NYSGA 2015
Geology of the Northeastern Adirondack Mountains
and Champlain–St. Lawrence Lowlands of
New York, Vermont and Québec
New York State Geological Association, 87th Annual Meeting
State University of New York at Plattsburgh,
Plattsburgh, NY
12–13 September 2015
87th Annual Meeting – New York State Geological Association
2015 NYSGA Organizing Committee
David Valentino (NYGSA Executive Secretary)
Mary K. Roden-Tice (2015 Meeting Coordinator)
David Franzi (2015 Guidebook Editor)

A1: THE BENNIES BROOK SLIDE - A WINDOW INTO THE CORE OF THE MARCY ANORTHOSITE
Jeff Chiarenzelli
Department of Geology
St. Lawrence University
Canton, New York 13617
Marian Lupulescu
New York State Museum
Research and Collections
3140 CEC, Albany, New York 12230
Sean Regan
Department of Geosciences
University of Massachusetts
Amherst, MA 01003
Dave Valentino
Department of Atmospheric and Geological Sciences
SUNY Oswego
Oswego, NY 13126
Doug Reed
238 Little Colfax Rd.
Cambridge, NY 12816
INTRODUCTION
The Marcy anorthosite massif forms the core of the Adirondack High Peaks region and is the geographic center of the associated topographic dome. On the State Geological Map (Adirondack Sheet, 1:250,000) it is depicted as a single rock type consisting of meta-anorthosite. However, this is a simplification and, in detail, numerous lithologies occur within the massif and display complex relations. This trip is designed to introduce participants to a 1.5 km, near continuous, linear landslide exposure which follows the course of Bennies Brook on the northern flank of Lower Wolf Jaw Mountain. Saturated conditions associated with Hurricane Irene in August of 2011 resulted in the most recent of landslide events along the course of Bennies Brook and additional exposure. The features unearthed here provide an unprecedented view into the interior of the massif and demonstrate its geological complexity.

Departure Point

Meeting Point
I will meet trip attendees at the cairn where Bennies Brook meets the Southside Trail at 10 am.

STOP 1. Intersection of Bennies Brook Slide and South Side Trail.
→ Features: 1) typical Marcy-type anorthosite; 2) garnet coronas; 3) small ductile shear zones; 4) sub-meter scale north-trending joints
The South Side trail follows John’s Brook and was severely damaged during the flooding associated with Hurricane Irene. The trail crosses the slide a few hundred meters above the confluence of Bennies Brook with Johns Brook. A stone cairn marks the intersection (Figure 6). If you don’t see the stone cairn you are at one of the other nearby slides in drainages on the north side of the range and this guidebook won’t be of much use. If you look to the north towards Johns Brook you can see a variety of woody debris and boulders from the slide (Figure 7).
Exposed in the bed and banks of Bennies Brook is typical Marcy type anorthosite (Figure 8). Here it consists mostly of 90% or more of darker andesine, some reaching lengths of 5 cm or more; mere pikers compared to what is to come. The lighter colored matrix is also plagioclase, much of which shows bent twins and other features indicative of some type of strain (recrystallization or auto brecciation?) during or just after final emplacement. Other minerals include pyroxene (often of both persuasions), oxides (magnetite and/or ilmenite), and trace minerals including biotite, hornblende, and garnet. Close examination reveals that much of the garnet rims mafic minerals and provides evidence of high-grade metamorphism. Throughout the outcrop proportions of minerals and grain-size vary.
Here cm-wide shear zones are developed. Some concentrate and deform garnets and pyroxenes and this will be much more evident and convincing as you proceed upwards. They are generally subvertical and trend about 340o here. A set of parallel joints are developed slightly oblique to the ductile shear zones and trend 352o 90o.

STOP 2. Heavy Minerals in Bennies Brook and Garnet Coronas
→ Features: 1) garnet and heavy mineral-rich sand in brook; 2) garnet coronas There are deposits of reddish and black sand exposed in the bed of Bennies Brook, often forming layers or ripples or other bed forms (Figure 9). These are derived primarily from within the drainage basin and show the relative abundance, and density, of garnet, pyroxene, and oxides. A sample of this sand has been collected and will be characterized by SEM and detrital zircon studies.
Wandering up from this location for the next 50 m or so will show some excellent examples of garnet coronas. Coronas are developed upon mafic minerals such as pyroxene and olivine, as well as, on oxide minerals (Figure 10). Some form delicate single grain necklaces or chains. In many cases they may represent the only metamorphic reaction visible in the rock. Davidson and Van Breeman (1988) showed similar zircon coronas on baddeleyite in Grenville mafic dikes. The zircon gave ages 100 million years younger than the baddeleyite cores documenting both igneous crystallization (1150 Ma) and Metamorphism (1050 Ma). In shear zones garnets are deformed and particularly abundant indicative of deformation and high-grade conditions and the abundance of the residual components needed for it to form (i.e. Fe, Al, Si).

STOP 3. Altered Anorthosite and Brittle Structures
→ Features: 1) alteration zones in anorthosite; 2) closely spaced fractures and brittle faults; 3) tiny carbonate veins
While typically fresh, and well preserved, some anorthositic rocks have experienced numerous lower grade alteration events. While the exact nature and timing of events is currently unknown many appear to be associated with low-temperature hydrothermal alteration associated with areas of high fracture density and brittle faults. Typically bluish grey when fresh, anorthosite often appears chalky white to pale green where altered and rusty where also weathered (Figures 11 and 12). In these zones andesine is nearly completely altered to epidote, pumpellyite, clinozoisite, chlorite, carbonate, and other secondary minerals.
At this locality there are composite alteration zones and veins that range in color from white to pink to pale green (Figure 11). They appear to be in close proximity and best developed where brittle faults occur (Figure 12). In addition, to these features the rock is heavily laced with thin (mm-scale) white calcite veins. Secondary calcite in microfractures is ubiquitous within the anorthosite (Morrison and Valley, 1988). A similar mineralized alteration zone up to several meters wide, cutting anorthosite, has been studied by the senior author and colleagues at SUNY Potsdam.

STOP 4. Xenoliths in Anorthosite
→ Features: 1) xenoliths; 2) white, finer-grained anorthosite
In this stretch of the slide there are a number of xenoliths. This is not uncommon for the anorthositic rocks to include well documented metasedimentary country rock lithologies (Valley and Essene, 1980) and also igneous xenoliths (McLelland). Here their origin is debatable, although some display a deep green clinopyroxene, possibly diopside (Figure 13). Many of the xenoliths are highly disrupted and partially assimilated into the anorthosite. In some areas the anorthosite appears to be chilled against larger xenoliths resulting in white, fine-grained regions in the anorthosite where in contact with the xenoliths. However, white finer-grained anorthosite also appears to have intrusive relationship with fresh anorthosite as well.

STOP 5. Dikes and Shear Zones
→ Features: 1) intrusive granitic dikes; 2) narrow ductile shear zones
In this section of the slide, as elsewhere, you will note a number of straight-walled and narrow 5-10 cm granite and pegmatite dike-like intrusions into the anorthosite. On closer inspection, shearing and mylonitic fabrics parallel to the boundaries of dikes are noted (Figure 14). Quite often the surrounding anorthosite is completely unaffected by the strain. Most of these shear zones trend nearly 110°, coincident with the major east-west trending Shawinigan-aged Piseco Lake Shear Zone in the southern Adirondacks. Where observable the shear zones display left-lateral kinematics. Many such dikes can be traced for 10’s of meters across the slide.

STOP 6. Complex Shear Zone
→ Features: 1) gabbroic pegmatite; 2) sinistral shear zone modified by late dextral motion; 3) relative age relations
At this stop a spectacular example of complex shearing, lithologic variation, and relative age relationships can be seen. Shown in Figure 15 there is a contact between Marcy-type anorthosite (left-hand side) and gabbro (right hand side). This contact was intruded by a gabbroic pegmatite with crystals approaching 10 centimeters in length. All three of these rocks were intruded by a thin aplitic granite trending 110°. After intrusion both the gabbroic pegmatite and aplite show strong left-lateral displacement parallel to the aplite margins. After cessation of left-lateral motion the rocks were affected by right -lateral shear (200°), deflecting the contact and the previously developed left-lateral shear fabric. Where observed elsewhere dextral shear also follows earlier sinistral shearing at 110° but also appears to have operated at high-grade metamorphic conditions.

STOP 7. High-grade Sinistral Shearing in Anorthositic Rocks
→ Features: 1) large sinistral shear in anorthosite; 2) pyroxene kinematic indicator; 3) shear in granitic pegmatite
As observed at the previous location and elsewhere on the Bennies Brook slide a large number of small sinistral shear zones developed along lithologic discontinuities occur. In most cases the strain is taken up in the dike or pegmatite and is parallel to the dike walls and often is completely absent in the anorthositic wall rocks. At this location you will see a ∼2m shear in anorthositic rocks indicating deformation of sufficient magnitude to affect a relatively rigid lithology (Figure 16). Also in the same shear zone you will see evidence of the high-grade nature of the shear zone in the form of sinistral pyroxene “fish-shaped” kinematic indicators (Figure 17). Finally, you will see evidence for opening of fractures approximately perpendicular to dike margins, followed by shearing parallel to the contact (Figure 18). This appears to indicate a nearly 90° change in stress direction after anorthosite intrusion.
These features provide clues to the overall kinematic history of the High Peaks region, when deformation occurred, and the relative age of events. Here they are clearly exposed without ambiguity. The trend of the shear zone is 114° 82° and a brittle fault is developed parallel to its margin. A well- developed lineation in the ductile shear zone has the orientation of 115° 16°.

STOP 8. Neoproterozoic dikes
→ Features: 1) Neoproterozoic Dikes; 2) Magmatic layering in anorthosite
Straight walled Neoproterozoic dikes are common in the Adirondack Region but rarely exceed a meter or two in thickness. They are believed to be part of the extension and mafic magmatism associated with the opening of the Iapetus Ocean. The dikes cut Grenville rocks over a large area from the Appalachian inliers to Newfoundland, with ages ranging between ∼570-640 Ma. Here a variety of dikes ranging in size from a few cm’s to a few meters are found. Most are vertical and northeast (∼40°). Some appear to be composite (Figure 19) and higher up in the slide relatively rare felsic trachytic dikes intrude the basalts, allowing the possibility of U-Pb dating.
In many areas anorthositic rocks subtle but regular changes in grain-size and mineral proportions are readily observable with the naked eye. Some are so subtle that only a vague meter-scale alteration of color demarks them (Figure 20). Others have finer layering and contain considerably more mafic minerals. While these layers are nowhere nearly as well developed as in mafic-ultramafic intrusions, they are believed to represent original magmatic, perhaps cumulate, layering with the anorthosite.

STOP 9. Ferrodiorite Contact
→ Features: 1) giant andesines; 2) raft of coarse Marcy type anorthosite; 3) ferrodiorite contact with anorthosite
The anorthosite within the massif is often a true pegmatite in terms of its grain-size and exceptionally large andesines can be found as xenoliths and rafts in other lithologies. At this locality large andesines up to 20cm or more can be found included in anorthosite of more moderate grain-size (Figure 21). In addition, a large raft of exceptional coarse anorthosite can also be seen (Figure 22). The raft has minimum dimension of 2 x 3 meters and consists of large andesine crystals with few other minerals. In an upcoming stop we will see exceptionally large pyroxenes as well.
On the upper portions of the Bennies Brook slide ferrodiorite or jotunite becomes a dominant rock type intruding the anorthosite and related rocks (Figure 23). Travelling up the slide another 20 meters or so reveals ferrodiorite as a network of thin, box-like dikes around anorthosite blocks and as large intrusive sheets. It is invariably chilled against the anorthosite and has scalloped or complex margins. In some cases andesine megacrysts can be seen increasing in abundance towards the contact and are often aligned along their long axis. The ferrodiorite is demonstratively one of the last intrusive phases as it cuts gabbroic pegmatites which have been previously deformed (Figure 24), indicating intrusion after ductile deformation.

STOP 10. Pyroxenite Block
→ Features: 1) numerous xenoliths; 2) pyroxenite block; 3) composite shear zone
As previously noted a wide variety of xenoliths can be found in the anorthositic rocks including metasedimentary country rocks of the Grenville Supergroup (Valley and Essene, 1980) and those of igneous parentage. Those of igneous origin are generally assumed to be related to the anorthosite and related rocks. At this locality a large meter-scale block of coarse-grained pyroxenite can be observed (Figure 25). If the anorthosite did develop by fractional crystallization from gabbroic magma, then it is lacking in requisite mafic minerals. One way to explain this dearth of Fe and Mg-bearing silicates, is to separate the anorthosite by density differences. Perhaps this pyroxenite was entrained as a rising mass of anorthosite made its way higher into the crust?
Also at this locality is a second photogenic example of a composite shear zone with sinistral movement modified by a later dextral shear (Figure 26). This example is developed in a gabbroic pegmatite cutting anorthosite. The sinistral shear trends ∼110°, while the dextral shear, larger of the two in this case, trends ∼350°. This confirms the sequence of deformation described previously and suggests both sinistral and dextral components of late deformation affected the rocks exposed in the High Peaks.

STOP 11. More Ferrodiorite
→ Features: 1) ferrodiorite defining block structure in anorthosite; 2) giant pyroxenes; 3) off-set of chill margin
At this point on the slide ferrodiorite is a major rock type and it occurs on all scales from small dikes (2-3 cm across) to large sheets intruding anorthosite. At several locations irregular dikes of various orientations and sizes and shapes define anorthosite blocks in a jig-saw fashion (Figure 27). Also exposed is a meter-scale xenolith with giant pyroxenes up to 20 cm or more in diameter (Figure 28). It is engulfed in anorthosite, and both are cut by ferrodiorite. An offset contact between anorthosite and an andesine xenocryst-rich ferrodiorite or ferrogabbro is also exposed. The offset appears to be brittle and right lateral (Figure 29). However, the origin of the rock at the contact can be debated. Is it a contact chill margin or a layered dike that intruded along the contact or the wall of a magma chamber? We will let your group argue over how best to interpret and test the origin of this enigmatic exposure.

STOP 12. The End
→ Features: 1) magmatic layering in xenolith; 2) composite bimodal Neoproterozoic dike
This is the final stop just before the splitting of the slide into three individual forks of different ages. Here the route steepens as you approach the head wall and it is an excellent place to turn around and descend. Two unique features are exposed in this section of the slide and include a layered mafic intrusion and a composite dike.
The layered mafic intrusion shows rhythmically cm-scale alteration in the proportion of mafic minerals (Figure 30). Its overall composition suggests it may be a variety of ferrodiorite; however, further work is needed to confirm this. Nonetheless it is striking in its difference from the bulk of the ferrodiorite and other rocks and testifies to the variety of igneous textures possible in this situation.
A number of small basaltic dikes occur at this elevation on the slide. One reaches a width of 2m and has an orientation of 50° 90°. What is unique here, though, is a few small felsic dikes (Figure 31) that appear to concordantly intrude the basaltic dikes. This suggests that they are broadly coeval. A similar trachytic dike intruding in a swarm of Neoproterozoic basaltic dikes near Dannemora yielded zircon. This zircon gave a concordant U-Pb zircon age of 643±4 Ma. Work is in progress to date this dike as well.

A2: PRECAMBRIAN GEOLOGY OF THE EAGLE LAKE QUADRANGLE, ESSEX COUNTY, NY
Sean P. Regan, Phillip S. Geer, Gregory J. Walsh, and John N. Aleinikoff
United States Geological Survey
Graham B. Baird
Department of Earth and Atmospheric Sciences
University of Northern Colorado
Greeley, CO 80639
Peter M. Valley
Weatherford Labs
Houston, TX 77086
Michael L. Williams and Michael J. Jercinovic
Department of Geosciences
University of Massachusetts
Amherst, MA 01003
Timothy W. Grover
Castleton State College
Castleton, VT 05735
INTRODUCTION AND REGIONAL BACKGROUND
The Grenville Province of eastern North America has long been interpreted as the roots of an ancient orogenic system that formed approximately 1.0 billion years ago (Hoffman, 1988; McLelland et al., 2010). The entire region formed and evolved during punctuated phases of crustal growth during the Mesoproterozoic involving a variety of tectonic processes including accretion, arc magmatism, back-arc development and collapse, and unique to the Grenville, anorthosite-mangerite-charnockite-granite (AMCG) plutonism. These processes occurred over > 300 my and are collectively referred to as the Grenville Orogenic cycle (Fig. 1; McLelland et al., 1996; Rivers, 2008). The northeast Grenville Province has been shown to have a long history, extending back to ca. 1.35 Ga (McLelland et al., 1993), and contains evidence of long-lived active margin processes (Dickin and McNutt, 2007; Chiarenzelli et al., 2010a). However, the most extensive regional deformation has been interpreted as having formed during a culminating collision at ca. 1.07 Ga (Ottawan orogeny), which resulted in the final assembly of the supercontinent Rodinia (McLelland et al., 1992; 1996; 2001a; 2010; 2011; Rivers et al., 2008). The Grenville Province offers an excellent opportunity to study the growth and modification of mid to lower continental crust during these tectonic processes within the context of a long-lived active margin (Mezger, 1992).
The Adirondack Mountains of northern New York are a large domical uplift and are subdivided into the northwest Adirondack Lowlands and southeast Adirondack highlands separated by the Carthage- Colten shear zone (Streepey et al., 2001; Selleck et al., 2005). The Adirondack Highlands are dominated by granulite facies (meta)igneous rocks in contrast to the Lowlands, which are dominated by supracrustal rocks metamorphosed to amphibolite facies. Another important difference is the timing of deformation in each region. In the Lowlands the majority of deformation was the result of Shawinigan orogenesis (ca.1.17 Ga; Fig. 1b; Wasteneys et al., 1999; Heumann et al.,2006; Chiarenzelli et al., 2010b; Wong et al., 2011a), and evidence for Ottawan deformation (ca. 1.07 Ga; McLelland et al., 2001a) is lacking (Baird and Shrady, 2011). In contrast, the Adirondack Highlands are known to have experienced both Shawinigan and Ottawan phases of the Grenville Cycle. Here, differentiating the effects of these two phases remains a major challenge that is critical for understanding the role and importance of the Ottawan orogeny in the Grenville Province as a whole (Rivers, 2011).
This field trip focuses on the general Precambrian geology of the Eagle Lake quadrangle (ELq) in the eastern Adirondack Highlands. We will discuss recent 1:24,000 scale mapping of the ELq and the historical mining area Hammondville (Geer et al., 2015). The new mapping represents work by Regan, Geer, and Walsh. Provisional mapping by Walton (1960) has proven quite helpful, and provided a foundation for the new work. We will present a new structural nomenclature established during mapping that is regionally transformative and will hopefully be a useful guide for future analytical work. New forward petrologic modeling of the metamorphic assemblages that define a major structural fabric, geochronologic data that constrain the timing of both granite petrogenisis and anatexis of metasedimentary lithologies, and geochemical data from the historical mining district of Hammondville.
The Adirondack Highlands are predominately underlain by granulite facies (meta) igneous lithologies (McLelland et al., 1996). The most voluminous plutonic rocks exposed belong to the anorthosite-mangerite-charnockite-granite (AMCG) suite, which underlies approximately 60 % of the Adirondack Highlands (Hamilton et al., 2004). The petrogenisis of AMCG suites has been debated for decades, but detailed U-Pb geochronologic work presented in McLelland et al. (2004) and Hamilton et al. (2004) shows that granitoid lithologies (1158 Ma) are statistically older than anorthositic and other mafic members of the suite (1154 Ma), but overlap in age (Figure 5e). These data were interpreted as showing that AMCG suites are “… coeval, but not necessarily comagmatic” (McLelland et al., 2004). Geochemical and isotopic analysis performed by Regan et al. (2011) showed that coronitic metagabbros of the Adirondack Highlands (McLelland and Chiarenzelli, 1988) are permissible as the parental magma for the anorthosite. Quartz-bearing end-members were interpreted to have been derived from partial melting of lower continental crust where anorthosite formed via fractional crystallization of an asthenospherically derived gabbroic magma (McLelland et al., 2004; Bickford et al., 2008; Regan et al., 2011). The largest body of anorthosite is the Marcy Massif, which underlies the high peaks region of the Adirondack Mountains (Buddington, 1939). This suite of rocks has been interpreted as the result of lithospheric- mantle delamination at the waning phases of Shawinigan orogenesis (McLelland et al., 2004). Therefore, the deformation and metamorphism present throughout AMCG rocks has been interpreted as Ottawan in age.
The Lyon Mountain Granite Gneiss (LMG: nomenclature after Postel, 1952) is the youngest igneous rock exposed within the Adirondack highlands (Buddington, 1939; Postel, 1952; McLelland et al., 2001a,b,c; Selleck et al., 2005; Valley et al., 2009; 2010). It is a leucogranite that rims the Adirondack Highlands and is host to low-Ti magnetite deposits interpreted as forming via kiruna-type Fe – oxide, copper, gold (IOGC) mineralization processes (McLelland et al., 2001b,c; Valley et al., 2009). The LMG contains no consistent shape-preferred orientation (except in the East Adirondack shear zone; see Grover et al., trip B-2), and very little evidence of regionally significant subsolidus deformation, although the nature of layering is currently being debated (McLelland et al., 2001a). The LMG has been interpreted as accompanying orogenic collapse at the tail end of Ottawan orogenesis (McLelland et al., 2001a; Selleck et al., 2005).

