined by P to S wave conversions, abruptly decreases by 10 to 15 km from west to east approximately below this contact in southern New England (Li et al., 2018).

Dirt Road on right is Torrey Hill Road.
This road leads to the Reed Brook Preserve managed by the Nature Conservancy. A long and winding trail leads uphill to a large ultramafic lens with rare plants. Some large boulders of ultramafic rock are visible in the road a short distance from the paved road.

STOP 3. Rowe Schist, Mafic and Ultramafic lenses.
After crossing the bridge, continue straight on foot along the dirt road. Go uphill and across the railroad tracks. Turn right and avoid the drive ways to follow Tunnel Road on the far side of the tracks through the woods to get to Stop 3 outcrops. Lots of outcrop of Rowe Schist in woods upslope to left along the road.

STOP 3A. Rowe Schist
The large outcrop of Rowe Schist next to the road contains abundant graphite and large garnet porphyroblasts. Note the steeply-dipping well-developed foliation and prominent down-dip lineation.

STOP 3B. Mafic Schist
The prominent outcrops on the left and right of dirt road are composed of mafic schist and gneiss, likely derived from oceanic crust caught up in the subduction zone between the Moretown terrane and the down-going Laurentian margin rocks.

STOP 3C. Highly altered Ultramafic Rocks of the Hoosac Tunnel Soapstone Quarry.
The large blocks on the right side of the dirt road are waste from a soapstone quarry and talc mine. Soapstone was quarried between 1885 to 1895. In 1909 an adit was drilled in the quarry to mine talc. The operation did not last long. This fine-grained rock contains talc, actinolite, garnet, antigorite, and calcite. Ultramafic lenses are common in the Rowe Schist and Moretown Formation near their contact, and are probably derived from altered mantle peridotite.

Outcrop of Moretown Formation.

STOP 4. Moretown Formation.
This enormous outcrop is a convenient place to see the Moretown Formation in good light. Most of the light-colored rock is quartz, plagioclase, muscovite, biotite granofels with a distinctive pin stripe texture. There are some quartz-rich layers, one of which provided one of our many Moretown Formation detrital zircon samples. Dark layers are mafic dikes and sills with diverse geochemistry (Kim and Jacobi, 1996; Karabinos et al., 2017). The Moretown Formation is Cambrian in age (Macdonald et al., 2014; Karabinos et al., 2017). Based on geochemistry, the undated mafic intrusions most likely include Cambrian to Early Ordovician arc rocks, some of which may have formed during back-arc rifting, and Silurian or possibly Devonian rift-related rocks coeval with the initiation of the Connecticut Valley basin. Detrital zircons from the Moretown Formation from northern Vermont to southern Massachusetts show prominent 550 to 650 Ma peaks typical of Gondwanan-derived sediment (Figs. 1C and 4) Macdonald et al., 2014; Karabinos et al., 2017). There is a remarkable lack of Grenville peaks in the zircon populations.

Moretown Formation on left

STOP 5. Hawley Formation.
Park on right before bridge. Proceed on foot along Middle Road across bridge and turn left on pleasant dirt road. Make your way down to the Chickley River to the left. Note that this is private property and that permission is needed to access this outcrop.
The mafic gneiss exposed here is one of several common rock types in the Hawley Formation. Jon Kim studied these rocks for his doctoral dissertation and Kim and Jacobi (1996) is an excellent account of the geochemistry of this diverse unit. Structures widely interpreted as pillow basalts are present is some parts of this long exposure. The Hawley Formation is part of the Shelburne Falls arc and its Ordovician age is tightly constrained by a TIMS date of 475 Ma on the intrusive Dell Metatrondhjemite and essentially coeval detrital zircons in metasediments (Fig. 6; Macdonald et al., 2014). At least some of the mafic rocks (Kim and Jacobi, 1996) and perhaps most, if not all, of the metasediments formed during an episode of back-arc rifting. An anoxic environment, consistent with the restricted circulation of a rift basin, is required to explain the highly graphitic schist present in the western and eastern exposures of the Hawley Formation. Notable sulfide and manganese deposits in the Hawley Formation are also consistent with intra-arc rifting/splitting.
Our tectonic model described above is strongly driven by the fact that the detrital zircon populations in the Hawley Formation samples indicate both Laurentian and Gondwanan sources. The Gondwanan source is easy to explain because the Shelburne Falls arc formed on the Moretown terrane. The Laurentian detrital zircons component indicates that by about 475 Ma, the Moretown terrane and the Shelburne Falls arc were close enough to Laurentia to receive detritus from the Grenville rocks so prominent along the Laurentian margin.