A2: PRECAMBRIAN GEOLOGY OF THE EAGLE LAKE QUADRANGLE, ESSEX COUNTY, NY
S.P.Regan, P.S.Geer, G.J.Walsh, J.N.Aleinikoff, G.B.Baird, P.M.Valley, M.L.Williams, M.J.Jercinovic, T.W.Grover

Ticonderoga McDonalds Parking Lot
S.P.Regan, P.S.Geer, G.J.Walsh, J.N.Aleinikoff, G.B.Baird, P.M.Valley, M.L.Williams, M.J.Jercinovic, T.W.Grover

STOP 1. Roadcut on NY-74, west of Ticonderoga, NY
The easternmost rock is a poikiloblastic garnet amphibolite that contains garnet-clinopyroxene, plagioclase, and hornblende. Fe-Ti oxides are present as well, though 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 and appears to both traverse and wrap around the garnet poikiloblasts, suggesting garnet growth was syn-kinematic.
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% and preferentially contains garnet, while biotite is concentrated in the melanosome. Though some texturally late leucosome contains biotite, this can be shown to grade into garnet-bearing leucosome. 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 a number of garnet rich boudins occur, 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.
Mineralogy of the pegmatites includes feldspar, quartz, and biotite. Most feldspar appears to have exsolution lamellae indicating alkali feldspar, but locally plagioclase has been observed. Due to the coarse grain size (upwards of 30 cm for feldspar and quartz and 50 cm. for biotite) the exact lithology has not been constrained, but it is thought to be an alkali-feldspar granite. Locally associated with the pegmatites in the host migmatite are cm. books of graphite.
Fabric orientation is essentially parallel among the garnet amphibolite, the migmatitic paragneiss, and the transitional contact zone between the two. The foliation in the garnet amphibolite and eastern portions of migmatitic paragneiss is approximately 062°, 43° but changes to about 103°, 45° by the westernmost end of the outcrop. Lineation is reasonably parallel to the outcrop face and ranges from a horizontal to 30° plunge, tending 100°. The lineation is defined by aligned hornblende, biotite, or sillimanite, depending on location within the outcrop. Asymmetric tails on garnet poikiloblasts in the garnet amphibolite suggest south side (top) to the west transport.
To constrain timing of geologic events recorded in the outcrop, samples of the migmatitic paragneiss and of pegmatite were processed for U/Pb zircon geochronology. For the paragneiss, given the scale of leucosome and melanosome interlayering, it was not possible to separate the two upon processing, so separated zircon is from both components. Such zircons are euhedral to subhedral, stubby to elongate, yellow to amber to brown in color, and are 100 µm to 500 µm in length.
Cathodoluminescence (CL) of polished grain cross sections show that zircons have a wide variety of internal zoning patterns from oscillatory to irregular zoning (Figure 11 inset). No consistent core-rim zoning pattern (or of any other type) could be generalized for the separated zircon grains. Results of 21 SHRIMP-RG analyses of all types of zircons and all CL zones reveal what it interpreted as a unimodal age distribution. 16 of 21 analyses define a chord with an upper-intercept of 1186 ± 25 Ma, which is interpreted to be the age of migmatization for the paragneiss (Figure 11).
Zircon from one pegmatite body is euhedral to subhedral, equant to stubby, and light pink to purple in color. Zircon size is poorly constrained as most grains are fragments of grains greater than approximately 500 µm. Cathodoluminescence of polished grain cross sections reveal all grains possess oscillatory or sector zoning (Figure 12 inset). No convincing core-rim relationships were found. Fourteen SRHIMP-RG analyses produce what is interpreted as a unimodal distribution of ages. Six if these are nearly concordant and produce a 207Pb/206Pb weighted mean age of 1051 ± 25 Ma that is thought to best represent the age of pegmatite formation (Figure 12).
S.P.Regan, P.S.Geer, G.J.Walsh, J.N.Aleinikoff, G.B.Baird, P.M.Valley, M.L.Williams, M.J.Jercinovic, T.W.Grover

STOP 2. Transposed Grenville Supergroup
This is an exposure of metamorphosed and folded meta-siliciclastics of the Grenville Supergroup (Davidson, 1998). Notice the very well layered nature of the rock, the related compositional variability, and the gradational transitions between various layers. Also note several asymmetric folds that are particularly easy to see within leucosome. This suite of rocks saw the following events: 1) deposition, 2) partial melting (M1), and 3) folding of this migmatitic gneissosity, and the development of axial planar fabric (S2). These relationships suggest that anatexis and high-grade metamorphism likely accompanied S1, which is interpreted to be Shawinigan based on the geochronology discussed at stop 1. This suite of rocks is common throughout the ELq, and represents some of the oldest rocks preserved within ELq.

STOP 3. Contact between mangeritic gneiss and the Grenville Supergroup.
Several small outcrops are present to the southwest along Old Furnace Road. We will start by the parked cars. In stark contrast to the rocks seen at the previous outcrop, this rock is fairly homogenous. It is a mangeritic gneiss, with hbl ± grt composing the tails around primary opx (McLelland et al., 1980).
Do you see evidence of more than one phase of deformation within this lithology? As we move south, we will see a quick transition back into the Grenville Supergroup. This produces an interesting map pattern, a dome structure that is cored by monocyclic mangeritic gneiss, engulfed by meta-siliciclastic rocks of the Grenville Supergroup, with the predominate fabric running parallel to the axial trace of (F2?). However, we know from the previous stop that the predominate fabric we see in the paragneisses is an aggregate of two planar fabrics. The big question is; can we see S1 in the mangeritic gneiss?

STOP 4. Hill No. 8
Hill No. 8 is in the northern part of the historic mining town of Hammondville. On the north eastern base is the remains of a great furnace used to melt and separate the ore mined here in the late 19th century (next stop). The remaining slag is the blue stone used on many of the roads in the area that you can see on the drive in and out.
Hill No. 8 was the northern most extent of the mining operations of Hammondville and included two major pits: the Hammond Pit and Number 8 Mine. The number 8 mine is the deepest in Hammondville, reaching depths of up to 1000 ft (Penfield museum, personal correspondence), however the exposure in the Hammond pit is better and more easily accessible.
Magnetite seams in the Hammond pit range from centimeter scale to roughly 2 meters in thickness. The seams display a discrete wavy contact with the host LMG and split into multiple seams in places. They are concordant to the local fabric (038°, 27°) which is much stronger within the vicinity of mineralization. The ore is >70% Fe2O3 with slightly elevated REE levels. A thin section of a seam contained beautiful S-C fabrics indicating a reverse sense shear with top to the southeast, Further investigation into the relationship between the timing of deformation and magnetite mineralization is underway. The host LMG immediately surrounding mineralization seems to be very gneissic and contains albite, quartz, amphibole, and pyroxene. It has an Fe2O3 content of less than 7%, Na2O of more that 6.5% and REE levels slightly higher than that of the ore.
A sample of LMG was also collected roughly 100 m from the mine on the western side of Hill #8 and is elevated in REEs and contains a potassium content an order of magnitude higher than the immediate host. It also records subsolidus deformation, perhaps related to the massive magnetite deposits, however the reverse shear is top to the NW. Again, constraining timing for this is underway.

STOP 5. Old Furnace.
Old Furnace locality. Depart by turning left onto Old Furnace Road.

STOP 6. Schofield magnetite mine.
WARNING: Route 74 is a busy road with lots of heavy truck traffic. Walk single file and stay well to the side of the road.
Walk west down the road for approximately a quarter mile until you reach a brown and yellow sign marking the boundary of the Adirondack Park. Proceed into the forest for approximately 75 meters to the base of a cliff.
This stop examines the relationship between magnetite ores and the Lyon Mountain granite (LMG) that hosts the ore. Most of the main ore seam has been mined out and is covered by large blocks of mine waste. However there are places where the relationship between the ore and the host granite can be observed. The main ore seam starts here at the base of the cliff and gradually climbs up the hill to the east parallel to the main foliation in the granite. The host rock on either side of the ore seam is albitized granite and is the result of fluid alteration. The amount of albite in the rock gradually diminishes as distance from the mineralized ore zone increases and perthitic feldspar becomes common. The ore is comprised of magnetite and apatite ± quartz. The apatite is generally reddish in outcrop and is extremely enriched in light rare earth elements and high field strength elements (especially Y, Ce and Nd). The mineralogy of the “granite” varies from perthitic feldspar + quartz + magnetite ± clinopyroxene, biotite, and amphibole in the least altered rocks, to quartz + albite + magnetite ± apatite and chlorite in the most altered zones. The contact between the ore and the granite is generally sharp but a large boulder here at the base of the cliff shows numerous mm-thick layers of magnetite. Zircon grains from the ore have been dated at 1000.9 ± 9.2 (2σ) (Valley et al., 2009).
Proceed uphill along the ore seam approximately 75 meters to the east by following an animal trail and scrambling over boulders. Here the ore is perfectly exposed in a large out of place slab and is where the geochronology sample was collected. Even though a reliable age for the host LMG could not be obtained at Skiff Mountain, the age of the zircon in the ore here is at least 34 my younger than that of the granite. This is based on the age of zircon rims in the LMG that was successfully dated near here at Eagle Lake (see EAGLE LAKE STOP; Figure 13).
Continue following the ore seam uphill to the east. Watch for a small outcrop of hematized breccia. The ore seam will disappear under mine waste rock and forest. At this point head directly uphill until reaching a relatively flat area. You will see what appears to be a ditch, sometimes filled with water, backed by a 3-4 m high outcrop. This is the top of the upper ore seam. Follow this relatively flat area to the west. At 43.886056, 73.637028, the top of the upper ore seam is visible. The ore seam is approximately 0.5 m thick with a second 2-3 cm thick vein just above the main ore. The adjacent granite is typically lacking in disseminated magnetite in the vicinity of the veins, but disseminated magnetite reappears where veins are lacking and distal to the ore vein. Near the top of this outcrop small lozenge shaped quartz-feldspar-magnetite pegmatites are present (∼20-30 cm long). Follow the upper ore seam west for approximately 60 m. Here 1 m below the ore body, there are two amphibolite layers 10 to 20 cm thick. Recent geologic mapping suggests that these are mafic “screens” that were incorporated in the LMG during intrusion. Zircon grains from one of these layers have been dated at 1046 ± 11 Ma (2σ) (Valley et al., 2011) with rare zircon cores that are ∼1150 Ma. These layers and the host LMG have been overprinted by Na fluid alteration providing a maximum age for Na fluid alteration. The increase in albitization around the ore suggests Na metasomatism is coeval with zircon growth and ore mineralization. It is probable that the U/Pb age from zircon in the amphibolite layer is the product of metamorphism or possibly a fluid alteration event that is older than the fluids associated with the growth of zircon in the ore at Skiff Mountain.

Parking for Stop 7

STOP 7. The Lyon Mountain Granite Gneiss (after Postel, 1956)
From the cars, walk east about 0.3 miles to an outcrop on the right (south) side of the road. At this locality a granitic pegmatite dike crosscuts the fabric of the LMG. The LMG has been altered by K-rich fluid alteration that has been overprinted by minor Na alteration. The crosscutting dike has also experienced minor Na alteration. Both the dike and the granitic hosts have extreme concentrations of potassium for granitic rocks (∼8 wt.% for both rocks). It is not clear if the dike is high in potassium because it intruded an already metasomatized rock and thus was enriched in K2O by a relatively closed system, or if both the dike and the host were altered by potassium bearing fluids together in an open system. Both the dike and the granite experienced subsequent minor sodic alteration. The dike is comprised of coarse microcline, quartz, and biotite, with minor plagioclase, zircon, apatite, and magnetite ± clinopyroxene and muscovite. Plagioclase is secondary and is present interstitially and in “patch” perthite which crosscuts microcline grains.
A sample from the dike was collected for U/Pb zircon geochronology and was dated by Secondary Ion Mass Spectrometry (SIMS). Zircon crystals from this sample are elongate (300-500 µm long), clear with patchy zonation in BSE images, and typically contain large inherited cores of both AMCG (∼1150 Ma) and LMG (∼1060 Ma) age. The zircon rims from grains with relict cores and grains without cores have a concordant age of 1030.4 ± 1.8 Ma (2σ) (Valley et al. 2011). The aforementioned field relations imply that Na alteration has to be younger than the U-Pb zircon age of the dike. They also suggest that the forces responsible for fabric development within the LMG must have ended by this time.
A sample of the host rock was collected within in a few meters of the dike discussed above in order to constrain the timing of Skiff Mt. LMG intrusion. Zircons separated from the sample are euhedral to subhedral, stubby to elongate, typically 200-500 µm long, and are amber to brown in color.
Cathodoluminescence (CL) of polished grain cross sections reveal that most zircons have oscillatory zoned cores and prominent dark rims with subtle and limited oscillatory or patchy zoning (Figure 14). A few grains appear to possess a core, mantle, and rim. In such grains, rim CL is consistent with the rims of other grains, with the mantle and core both having oscillatory zoning. Twenty U/Pb SHRIMP-RG analyses of all zones demonstrate that the cores (or mantle, if three zones are present) are 1086 ± 34 Ma (2σ 207Pb/206Pb weighted average, 5 analyses; Figure 14), while rims cluster at 1043 ± 9 Ma (2σ) 207Pb/206Pb weighted average, 5 analyses, Figure 14). For grains with three zones, only two cores were analyzed and produced 2σ 207Pb/206Pb ages of 1169 ± 38 Ma and 1138 ± 28 Ma. These older core ages are interpreted to be inherited, while the c. 1086 Ma age is interpreted to be the magmatic crystallization age of the Skiff Mt. LMG body, with the c. 1043 Ma rim age representing the timing of fluid alteration. Though the timing of fluid alteration age is slightly older than the age of the altered pegmatite dike at the limits of error, we believe that pegmatite dike intrusion and fluid alteration were in close temporal proximity.

STOP 8. Highly deformed augen gneiss
We are in the eastern portion of the quadrangle, which is dominated by large amplitude, upright folds that contain a calculated β-axis of 10 to 112 (Figure 6). There is no identified axial planar fabric associated with this fold generation (F3). The rock exposed here is a strongly deformed augen granitic gneiss (Yggn on new 7.5’ quadrangle map). Any one exposure typically contains both pegmatitic (mesoperthitic) varieties that are weakly deformed as well as biotite-megacrystic varieties that are strongly deformed and contain >5% modal garnet (vol%). The latter unit is isoclinally folded (F2), and contains a very strong S2 here. Can we see a lineation? Is there evidence of folding at the outcrop scale (F1)?

Parking for Stop 9

STOP 9. Penfield Pond – granite pegmatite cross cutting amphibolitic gneiss
We are in the hinge of an F3 fold. We have walked through the LMG the entire way here. This outcrop contains amphibolitic gneiss, which is locally migmatitic, and a concordant lens of mesoperthite granite. The mesoperthite granite is typically associated with the augen gneiss seen at Stop 8, despite the lack of tectonic fabric within the granite. At one location in the outcrop (Fig. 5a,b), the granite cross cuts the amphibolite gneiss at a low angle. This truncation suggests that migmatization predates the mesoperthite granite, and that migmatitic layering is transposed to be nearly parallel to S2, but may have occurred during S1.

A3:CURRENT RESEARCH IN STRUCTURE, STRATIGRAPHY, AND HYDROGEOLOGY IN THE CHAMPLAIN VALLEY BELT OF WEST- CENTRAL VERMONT
Jonathan Kim
Vermont Geological Survey
Montpelier, VT 05620
Keith Klepeis
Dept. of Geology
University of Vermont
Burlington, VT 05405
Peter Ryan
Middlebury College
Middlebury, VT 05753
Edwin Romanowicz
Center for Earth and Environmental Science
SUNY at Plattsburgh
Plattsburgh, NY 12901
INTRODUCTION
Over the past four years, the Vermont Geological Survey and professors and undergraduate students from the University of Vermont, Middlebury College, and SUNY at Plattsburgh geology departments have formed a multidisciplinary fractured bedrock consortium. This consortium integrates varying expertise and resources to comprehensively address applied geologic issues in Vermont, such as groundwater quality (i.e. radionuclides, arsenic, nitrates, fluoride, and manganese), groundwater quantity of domestic and public wells, groundwater-surface water interaction, and shallow geothermal energy. The purpose of this trip is to visit field sites in the Champlain Valley Belt of west-central Vermont that illustrate our group’s current research efforts in fractured bedrock hydrogeology. At each site, we will discuss how structural geology, stratigraphy, and hydrogeology (including geophysical well logging) bear on a specific environmental issue. This trip will not only visit classic sites such as the Champlain Thrust at Lone Rock Point and the Hinesburg Thrust at Mechanicsville, where we will discuss refined structural chronologies, but also locations that exhibit a strike-slip fault zone in the Winooski River Spillway (Williston), a well-described wrench fault site in Shelburne, phosphorite layers that explain elevated radioactivity in the bedrock aquifer (Milton), and a site in Hinesburg where field mapping of fractures has been correlated with those in geophysical logs. The following Bedrock Geology of Vermont, Field Area Geology, Structural Geology, Metamorphism, and Geochronology sections are modified from Kim et al. (2011).

A3:CURRENT RESEARCH IN STRUCTURE, STRATIGRAPHY, AND HYDROGEOLOGY IN THE CHAMPLAIN VALLEY BELT OF WEST- CENTRAL VERMONT

Meeting Point - Colchester Park and Ride Lot

STOP 1. Elevated Radioactivity in Groundwater from the Clarendon Springs Formation, Milton, Vermont
Introduction Elevated radionuclide levels were reported in groundwater from ∼30% of bedrock wells tested in the Clarendon Springs Formation in the towns of Colcheser Quadrangle.
Lithology Massive gray dolostone of the Late Cambrian Clarendon Springs Formation.
Structure The Muddy Brook Thrust (MBT), which is described below, is located in the valley to the east. In this area, the dominant fracture set strikes NE, dips steeply, and is associated with cross faults that offset the MBT and lithologies on both sides (Kim and Thompson, 2001).
Tectonic/Stratigraphic Context This site is at the north end of the Hinesburg Synclinorium. The shallowly east- dipping MBT carried black phyllites with thin dolostone interlayers over the Clarendon Springs Formation dolostones during the Taconian Orogeny. The MBT is west of and probably synchronous with the Hinesburg Thrust (Kim and Thompson, 2001).
Hydrogeology and Groundwater Geochemistry Groundwater produced from the Clarendon Springs Formation in Milton and Colchester is known to contain elevated uranium, radium, and alpha radiation. Radioactivity is high enough that the average alpha radiation from 131 domestic bedrock wells tested in a 12 km2 area is 32 picocuries per liter (pCi/L), and 22 % of these wells produce concentrations of alpha radiation above the EPA’s maximum contaminant level (MCL) of 15 pCi/L; furthermore, the three most contaminated wells, which occur within a kilometer of each other, average 841 pCi/L. Dark grey to black phosphorites (with 7 to 37 % P2O5) occur throughout the region that contains elevated radionuclides in well water; in outcrop, these phosphorites occur in two main forms: (1) subangular, pebble-sized clasts in a dolostone matrix, sometimes with imbricated clasts that suggest deposition from a current, and (2) wispy beds of dark grey phosphorite (possibly “hardground”). Breccias are more common than wispy sedimentary layers, and both indicate that the concentrated phosphate is syndepositional in origin. McDonald (2012) cited a model involving upwelling sea water and reducing biochemical conditions as factors responsible for deposition of phosphorite with elevated U.
Rock samples collected from outcrops located upgradient of the most highly contaminated wells exhibited the following: (1) the phosphorites contain 80 to 430 mg/kg uranium while the dolostone matrix contains less than 10 mg/kg U; (2) the phosphorite mineral — as determined by powder XRD — is fluoroapatite, and trace amounts of autunite also occur in some phosphorites; (3) a gamma ray survey of a 160 m deep bedrock well documents the interbedding of U-rich phosphorite beds and dolostone beds throughout the Ccs in Milton-Colchester, perhaps alternating cyclically. U-rich phosphorites produce a gamma ray signal of 500 to 3400 cps whereas U-poor carbonates have a signal < 50 cps gamma radiation; and (4) SEM-EDS element maps (Bachman, 2015) show that phosphorite clasts and layers lacking significant post-depositional deformation contain the highest levels of uranium. SEM-EDS also indicates that U occurs in two broadly-defined mineralogical forms: (1) diffusely distributed U in cryptocrystalline fluoroapatite, and (2) perhaps more importantly, concentrated U in microcrystalline minerals (e.g. autunite, coffinite, brannerite) scattered throughout the phosphorite. This may suggest that uranium was initially substituted for Ca in fluoroapatite — during crystallization on the seafloor (consistent with literature reports)—, but then was incorporated into anhedral secondary minerals when (at least some of the) fluoroapatite dissolved. This could have occurred any time from very early diagenesis to dolomitization. Weathering of pyrite may trigger U release by locally lowering pH — evidence for this is the occurrence of autunite [Ca(UO2)2(PO4).nH2O] in rusty Fe hydroxide matrix at the boundary of weathered pyrite and adjacent, unweathered fluorapatite.

STOP 2. Champlain Thrust at Lone Rock Point
Introduction This stop description was modified from West et al. (2011). The Champlain thrust fault at Lone Rock Point is arguably the iconic geologic feature in Vermont and perhaps the finest thrust fault exposure in eastern North America. The “older-on-top-of-younger” relationship exposed here is a fundamental indicator of thrust faulting. Hitchcock et al. (1861) was the first to recognize that the contact relationships exposed at Lone Rock Point are the result of major regional faulting.
Lithology Early Cambrian massive dolostone of the Dunham Formation structurally overlies Late Ordovician Iberville Formation black shales (Figure 5).
Structure An interesting feature at Lone Rock Point is the interplay between brittle deformation and ductile flow mechanisms. Brittle deformation involves the breaking of material along discrete surfaces, which can be fractures, veins or, if they accommodate slip, faults. These two styles are not completely independent of one another, and commonly occur together to accommodate shortening. The material type and the conditions under which deformation occurs typically dictates the types of deformation processes. As such, the type of parent lithology (what type of rock was present prior to deformation) is a critical influence on style of deformation. At Lone Rock Point two significantly different rock types are juxtaposed with the strong Dunham dolostone thrust above the weak Iberville shale (Fig. 5A). Deformation is not restricted to slip along the fault plane, and can be dived into two domains as described by Stanley (1987). These are composed of an inner fault-zone including the fault surface at the base of the Dunham dolostone and a proximal region consisting of broken limestone and highly contorted shale. The outer fault-zone occurs in the Iberville Shale and consists of a high concentration of veins, subordinate faults, and tightly folded compositional layering. By recognizing various features within both the Iberville Shale and the Dunham dolostone we can establish various mechanisms of deformation and compare the way these two units accommodate shortening. For this reason, the exposure of the Champlain Thrust fault at Lone Rock point is an excellent location to teach fault-zone processes and the influence of material properties on deformation.
The following sections briefly describe key features that document deformation styles and the motion history of Lone Rock Point.
Structural Slickenlines. The basal surface of the Dunham Formation is the slip surface on which a significant amount of displacement has occurred. This surface contains corrugations referred to as fault mullions with wavelengths on the order of half a meter. These features, as depicted in Figure 8 can form with crest lines parallel to the motion of the fault and, therefore, can help constrain motion direction. In addition to fault mullions, striations can occur from the scraping and gouging of the fault surface by resistant objects. These also help constrain the motion of the Champlain Thrust fault (Figure 5B).
Original Bedding. Despite the highly deformed nature of the Iberville Formation, compositional layering is visible as resistant quartz rich layers. These layers are frequently folded and faulted into small isolated pods surrounded by soft clay rich rock. Although dolostones of the Dunham Formation are massively bedded, depositional surfaces are intact and relatively undisturbed (Figure 5A). Veins and Mineral Slickenlines. Veins form from the deposition of material from solutions that fill voids in rocks. The calcite veins at Lone Rock Point display complex geometries, including folding, faulting, and shearing. The shear veins, in particular, can be good indicators of the motion history of deformation and frequently form lineated surfaces known as mineral slickenlines. The formation of mineral slickenlines generally involves the infilling of void space with material created by offsets on the fault surface (Figure 5C).
Cleavage. Cleavage planes are formed from the preferential alignment of mineral grains due to flattening, which is accommodated by dissolution and the removal of soluble material. Cleavage is common in shales of the Iberville Formation, and are generally is oriented at low angles to the fault surface. However, as the distance increases away from the fault, the orientation of cleavage planes tends to steepen. This rotation of cleavage planes can be used to infer the sense of motion on the thrust surface.
Fractures. Unlike the intensely deformed Iberville Formation, massive dolostones of the Dunham Formation are relatively intact and show little evidence of internal deformation (i.e., no visible cleavage). However, fractures (breaking of the rock along discrete planes without differential motion), are common in this more competent unit. New fracture data will be presented for the hanging wall.
Subordinate Faults. Many small scale faults can be observed within the footwall rocks of the Iberville Formation that offset folds, cleavage, and veins. These faults do not continue into structurally overlying dolostones of the Dunham Formation. These small scale faults in the Iberville tend to rotate into the direction of the fault motion with proximity to the fault (Stanley, 1987). Because the Champlain thrust fault at Lone Rock Point contains numerous features that help constrain the transport direction along the fault plane, an exercise can be constructed to identify as many of the features that contain motion information (i.e. fault mullions, gouges, and mineral slickenlines). These features generally indicate displacement and tectonic transport in a west-north-west direction. Teaching considerations include the following questions: How does the geometry of folded compositional layering change with proximity to the fault? How does the orientation of the dominant cleavage change with proximity to the fault surface? Is displacement constrained to slip along the fault surface only? What is the general temporal progression of deformational style using cross-cutting relations? What fundamental influence does the type of bedrock have on deformation style? These questions address the fundamental aspects of thrust faulting and the inherent relationship between initial lithology and deformation mechanisms.
New Work on Foot Wall Structures
The following has been modified from Strathearn et al. (2015). Although the Champlain Thrust has been studied previously at Lone Rock Point, the multiple generations of ductile and brittle structure in shales of the footwall have never been systematically defined. We present the following relative chronology of structures:
1) Formation of bedding planes (S0), characterized by thin layers of carbonate within black shale.
2) Formation of rootless isoclinal folds (F1) of brittle carbonate layers and the development of an spaced pressure solution cleavage (S1) that parallels the axial planes of the folds.
3) The S1 cleavage is deformed into asymmetric S-C shear bands that merge into parallelism with, and are cut by intraformational thrusts. The thrusts form oblate, eye-shaped structures that are stacked on top of one another forming thrust duplexes. A second cleavage (S2) defines a part of the S-C fabric and is intensified in thrust zones. Calcite slickenlines on fault surfaces plunge to the SE and NW and slip directions fan up to 40 degrees with respect to one another in different thrust horses
4) Formation of sets of upright, north (F3) and east-striking (F4) folds of S2 warping the CT.
5) Formation of conjugate sets of normal faults that record top-down-to-the–north and -south kinematics.
6) Formation of the steeply-dipping fracture sets (N-S and E-W striking) that cut across competent lithologies.
Hydrogeology and Groundwater Geochemistry In a case study examined on this field trip (Hinesburg Thrust; Kim et al., 2014a), the hanging wall is sometimes responsible for producing groundwater with elevated radionuclides. In the case of the Champlain Thrust, the culprit is the footwall. Of 35 tests for fluoride (note: F can cause bone disease), 37 % (13/35) exceeded the Vermont recommended F level in public water systems (0.7 mg/L). [Relative to the EPA MCL of 4 mg/L, however, only 3/35 wells were above the F threshold]. Sodium is elevated in ∼ 45 % of footwall wells (13/29) relative to EPA's 20 mg/L Drinking Water Equivalency Level (guidance level), and the average Na concentration in footwall wells (82 mg/L) is four times greater than the DWEL. Both issues are related to the behavior of illite in this black shale-influenced bedrock aquifer system. Regarding fluoride, diagenetic illites typically contain 0.2 to 0.5 % F in isomorphous substitution for OH (Thomas et al., 1977), so dissolution of illite or exchange of Cl- or OH- for F- are potential F sources in groundwater.
Dissolution of apatite (4 % F) could also be a source, although it is far less abundant and likely less reactive than illite. Other potential F minerals (e.g. fluorite, titanite) are not likely F sources in this system. Regarding elevated Na, we observe a weak but positive correlation of Na and Ca in solution, an occurrence that likely relates to Na-Ca ion exchange. When Ca is released to solution upon weathering of calcite (nearly ubiquitous in these Ordovician black shales), the higher-charge, less-hydrated Ca+2 cation is more strongly attracted to cation exchange sites (e.g. on illite) than is Na+1, and the exchange reaction releases Na+1 into solution. This also is cited as the cause of high Na in black slate-influenced wells of the Taconics of southwestern Vermont (Ryan et al., 2013). Another element worth noting in footwall wells of the Champlain thrust is arsenic, which exceeds 10 ppm in 3 % of wells tested for As and exceeds 5 ppm in 10 % of wells (3/29); by comparison, in Taconic slates, 22 % of bedrock wells (52/236) exceed 10 ppb and 24 % exceed 5 ppb. Deeper anoxia in the Taconic seaway relative to open-shelf Iberville and Stony Point shale depositional environments is likely the cause of greater amounts of pyrite and arsenic in Taconic black slates.

STOP 3. Strike Slip Fault Zone in the Winooski River Gorge, Williston/Essex, Vermont
Introduction The Winooski River flows through a bedrock gorge downstream of a hydroelectric dam at the Essex/Williston border. During bedrock mapping in the Town of Williston (Kim et al., 2007), it was observed that most of the bedrock channels were northeast-striking fracture intensification domains (FIDs). Further investigation revealed a more complicated scenario.
Lithology Massive gray dolostone of the Late Cambrian Clarendon Springs Formation.
Structure Preliminary mapping and analysis of the brittle structures in the Winooski River Gorge reveals that the river channels are northeast- striking and steeply-dipping fracture intensification domains (FIDs) that are also strike-slip fault zones. Throughout the river gorge, these faults are intersected by steeply-dipping, north-northeast striking Riedel shear zones. Riedel shears usually form at a 10-20 degree angle to the main strike slip fault, in clockwise and counterclockwise directions, respectively, to right lateral and left lateral displacements (e.g Twiss and Moores, 2006). The Riedel shears in this gorge are consistent with left lateral displacement on the main strike-slip faults.
Tectonic/Stratigraphic Context This body of Clarendon Springs Formation is bounded to the east by the Ordovician Muddy Brook Thrust Fault and to the west by the presumed Mesozoic down-to-the-east St. George normal fault. Hydrogeology and Groundwater Geochemistry Bedrock geochemical analysis indicates some interesting differences in Ccs composition here compared to the phosphorite-bearing localities in Milton-Colchester and Highgate (McDonald, 2012). P2O5 here in the Winooski gorge is < 0.5 % whereas P2O5 ranges from to 7.4 to 36.9 % in phosphorite-rich dolostone and phosphorite layers or clasts. Uranium is also low in the Ccs in the Winooski River gorge (1.3 to 8.3 ppm) relative to Milton-Colchester (where it reaches 432 ppm). No phosphorite clasts or layers are observed in the Winooski River gorge, explaining the low amounts of P and U at this locality. Geochemically-notable are two samples of black wispy clasts coated with Fe hydroxide that contain 120 and 283 ppm As (for comparison, As < 16 ppm at Milton-Colchester outcrops). These black wispy clasts occur in a dolomitic breccia, and yet while this facies appears similar in outcrop to some of the phosphorite occurrences at Milton-Colchester, the black wispy clasts in Winooski gorge contain < 0.5 % P2O5 and < 10 ppm U. In terms of major elements, the Winooski gorge Ccs black clasts are dominantly Si, Mg and Ca; otherwise, only elevated iron (5.7 and 10.1 % Fe2O3, respectively) and arsenic (120 and 283 ppm, respectively) are anomalous relative to other dolomitic Ccs rocks. One possible interpretation: if the black color is from mature organic matter, these wispy black clasts may represent an organic-rich shallow marine or subaerial layer that was incorporated into a cave breccia. The As-bearing Fe hydroxide is likely remnants of weathered disseminated pyrite.
The municipalities in the vicinity of Winooski gorge are part of the Champlain Water District and their drinking water is supplied from Lake Champlain. We know of no chemical data for bedrock wells in this area, so it is unknown how the composition of this belt of Clarendon Springs Formation may affect groundwater.

STOP 4. Wrench Faults and Fractures at the Shelburne Boat Access, Shelburne, Vermont
Introduction This site was used by Rolfe Stanley of the University of Vermont as a teaching outcrop for his structural geology classes for decades and is currently used by Keith Klepeis. Stanley (1974) wrote a Geological Society of America Bulletin article that details the structural history of this site.
Lithology Early Cambrian ferruginous quartzite of the Monkton Formation. Stanley (1974) reported that the quartzite at this site is composed of 93% quartz, 5% hematite, 1% microcline, and traces of zircon, plagioclase, detrital staurolite, and dolomite.
Structure As outlined by Stanley (1974), there are three major structural features at this site, which are:
1) High – Angle Wrench Faults:
A) Generation 1 Northeast to west striking, moderately-steeply dipping faults (Figure 7).
B) Generation 2 North-northeast to north-northwest striking, steeply dipping faults that offset generation 1 faults (Figure 7).
2) En Echelon Fractures: Dextral and sinistral en echelon fracture arrays that are filled with quartz. The trend of the en chelon arrays corresponds to the strike of a particular fault. These arrays are only associated with the first generation of high angle faults (see fracture arrays on Figure 7).
3) Fractures: The four major fractures sets (contoured maxima) at this field site are:
A) East-west striking, steeply north-dipping,
B) North-south striking, steeply west-dipping,
C) Northwest striking, steeply southwest-dipping, and
D) Northeast striking, steeply northwest dipping.
Set A is associated with the first generation wrench faults whereas sets B, C, and D overprint set A and are associated with the second generation wrench fault generation.
Tectonic/Stratigraphic Context Stanley (1974) proposed that the east-west striking and steeply dipping wrench faults were related to the Shelburne Bay cross-fault, which offsets the Champlain Thrust. Both sets of wrench faults were presumed to be related to the Devonian Acadian Orogeny. Kim et al. (2011; 2014) described dome and basin fold patterns in the Champlain Valley Belt that were the result of the superposition of fold sets with steeply dipping ∼north-south and ∼east-west striking axial surfaces. Could the two sets of wrench faults and their associated fractures be related to these folding events?
Hydrogeology and Groundwater Geochemistry We are using this study as a detailed context for fracture analysis throughout the Champlain Valley Belt. These brittle structures may influence groundwater flow in the bedrock aquifer.

STOP 5. Mount Philo State Park in Charlotte, Vermont
Introduction This stop description, which was modified from Kim et al. (2011), is designed to give an overview of the four major tectonic zones in the Lake Champlain area of west-central Vermont and eastern New York, which are, from west to east, and from structurally lowest to highest: 1) Autochthon, 2) Parautochthon, 3) Hanging Wall of the Champlain Thrust, and 4) Hanging Wall of the Hinesburg Thrust. A major thrust fault separates the Autochthon from the Parautochthon on the west side of Lake Champlain.
Lithology At the top of Mt. Philo, you are standing on the upper member of the Monkton Formation, which is a ferruginous quartzite. The lower dolomitic sandstone member of the Monkton Formation is exposed near the Champlain Thrust below.
The autochthonous rocks on the west side of Lake Champlain include Mesoproterozoic metamorphic rocks of the Adirondacks (see Yu in Figure 3) that are uncomformably overlain by sedimentary rocks of the Beekmantown Group (Isachsen and Fisher, 1970). The Autochthon is structurally overlain by by: 1) Late Cambrian - Middle Ordovician weakly metamorphosed sedimentary rocks of the Parautochthon, 2) weakly metamorphosed sedimentary rocks of the Hanging Wall of the Champlain Thrust; and 3) low grade (chlorite-sericite to biotite) rift clastic metasedimentary rocks of the Hanging Wall of the Hinesburg Thrust.
Structure At this lookout, the Champlain Thrust is ∼160 vertical feet (50 m.) below your feet, where the dolomitic lower member of the Monkton Quartzite is in tectonic contact with the Middle Ordovician Stony Point Shale. From Mt. Philo, one can see the geomorphic expression of this thrust to the south as the slope breaks (steep = Monkton and gentler = Stony Point) on Buck and Snake mountains, and to the north on Pease Mountain.
Tectonic/Stratigraphic Context At the Lone Rock Point exposure of the Champlain Thrust in Burlington (e.g Stanley, 1987)(Stop 2), this thrust juxtaposes the Lower Cambrian Dunham Dolostone with the Iberville Shale; this suggests the presence of an along-strike ramp that climbs ∼2000’ (610 m.) up section between Burlington and Mt. Philo. The stratigraphic throw of the Champlain Thrust is ∼9000’ (2743 m.) at Lone Rock Point and ∼6000’ (1830 m.) at Mt. Philo (Stanley, 1987). Total displacement along the Champlain Thrust ranges from 34-62 miles (55-100 km.) (Stanley, 1987; Rowley, 1982).
Apatite fission track work by Roden-Tice (2000) indicates that the exhumation of the Adirondacks was complete by the Cretaceous. This uplift may have reactivated faults between the Champlain Valley and Adirondacks.
Hydrogeology and Groundwater Geochemistry Based on domestic well logs, average yields from the Monkton Formation in the Town of Charlotte are very favorable (Springston et al., 2010). Although we do not know of any bedrock groundwater wells that penetrate through the Monkton Formation into the underlying shales, Charles Welby (pers. Comm., 2009) said that numerous wells do. Such wells are often characterized by unpleasant amounts of hydrogen sulfide (rotten egg smell) gas. A Middlebury College student will be investigating the groundwater chemistry of the hanging and foot wall aquifers of the Champlain Thrust during 2015-2016.