Hallockville Pond. Lunch stop

STOP 6. Hallockville Pond Gneiss.
The Hallockville Pond Gneiss is a 475 Ma tonalite (Karabinos and Williamson, 1994; Macdonald et al., 2014) that intruded the Moretown Formation. The intrusive contact is not exposed at this stop, but it is visible nearby on the small hill 0.5 km northwest of the intersection of Hallockville and King Corner Roads in Dubuque State Forest. The arc geochemistry of this pluton (Karabinos et al., 1998) together with the observation that it intruded the Moretown Formation, provides compelling evidence that the Shelburne Falls arc was constructed on the Moretown terrane. The Hallockville Pond Gneiss contains a well-developed foliation that was overprinted by a younger crenulation cleavage. The older fabric, and possibly both, formed during the Taconic orogeny. 20 km to the south the Middlefield Granite at Glendale Falls in Middlefield, MA, is a 447 Ma weakly-foliated granodiorite (Karabinos and Williamson, 1994; Macdonald et al., 2014) that intruded both the Rowe Schist and the Moretown Formation at their contact, but the pluton is not offset at this major terrane boundary. Thus, the suture zone formed before this 447 Ma intrusion. These observations suggest that the strong deformation fabric in the Hallockville Pond Gneiss formed during the Taconic orogeny, and that the Rowe Schist-Moretown Formation fault contact was not greatly modified during the Acadian orogeny.

Outcrop of Collinsville Formation on right in arc of road.

Outcrop of Collinsville Formation on both sides of road.
A sample from this outcrop gave a 475 + 1.4 Ma TIMS date (Karabinos et al., 1998).

Trailhead for Mahican-Mohawk Trail
Alternate way to Stop 7 for those with very low clearance vehicles.

Turn right on Wilcox Hollow Road.
Dirt road can be rough after severe storms. Use alternate route and hike down Mahican-Mohawk trail if desired.

STOP 7. Collinsville Formation in the Shelburne Falls dome.
Park without blocking dirt roads. Continue down to power line. Look for small path on right that leads down to the Deerfield River. Long outcrop along the river continues upstream. Pay attention to rising water warnings; this outcrop is just below a hydroelectric dam.
This accessible exposure of the Collinsville Formation includes both tonalite and mafic gneiss. It is great fun to examine and argue about the intrusive relationship between the felsic and mafic rocks. The rocks are strongly foliated, lineated, and folded. The Collinsville Formation makes up the cores of the Shelburne Falls, Goshen, Granville domes. Core rocks from both the Shelburne Falls and Goshen domes are approximately 475 Ma (Karabinos et al., 1998) and surrounded by much younger Devonian metasediments. The contact is not an unconformity, but rather a major extensional shear zone, tops to the south, similar to that described for the Chester dome (Karabinos et al., 2010). Another fascinating feature about this part of the western New England dome belt is the occurrence of potentially ultrahigh-pressure garnet bearing rocks from the Goshen dome just south of the Shelburne Falls dome (Peterman et al., 2016).