STOP 6. Hinesburg Thrust at Mechanicsville, Vermont
Introduction Elevated radionuclide levels have been reported in groundwater from numerous bedrock wells completed in the Hanging Wall of the Hinesburg Thrust and drilled through this thrust (Kim et al., 2014a). This stop description was modified from Kim et al. (2011).
Lithology The Hinesburg Thrust at Mechanicsville places overturned light green and gray metapsammitic schists and quartzites of the Early Cambrian Cheshire Quartzite (argillaceous member) on top of a slice of deformed dolomites, limestone and marble of the Ordovician Bascom Formation (Thompson et. al., 2004). The hanging wall rocks display a lower greenschist facies metamorphic mineral assemblage that includes chlorite, quartz, sericite, and biotite. The footwall is dominated by a chloritic and graphitic carbonate mineral assemblage, indicating that the Hinesburg Thrust is marked by an abrupt change in metamorphic grade as well as a change in lithology.
Structure The thrust fault itself is defined by a narrow zone (<10 cm thick) of cataclastic dolomite at the top of the Bascom Formation (Figure 8). The dolomites below the fault are highly fractured and sheared. Ductile shear bands (C’ type) deform a penetrative phyllitic cleavage (S1) up to at least a meter beneath the fault. The shear bands are variable in orientation, but most dip gently to the northeast (298 31 NE). Quartz rods define a mineral stretching lineation on C’ shear planes. This lineation plunges gently to the east and southeast (093 02). The asymmetry of the shear bands yields a consistent top-to-the-northwest sense of displacement on the Hinesburg Thrust.
Above the Hinesburg Thrust at Stop 6 the metapsammite schists and quartzite layers of the Cheshire Fm are mylonitic (Strehle, 1985; Strehle and Stanley, 1986). The oldest structures preserved include a compositional layering (S0) defined by alternating bands of micaceous and quartz-rich laminae and >3 cm thick quartzite beds (Figure 8). This layering, which most likely represents sheared, stretched, and recrystallized bedding planes, is deformed into a series of tight-isoclinal, reclined-recumbent folds (F1). A fine grained mylonitic foliation (S1) defined by the alignment of graphite, mica, and recrystallized quartz parallels the axial planes of the F1 folds (Figure 8). On S1 surfaces, a penetrative mica and quartz mineral lineation (L1) plunges gently to the east and southeast. Locally, and especially within thin quartzite bands, the L1-S1 fabric is deformed by a series of shear bands (C’ type) similar to those in the foot wall rocks (339 13 NE). Both sets yield a similar top-to-the-northwest sense of shear parallel to the L1 mineral lineation (106 12). The S1 foliation also parallels the surface of the Hinesburg Thrust and is interpreted here to reflect early ductile thrusting at depth prior to the final emplacement of the Cheshire Fm onto the Bascom Fm along the semibrittle Hinesburg thrust fault.
The structural features observed in the mylonitic rocks above the Hinesburg Thrust display an elegant interplay between ductile deformation, in the form of folds and cleavages, and brittle deformation, in the form of veins. Throughout the outcrop the mylonitic S1 foliation locally is cross cut by a set of quartz veins (V1) that are tightly folded within the F1 folds, indicating that they formed during folding, probably as a result of pressure solution and fluid transfer processes. Cross cutting both the S1 cleavage and the V1 veins is a second set of asymmetric quartz tension gashes (V2) that localized within thick (>30 cm) quartzite layers (Figures 8, 9a). The tips of the asymmetric veins penetrate into the mylonitic schist surrounding the quartzite layers. A close inspection of the V2 veins (both on the outcrop and in thin section) indicates that they are sheared and protomylonitic (look for the milky white, recrystallized appearance and the presence of quartz ribbons). These characteristics contrast with a younger set of quartz tension gashes (V3) that cross cut the V2 set in the same quartzite layers (Figure 8). The V3 vein set is only weakly deformed, less recrystallized than the V2 veins, and are mostly symmetric to slightly asymmetric. A black, quartz-poor pressure solution selvage surrounds the veined quartzite layers (Figure 9a) strongly suggesting that the vein material was locally derived and that dissolution and fluid migration depleted these zones of silica during progressive deformation. These relationships indicate that crystal-plastic deformation alternated with brittle deformation as the superposed sets of tension gashes formed.
Both the V3 and V2 vein sets, as well as the F1 folds, S1 cleavage, and quartzite layers, are all deformed into a series of northwest-vergent asymmetric folds (F2) of variable tightness (Figure 18). The fanning of the V2 vein sets around fold hinges is a good indicator that they are folded (Figure 19c, 19d). The tightest folds are recumbent and tend to occur nearest the Hinesburg Thrust close to the base of the mylonitic section. Farther above the thrust, the F2 folds tend to be more open and upright to gently inclined. This increase in fold tightness and orientation suggests that the folds record an increase in finite strain downward toward the Hinesburg Thrust.
A spaced crenulation cleavage (S2) parallels the axial planes of the F2 folds and also displays variable dips (Figures 8, 9d). In addition to recording a strain gradient, the variability in axial plane and cleavage orientation with increasing fold tightness provides kinematic information. The rotation of fold axial planes and S2 to the northwest as fold tightness (and finite strain) increases, yields a top-to-the-northwest sense of shear identical to that indicated by the shear bands (Figure 9d). This relationship indicates that the F2 folds reflect progessive deformation during the same ductile thrusting event that produced the S1 mylonites and F1 folds.
The following model, which is based on sketches of features at Mechanicsville, explains the evolution of the veins and the F2 fold structures. See if you can find features on the outcrop that record each of these stages:
Stage 1 (Figure 9a).
En echelon arrays of quartz veins (V2) open perpendicular to the direction of maximum stretch (X) of the instantaneous strain ellipse (ISE) in quartzite layers.
Stage 2 (Figure 9b)
After the V2 veins finish forming, noncoaxial shear zones localized by the rheological contrast between the shale and the quartzite causes the parts of the vein tips that extend into shale to deform and rotate to the left. This process causes the veins to become asymmetric. A new set of veins (V3) open perpendicular to X-direction of instantaneous strain ellipse (ISE). A comparison of instantaneous and finite strain ellipses and the asymmetry of the two vein sets yield a top-to-the-NW sense of shear, identical to that recorded by shear bands in the mylonitic matrix. Note that this process differs than that which forms sigmoidal veins in brittle shear zones where the veins continue to open during shearing (Figure 10). In this latter model, a comparison of instantaneous and finite strain ellipses yields a top-to- the-SE (normal) sense of shear. This is because, in this latter case, the tips of V2 veins are younger than their centers and so the former record instantaneous strains (ISE, Figure 10) and the latter record finite strains (FSE, Figure 10). We ruled out this model because, given the top-to-the-NW sense of shear at Mechanicsville, it would produce V2 and V3 vein asymmetries opposite to those observed (Figure 19). In the Mechanicsville model, the vein is required to form quickly and finish opening before ductile shear begins, yielding an asymmetry similar to that of a shear band.
Stage 3 (Figure 9c)
As the rotation of the veins during noncoaxial shear continues, the F2 folds begin to form along with an axial planar crenulation cleavage (S2). The F2 axial planes and S2 cleavage initially form at 45° to the quartzite layers and then rotate to the northwest toward the shear plane (defined by S1). The V3 vein sets also begin to rotate toward the northwest as noncoaxial shear continues.
Stage 4 (Figure 9d)
As noncoaxial shear continues the F2 folds continue to rotate and tighten, recording a progressive increase in finite strain durign ductile thrusting. The S2 crenulation cleavage rotates to the northwest along with the folds. The V2 veins exhibit a characteristic fanning geometry around fold hinges, indicating that they also rotated during folding. The F1 and S1 structures are transposed parallel to S2. The rotation of F2 axial planes to the left with increasing fold tightness yields a top-to-the-northwest sense of shear, which is the same as that indicated by the asymmetric veins and shear bands.
The Hinesburg Thrust surface, and all other structures above and below it, are corrugated by two orthogonal sets of gentle, upright folds, forming a dome and basin pattern with a wavelength on the order of a few meters (Figure 10). This fold geometry mimics a kilometer scale dome and basin interference pattern formed by north-plunging (F3) and east and west plunging (F4) folds across the field trip area. These orthogonal fold sets are among the youngest ductile structures preserved at Stop 6. On the thrust surface itself two orthogonal crenulation lineations (L3, L4) mark the presence of the corrugation folds (Figure 11). A regional correlation of similar structures across the field area indicates that the orthogonal fold sets everywhere postdate thrust sheet emplacement and imbrication on the Champlain, Hinesburg and Iroquois thrusts. Earle et al. (2010) suggested that the two folds sets formed together as a result of a constrictional style of deformation during the Acadian orogeny, possible reflecting the reactivation of inherited basement faults or lateral thrust ramps. However, it is also possible that the fold sets formed in sequence as separate events.
Hydrogeology and Groundwater Geochemistry Elevated levels of alpha radiation (> 15 pCi/L as gross alpha) were observed in 38 % (12/31) of bedrock wells drilled into the hanging wall, including wells that penetrated the thrust into carbonates below; for comparison, no wells [0/21] in the carbonate-dominated footwall west of the thrust front exceeded 15 pCi/L (the EPA MCL). The source of the elevated radioactivity was evaluated by testing groundwater from hanging wall and footwall bedrock wells and by analyzing compositions of bedrock from local outcrops. The chemical composition of groundwater in hanging wall and footwall aquifers is mainly controlled by whole-rock geochemistry. Hanging wall groundwater is enriched in alpha radiation, K, Cl, Ba, Sr and U, whereas footwall groundwater is enriched in Ca, Mg, and HCO3. These signatures reflect the distinct compositions of phyllite-dominated bedrock in the hanging wall compared to Ca-Mg-CO3-rich limestones and dolostones of the footwall. Elevated alpha radiation and U in wells that produce from footwall carbonates below the thrust recorded compositions that are intermediate to hanging wall and footwall end-members, indicating that alpha and U are transported in groundwater downward through the thrust via fractures into the footwall below (Kim et al., 2014a).

STOP 7. Multiple Structural Generations and Hydrogeology at Geprags Park, Hinesburg, Vermont
Introduction The Town of Hinesburg drilled three bedrock wells in Geprags Park in 1996. Although these wells were completed as future public water supplies, surface water contamination in one of these wells precluded them from ever being used. The Vermont Geological Survey and SUNY at Plattsburgh conducted comprehensive geophysical logging of these wells during 2012-2014 using temperature, conductivity, gamma, caliper, and accoustic televiewer tools. During 2014, Hinesburg drilled two new wells at a nearby site, which is also located on Shelburne Falls Road, but closer to the intersection with Route 116. These wells were drilled to increase water production for the town. One of these new wells (#48477) was logged with the previously mentioned tools. One well in Geprags Park (#7797) and another on the Wainer property were logged using a heat pulse flowmeter. Structures and lithologies were also mapped in the field to compare with those in the wells.
Lithology In order of decreasing age, Early Ordovician Shelburne Formation (marble and dolostone), Late Cambrian Clarendon Springs Formation (dolostone) and Danby Formation (dolomitic sandstone), and Middle Cambrian Winooski Formation (dolostone).
Structure All lithologies are folded by the north-plunging Hinesburg Synclinorium, which is thought to have formed during the Ordovician Taconian Orogeny. In the western half of Geprags Park, this synclinorium is truncated by the north striking and steeply east-dipping St. George Fault, a down to the east normal fault of presumed Cretaceous age (Figure 2). The dominant fracture set in this area is axial planar to parasitic folds in the Hinesburg Synclinorium (Figure 12). Isolated east-west trending fracture zones also occur (Figure 13). The integrated geophysical logs for Geprags Park well #7797 showed that this well was completed through the St. George Fault.
Tectonic/Stratigraphic Context This site sits in a similar tectonic context to Stop 3, in the hanging wall of the Champlain Thrust just east of the St. George Fault.
Hydrogeology and Groundwater Geochemistry Using temperature, conductivity, caliper, and acoustic televiewer logs for the four wells, we were able to delineate water producing zones. A summary cross section for Geprags Park only was drawn by Kim et al. (2014b) In addition, the attitude of structures intersecting each well were calculated from the acoustic televiewer logs. Using structural data from bedding and fractures acquired in the field as a template, we were able to accurately categorize the borehole structures. Finally, we were able to determine which structures were producing groundwater in each well (Kim et al, (2015). All of our data has been presented to the Consulting Hydrogeologist for the Town of Hinesburg.

A4: ANATOMY OF AN ORDOVICIAN VOLCANIC ARC AND FOREARC BASIN IN THE SOUTHERN QUÉBEC APPALACHIANS
Alain Tremblay
Département des Sciences de la Terre et de l’Atmosphère
Université du Québec à Montréal
Montréal, Qc, H3C 3P8
INTRODUCTION
This field trip focuses on the lithological characteristics and structural features of an Ordovician forearc basin sequence, the Magog Group, and adjacent volcanic arc, the Ascot Complex, from the oceanic domain (Dunnage Zone) of the southern Québec Appalachians. A particular emphasis will be put on stratigraphic relations within and between both units, which will serve as a template in order to better understand volcanic and sedimentation processes that has been coeval with the Taconian orogeny. We will visit key outcrops in the Sherbrooke area of southern Québec.
The fieldtrip incorporates numerous discussions with scientists involved in the geology of southern Québec over the last 30 years. I wish to express my gratitude to the late P. St-Julien for introducing me to the area and sharing his knowledge during earlier field seasons, P. Cousineau for his sedimentological expertise on the Magog Group, and to the southern Québec Appalachians team of graduate students (B. Lafrance, S. Castonguay, S. Schroetter, P. Pagé, C. Daoust, S. de Souza, P.-E. Mercier and M. Perrot) for their passion for field geology. It must be remembered, however, that this document represents my own vision of the southern Québec Appalachians, and that other interpretations exist. I hope that this short fieldtrip will be a forum for discussion of volcanic arc and related sedimentary basins genesis along the margins of active collision zones.

A4: ANATOMY OF AN ORDOVICIAN VOLCANIC ARC AND FOREARC BASIN IN THE SOUTHERN QUÉBEC APPALACHIANS

STOP 1. The uppermost fine-grained sequence of the Magog Group.
Location: Take Jean-Paul-Perrault Street, and drive to HGW 410N. Follow HGW 410N for approximately 1 km and take the exit to HGW 10W (toward Montreal). Drive for ca. 7 km and take Exit 133 to St-Rock Road. Turn left on St-Rock Road, and then right on the exit back to HGW 10E. Park along the access ramp, the outcrop is the roadcut along both sides of the ramp.
Field description: This outcrop exposes a well-bedded sequence of siltstone and black slate that are typical of the uppermost stata of the St-Victor Formation in southern Québec. Relationships between the bedding and the regional schistosity indicate that this exposure is located along the hinge of a major regional fold that is the St-Victor synclinorium.
The source of the St-Victor Formation is believed to be siliciclastic and fine-grained felsic volcanic rocks (Cousineau, 1990). Detrital chromite and serpentinite indicate an ultramafic component in the source area (St-Julien, 1987). Cousineau (1990) concluded that this deposit originated from two sources: an uplifted accretionary prism (the St-Daniel mélange) to the West, and a volcanic arc (the Ascot Complex) to the East.

STOP 2. The conglomeratic marker horizon.
Location: Drive back to HGW 10E toward Sherbrooke. Follow the highway for ca. 8 km and take Exit 141 for Monseigneur-Fortier Blvd (next exit after HGW 410S). Turn left on Monseigneur-Fortier Blvd, right on Lionel-Groulx Blvd and then right again on Arnold-Pryce Road. Drive approximately 100 metres and park (there is a private road to the right). The outcrop is a roadcut on your right along the bike path.
Field description: This outcrop is a typical exposure of conglomerate and arkosic sandstone that make a stratigraphic marker horizon in the Sherbrooke area (Figure 6). The conglomeratic facies constitutes ca. 50% of this unit. It contains cm- to dc-sized, subangular to well-rounded clasts and blocks of volcanic, granitic and sedimentary rocks. Some granitic blocs can be up to 1 metre in diameter (Figure 7). Volcanic rocks fragments are predominant. This conglomerate forms a metres-thick lenticular horizon within feldspathic sandstones. The conglomerate and sandstone belong to the former lower Silurian (?) Sherbrooke Formation which was interpreted as a basal unconformity by St-Julien (1963), but has been since re-interpreted and included into the Magog Group by St-Julien et al. (1972) and St-Julien and Hubert (1975). Preliminary results of a regional U-Pb detrital zircon dating study suggest, however, that it effectively belongs to the Silurian (Perrot et al., 2015; Perrot, in progress).

STOP 3. Typical sandstone-rich turbidites of the Magog Group
Location: Drive back to HGW 10E (55N) and follow the exit for HGW 610E. Drive ∼ 2 km and take Exit 3, right after crossing the river, to St-François Street. Turn left. There is a parking lot at ca. 100 metres, park there. The outcrop is a huge roadcut along the access ramp to HGW 610E.
Field description: This is an exposure of the typical sandstone-rich turbidites of the Magog Group. It comprises interbedded quartzo-feldpathic sandstone, siltstone and mudslate. Sandstone beds are characterized by various primary structures such as graded-bedding, cross-laminations and ripples. Intraformational breccias are locally visible. The outcrop corresponds to the hinge of an overturned anticline (Figure 8). Fold-related faults are found and the presence of bedding-parallel shear zones indicates that flexural-slip folding was dominant.
This outcrop is at the same stratigraphic level than the 09SV02 sample of de Souza et al. (2014), which yielded the following age results by U-Pb zircon dating: 41% of the zircons yielded two Ordovician ages (480 ± 12 Ma; 459 ± 23 Ma) and 31% yielded Silurian to Early Devonian ages (Figure 9D). Nine of these analyses yielded Silurian-Early Devonian concordant ages ca. 90 percent to 105 percent, determining an age cluster from which was calculated a weighted mean 238U/206Pb age of 424 ± 6 Ma (Figure 9D).

STOP 4. The lowermost sequence of the Magog Group
And the contact with the Ascot Complex phyllites.
Location: Drive back to HGW 610E and follow it until the intersection with Road 112 (approximately 7 km). At the intersection with Road 112, turn around the traffic circle and drive back HGW 610 toward the West. Drive approximately 800 metres and park (there is a service road on your right). The outcrop consists in a series of roadcuts along the highway.
Field description: In the Sherbrooke area, the lower most stratigraphic unit of the Magog Group is made up of volcanic conglomerate and sandstone. On the basis of regional mapping, it was interpreted as unconformably overlying the sedimentary rocks of the Ascot Complex (Tremblay, 1992b). This outcrop is a new roadcut that exposes this contact. We will walk the section from East to West (from the Ascot Complex phyllites toward the contact with the Magog Group).
The phyllites of the Ascot Complex consist here of black and rusty argillite with occasional but typical, cm-sized nodules of pyrite. The deformation of these rocks is complex; three generations of folds and fabrics can be seen. The contact with the Magog Group corresponds to a diffuse zone, approximately 1 metre-wide, in which well-bedded sandstone strata (typical of the Magog Group) progressively appear. The main point to examine here is the deformational contrast between the Ascot phyllites and the Magog Group strata: is there a missing structural fabric in the latter?
Stratigraphically higher (westward) along the section, the outcrop exposes a volcanic, polygenic conglomerate (Figure 10) which constitutes more than 50% of the basal unit of the Magog Group in the Sherbrooke area (Tremblay, 1992b). Cm-scale clasts consist of felsic volcanic and pyroclastic rocks, less commonly of pelite, granite and siltstone. Regionally, this sequence is interlayered with quartz-feldspar sandstone and lithic tuff. In the Stoke Mountains area (approximately 10 km to the northeast), this same sequence unconformably overlies the felsic volcanic rocks of the Stoke domain (Mercier, 2013), and yielded preliminary U-Pb detrital zircon ages of ca. 435 Ma (Perrot, in progress).

STOP 5. The Ascot Complex
Interlayered felsic and mafic lavas of the Stoke domain.
Location: Continue on HGW 610W and take Exit 7 to 12e Avenue. Turn left onto 12e Avenue and then back onto HGW 610E. Drive the highway until Road 112. Turn right on Road 112, and then left on Gastin Street after 150-200 metres. Turn left on Bibeau Road until the end (ca. 300 metres) and park. The outcrop is an abandoned access ramp, ca. 100 metres to the east of the parking lot.
Field description: This outcrop nicely exposes a series of felsic and mafic volcanic rocks and schists of the Stoke domain (Figure 6). In terms of REE profiles (Figure 11), the felsic volcanics are very homogeneous. They are LREE-enriched and show a typical negative Eu anomaly that is attributed to the fractionation of plagioclase (Tremblay et al., 1989a). The mafic rocks are tholeiitic basalts with a typical MORB-type REE composition (Figure 11).
On the outcrop, the volcanic rocks are strongly sheared and hydrothermalized. The northern wall of the ramp exposes a ca. 5-10 metres-wide ductile shear zone marked by sericite-rich ≪paper≫ schist on the hanging wall, and chlorite-carbonate laminated schist on the footwall. 40Ar/39Ar muscovite dating of a felsic schist in the vicinity of this outcrop (along road 122) yielded plateau ages of ca. 379 Ma (Figure 12), which corresponds to the age of regional metamorphism (Tremblay et al., 2000).

STOP 6. The Ascot Complex – the granitic synvolcanic intrusion.
Location: Drive back to Road 112 and turn right. Drive for 1.2 km and park on your right.
Field description: The volcanic rocks of the Stoke domain are intruded by a granitic massif, the Ascot Complex pluton, interpreted as a plutonic equivalent of the extrusive sequence (Tremblay et al., 1994). The Ascot Complex pluton is a well-foliated granitic rock showing the same deformational style and fabrics as the adjacent volcanic rocks. It is a coarse-grained rock (4-7 mm in diameter) of deformed crystals of quartz, plagioclase, K-feldspar, chlorite, muscovite and epidote with minor amounts of biotite, sphene, zircon, calcite and pyrite. K-feldspar is only locally visible. Granophyric texture is common. Chlorite/muscovite ratios are variable and are mostly the product of metamorphic recrystallization and alteration of primary minerals.
In a comparative geochemical study between granitic boulders found in conglomerates (Stop 2) of the Magog Group and the Ascot Complex intrusive rocks, Tremblay et al. (1994) have shown that both are quite similar (Figure 13), and that their petrographical and geochemical characteristics suggest that both types of granitoids are peraluminous, anatectic granitic magmatic pulses originating from destructive plate magmatism.

STOP 7. Hanging wall and footwall of the La Guadeloupe faul
Sheared metasedimentary rocks.
Location: Continue on Road 112 for approximately 2.5 km and turn right on Biron Road. Drive it for 2.5 km and park at the intersection with Provost Street; stop 7 is a roadcut at that intersection.
Field relations: In the Stoke Mountains area, the progressive development of granitic mylonites related to the La Guadeloupe fault is nicely exposed in the Big Hollow brook section (Figure 14; Tremblay et al., 2000), which has been used as a type section for the 40Ar/39Ar dating of the regional Acadian metamorphism. Isotopic analyses of sheared rocks of the Big Hollow section clearly demonstrated that Acadian peak metamorphism in the southern Québec Appalachians can be tightly constrained in the 380-375 Ma time range (Figure 15).
This stop illustrates the intensity of shear deformation that can be observed in metasedimentary in the hanging wall and footwall, respectively, of the La Guadeloupe fault.

STOP 8. Hanging wall and footwall of the La Guadeloupe fault
Sheared granitic rocks.
Location: Drive another 800-900 metres along Biron Road and turn right onto Mont-Blanc Street. The outcrop is on your right approximately 100 metres from the intersection with Biron Road.
Field relations: In the Stoke Mountains area, the progressive development of granitic mylonites related to the La Guadeloupe fault is nicely exposed in the Big Hollow brook section (Figure 14; Tremblay et al., 2000), which has been used as a type section for the 40Ar/39Ar dating of the regional Acadian metamorphism. Isotopic analyses of sheared rocks of the Big Hollow section clearly demonstrated that Acadian peak metamorphism in the southern Québec Appalachians can be tightly constrained in the 380-375 Ma time range (Figure 15).
This stop illustrates the intensity of shear deformation that can be observed in granitic rocks in the hanging wall and footwall, respectively, of the La Guadeloupe fault.

STOP 9. Contact between felsic volcanic rocks of the Eustis domain and black phyllites.
Location: Follow Moulton Hill road until the intersection with College Street (Road 108), just after another bridge over St-François River. Turn right onto College Street and then (downtown) left on Road 143. Drive for approximately 1.5 km and park (there is a small dirt road right after the second traffic circle for safe parking). Walk back to the traffic circle, the outcrop is a huge roadcut on the left.
Field description: This is a new outcrop that exposes the northwestern contact between the volcanic rocks of the Eustis domain and the laminated phyllites of the Ascot Complex. Based on constrasting lithological facies and on geochemical characteristics of the volcanic rocks, the contact between the different domains of the Ascot Complex and the surrounding sedimentary rocks has been originally interpreted as faults (Tremblay, 1992b). However, such a contact is superbly exposed on this outcrop (Figure 16) and, obviously, it does not correspond to a tectonic contact but is rather depositional, and probably represents a disconformity between a ≪basement≫ made up of an heterogeneous series of volcanic rocks (i.e. the domains of the Ascot Complex) and a sequence of manganiferous, fine-grained siltstone grading up into deep-marine mudstone representing an in situ seafloor sedimentation. If correct, such an interpretation raises important questions regarding the origin and the accretionary history of the Ascot Complex.
On this outcrop, cm-thick layers of brown-colored siltstone and sandstone, and more rarely, limestone are visible in the phyllites. Locally preserved cross-laminations and graded-bedding structures indicate that stratigraphy is overturned, and that the phyllites overly the volcanic rocks. Fabrics and structures visible in the phyllites are quite complex and three generations of folding are visible (Figure 17). This structural complexity has been attributed to the proximity of the La Guadeloupe fault (Tremblay and St-Julien, 1990; see next Stop) which is located less than a hundred metres to the SE.

STOP 10. Quartz-sericite schist of the Eustis domain.
Location: Continue on Road 143 for approximately 5 km and turn right onto Road 108 (Capelton Road). After ca. 1.5 km, turn right on Du Boisé Street. Drive ca. 100 metres and park.
Field description: The felsic pyroclastic rocks of the Eustis domain are lithologically and chemically similar to those of the Sherbrooke domain, suggesting that both domains are possible lateral equivalents. D2 deformation is however stronger in the Eustis domain than elsewhere. S2 is there a mylonitic foliation, axial-planar to F2 folds which are crosscut by anastamosing ductile genetically related to the La Guadeloupe fault (Tremblay and St-Julien, 1990). This outcrops is structurally located a few tens of metres in the footwall of the La Guadeloupe fault and exposes a typical example of quartz-sericite schist of the Eustis domain. Note the intensity of structural fabric development and the down-dip plunges of the mineral lineation.

A5: SEDIMENTOLOGY AND STRATIGRAPHY OF THE CAMBRIAN-ORDOVICIAN POTSDAM GROUP (ALTONA, AUSABLE AND KEESEVILLE FORMATIONS), NORTHEASTERN NY
Dave Lowe
Department of Earth Sciences,
University of Ottawa
Ottawa, ON, Canada
Ryan Brink, Charlotte Mehrtens
Department of Geology
University of Vermont
Burlington 05405
Geological Setting
The Cambrian to Lower Ordovician Potsdam Group is the lowermost unit of the Paleozoic platform succession overlying rocks of the 1 – 1.5 Ga Grenville orogen within the Ottawa Embayment and Quebec Basin, two semi-connected fault-bounded basins that form an inboard salient of the St. Lawrence platform succession in eastern Ontario, northern New York State and western Quebec, over an area of approximately 26,000 km2 (Figure 1). Potsdam strata was deposited from the Late Early Cambrian until the Early Ordovician at subequatorial latitudes between 10° – 30° south approximately 200 – 400 km inboard of the south-facing Laurentian margin, during the late postrift and subsequent passive margin phase associated with the rifting and breakup of Rodinia (Torsvik et al. 1996; McCausland et al. 2007, 2011; Lavoie, 2008; Allen et al., 2009; Figure 2). It is mainly a siliciclastic unit that, in general, exhibits an upwards progression from terrestrial to marine deposits (Sanford and Arnott, 2010; with some exceptions, including the Altona Formation and Riviere Aux Outardes Member, see below). This terrestrial to marine progression is shared by stratal equivalents of the Potsdam Group on the Laurentian margin and in other North American cratonic basins that collectively form at the base of the Sauk Megasequence, a transgressive Megasequence that records eustatic rise across much of Early Paleozoic Laurentia preceding the Taconic Orogeny (Sloss, 1963; Lavoie, 2008).
There is no consensus from the existing Potsdam literature to suggest whether the Ottawa Embayment and Quebec Basin were pre-existing basins that were passively filled during the Early Paleozoic or were active rift or tectonically-reactivated basins. Also poorly understood is the relationship between these basins to structures within the rift margin to the east. However, features of the Potsdam including its highly irregular isopach distributions, presence of internal unconformities, variable provenance, localized soft-sediment structural deformation, in-situ brecciation, and debris flows adjacent to faults all suggest that tectonism played a role in Potsdam sedimentation, probably by the reactivation of inherited zones of tectonic weakness and pre-existing faults that developed in the Proterozoic (e.g. Wiesnet and Clark, 1966; Wolf and Dalrymple, 1984; Landing et al., 2009; Sanford and Arnott, 2010). In fact, the Ottawa Embayment and Quebec Basin lie in an area of prolonged rifting referred to as the Ottawa-Bonnechere Graben, a probable aulocogen that became active in the Neoproterozoic and was reactivated episodically until the Mesozoic and even until more recent times (Kay, 1942; Kumarapeli and Saull 1966; Kamo et al. 1995; Roden-Tice et al. 1999; Aylesworth et al. 2000; Malka et al. 2000; Bleeker et al. 2011); though the exact timing of episodic tectonic events is poorly constrained. Also poorly understood from existing literature is the approximate configuration of the basins and adjacent highlands during the Cambrian-Ordovician and the original extent of Cambrian-Ordovician strata due to post-Ordovician faulting, Mesozoic hotspot migration and unroofing of the Adirondack Dome (Kay, 1942; Fisher, 1968; Isachsen, 1975; Crough 1981; Roden-Tice et al. 1999).
The Potsdam Group is made up of four formations: the Altona, Ausable, Hannawa Falls and Keeseville Formations (Figure 3), which are described in more detail below. This trip will focus on the Potsdam strata in from the southern Quebec Basin and westward along the axis of the Oka-Beauharnois Arch that separates the Quebec Basin and Ottawa Embayment. In this area the Altona, Ausable and Keeseville Formations are present and exposed (Figure 4).

A5: SEDIMENTOLOGY AND STRATIGRAPHY OF THE CAMBRIAN-ORDOVICIAN POTSDAM GROUP (ALTONA, AUSABLE AND KEESEVILLE FORMATIONS), NORTHEASTERN NY

STOP 1. Murtagh Hill, NY: The basal Altona Formation
Walk up the dirt road until power lines are visible ahead and then veer into the woods on the left. BE CAREFUL OF SYRUP PIPING AND SUPPORTING GUIDE WIRES!! Proceed ∼25’ to the small stream and walk upstream to Stop 1. NO HAMMER OUTCROP (there is plenty of float available to sample).
This outcrop (Figure 10) exposes the stratigraphically lowest horizons of the Altona Formation in the region. The 3.5 meter thick exposure of coarse-grained, poorly sorted arkosic sandstones of Lithofacies 6 can be observed in close proximity to Precambrian basement rocks. Beds have non-planar (erosional?) bases and may exhibit graded bedding. Planar cross bedding can be present.
Based on grain size, graded bedding and the stratigraphic position of this lithofacies above basement and below arenaceous dolostones it is interpreted as recording sheetflood deposition into the marginal marine setting.

STOP 2A. NO HAMMER!
The first stop exposes Lithofacies 1 of the Altona Formation, between 22 to 24 meters in the Altona composite section (Figure 11). Although covered in moss and lichen, patient viewing will reveal a variety of ripple morphologies that are very characteristic of much of the Altona sandstones. Lithofacies 1 is a very fine to medium grained, moderately well-sorted sandstone containing a variety of sedimentary structures including upward bundling ripple cross laminations, combined flow ripple cross laminations, and swaley cross stratification (Figure 9). Along with these structures, planar cross stratification and dwelling burrows have been identified. Planar cross stratification is up to 1 m thick, pinches out laterally over 10 m, and is generally topped by 2D or 3D weakly asymmetrical ripple cross lamination. Swaley cross stratification up to 2 m thick passes into upward bundling cosets of ripple cross laminations < 0.5 m thick. Individual ripples consist of both form concordant and form discordant morphologies with wavelengths of up to 10 cm and amplitudes < 8 cm. Sandstones are dolomitized, reacting weakly to dilute HCl.
Lithofacies 1 is interpreted to represent sediment that accumulated on a wave-dominated upper to middle shoreface that was periodically inundated by storms. This interpretation is based on the combination of sedimentary structures including upward bundling ripples, swaley cross stratification, combined flow ripple cross lamination and planar cross stratification.
Walk upriver and cross the small access stream before ascending up to the field (while it is possible to walk in the shallow channel and view the stratigraphy, for efficiency with a large group we will ascend up to the field, walk northward and descend to the river at various points. WARNING: old, difficult to see barbed wire is present in the woods. Be careful!).

STOP 2B. Mudstone
This portion of the Altona stratigraphy, between 62-65 meters in the Altona composite section (Figure 12), exposes Lithofacies 4, a heterolithic unit dominated by mudstone but containing sandstone, siltstone, and dolostone. Note the interbedded nature of the sandstones and mudstones as well as the sedimentary structures found within the sandstones. Mudstones are red and grey in color with mm scale laminations. Trace fossils are present but rare. Sandstones are poorly sorted, fine to coarse-grained, and contain planar cross beds and ripple cross lamination. Slightly thicker and finer-grained sandstone beds contain structures resembling hummocky cross stratification (HCS). Planar cross beds have erosional bases, are occasionally graded, and change thickness laterally over ∼50 meters from 4 to 30 cm. Ripple cross laminations are present with cosets less than 10 cm thick and are either form concordant or discordant. Siltstones contain 1 cm thick ripple cross lamination and locally there are 1–2 cm thick lenses or nodules of dolostone.
Lithofacies 4 is interpreted to record deposition of offshore muds near the transition from the lower shoreface to offshore zones. The intercalated beds of sand are interpreted as tempestite deposits; episodically high energy deposition in an overall lower energy, further offshore setting. Transitions between lithologies are abrupt; there are no mud-rich dolostones, sandstones or muddy siltstones that would be expected on a graded shelf. This suggests the possibility that the deposition of mud represents a salinity controlled discharge-related event from hypopycnal flow.

STOP 2C. Dolostone bench
A prominent dolostone bench forms a small waterfall between 67-69 meters in the Altona composite section (Figure 13). Dolostone beds may contain disseminated sand-sized clastic material but the majority of beds are carbonate. Thin section analysis has not revealed fossils, sedimentary structures or fabrics that can be used to interpret paleoenvironments as there is evidence of multiple dolomitization events. The carbonate material may be shallow water (photic zone) algal in origin or it may be a chemical precipitate (“whitings”). The shoreline setting suggested for the clastic lithofacies of the Altona suggests that the near shore clastics may have passed offshore into more clear water carbonate ramp deposits. During sea level highstands, clastic material would be trapped in terrigenous fluvial systems which allowed carbonate deposition to commence. The interbedding of sandstone/mudstone and dolostone would record discharge events rather than sea level fluctuations. Dolostone beds containing clastic material frequently exhibit graded bedding of the sand, which may represent punctuated mixing during storm events. Dolostones containing fine sand disseminated throughout the carbonate may record eolian input of clastics.

STOP 2D. Poorly exposed uppermost Altona
This portion of the Altona stratigraphy, 81-84 meters in the Altona composite section, is poorly exposed but includes Lithofacies 5, a poorly sorted, feldspathic fine to medium-grained sandstone (Figure 14). 50 cm thick beds appear to be graded with faint ripple cross laminations. Due to its graded nature, stratigraphic position interbedded with dolostones, and stratigraphic position (overlain by Lithofacies 6), this lithofacies is interpreted to have been deposited under conditions of waning flow velocities associated with discharge events in the adjacent terrestrial fluvial system (Ausable Fm.). This lithofacies is interpreted to represent “event beds”: siliciclastic sediment washed onto the carbonate ramp following a runoff event (storm). This lithofacies may represent the nearshore component of the event beds represented by the graded planar cross bedded sandstones of Lithofacies 4.
Ascend up to the fields and walk south back to the vehicles, which need to turn around to retrace the route back to the Military Turnpike.

STOP 3. Great Chazy River, North Branch Near Ellenburg: The Ausable Formation
Park at the angler’s parking along Cashman Road. Walk south almost to the intersection of Cashman Road and Rt 190 (Star Road) and work your way down to the river. The section is continuous on the east side of the river from beneath the bridge southwards for some distance upriver. The section at this stop is along a river and waterproof footwear may be needed, depending on the current water level. Watch your step, as rocks are strewn across the riverbed and may be slippery.
This stop showcases the lithofacies and architectures of braided fluvial strata of the Middle Cambrian Ausable Formation, stratigraphically beneath the Keeseville but above the Altona. We are now on the foot wall of the east-dipping normal Havelock fault upon which lower potions of the Potsdam succession are exposed on the axis of the Oka-Beauharnois Arch (Figures 1 and 4), but nevertheless the Altona –Ausable contact is probably several hundred meters in the subsurface at this location. Strata here records deposition of pebbly arkose by braided rivers in a humid climate setting. The braided rivers flowed roughly SE and had nearby hinterlands 70 – 150 km to the NW. These rivers drained into a ∼N-S trending sub-basin that covered ca. ∼10,000 - 15,000 km2 in this region, resulting in the deposition of ∼300 – 400 m of braided fluvial arkose that occurs from Montreal south to Plattsburgh (Figure 6a).
Most of the section consists of ∼5 – 15 cm thick cross-stratified sets of moderately sorted upper medium to very coarse-grained pebbly sandstone formed by 2D or 3D dunes in shallow high energy steady unidirectional currents; although thicker (∼15 – 45 cm) planar-tabular sets are interpreted to have been deposited by simple downstream-accreting solitary bars (i.e. unit bars, see Bridge, 2003; Reesink and Bridge, 2011) (Figure 15). Architectural analysis reveals that dunes and unit bars coalesced to form larger-scale accretional elements (bounded by the red lines in Figure 15) where in which sets and cosets bounded by accretion surfaces dip at low angles in the downstream direction. These accretional elements are interpreted as the deposits of low relief compound braid bars in shallow braided rivers, ubiquitous components of modern and ancient braided fluvial environments. Other features include 4 – 10 cm thick pebble conglomerates that record energetic bar-top deposition and ∼ 1.5 – 6 cm thick argillaceous fine-grained sandstone beds and associated local rip-up clasts. In thin section the latter beds are shown to contain preferentially degraded feldspars, swollen biotite and eluvial clay and silt (Figure 16) recording preferential weathering (mostly by hydrolysis) and illuviation, and are therefore interpreted as overbank floodplain deposits.
Relatively thick (∼1.3 – 2.5 m) repetitive successions usually capped by thin overbank deposits are recognized here and elsewhere (bounded by heavy black lines in Figure 15) and interpreted as channel belt successions. At this outcrop the at least one entire channel belt succession is exposed, over- and underlain by parts of others.

STOP 4A. High Falls viewpoint
This very thick section exposes a sharp but conformable contact between two distinct parts of the Upper Cambrian quartz arenitic Keeseville Formation (Figure 17). The lower part was deposited in a coarse-grained sandy braided fluvial setting and exhibits large-scale cross-stratification formed by large dunes and unit bars with ∼westward paleoflow. Although in the past this lower unit has been correlated to the Ausable Formation due to similarity in sedimentary lithofacies, its detrital composition, detrital zircon ages and its westward paleoflow instead suggest that this is a part of the Keeseville Formation. The upper part of the section exhibits a planar-tabular character. Based on lithofacies analysis it is interpreted to have been deposited mostly by unconfined terminal ephemeral sheet floods in an arid or semi-arid terrestrial environment (i.e. ephemeral fluvial). Each sheet flood deposit was largely deflated and reworked by the wind during the long recurrence intervals between flooding, which resulting in a generally tabular architecture with abundant windblown deflation lags. In places, particularly near the top of the section, upstream,-migrating scour-filling features are exposed and interpreted as supercritical chutes-and-pools or cyclic steps. These features, particularly cyclic steps, are better exposed and thus described at stop 6B.
The sharp contact separating the braided and ephemeral fluvial strata records a shift in climate of from humid to semi-arid. This climate shift is correlated to a global Late Cambrian climate fluctuation at around ∼494 Ma recorded by numerous isotopic and lithofacies indicators and attributed to orbital forcing (Cherns et al., 2013).
Return to the parking lot, where we will proceed back to Jerdon Road and make our way to the section above the High Falls that exposed the ephemeral fluvial unit.

STOP 4B. Above High Falls
After leaving the High Falls parking lot, turn left (south) on to Jerdon Road. Then turn left (east) on to Pulp Mills Road, cross the bridge and turn right (south) on to La Plant Road. Park along La Plant Road and walk back to Pulp Mill Road and turn left (west). Cross the bridge and then make a right and walk north along a beaten path towards the Chateguay River. The outcrop is continuous along both sides of the river. From this (the east) side the finer detail of the ephemeral fluvial deposits can be seen, while spectacular large-scale features, including cyclic steps, can be seen by looking across to the other side of the river.
Most of the section along the west side of the river consists of planar-stratified medium to very coarse grained sandstone (Figure 18). Many if not most planar laminations are inversely graded and record deposition by the migration of wind ripples. Many planar layers are capped by coarse- to very coarse-grained deflation lags. Bedding surfaces are either featureless or expose low relief regular or irregular features suggesting adhesion and/or microbial processes. Sand Grains, although coarse, are rounded with common pits and bulbous edges, suggesting aeolian processes. Eluvial silt and clay is common along grain margins, indicating illuviation of fine-grained windblown material. Scours and low angle cross-stratified sets, probably formed by antidunes under energetic sheet floods, are common (as in Figure 18). Altogether this association records episodic deposition by terminal sheet floods (i.e. terminal splays) followed by long periods of aeolian deflation and reworking in an ephemeral fluvial environment. Most of these features can be observed close-up while walking north along the exposed sections. Farther north, pink colored apatite-cemented layers occur in lower parts of the section. These preserve high primary intergranular volumes and prohibit silica overgrowths present elsewhere (Figure 19) and thus are early intergranular cements. The origin of the phosphate is unknown, but was most likely derived from phosphate-bearing brines from underlying basement rocks.
As we walk back along the outcrop to Pulp Mill road, a large scour-filling feature can be seen on the other side of the river. This feature (outlined in Figure 20) is interpreted as a cyclic step deposit. Cyclic steps form under high energy and shallow supercritical flows in which the formation of hydraulic jumps permits the buildup and upstream-migration of cross-stratified sandstone (see Figure 20). This scour records the upstream-migration of the erosional trough of a cyclic step bedform. On the upstream side of the trough the flow accelerated causing erosion, while further downflow the flow expanded under a stable hydraulic jump resulting in stoss-side deposition (e.g. Cartigny et al. 2014). The sets are bounded by erosion surfaces that record natural fluctuations in the quasi-stable hydraulic jump during the same flood event. The final position of the cyclic step trough and its subsequent fill under subcritical flow conditions is recorded by the symmetrical scour immediately upstream of cyclic step sets, which contains downstream-migrating dune cross-stratified sandstone (Figure 20). Thus this features records high energy sheet flood conditions, perhaps during an atypically voluminous high energy flood or a flood closer to the source of overland drainage (such as a nearby isolated storm cell or front). The erosion and filling of cyclic step troughs (the scour) are crucial to the preservation of these features from subsequent aeolian reworking/deflation.
Following some discussion/observation, views of the east side can be made from the Pulp Mill Bridge and by a small beaten path on the other side of the bridge. But USE CAUTION the bank is VERY STEEP and quite high on this side. DO NOT GET TOO CLOSE TO THE EDGE! From here we can examine the large-scale stratal architecture of these deposits that previously we examined close-up. Most of the upper section is made up of the planar-stratified aeolian-reworked terminal splay deposits we examined close-up, while some parts below are interpreted as scour-filling antidune cross-strata or cyclic step sets (Figure 21).
Ascend the path back to Pulp Mill Rd and return to the vehicles parked on La Plant Road. We must turn the vehicles around and proceed back to the T junction with Pulp Mill Road.

STOP 5. Rainbow Quarry: The Keeseville Formation
Park at or just beyond the quarry entrance. The best section is exposed in the southwest corner of the Quarry. Here relatively thick (∼4-6 m) sets and cosets of cross-stratified medium grained sandstone are present and are interpreted to have been formed by the migration of large aeolian dunes in a generally dry aeolian erg setting (Figure 22). Close examination of the cross beds reveals inverse graded grain flow deposits in foreset strata and inversely-graded wind ripple laminations in the toe sets. At this location and at a similar quarry ∼1.6 km southwest of here the aeolian dune sets are separated by generally planar stratified sandstone underlain by horizontal surfaces erosion that most likely record dry to damp interdune erosion and deposition of interdune sand flats (see Hagadorn et al. 2008 for more details). Aeolian erg strata here are interpreted to be coeval to the ephemeral fluvial strata at the top of the previous stop (Chateauguay Chasm) and most likely accumulated by the deflation and downwind transport of the ephemeral fluvial splay deposits at the Chateauguay Chasm. This correlation is supported by similarities in detrital zircon ages.