C4: MOUNT GREYLOCK AS A COSMOGENIC NUCLIDE DIPSTICK TO DETERMINE THE TIMING AND RATE OF SOUTHEASTERN LAURENTIDE ICE SHEET THINNING
Christopher T. Halsted
Rubenstein School of Environment and Natural Resources
University of Vermont,
Burlington, VT 05405
chalsted@uvm.edu
Jeremy D. Shakun
Dept. of Earth and Environmental Sciences
Boston College
Chestnut Hill, MA 02467
jeremy.shakun@bc.edu
P. Thompson Davis
Dept. of Natural and Applied Sciences
Bentley University
Waltham, MA 02454
pdavis@bentley.edu
Paul R. Bierman -- pbierman@uvm.edu
Lee B. Corbett -- Ashley.Corbett@uvm.edu
Dept. of Geology
University of Vermont
Alexandria J. Koester
Dept. of Earth, Atmospheric, and Planetary Sciences
Purdue University
West Lafayette, IN 47907
This field trip to Mt. Greylock serves several purposes: 1) discuss and review the glacial history of Mt. Greylock and the surrounding region, 2) describe the motivation behind an NSF-funded project to understand the timing and rate of ice thinning in New England and adjacent areas, 3) explain the ‘dipstick approach’, the primary method being used to reconstruct ice thinning, and demonstrate the field work involved, 4) examine field sampling sites from this project and other glacial features on Mt. Greylock, and 5) discuss implications of initial results for paleoclimate and paleo-sea level change during the last deglaciation. Those interested can participate in a walk along the Appalachian Trail of approximately 2.7 miles (although trip participants are encouraged to lengthen their walk by ∼2 miles and return back to the Greylock summit via the AT if desired).
Before beginning this field trip, it is important to reflect on why the fields of glacial geomorphology and ice sheet reconstructions are areas of interest. Understanding ice sheets and their interactions with other components of the climate is essential to climate change mitigation and adaption discussions, but present projections of future ice sheet stability contain substantial uncertainties (e.g., Stocker et al., 2014; DeConto and Pollard, 2016, Hansen et al., 2016). These large uncertainties are due primarily to a limited understanding and poor model representation of ice sheet decay mechanics, particularly marginal processes such as ice-stream flow, ice-shelf buttressing, and subglacial meltwater drainage (Kleman and Applegate, 2014; Stokes, 2017). Modern observations of ice sheets in Greenland and Antarctica capture only a brief window in their history, and important basal processes are nearly impossible to observe. Reconstructing and studying the behavior of ice sheets during past periods of climate change offers a potential solution to observational and computational limitations.
The largest Northern Hemisphere ice sheet at the Last Glacial Maximum (LGM), the Laurentide Ice Sheet (LIS), is one of the most-investigated paleo-ice sheets, as its size and mixture of terrestrial and marine margins provide a unique opportunity to study the long-term behavior of ice margins analogous to modern-day locations in Greenland and Antarctica (Kleman and Applegate, 2014; Margold et al., 2015; Stokes et al., 2016; Stokes, 2017). Additionally, the near-complete retreat of such a large ice mass in only ∼15-kyr illustrates the sensitivity of ice sheets to changing climatic conditions (Abe-Ouchi et al., 2013; Ullman et al., 2015).
The nature of interactions between the LIS and the rest of the climate system during the last deglaciation (∼21- 11 ka) are the source of much debate in the paleoclimate community. The last deglaciation was marked by a series of abrupt climate change events that are particularly apparent in proxy records from the North Atlantic and Greenland (e.g. Clark et al., 1999; Thornalley et al., 2010). A mechanism for similar abrupt events observed further back in the paleoclimate record (Dansgaard et al., 1993; Johnsen et al., 2001) was originally proposed by Broecker et al. (1985). They argued that changes in the strength of the Atlantic Meridional Overturning Circulation (AMOC) caused the North Atlantic to rapidly shift between two quasi-stable modes of climate. When the AMOC was strong, North Atlantic average annual temperatures were high (interstadial), and when the AMOC was weak, annual North Atlantic temperatures dropped (stadial). Clark et al. (1996) noted that a similar mechanism may have caused the abrupt warming and cooling events observed during the last deglaciation, and eventually hypothesized that the direction and volume of the LIS meltwater flux may have had a direct influence on the strength of the AMOC (Clark et al., 2001).