STOP 6. Lasalle Dam, Altona NY: The Keeseville Formation
Park at the Feinberg Park and walk south along Devils Den Road across the bridge. The outcrop is along both sides of the river, and the section is continuous upriver to the abandoned McGregor Powerhouse. Following some discussion on the south side of the river, the north side can be accessed directly from Feinberg Park.
This section exposes a part of the Upper Cambrian Keeseville Formation which underlies the surrounding area in Altona north of the Flat Rock State forest (Figure 4). Here the Keeseville consists of a sheetflood-dominated ephemeral fluvial facies association that was deposited on a Late Cambrian coastal plain inboard of the paleo-shoreline located at or somewhere south of the Ausable Chasm section, ca. 40 km to the south, which was mostly a coastal sabkha throughout the Late Cambrian. A good modern approximation for such a setting might be the coastal plain of the Skeleton Coast of Northern Namibia in which ephemeral rivers discharge directly into the Atlantic forming ephemeral fluvial “deltas”.
Most of this section consists of beds of slightly inclined (≤10°) or horizontal planar-stratified medium grained quartz arenite usually truncated by low angle or horizontal surfaces (Figure 23). Internally these beds consist of finely-interlaminated shallow water (depth-limited orbital ripples, upper plane bed, rare current ripples) and aeolian (wind ripple and adhesion) lithofacies recording alternating wet and dry conditions. Coarse-grained deflation lags are common and desiccation cracks are also present but rare. These beds are interpreted terminal splay lobes in an ephemeral fluvial setting, and were built up over time by the accretion of episodic flood deposits. Flood deposits were largely reworked by the wind between episodic floods, resulting in common aeolian and adhesion features. Features such as oscillation ripple marks, parting lineations and adhesion ripples and warts immediately catch the eye on the bedding planes of these terminal splay lobes but belie the aeolian influence on deposition. Keep an eye out for less obvious features, such as inversely-graded laminations or low relief step-like features on bedding planes attributed to wind ripple stratification, coarse-grained deflation lags, or featureless planar laminations and surfaces most probably recoding the adhesion of windblown sand onto damp surfaces (i.e. adhesion stratification: Hunter, 1980; Kocurek and Fielder, 1982). Infiltrated matrix is also present in these strata and can be seen in thin section, and record the infiltration of clays into the vadose zone by illuviation, probably during rainstorms.
Upper medium to coarse-grained antidune cross-stratification is also present at this section and record high energy, probably more proximal, waning shallow sheet flood conditions. Antidune strata are recognized by low angle symmetrical trough-cross stratified bottom sets and convex-up low angle symmetrical formsets with wavelengths in the order of ∼25 cm – 1 m (Figures 23 and 24). It has a very similar structure to HCS, something noted from numerous fluvial deposits (Cotter and Graham, 1991; Rust and Gibling, 1990; Fielding, 2006) and also from experimentally produced antidune cross-stratification (Alexander et al., 2001; Cartigny et al., 2014). However, several features including the small thickness and wavelengths of sets and formsets, its 2D form on bedding planes, its relatively coarse grain size and lack of association with bioturbated lower shoreface deposits rules out classic HCS.
Also present near the base of the section are erosionally-based and crudely channelized trough cross-stratified sandstone formed by the migration of subaqueous dunes. These strata record the filling of an erosional distributary channel (like an arroyo or wadi) that had previously funneled energetic flood waters to terminal splays, probably during the late stages of lower energy floods leading to the abandonment of this particular distributary channel.
Isolated Proticnites trackways formed by arthropods have been observed at this location, but are rare.

A6: QUATERNARY DEGLACIATION OF THE CHAMPLAIN VALLEY WITH SPECIFIC EXAMPLES FROM THE AUSABLE RIVER VALLEY
David A. Franzi
Center for Earth and Environmental Science
SUNY Plattsburgh
Plattsburgh, NY 12901
David J. Barclay
Geology Department
SUNY Cortland
Cortland, NY 13045
Rebecca Kranitz and Kelly Gilson
Deglacial events and paleoenvironments in northeastern New York were influenced by the high relief of the Adirondack Mountains and predominance of north to northeast-draining river valleys (Figure 1). Landforms and sediments within these valleys record sequences of proglacial lakes that existed between the retreating margin of the Laurentide Ice Sheet (LIS) and drainage divides farther south. In general, these proglacial lakes evolved from small headwater impoundments to lowland lakes as the LIS exposed connecting spillways, and each change in lake level caused concomitant changes in tributary stream channels as river systems adjusted to the drop in local base level. Because of the interdependence of these water bodies and the adjacent LIS margin, the chronology of these lakes forms a key record for constraining the timing and style of deglaciation of this area.
In this trip we will examine the field evidence for these lakes and consider the implications of their chronology. In particular, we will focus on the methods used to reconstruct the sequence of proglacial lakes in the Ausable River valley and the constraints that this lake chronology places on a putative interval of post-LIS alpine glaciation in the Adirondacks. These deglacial events and deposits have influenced the evolution of modern drainage patterns, and we will also consider the geomorphology of Ausable Chasm and the nearby Ausable River delta at the shoreline of Lake Champlain as they relate to late Pleistocene and Holocene landscape evolution.

A6: QUATERNARY DEGLACIATION OF THE CHAMPLAIN VALLEY WITH SPECIFIC EXAMPLES FROM THE AUSABLE RIVER VALLEY

Meeting point

Alternate Meeting Point

STOP 1. Chapel Pond, Keene Valley, NY
Chapel Pond Pass (elev. 505 meters a.s.l.) served as the outlet for proglacial Lake Chapel. We shall use this scenic spot to present a brief summary of the field trip, regional geography and deglacial lake chronology of the Ausable Valley.

STOP 2. Styles Brook Mouth, Jay, NY
We shall stop briefly near the mouth of Styles Brook to discuss Tropical Storm Irene, the damaged it caused, the restoration efforts and natural recovery of the fluvial system. We may also use this opportunity to consolidate vehicles to reduce parking congestion at the next stop.

STOP 3. Styles Brook Valley Bottom, Jay, NY
We will access the Styles Brook study area via a trail that crosses private property. Time constraints prevent us from viewing the entire study reach but the flood damage and late Pleistocene sediment exposures at this location are typical.

STOP 4. Jay Covered Bridge, Jay, NY; Discussions and Lunch Stop
The covered bridge in Jay lies adjacent to a knickpoint formed by a metanorthosite outcrop in the East Branch channel. We shall take advantage of this scenic location to enjoy lunch and discuss our paleogeographic reconstructions and implications of Styles Brook stratigraphy and landforms for alpine glaciation.

STOP 5. Clintonville Pine Barren, Clintonville, NY
The Clintonville pine barren is situated on a large deltaic sandplain built by the Ausable and Little Ausable rivers into an embayment of glacial Lake Coveville in the lower Ausable Valley (Figures 4B and 4C). The Little Ausable River was originally a tributary of the Ausable River but the lowering of glacial Lake Vermont water level to the Fort Ann stage initiated downcutting into the Coveville delta deposits and resulted in its diversion across the Ausable drainage divide to its present course (Figure 4D). The Clintonville pine barren is a fire-dependent ecosystem, consisting primarily of pitch pine with an understory of heath plants that has adapted to the low nutrient and drought-prone soils of the deltaic sandplain.

STOP 6. Clintonville Stream Terrace Deposits, Clintonville, NY
The upper portion of the north face of this small gravel pit exposes about 4m of medium to thickly bedded pebble gravel, sand and minor silt that is part of a stream terrace graded to the lower glacial Lake Fort Ann delta at Keeseville. Individual beds range from about 0.1 to 1.0 meters thick and are traceable for several meters across the exposure. The middle approximately 2–3 meters of the section contains a remarkable set of climbing dunes, the upper and lower bedding surfaces of which dip gently upvalley.
The stream terrace gravel and sand disconformably overlies a discontinuous (0.0–0.7 m thick) layer of reddish brown rhythmically laminated silt and clay (varves). The undisturbed clay section is inaccessible but estimates from photographs indicate that it contains at least 20 silt-clay couplets. The west wall of the pit is oriented perpendicular to the paleocurrent and parallel to the valley slope. The rhythmite section in this exposure is thicker (up to a meter thick) and contains soft sediment deformation and slump structures. The rhythmites are draped over 5–6m of sandy, light brownish gray diamicton. The sandy texture of the diamicton is typical of surface till deposits in the region but is markedly dissimilar to the silty, dark gray diamictons observed in exposures at Styles Brook.

STOP 7. Lake Vermont Lacustrine and Deltaic Deposits, Keeseville, NY.
This exposure lies along the south bank of the Ausable River approximately 0.5 mi upstream (south) of the Keeseville Industrial Park (KIP) section that was described by Diemer and Franzi (1988), Franzi et al. (2002) and Franzi et al. (2007) and contains a similar record of lacustrine and deltaic sedimentation. The basal portion of the section consists of approximately 3 meters of dark gray silt and clay rhythmites that probably record pro-deltaic varved clay sedimentation in the Coveville phase of Lake Vermont. Water depth here at this time exceeded 80 meters. Deltaic sedimentation at this site began following the glacial Lake Iroquois breakout when glacial Lake Vermont was lowered to the Fort Ann level. The ice front would have receded about 30 km north to the Cobblestone Hill ice margin over this time interval at an average retreat rate of approximately 0.44 km/yr (Franzi et al., 2007). The locus of deltaic sedimentation shifted nearly 10 km downvalley (eastward) during the breakout (Figures 4C to 4D) from the Coveville-level Clintonville delta at approximately 205 meters a.s.l. to the lower Fort Ann-level delta in Keeseville at approximately 155 meters a.s.l.
The deltaic section at this site exposes approximately 10 meters of lower to middle delta slope foreset and bottomset beds. Excellent examples of ripple-drift cross lamination are commonly exposed.

STOP 8. Ausable Chasm, Keeseville, NY.
Ausable Chasm is a post-glacial gorge cut deeply into the Cambrian-aged Keeseville Member or, to a lesser extent, the Ausable Member of the Potsdam Formation. Landing et al. (2007) note three items of interest in the Cambrian section; 1. the occurrence of scyphomedusae impressions 2. siliciclastic microbial structures and 3. a low-diversity suite of trace fossils. These features are typically restricted to medium to fine grained quartz arenites and rare mudstones in the exposures immediately upstream from the US 9 bridge over the chasm, which are interpreted to represent shallow to emergent littoral sand flat facies (Landing et al., 2007). The scyphomedusae impressions from these outcrops were recently described by Hagadorn and Belt (2008). The outcrops also contain extensive horizontal trackways of Climactichnites and Protichnites (Landing et al., 2007).
A delta complex built by the Ausable and several smaller streams to the northwest overlie the Potsdam sandstones at this location (Figure 4E) (Denny, 1974; Franzi et al., 2007). Forced regression of the Champlain Sea shoreline accompanied isostatic uplift of the region causing the Ausable River to cut through the unconsolidated deltaic sediments and the underlying Potsdam Formation. Incision through the bedrock was probably accomplished by a combination of gradual incision and knickpoint migration and was influenced by the well-developed joint system in the sandstone, resulting in a rectangular channel pattern through the chasm.

STOP 9. Spring Sapping Channels in Champlain Sea Deltaic Deposits.
Spring sapping channels like the two at this site are commonplace in the deltaic deposits surrounding Wickham Marsh. In 1985 the heads of these channels were located several tens of meters farther downstream and their headcuts stood 2 meters high with slight to moderate spring flow near their bases. Since then, the channel heads have migrated headward and the height of the headcuts decreased. We shall discuss the significance of spring sapping as a channel-forming process in the Champlain Sea deltaic deposits and its possible role in the channel avulsion that led to the abandonment of the Wickham Marsh delta lobes and the creation of the modern Ausable delta in Lake Champlain (Figures 11A and 11B).

STOP 10. Wickham Marsh.
Wickham Marsh is a bayhead wetland on Lake Champlain that is separated from the main waterbody by a barrier beach, although nearly all of the beach is presently buried by a railroad embankment. The marsh’s basin was cut as the Ausable River and smaller tributaries adjusted to falling baselevel during the latest stages of the Champlain Sea (Figures 11A and 11B) and possibly early Lake Champlain (Denny, 1974). Spring sapping might have contributed to the development of the tributary stream networks. The lowest deltaic terrace presently lies about 4 meters below the level of Lake Champlain and represents the last deltaic surface formed before the avulsion of the Ausable River to its present course.

STOP 11. Ausable Delta on Lake Champlain.
This scenic spot provides an opportunity to conclude the trip on the modern Ausable delta and summarize the deglacial history of the region.

A7: NON-POINT SOURCE POLLUTION IN THE LITTLE CHAZY RIVER OF NEW YORK
Robert D. Fuller and David A. Franzi
Center for Earth and Environmental Science
State University of New York
Plattsburgh, NY 12901
Steven Kramer and Eric Young
Miner Agricultural Research Institute
Chazy, NY 12921
The Little Chazy River watershed (Fig. 1) drains into Lake Champlain, which is an oligotrophic to mesotrophic water body with low to moderate levels of phosphorus and nitrogen, the primary nutrients for primary productivity and principal determinants for associated water quality issues. Major sources of nutrients in Lake Champlain include point sources such as municipal sewage treatment plants and non-point sources including urban runoff and agricultural inputs. Extensive dairy operations in the Lake Champlain basin produce large quantities of manure, which is applied back to soils and can potentially become a major source of nutrients to nearby surface waters (Fig. 2). Consequently, reduction of nutrient inputs to Lake Champlain, which promotes healthy and diverse aquatic ecosystems and provides for sustainable human use and enjoyment of the lake, is a priority.
SUNY Plattsburgh and the William H. Miner Agricultural Research Institute (Miner Institute) used a high-resolution monthly synoptic water-sampling strategy to examine the spatial and downstream distribution of nutrient concentrations (i.e. nitrate-N, total phosphorus and soluble reactive phosphorus) in the Little Chazy River and its principal tributaries for five full years between January 2008 and December 2012 (Fig. 1). The Little Chazy River watershed is typical of medium-sized (basin area = 145 km2), rural watersheds in the region, possessing a broad range of watershed issues and concerns reflected throughout the Champlain lowlands. Agriculture accounts for approximately 17% of land cover in the watershed, compared to approximately 24% of land cover county-wide, most of which is concentrated near the lake shore east of Chazy village. The watershed was identified on the 1996 Lake Champlain Basin Waterbody Inventory/Priority Waterbodies List Report as a class C river with possible impaired fish survival (NYSDEC, 2001). Karim (1997) cited on-site septic problems, particularly in Chazy and West Chazy, and high levels of livestock and crop agriculture as water pollution concerns in the watershed. The Little Chazy River has the highest median nutrient concentrations and median unit nutrient load (nutrient load per unit area of watershed) of any New York tributary to Lake Champlain that is monitored as part of the Lake Champlain Long-term Water Quality and Biological Monitoring Program (Vermont DEC and New York DEC, 2002).

A7: NON-POINT SOURCE POLLUTION IN THE LITTLE CHAZY RIVER OF NEW YORK

STOP 1. Miner Dam, Town of Chazy, NY
The Miner Dam was constructed near the headwaters of the Little Chazy River circa 1913 and was abandoned shortly thereafter due to excessive leakage. These headwaters traverse the Altona Flat Rock Pine Barrens, consisting largely of jack pine, pitch pine and lichens growing on a thin mantle of organic soil over sandstone bedrock. The area was scoured by a catastrophic breakout flood at the end of the last ice age, approximately 13,000 years ago. Lacking a substantial mineral soil, these ecosystems are oligotrophic, with extremely low nitrate-N concentrations (Fig. 4A), particularly during the summer growing season (Fig. 6). Nitrate is probably removed by biological uptake, both in the soil and in the impoundments above the dam. Phosphorus concentrations are also low (Fig 4B), and are largely in an organic or particulate bound form (Fig. 5).

STOP 2. Agricultural Tributary Sample site, Town of Chazy, NY
This site (StratAgTrib) illustrates the importance of agricultural inputs and their influence on nutrient concentrations in the Little Chazy River (Fig. 2). Immediately upstream from this site, nitrate concentrations rose as the River traversed dairy agricultural operations on well-drained, sandy glacial till and outwash parent materials. Numerous agricultural tributaries and ditches in the Little Chazy River receive inputs from tile-drained fields which receive manure applications, typically 2-3 times per year (Fig. 2). Nutrient concentrations in these tributaries are typically high, but low flows typically cause only minor increases in nutrient concentrations in the Little Chazy River. However, their effects are cumulative.

STOP 3. Wood Sample site along the Little Chazy River, Town of Chazy, NY
The Wood sample site is more typical of the lowland reach of the drainage basin, with soil parent materials composed of fine-textured glacial lacustrine sediments. River gradients are typically low, and many of the agricultural fields have subsurface drainage. Evidence of river meanders is common, and during high flows extensive streambank erosion can occur. Throughout this reach, total phosphorus concentrations increase slowly, driven largely by the soluble reactive phosphorus fraction. Nitrate concentrations, however, remain relatively constant throughout this reach, suggesting that the magnitude of different flow paths may vary for the two nutrients. Downstream from this site, the River enters some small impoundments in which nitrate concentrations decline, particularly during low flow, suggesting either conversion to other forms and subsequent loss or biological sequestration.

STOP 4. USGS Sample site along the Little Chazy River, Town of Chazy, NY
The U.S. Geological Survey maintains a discharge monitoring gauge at the USGS sample site, providing continuous flow data. Episodic sampling sampling was conducted at this site during three successive storm events in the fall of 2010. During each of these events, maximum concentrations of both nitrate, total P and soluble reactive P were observed on the rising limb of the hydrograph, while precipitation was still falling. By the time the maximum flow was observed, concentrations had declined to baseline levels. Loadings were near maximum levels at moderate flows on the rising limb due to high concentrations, and during the peak of the event due to high flows. The relationship between discharge and load shows a marked hysteresis.
Immediately upstream from this site is the outflow from the publicly owned treatment works for the village of Chazy. Sampling immediately upstream and downstream of this site showed only slight increases in concentration of either N or P fractions, suggesting that this source is small relative to other sources, such as those arising from agricultural practices. Immediately downstream, both nitrate and P fractions increase rapidly, as the lower stretch of the drainage basin has extensive agriculture, almost all with subsurface drainage.

STOP 5. R15 Agricultural Field at Miner Institute, Town of Chazy, NY
The R15 agricultural field is operated by the William H. Miner Agricultural Research Institute, a fully functional research dairy farm. Miner Institute conducts research on dairy operations as well as their impacts on the environment. The R15 field has been tile-drained since 1975, and includes extensive shallow groundwater sampling devices, a deep groundwater monitoring well, and a meteorological station. At this site, we will conduct a demonstration of “smoking” a tile line. This demonstrates the connectivity of subsurface tiles (approximately 1 m deep) to the surface through soil macropores which include worm tunnels as well as structural cracks. This provides a potential route for direct and rapid transit of pollutants applied on the surface, such as those through manure application, to tiles and surface waters.

STOP 6: Edge of Field Study Site, Town of Chazy, NY
The USDA-NRCS edge-of-field (EoF) monitoring program is designed to quantify impacts of various nutrient management practices on runoff water quality. Data collection and analysis for EoF projects utilize a small paired watershed approach. Typically, one best management practice is evaluated in comparison to a reference condition. The objective of our project is to quantify water quality impacts associated with the practice of controlled subsurface tile drainage (CD) as compared to the standard practice of freely-drained subsurface tile drainage (FD). The practice of CD uses inline water control structures (AgriDrain, Adair, IA) to artificially raise the drainage outlet elevation during the non-growing season. Studies show that CD can reduce annual drainage flow volumes by as much as 80% compared to FD. The two crop production fields were selected due to their close proximity, similar soil types, and for their ability to be instrumented properly to capture both surface and subsurface runoff water. The fields are currently in 2nd year production of corn for silage. Subsurface drainage tile (slotted plastic drain tile, 10-cm i.d.) was installed via a tile-plow at an average depth of 1.2 m below the soil surface during November 2014 (10.7 m spacing between tile lines within each field). Subsurface drainage water from each field’s lateral lines flows to a main outlet pipe (schedule 40 PVC, 20-cm i.d.) that drains to individual concrete manholes (1.2-m i.d. and 2.4 m deep) where flow is gauged and sampled. Drainage water flows are measured by routing water through a plastic barrel equipped with a v-notch weir and stilling well that houses a pressure transducer-data logger to capture water elevation changes related to flow differences. Drainage flows are gauged by rating curves developed between measured flows and water height inside the barrels. Variation in nutrient and sediment concentrations will be quantified by taking multiple samples over an event using flow-programmable autosamplers. Overland flow is measured from each field by routing surface water to a fiberglass flume (Plasti-Fab, Tualatin, Oregon) using a combination of natural swales, field contouring and an earthen, clay berms. Event-based and seasonal mass loads of total N, nitrate-N, ammonium-N, total P, dissolved (<0.45µm) reactive P (orthophosphate), unreactive P (particulate-organic P), and total suspended solids will be calculated. Baseline monitoring will occur for two years after which CD treatment will be implemented and monitored over four years. Differences in mass nutrient loads between FD and CD will be evaluated using analysis of covariance. Agronomic data is also being tracked (i.e., yields, fertilizer and manure inputs and nutrient uptake) and will be utilized to calculate N and P mass balances each year of the study. A field-scale water quality model (APEX) will also be parameterized and calibrated as part of the project. Results will help quantify tradeoffs between drainage intensity, surface/subsurface hydrology, and nutrient loss dynamics at the field scale.