C4: MOUNT GREYLOCK AS A COSMOGENIC NUCLIDE DIPSTICK TO DETERMINE THE TIMING AND RATE OF SOUTHEASTERN LAURENTIDE ICE SHEET THINNING
C.T.Halsted, JH.D.Shakun, P.T.Davis, P.R.Bierman, L.B.Corbett, A.J.Koester

Meeting Place
C.T.Halsted, JH.D.Shakun, P.T.Davis, P.R.Bierman, L.B.Corbett, A.J.Koester

STOP 1. Park at Stoney Ledge.
Sperry Rd makes a loop here, please find parking off the road. Stoney Ledge offers fantastic views of Mount Greylock and The Hopper, a glacial cirque on Greylock’s west flank. Exposed bedrock offers the potential to investigate glacial striations!
C.T.Halsted, JH.D.Shakun, P.T.Davis, P.R.Bierman, L.B.Corbett, A.J.Koester

STOP 2. Arrive at Mount Greylock summit.
Park in designated parking lot and meet outside Bascom Lodge. Bascom Lodge provides restrooms, food for purchase, and free clean drinking water. An in-depth discussion of regional glacial geomorphology, the chronology of the Laurentide Ice Sheet in the area, and the motivations and preliminary results of an NSF-funded ice sheet reconstruction project will take place here. For those not participating in the walk on the Appalachian Trail, the summit marks the end of the trip. Auto roads depart the summit to the north and south (we will drive up the southern road), please drive carefully!

STOP 3. Two boulders sampled for project
Briefly examine two glacially-transported boulders deposited on the Greylock ridgeline. These boulders were sampled for the NSF dipstick project.

STOP 4. Mt. Williams summit.
This is not truly a separate mountain, as it is part of the Greylock ridgeline, but it is a local high-point that offers fantastic views of the surrounding area. Two samples were collected at this location for the NSF dipstick project, one bedrock and one boulder. From here, trip participants can choose to either return via the AT to the Greylock summit or continue down the AT for another half mile, eventually meeting the northern Greylock auto road (Notch Rd). The trip leader will arrange to have a van meet at Notch Rd and transport participants back up to the summit, although participants are welcome to arrange their own transport if desired.

C5: HEAVY METAL CONTAMINATION FROM ILLEGAL BURN PILES IN AN ECOLOGICALLY
SENSITIVE SITE IN WEST HAVEN, VERMONT
Helen Mango -- helen.mango@castleton.edu
Mary Droege -- mary.droege@castleton.edu
Department of Natural Sciences
Castleton University
Castleton, VT 05735
J. Murray McHugh -- mmchugh@tnc.org
The Nature Conservancy
Southern Vermont Office
348 Bentley Ave.
Poultney, VT 05764
Michele Hluchy -- fhluchy@alfred.edu
Geology and Environmental Studies
Alfred University
Alfred, NY 14802
The Nature Conservancy owns and manages the Helen W. Buckner Memorial Preserve in West Haven, Vermont. The 3827-acre preserve has one of the highest levels of biodiversity in the state. It is located on the border of New York State, bounded on the south by the Poultney River and on the west by the South Bay of Lake Champlain (Fig. 1). It is an environmentally sensitive and important tract of land, providing habitat for eleven rare/uncommon animal species (including the timber rattlesnake), eighteen rare/uncommon plant species, and ten rare/uncommon natural community types. It encompasses undeveloped lake shoreline, wetlands along the river, floodplain, Clayplain and upland forest, marshes, and cliffs. It includes Austin Hill, composed of Precambrian (Middle Mesoproterozoic) Adirondack biotite tonalite gneiss, and bounded on the west by Cambrian/Ordovician dolostone and quartzite (Fisher, 1984; Ratcliffe et al., 2011).
Illegally dumped and burned garbage piles have existed for years on the access road (Galick Road) along a stretch of the Poultney River. Analysis of ash and soil samples from different burn piles over several years shows elevated levels of many contaminants, including arsenic, copper, lead, zinc, chromium and cobalt, often exceeding background levels by orders of magnitude. The area is seasonally flooded, resulting in both erosion and deposition. As a result of our research, the Vermont Department of Environmental Conservation (DEC) remediated the most contaminated site and installed a parking area and boat launch (allowing for patrol by state law enforcement). However, dumped and burned garbage continues to be a problem at this site.

STOP 1A: Site of Major Garbage Burn Piles
This is the site of Burn Piles 1 and 2. Although geochemistry does not lend itself well to field trips, there are several things worth observing here. As a result of this study, the State of Vermont hired a geologic consulting company to assess the site, remove contaminated soil, place boulders to prevent vehicles from driving further along the river bank, re-grade the burn pile site as an official “boat launch” and create the parking area. Making this an official boat launch allowed Fish and Wildlife law enforcement to patrol the area and issue citations. (Previously, it was up to the town of West Haven to supply law enforcement, which, despite the good efforts of the local sheriff, was completely inadequate to keep up with the illegal/questionable use of the site.)
It is highly likely that there will be a reasonably fresh burn pile, and while we might not see the variety of materials that have been dumped and burned in the past, this will give trip participants a sense of the scale of activity. We can also observe the current river level, and look for evidence of flooding (which periodically erodes burn piles, sending their contents and the contaminated soil around and beneath them, directly into the river and shortly thereafter into Lake Champlain.

STOP 1B. Clay Plain Soils and Reforestation
The Nature Conservancy’s Southern Lake Champlain Valley office launched an ambitious 11-year Clayplain forest restoration project in 2004 in multiple sites in West Haven and Benson, Vermont. A decision to use only local seed source lead to the establishment of a small native woody plant nursery called the Champlain Valley Native Plant Nursery. 76,000 native trees and shrubs were planted between 2004 and 2014 on more than 200 acres. Funding was provided by USDA’s Wetland Reserve Program, Wildlife Habitat Incentive Program and USFWS Partners For Fish and Wildlife. Survival of the plantings has been mixed, but adjacent to the parking area on Galick Road, which was planted in 2005, survival and growth has been excellent.

STOP 2. Site of “Background” Samples and View of Cliffs and Talus Slope
This is a very brief stop to indicate the location of the “background” samples used in this study. The trailhead is just in the woods next to a large field that is kept open for field-nesting birds, amphibians, and other wildlife. The 2014 background sample was obtained from the very edge of the field, and the 2015 sample was obtained just into the woods, both approximately 100 m from the road.
A little further along the trail (50 m) is an excellent view of the cliffs of Middle Mesoproterozoic gneiss and the talus slope below.

Kiosk - Information Panel
At the kiosk, go left to take the trail that goes through the woods along the base of the cliff. The trail to the right goes along the edge of the meadow; we will return by this trail.

STOP 3. Mesoproterozoic Granitoid Gneiss-Granulite
The trail is right next to the cliffs here.
This is a brief stop to examine the biotite tonalite gneiss. The rock contains mostly quartz and plagioclase feldspar, with the result that where the foliation is subtle the rock could more appropriately be called a granulite. Minor biotite and pyroxene are visible in hand sample and microscopically. Farther along the trail there is an outcrop that exhibits folded foliation. The rock is very hard, creating the cliffs, and weathers in large blocks, creating the talus slopes that provide critical habitat to a number of species, including the timber rattlesnake.

STOP 4. Ward Marsh Overlook
Ward Marsh is visible to the left (east). In addition to the rattlesnake research discussed above, the Buckner Preserve is home to what has recently been recognized as the largest populations of golden-winged warblers, bluewinged warblers and their hybrids in New England. The Nature Conservancy is partnering with Audubon Vermont to monitor the populations and manage areas of the preserve to maintain and improve habitat for these birds.
To the southeast (to the right of Ward Marsh from this vantage point) is Skene Mountain. It consists of Cambrian and Ordovician sedimentary rocks. At the base is Potsdam Sandstone, overlain by Ticonderoga Formation dolostones and above that the Whitehall Formation. The Whitehall Formation consists of interlayered dolostones and limestones; Skene Mountain is capped by the Skene Member, a very hard, massive dolostone.

Outcrops showing gneissic fabric
Outcrops showing gneissic fabric are plentiful as the trail heads downhill through a hemlock forest.

STOP 5. LUNCH And View
This overlook provides an excellent view of the Poultney River near its mouth and the South Bay of Lake Champlain. The hill directly south across the Poultney River is made of Hague Gneiss and quartzite of the Lake George Group. Skene Mountain is visible to the east.