A8: GEOBOTANY OF GLOBALLY RARE SANDSTONE PAVEMENT PINE BARRENS, CLINTON COUNTY, NY
Jacob N. Straub and Rachel E. Schultz
State University of New York at Plattsburgh
Geobotany is the study of plant distributions, which often are determined by the interaction of plants with the underlying geology of an area. Historic and contemporary geologic and climatological influences have shaped the sandstone pavement pine barrens botanical community, which is considered rare in New York and the world (Edinger et al., 2014). The pavement barrens of Clinton County, New York were created over 12,000 years ago when glacial Lake Iroquois drained into glacial Lake Coveville (Franzi and Adams 1993). Plants have been able to survive in this harsh, nutrient poor ecosystem due to specialized characteristics and interactions with other organisms. The Altona Flat Rock sandstone pavement pine barren is fire-dependent, and the last stand-replacing fire burned over 1,200 ha in 1957 (Gooley, 1980). Fire acts as an ecological mechanism to restart succession in this type of community. An outstanding quality of this pine barren is its location at the northern extreme of pitch pine and southern extreme of jack pine, and both species depend on fire as a mechanism for reproduction. Another unique community found in the pavement barrens is the perched bog. These acidic peatlands are separated from the water table due to the sandstone bedrock and host unique plant species such as the carnivorous pitcher plant (Edinger et al., 2014). On this field trip we will visit the Altona Flat Rock pine barrens owned by the W.H. Miner Agricultural Research Institute and the Gadway Pine Barrens owned by the Nature Conservancy.

A8: GEOBOTANY OF GLOBALLY RARE SANDSTONE PAVEMENT PINE BARRENS, CLINTON COUNTY, NY

Meeting Point

STOP 1. Altona Flat Rock Pavement Barrens, Chazy, NY
The Altona Flat Rock pine barren is located in the township of Altona, Clinton County, and is primarily owned by The W. H. Miner Agricultural Research Institute. Altona Flat Rock is the largest of the sandstone pavements in northeastern New York (Figure 1). The central portion of Altona Flat Rock is drained by Cold Brook, a headwater tributary of the Little Chazy River that originates near the Dead Sea.
The Altona Flat Rock was heavily influenced by a large ice storm in 1998 (Irland 1998). An assessment of overstory tree composition at this site in 1999 revealed it was dominated by an admixture of jack pine (Pinus banksiana), pitch pine (P. rigida), and red maple (Acer rubrum). An outstanding quality of this pine barren is its geographic location at the northern limit of pitch pine and southern limit of jack pine (Burns and Honkala 1990). Both species depend on fire for reproduction (Yorks and Adams 2003). This particular sandstone pavement pine barren is fire-dependent and considered a rare ecosystem by the New York Natural Heritage Program (Edinger et al. 2014). The last stand-replacing fire to affect this pine barren occurred in 1957 and it burned over 1,200 ha. In this type of ecosystem fire acts as an ecological mechanism to set the successional state back to a less developed state (Noble and Slatyer 1980).

STOP 2. Gadway Pine Barrens, Mooers Forks, NY
The Gadway Sandstone Pavement Pine Barren is located in Clinton County, near the border with Quebec (Figure 1). The Gadway barren is an outstanding example of a sandstone pavement barren, a globally rare ecological community with fewer than 20 sites in the world and fewer than five sites in New York State. The Gadway barren is approximately 210 ha (520 ac) in size and is dominated by jack pine. With its serotinous cones, jack pine is one of the best examples of a fire-adapted species in New York.
The most recent example of a wildfire in the barrens occurred at Gadway on May 28, 2010. The day prior to the fire, there was lightning area which is believed to have caused this. With winds from a westerly direction, the main fire was a backfire spreading into the barren at a rate of less than 1 m/min. The forest ranger response was rapid and the fire burned only approximately 2 ha (5 ac).
In addition to the exploration of the plant communities at Gadway, we will also visit sites that expose fossils in the Potsdam Sandstone. According to Landing et al. (2007), “arthropod trackways (predominantly Diplichnites and Protichnites) as well as probable mollusk trackways Climactichnites and Plagiogmus, are well preserved on some bedding surfaces.”

B1: THE MAGNETITE-FLUORAPATITE ORES FROM THE EASTERN ADIRONDACKS, NEW YORK: CHEEVER MINE
Marian v. Lupulescu
Research and Collections
New York State Museum
Albany, NY 12230
Jeffrey R. Chiarenzelli
Department of Geology
St. Lawrence University
Canton, NY 13617
David G. Bailey
Geosciences Department
Hamilton College
Clinton, NY 13323
Sean P. Regan
Department of Geosciences
University of Massachusetts
Iron mining in New York State, a significant supplier of iron ore used in developing the resources and industries of the United States through the first part of the 20th century, has a long and interesting history. Iron ore was mined from Grenville-age metamorphic rocks in the Adirondack Mountains and Hudson Highlands, and from Silurian-age sedimentary rocks in central and western New York (Clinton-type).
Smock (1889) was the first author to present a complete report and the first classification of the iron ores “…which occurs in beds and deposits of workable size in the State of New York”. From his report we have brief descriptions, locations, and data on the ore quality, by-products, and output of the major mines operating in the late 19th century. Kemp (1897, 1908), Kemp and Ruedemann (1910), Nason (1922), Alling (1925), Gallagher (1937), Postel (1952), Buddington (1966), Hagner and Collins (1967), Baker and Buddington (1970), Foose and McLelland (1995), Valley et al. (2009), and Valley (2010) all made significant contributions to our understanding of the iron deposits in the Adirondack Mountains, but there is still no agreement as to their origin, and thus Alling’s statement from 1925 directed at individuals wanting to understand the “genesis of these important ore bodies in Northern New York may reach the conclusion that there is a hopeless disagreement among those who have studied these deposits” is still as valid today as it once was.
On this trip we will visit the Cheever Mine in the eastern Adirondack Mountains. Our goal is to present a brief classification of iron ore deposits in the Adirondacks and to briefly summarize pertinent aspects of the diversity, differences, and similarities between mines and prospects across the Adirondacks and ignite discussion on their origin. According to their mineralogical and geochemical compositions, the principal iron deposits and prospects in the Adirondacks (Figure 1) are 1) low Ti iron oxide – REEs (Mineville, Cheever, Palmer Hill, Rutgers, LaVake, and Arnold Hill mines), 2) low Ti iron oxide (Lyon Mountain and Benson mines), 3) low Ti iron oxide - boron (vonsenite) (Jayville and partially Clifton mines), and 4) Ti – Fe oxides (Tahawus). Based on their general features (Valley et al. 2009; Valley 2010) these ores are considered to belong to the IOCG (iron oxide – copper – gold) class of iron deposits (e.g., Hitzman et al. 1992) even though copper and gold mineralization is not known at these locations. One possible exception is the Lyon Mountain mines from which Callaverano and Zimmer (2009) reported on associated copper mineralization.
During this field trip we will visit the Cheever mine, Port Henry, Essex County, near Lake Champlain. Due to the elongate, sill-like geometry of the body it was worked at many locations along its length. To avoid confusion we will collectively refer to all the locations as belonging to the Cheever “mine”. On this trip we will examine the morphology of the magnetite – fluorapatite ore bodies, the mineralogy of the ore and the host (foot- and hanging-wall), the relation of the ore body to the neighboring coronitic metagabbro, and discuss the igneous origin of the deposit in light of age, major and trace element chemistry, and isotopic composition of the fluorapatites from the ore and host rock.

B1: THE MAGNETITE-FLUORAPATITE ORES FROM THE EASTERN ADIRONDACKS, NEW YORK: CHEEVER MINE

STOP 1. Cheever mine
At the Cheever mine, a 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 in the Mineville area. 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, and minor pyroxene, or quartz, albite, microcline, and pyroxene-rich units. 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 (Figure 2). Some pyroxenes display exsolution of tiny elongated lamellae of another pyroxene and ilmenite as the result of the sub-solidus re-equilibration during slow cooling. In some grains, both types of lamellae occur oriented along the two pyroxene cleavages. Rare grains of ilmenite and pyrrhotite pepper the rock. The coronitic metagabbro contains a layered facies with the same composition as the massive one.
The footwall rock which hosts the ore is granite that contains predominantly quartz and albite (Figure 3). Minor phases are grains of relict clinopyroxene (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. Further on, the rock displays a gneissic texture with alternating “bands” of quartz and albite, and pyroxene, amphibole, annite, magnetite, and 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 sustained also by microscopic study) and contain pyroxene, and rounded, in places 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, like those from Mineville (Figure 5a, b).
The magnetite-fluorapatite ore at the Cheever Mine forms 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 interraction 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.
Three samples of Cheever ore and host rock were selected for electron microprobe WDS analysis and in-situ U-Th-total Pb monazite geochronology. WDS compositional mapping was performed on the Cameca SX-50. Quantitative analysis of monazite was performed on the SX-100 ultrachron following procedures outlined in Allaz et al. (2013) at the University of Massachusetts, Amherst. Initially, thin sections were mapped for Mg, Y, Ce, Zr, and Ca or Fe (300 nA, defocused beam; Fig. 8). Close up maps of monazite were done with a fixed stage, using a beam rastering mode for sub-micron resolution, with the current at 200 nA. For large grains, stage mapping was employed due to the sufficient resolution at 2 micron-pixels. Preliminary geochronology will not be presented herein due to extremely low actinide contents.
As mentioned above, the ore contains magnetite, apatite, allanite, monazite, stillwellite, and other phases. All thin sections contain sharp boundaries between ore and the host albite-quartz rock with augitic pyroxene crystals. All WDS full section maps show abundant zircon in host rock, but no zircon in the ore. One thin section (Figure 9) contains abundant REE bearing phases interspersed with magnetite. Most notably are medium grained stillwellite (Ca, LREE, Th borosilicate) and coarse symplectic intergrowths of allanite and monazite. Symplectic monazite has very low actinide concentrations, and has been a challenge to date (as noted above). However, the low actinide content suggests a low temperature, post mineralization origin for the symplectite formation. The other two thin sections contain coarse monazite crystals, typically associated with, and sometimes rimming, apatite crystals. Recent grain mapping has revealed several high Th grains, which should yield robust and reliable counting statistics, and thus an acceptable age determination. This work is ongoing.

B2: DECIPHERING THE GEOLOGIC EVOLUTION OF THE EASTERN ADIRONDACKS: NEW FIELD, PETROLOGIC AND GEOCHRONOLOGIC DATA
Timothy W. Grover
Department of Natural Sciences
Castleton University
Castleton, VT 05735
Michael L. Williams
Amherst, MA 01003
James M. Mclelland
Colgate University
Hamilton, NY 13346
Sean P. Regan
Graham B. Baird
Department of Earth and Atmospheric Sciences
University of Northern Colorado
Greeley, CO 80639
Kimberlee A. French
Claire R. Pless
Models for the tectonic evolution of the Adirondack Mountains have advanced significantly over the past ten to fifteen years. Prior to this, most workers would suggest that the (ca. 1.06 Ga) Ottawan Orogeny was the dominant deformation-metamorphic event in the Adirondacks. Recently published results from geochronological studies based on high resolution, single grain, U-Pb zircon SHRIMP data (McLelland et al., 2001; Hamilton et al., 2004; Heumann et al., 2006; Bickford et al., 2008), along with geochemical and field based studies in the Adirondack Lowlands (Chiarenzelli et al., 2010; Baird and Shrady, 2011), and recognition of an ophiolite suite in the Lowlands (Chiarenzelli et al., 2011) has resulted in significant revision of models of the tectonic evolution of the Adirondacks. These data reveal a major mountain-building event, the Shawinigan Orogeny (∼1.2-1.16 Ga) - previously unrecognized in the Adirondack Highlands, resulted in widespread deformation and high grade metamorphism with accompanying partial melting. After the Shawinigan Orogeny, extensive anorthosite-mangerite-charnockite-granite (AMCG) magmatism (1.16. to 1.14 Ga ) is interpreted to reflect a period of lithospheric delamination that resulted in asthenospheric upwelling and decompression melting in the mantle (McLelland et al., 2010) that yielded gabbroic magmas parental to the anorthosites. Thermal effects of the gabbroic magmas provided heat to partially melt portions of the lower crust and produce the felsic magmas of the AMCG suite.
In addition, Selleck et al. (2005) proposed that the Lyon Mountain granite was emplaced at the end of the Ottawan Orogeny (ca 1.05 Ga) and that it accompanied extensional collapse following peak-Ottawan compression. Wong et al (2012) support this model with evidence for an extensional shear zone in the eastern Adirondacks and show, using in situ EMP monazite data that the latest, normal-sense motion along the shear zone occurred from approximately 1.05 to 1.0 Ga.
In summary, important developments include: (1) the recognition that the 1200-1160 Ma Shawinigan Orogeny played a major role in a multiphase orogenic history; (2) the extensive AMCG magmatic suite, previously considered to be anorogenic, is now thought to have occurred during post-Shawinigan lithospheric delamination; (3) the 1090-1050 Ma Ottawan Orogeny may overprint but does not obliterate the earlier Shawinigan features; (4) significant shear zones exist on the western and eastern sides of the Adirondack Highlands that record the early stages of collapse and exhumation of the Grenville orogen.
Distinguishing between Shawinigan, AMCG, Ottawan, and extensional stages in the deformation and metamorphic history is essential for P-T-D-t models and for constraining the Adirondack tectonic history in general. Our work in the eastern Adirondacks will delineate the nature and spatial extent of both the Shawinigan and Ottawan Orogenies. By nature we mean metamorphic grade, style of deformation, composition of igneous rocks, and importantly, age constraints on each of these processes.
The East Adirondack shear zone is an intriguing feature in the Adirondack Highlands, possibly correlative in timing and history with the Carthage-Colton shear zone on the west side of the Adirondacks. However, many questions remain before this late-stage deformation can be wholly integrated into the tectonic history of the Adirondacks. Our on-going research aims to address the following questions:
1) What is the width, length, and intensity of the East Adirondack shear zone?
2) What is the style of the shear zone, is it constructed of many narrow discrete zones (<10 m) or is it distributed more uniformly over a wide zone (> 100m)?
3) Is this an extensional or a compressional shear zone, or does it have a complex, multiphase, diachronous history similar to the Carthage-Colton shear zone?
4) What was the P-T path during shearing, did the latest extensional phase of shearing recorded by these rocks occur at different crustal levels and therefore different metamorphic grades?
5) How does the shear zone link to other high-strain zones in the Eastern Adirondacks and more broadly to the Carthage-Colton shear zone?
6) What is the timing of shearing and is there an inheritance relationship with deformation in the Shawinigan?
Our approach to answering these questions is multifaceted and includes field work, macro- and microscale structural analysis, petrologic analysis and modeling, and geochronology. Field observations are foundational for all the research that follows. Contacts and map scale structures are mapped; kinematic indicators such as deformed porphyroblasts or porphyroclasts and strong lineations are documented. Oriented polished thin sections are used for microscopic kinematic and textural studies. Extensive compositional mapping and elemental analysis provides the data for forward petrologic modeling using Theriak-Domino software (de Capitani and Petrakakis, 2010) in order to determine not only the P-T conditions of metamorphism but also the sequence of metamorphic reactions. In situ electron microprobe dating of monazite is used to constrain the timing of metamorphism and deformation (Williams and Jercinovic, 2012). Application of this technique is discussed elsewhere in the paper.
The stops on this trip are designed to illustrate the wide range of rock types present in the eastern Adirondacks including but not limited to metagabbos, deformed anothositic rocks, charnockites, and garnet-sillimanite gneisses. The geologic context of each stop will be discussed and new data, when available, will be presented. Lively discussion is encouraged. Of course this trip is not the first to visit many of these stops and additional information can be obtained from previous NYSGA and NEIGC field trips including McLelland et al. (2011) and Whitney et al. (2002).

B2: DECIPHERING THE GEOLOGIC EVOLUTION OF THE EASTERN ADIRONDACKS: NEW FIELD, PETROLOGIC AND GEOCHRONOLOGIC DATA
T.W.Grover, M.L.Williams, J.M.Mclelland, S.P.Regan, G.B.Baird, K.A.French, C.R.Pless

Meeting Point
T.W.Grover, M.L.Williams, J.M.Mclelland, S.P.Regan, G.B.Baird, K.A.French, C.R.Pless

Exposures of the “Great Unconformity”
Exposures of the “Great Unconformity” with the Cambrian Potsdam sandstone sitting unconformably on nearly vertically dipping Mesoproterozoic gneisses.
T.W.Grover, M.L.Williams, J.M.Mclelland, S.P.Regan, G.B.Baird, K.A.French, C.R.Pless

STOP 1. Isoclinal folds, Lyon Mountain granite, and AMCG rocks
Examine the granitic gneiss in the northern outcrops on the west side of the road. Note the numerous tight to isoclinal similar folds (Fig. 4). The fold axis trends almost due east and has a gentle eastward plunge. If you look closely at the nose of the fold you can see a new axial planar foliation has developed. If we call the folded foliation S1 then the new axial planar foliation would be S2. The foliation on the limbs of the axial planar folds would be a composite S1/S2 foliation. Our working hypothesis is that S1 developed during the Shawinigan Orogeny and S2 developed during the Ottawan Orogeny. However, this may be problematic if these granitic gneiss are part of the AMCG suite. We have collected oriented samples in this region to see if we can use monazite age data to constrain the timing of foliation development.
Further to the north along this outcrop is a reverse fault that is cut by a pegmatite. Where the pegmatite cuts the granitic gneisses it appears as if fluids from the pegmatite interacted with the gneiss and resulted in potassic alteration of the gneiss.
Proceed south, past a gap in outcrops to the next exposure. Here various phases granites, along with amphibolite that exhibits flow folding are intermixed. The white to gray granitic rock appears locally to be undeformed except for shear zones and is interpreted to be the Lyon Mountain Granite with its typically low (< 0.3 wt.%) MgO composition. The pink granitic rock contains few amphibolite enclaves and is thought to be an AMCG granite. McLelland et al. (2011) refer to this roadcut as an intrusion breccia involving white Lyon Mountain granite that engulfed, disrupted, and flow folded amphibolites. Portions of the Lyon Mountain granite have undergone various phases of alteration by sodic and potassic fluids resulting in elevated concentrations of those elements (Valley et al., 2011).

Outcrops of Potsdam sandstone

STOP 2 Lineated K-feldspar megacrystic gneiss and migmatitic paragneiss
This rock is a biotite-bearing, K-feldspar megacrystic rock with a pervasive linear fabric. Most of the K-feldspar crystals are strung out and completely recrystallized into finer grained mosaics (Fig. 5), however some large crystals are still present. The dominant lineation is gently plunging to the east.
Graham Baird recently analyzed zircons from this sample using the SHRIMP-RG at Stanford University. Zircon in the unit is euhedral to subhedral, elongate, amber colored, and ranges in length approximately 200-400 µm. Cathodoluminescence of polished grain cross sections reveal that all grains possess oscillatory zoning with no conclusive sign of metamorphic rims or common inherited cores (Fig. 6). Fourteen SHRIMP-RG analyses produce what is interpreted as a unimodal distribution of ages with one inherited core (1497 ± 78 Ma) and 3 anomalously young ages. Seven concordant and near concordant, well clustered analyses produce a 2σ   207Pb/206Pb weighted mean age of 1155 ± 15 Ma, interpreted to best represent the age of pluton crystallization (Fig. 6).
The 1155 Ma age is consistent with this granitoid being part of the AMCG suite. The strong, penetrative fabric suggests a phase of deformation after this time.
Further north along the outcrop the K-feldspar megacrystic granitic rocks are intermingled with migmatitic, biotite-bearing gneisses. The contact relationships between these two units are not clear cut. We have sampled both units in order to collect more geochonologic data using EMP analyses of monazite.

Mafic boudins in metasedimentary rocks

Pull out with folded granitic rocks near southern end.
Also a spring with potable water.

STOP 3. Garnet-sillimanite gneisses and coronitic metagabbro
Many field trips have stopped at this location over the years. McLelland et al. (1988a) cite evidence from 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. 7). The gneiss is a garnet-sillimanite-plagioclase-K-feldspar-quartz gneiss with a small amount of biotite. The garnet-sillimanite gneisses here are often referred to as khondalites. This name originated in India and refers to quartz-garnet-sillimanite gneisses with little to no biotite. Khondalites are often associated with quartzites and calc-silicate rocks and that is the case here. This mineral assemblage is consistent with upper amphibolite to lower granulite facies metamorphic conditions. The garnet-sillimanite gneiss is penetratively deformed with a well-developed foliation and a lineation that plunges gently to the east. 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. The gabbro is finely crystalline right at the contact with the gneiss (Fig 7c) and appears to get more coarsely crystalline towards the interior of the body suggesting a chilled margin formed at the contact between the gabbro and the gneiss. 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 metagabbros are found throughout the Adirondacks (Whitney and McLelland, 1973; Whitney and McLelland, 1983; Regan et al., 2011). Figure 8 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. The clouding in the plagioclase is due to numerous microscopic spinel inclusions. Another feature of these rocks that ilmenite is rimmed by Ti-rich hornblende. 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 coronitic metagabbro seen here, have a very strong penetrative fabric and in many locations are mylonites. These rocks contain the mineral assemblage garnet-clinopyroxene-plagioclase±orthopyroxene which is similar to the mineral assemblage found in the coronitic metagabbros. Kinematic indicators in these rocks suggest a west-directed thrust sense of motion. These observations suggest this period of deformation and metamorphism must postdate the emplacement of the AMCG rocks.
McLelland et al. (1988b) report a U-Pb zircon, multigrain age of 1144 ± 7 Ma for the metagabbro. This is interpreted as an igneous crystallization age and is consistent with the gabbro belonging to the AMCG suite. However, Aleinikoff (pers. comm.) has reported a preliminary SHRIMP age of ca. 1107 Ma. We dated monazite from the garnet-sillimanite gneisses following the techniques outlined in Williams et al. (2006), Williams et al. (2007) and applied in Williams and Jercinovic (2012). The technique involves locating all the monazite crystals in the thin section through full section mapping with the microprobe followed by detailed mapping of up to 30 monazite crystals in a thin section to identify different compositional domains within each monazite. Then the microprobe is used for complete spot analyses of different compositional domains in the monazite and ages are calculated from these data. We analyzed monazite from two thin sections of the garnet-sillimanite gneisses. The results are presented in figure 9.
The data broadly suggest three distinct periods of monazite growth. The oldest is set of ages yields a mean age of 1179 ± 9 Ma. We suggest that this age represents the timing of granulite facies metamorphism and fabric development in the garnet sillimanite gneisses. This age is consistent with metamorphism and deformation occurring during the Shawinigan Orogeny. The next population of analyses yields ages that cluster around 1151 Ma. We hypothesize that this represents a period of monazite growth driven by a thermal perturbation resulting from the intrusion of the gabbroic rocks. This age correlates well with the reported age of the gabbro. With the exception of one analysis yielding an age of approximately 1050 Ma, most of the remaining monazite ages are 1020 Ma or less. These ages are too young to correlate with the proposed timing of the peak of the Ottawan Orogeny (∼1090-1050 Ma). One possibility is that monazite grew during the period of 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 ages 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 fabric 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. We will discuss this further at Stop 8.

STOP 4. Mylonitic gabbroic anorthosite in a gravel pit
Since leaving stop 3, we have been passing through a unit of deformed and metamorphosed mafic rocks ranging in composition from gabbroic to anorthositic. Most of the rocks have a well developed foliation. Many contain porphyroclasts of plagioclase that have been variably deformed (Fig. 10). As mentioned earlier, shear sense indicators in these rocks suggest a west-directed, thrust sense of shear. The mineral assemblage in these rocks is typically garnet-augite-hornblende-plagioclase with little to no orthopyroxene or quartz. The presence of hornblende, a hydrous mineral, requires an H2O-bearing fluid, possibly introduced during deformation. Although we continue to search we have yet to find any monazite in these rocks.

Large outcrop of metagabbro with a white marble xenoltth.

Good exposures of dark green charnockite.

STOP 5. Pegmatite, charnockite, and mafic gneiss
This outcrop contains a beautiful exposure of a post-Ottawan pegmatite dike. The hanging wall of the dike is a mafic gneiss. Xenoliths of the mafic gneiss are in the upper third of the dike and show evidence of reacting with the magma (Fig. 11). If you look across the street you can see the continuation of the dike to the north. Most of the outcrop here is granitic gneiss and maroon-brownish charnockite. These rocks are thought to belong to the AMCG suite. If this is correct foliation in these rocks must have developed from a deformational event after ∼1155 Ma. Towards the south end of the outcrop there are cm-scale pyroxene crystals. This rock contains the high grade mineral assemblage garnet-augite-hornblende-plagioclase-quartz±orthopyroxene, indicating that these are upper amphibolite to lower granulite facies metamorphic rocks.

Outcrops of ca. 1.3 Ga tonalitic rocks
(McLelland and Chiarenzelli, 1990).

STOP 6. Garnet-sillimanite gneiss
The entire road cut consists of well foliated and lineated garnet-sillimanite-K-feldspar-plagioclase-quartz±biotite gneiss. This is another example of khondalite. The sillimanite in this rock forms a prominent lineation of ∼10/100. This is similar to all other lineations noted so far on this trip. The mineral assemblage is again consistent with upper amphibolite to lower granulite facies metamorphism.

STOP 7. Long outcrop of many different lithologies found in the Adirondacks
The rocks in this outcrop include many of the lithologies found throughout the Adirondack Highlands that have probably been tectonically juxtaposed. The foliation is dominantly NE striking and gently dipping to the southeast. As we move southward we will moving structurally down section. The north end of this outcrop is an impure marble that contains coarsely crystalline graphite, diopside, and quartz. The marble also contains numerous inclusions of a variety of different rock types. Observations such as these led Ebenezer Emmons in his 1842 New York State Monograph “Geology of the State of New York: Survey of the Second Geological District” to state “Of Adirondack rocks we know little of their origin except for the marbles which surely are igneous” (McLelland, pers. comm.). Climb on top of the outcrop for a clear view of a unmetamorphosed, Mesozoic age mafic dike (Fig. 12). There is a millimeter scale chilled margin on the dike. To the south the marble is in contact with a biotite-garnet-quartz-plagioclase±sillimanite gneiss with quartzofeldspathic leucosomes.
As you proceed south from the marble and paragneiss you will encounter a number of different rock types including, garnet-bearing, quartzofeldspathic gneiss, charnockitic gneiss, numerous marbles, calcsilicate rocks with orange grossular garnet and wollastonite, and hornblende-pyroxene plagioclase gneiss.
At the southern end of the outcrop (UTM NAD83: 626522 E, 4832092 N) are biotite-garnet-sillimanite-K-feldspar-plagioclase-quartz gneisses. On some foliation surfaces these rocks contain beautiful matchstick-sized sillimanite. The sillimanite crystals are aligned forming a strong lineation trending approximately 15/135. We are in the process of mapping thin sections from this exposure to look for monazite. We do have monazite data from similar rocks just to the south that we will discuss at the next stop.
OHM 2019: UTM coordinates given in the text are not consistent with map or road log.

STOP 8. Lineated sillimanite gneisses, mylonitic granitic rocks, and the East Adirondack Shear Zone
There is an outcrop paragneiss immediately to the south of the parking area on the east side of the road. This outcrop contains interlayered garnet-biotite-sillimanite gneisses with greenish calcsilicate lithologies. The well exposed foliation surfaces of the sillimanite-bearing gneisses commonly have well-lineated, coarsely crystalline sillimanite. The lineation is gently plunging to the southeast, similar to that in the last outcrop.
Follow the outcrop to the south remaining on the east side of Hwy 22 for the time being. Note the interlayer folds in the upper part of the roadcut on the east. This exposure illustrates an older foliation that was transposed into a new foliation (Fig 13). Perhaps this is an example of a Shawinigan S1 transposed into a younger S2.
Across the highway is an asymmetric mafic boudin (Fig. 14) with a long tail that continues to the south in the upper part of the roadcut. Although it is clear the boudin was involved in a period of ductile deformation, the roadcut 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 above are mylonitic, LS-tectonites with megacrystic K-feldspar porphyroclasts (Fig 15). 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 ziron age obtained using the SHRIMP-RG at Stanford, of the granite in figure 15 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 suggest that most of the strain in the rock is therefore related to Ottawan compression 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. From this they concluded that this rock experienced both compressive stresses and extensional stresses but the extension event was younger and overprinted the effects of the compressional event.
Wong et al (2012) also report U-Th-Pb electron microprobe ages from in situ monazite crystals. The data are shown in figure 16. Note the two groupings of peaks in the 1100-1000 Ma age range. The older peaks, shown in blue, cluster between 1080-1050 Ma. These data are from analyses of the outer cores of monazite crystals. They are interpreted to represent monazite growth during Ottawan compression and metamorphism. The younger peaks in red are less than 1050 Ma. These ages come from analyses of the tips of monazite crystals that are elongated in the extensional direction. Wong et al (2012) interpret this as monazite growth during extensional collapse following peak Ottawan compression. 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).
It is interesting to compare this 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 a strong Shawinigan, peak-Ottawan, and post-Ottawan signature and limited evidence for monazite growth during AMCG times despite AMCG rocks being found in the outcrop.

STOP 9. Mylonitic, migmatitic gneiss
This is a beautiful exposure of a garnet-biotite-plagioclase-K-feldspar-quartz ± sillimanite gneiss (Fig 17). Most of the white layers are interpreted as leucosomes that formed during anatectic melting via a biotite dehydration reaction such as:
Biotite + plagioclase + sillimanite + quartz → K-feldspar + garnet + melt.
Much of garnet in the leucosomes was probably produced via this reaction.
The strongly attenuated nature of the leucosomes, along with also remnants of pegmatite dikes that are now sheared out in the foliation plane (Fig. 17) attest to the significant strain recorded by these rocks. The rocks have the same strong southeast-trending, gently plunging lineation that we have seen in the last two stops. We suggest that some of the strain in these rocks may be the result of post-Ottawan extension collapse in the EASZ.
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 18 shows our preliminary U-Th-Pb electron microprobe monazite data from this outcrop. The data shown in figure 18 are from 9 monazite crystals in a single thin section. Our data show monazite growth from approximately 1080 Ma through 1000 Ma. The data are consistent with recrystallization during the Ottawan and post Ottawan extension and in good agreement with the results of Bickford et al. (2008). It is interesting to note the lack of any Shawinigan or AMCG ages when the rocks approximately one mile to the north have a strong Shawinigan signature. 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.

B3:GEOLOGY AND PETROLOGY OF THE MONT ROYAL PLUTON, MONTREAL
Nelson Eby
Environmental, Earth and Atmospheric Sciences
Lowell, MA 01854
The Cretaceous Monteregian Hills (MH) petrographic province consists of a group of small alkaline plutons, and associated dikes, extending along a linear trend from approximately 30 km west of Montreal, Canada, to approximately 60 km east of Montreal (Fig. 1). Mont Megantic, located approximately 150 km east of Montreal has been included in the MH province (because of its geographic location), but geologically Mont Megantic belongs to the Younger White Mountain petrographic province of New Hampshire, Maine and Vermont. The MH plutons are largely composed of silica-undersaturated alkaline rocks: carbonatite and other strongly silica-undersaturated lithologies occur at Oka, and nepheline normative gabbros and pyroxenites, nepheline-bearing diorites, and feldspathoidal-bearing syenites and monzonites occur at the other MH plutons. Slightly silica-saturated quartz-bearing syenites are found at Monts Brome and Shefford. Philpotts (1974) and Eby (1987) have summarized the geology and petrology of the MH province.
The White Mountain (WM) petrographic province (Fig. 1) consists of two groups of plutons (Eby, 1987), one group (Older WM) intruded during the Jurassic and the other (Younger WM) during the Cretaceous. In contrast with the MH province, all the Younger WM plutons, with the exception of Cuttingsville, are silica-saturated to silica-over-saturated and the dominant lithologies are gabbro, diorite, syenite, quartz syenite, and granite. Rhyolite and basalt are found at Ossipee, a classic ring complex. Of note is that the silica-under-saturated lithologies occur to the west of Logan’s Line (which represents the boundary between the Precambrian Grenville terrane and the folded Appalachians) while only silica-saturated lithologies are found to the east. The older WM plutons are dominantly composed of syenite, quartz syenite, and granite. Mafic rocks are generally sparse and are only abundant in the Pliny Range and Belknap Mountain complex (both contain diorite and monzodiorite). Silica-under-saturated rocks, nepheline syenite, occur at Red Hill and Rattlesnake Mountain, ME.

STOP 1. Voie Camillien-Houde traverse. Contact with country rock, pyroxenite and gabbro outcrops
Start the traverse across from lower entrance to the parking lot. The contact between the plutonic rocks and the Utica shale is exposed at this location. Also at this location the underlying Trenton limestone is in faulted contact with the shale. The shale has been metamorphosed to a hornfels and the limestone is slightly metamorphosed. Locally the limestone is intruded by camptonite dikes surrounded by narrow thermal aureoles. Return to the parking lot (look both ways before crossing) and proceed to the upper end of the parking lot and up the steps. Note the large xenolith of Utica Shale halfway up on the right side of the steps. A short walk beyond the steps leads to good exposures of gabbro with steeply dipping layering. Return to the parking lot and walk up the road. In places you will need to walk on the edge of the road. Be careful!
There is an almost continuous exposure of the gabbro-pyroxenite sequence which is cut by multiple felsic dikes. At the end of the traverse, opposite the entrance to the Cimetière Mont-Royal, part of the limestone screen is exposed. The intensely deformed metamorphosed limestone consists of a mixture of calcite marble and a greenish skarn of eipidote, tremolite, zoisite, and diopside.

STOP 2. Belvedere traverse. Pyroxenite and gabbro cut by various mafic and felsic dikes
This traverse is along the road leading to the Belvedere. The pyroxenites and gabbro are intruded by a variety of mafic and felsic dikes. The feldspathoidal monzonite and feldspathoidal syenite dikes are light gray in color, medium-grained, and vary in width from a few millimeters to a few tenths of a meter. The lithologies are distinguished on the basis of the alkali feldspar/plagioclase ratio, a distinction which is not apparent at the outcrop scale. Late stage medium- to coarse-grained quartz-bearing syenites, which weather to a distinctive salmon pink color, cut the feldspathoidal monzonites and feldspathoidal syenites. This traverse ends at the Belvedere. On a clear day there is a superb view of the Monteregian Hills plutons from this location.

STOP 3. Gabbro and pyroxenite
Large outcrops of gabbro in wooded area across from RCMP station.

STOP 4. Foliated leucogabbro
Foliated leucogabbro. Foliation is essentially vertical and the outcrop is cut by a felsic dike. The leucogabbro is distinguished on the basis of a lighter color index due to a greater amount of feldspar. Olivine is found in some specimens. Amphibole is much less abundant relative to the more mafic units.

STOP 5. Nepheline-bearing diorite
Nepheline-bearing diorite outcrops are found in the woods next to the grave site. The outcrops extend outside of the cemetery and at the time this field trip guide was written could be accessed through a break in the fence. Sample RY3 and RY5 (Table 1) were collected from this location. RY3 (Fig. 7) is medium-grained and shows a subparallel alignment of plagioclase which is interpreted as a flow texture. RY5 is similar, but coarser-grained.

STOP 6. Île Ste. Hélène diatreme breccia
The large hill on Île Ste. Hélène is composed of diatreme polymictic breccia (Figs. 19, 20). Many of the buildings (Fig. 21) on the island were built using the breccia and provide some of the best exposures (not to be sampled). The field trip stop is an abandoned quarry which was the source for some of the building material.
A large variety of clasts can be found in the breccia. These include fossiliferous shale and limestone clasts which contain Devonian brachiopods. The youngest exposed sediments in the St. Lawrence Lowland are lower Silurian in age and thus these clasts represent some of the missing overlying sedimentary column. Other clasts include Trenton limestone, Utica shale, and metamorphic basement clasts. The breccias exposed on Île Ste. Hélène do not contain an igneous matrix. However, the comminuted material does have an elevated Ti content and the presence of analcime, perovskite, and apatite suggest there is an igneous component. Hydrothermal activity subsequent to emplacement resulted in veins carrying calcite, analcime, apatite, and rare Ti-minerals.

B4: THE MIDDLE ORDOVICIAN SECTION AT CROWN POINT PENINSULA, NEW YORK
Bruce Selleck
Colgate University
Hamilton, NY 13346
bselleck@mail.colgate.edu
Charlotte Mehrtens
University of Vermont
Burlington, VT 05405
charlotte.mehrtens@uvm.edu
The Chazy, Black River and Trenton Groups are a well studied sequence of fossiliferous limestones, dolostones and sandstones in the Champlain Valley of New York, Vermont and southern Quebec. These rocks record shallow water cyclic sedimentation in the foreland basin of Laurentia prior to and during the initial stages of the Taconic Orogeny. This field trip reviews the stratigraphy of these units, as well as the evidence for reconstructing depositional environments and cyclic sea level changes.
THE CROWN POINT HISTORICAL SITE
This field trip takes place at the scenic Crown Point Historical Site (Figure 1), a no-hammer, no-collecting locality. Approximately 120 meters of section is exposed in a combination of excavations, ridges and shoreline exposures that dip ∼8° northwest. The overall stratigraphy (Figures 2, 3) from the Crown Point Formation (Chazy Group) to the Glens Falls Formation (Trenton Group) records the Middle Ordovician transgression seen throughout the Appalachians. Correlation of this stratigraphy exposed at Crown Point to the type Trenton in central New York State is provided by Prasopora spp. (Mehrtens and Barnett, 1979). Conodonts collected from strata to the north and east (Harris, pers. comm. and Harris, et al., 1995; Roscoe, 1973) provide additional age control. Detailed sedimentologic studies of these units have been published by Bechtel and Mehrtens (1995), Mehrtens and Cuffey (2003) and MacLean (1986) and Selleck (1988).

STOP 1. The Redoubt
The Crown Point Reservation Historical Site is a no hammer, no collecting outcrop.
Approximately 6 meters of burrowed, slightly dolomitic, thin to medium bedded bioclastic packstone and grainstone is exposed at this stop. Some beds are relatively well-sorted grainstones with sharp bases which are interpreted as tempestites. The same interpretation is made for the grainstone intraclast-rich bed visible in the low ledge at the southeast corner of the ditch. Approximately 3 meters above the base of the section abundant Girvanella algal oncolites are common allochems. Nuclei of fossil fragments are visible within some oncolites. In other beds fossils are relative abundant, and are best seen on slightly weathered bedding surfaces. Trilobite fragments, brachiopods, bryozoans, pelmatozoan fragments, nautiloids and large Maclurites magnus are present (exterior shells and opercula). Dolomite occurs in along shaley seams and in burrow fill.
The relatively high faunal diversity, general bioturbation, and storm-related sedimentation all point to a low energy shallow subtidal environment at depths slightly below fair-weather wave base. The abundant algal oncolites and discrete calcareous algal fossils (e.g., Hedstromia spp.) suggest depths well within the photic zone. A possible modern analogue is found in the mixed mud and sand shelf to the west of the emergent tidal flats of Andros Island, Bahamas, as described by Purdy (1963).
We interpret the wavy, irregular dolomite laminae as the result of dolomitization of lime mud, followed by compaction and pressure solution of calcite that produced irregular clay and dolomite-rich stylocumulate seams. Preferential dolomitization of burrows is due to contrasts in permeability of burrow-fill versus burrow-matrix sediment. The burrow-fill material retained permeability longer during diagenesis and allowed more pervasive dolomitization. In similar facies exposed on Bullwagga Bay (the west shore of the Crown Point Peninsula), modular limestone with shaley dolostone seams and stringers are present. The limestone nodules appear to have been cemented prior to significant burial compaction, whereas the shaley dolostone material was compacted around the cemented limestone. The early-cement limestones were resistant to dolomitization. This sort of fabric selective dolomitization is common in the Chazy and Black River Groups throughout the Champlain Valley.

STOP 2. Ridge of outcrop running NE from entrance gate to highway
WARNING: poison ivy is common along this ledge
Cross-stratified coarse bioclastic grainstones are well-exposed near the main gate along the entrance road and adjacent ridge. Nearly three meters of section form a prominent belt parallel to strike, extending from the entrance road to the main highway. Foreset cross-strata show bipolar dip directions. Angular quartz and feldspar grains are concentrated in some laminae, and along stylolites (Figure 10). The carbonate particles are dominantly sub-rounded abraded pelmatozoan plates with gastropod and brachiopod fragments. Large Maclurites fragments and grainstone intraclasts are present on the upper bedding surfaces of these ledges.
We propose that the environment of deposition of this facies was a shallow subtidal wave and current reworked bioclastic sand shoal. Active transport of abraded grains may have been accomplished by tidal currents (as suggested by the bi-directional cross-strata) or by storm-generated currents that produced complex, anastomosing patterns of cross-strata and intervening reactivation surfaces. The lack of burrowing and well-preserved whole-shell body fossils may be due to the inhospitable shifting sand substrate. A modern analogue for this environment would be the unstable sand shoals of the Bahamas Platform (Ball, 1967), as the scale and style of cross stratification are similar. Oxley and Kay (1959) report that similar strata in the northern Lake Champlain Valley are oolitic.

STOP 3. Low ledges adjacent to the entrance road (Picnic Pavilion ridge)
About 50 meters north of Stop 2
Brown-weathering, slightly shaley dolostone exposed here contains lenses and stringers of fossiliferous lime packstone and wackestone. The fauna is similar to that at Stop 1, with trilobites, small brachiopods and Maclurites common. The environment of deposition is assumed to be similar to that of Stop 1, however lacking evidence of storm activity. Note that some of the fossils are almost entirely encased in dolomite, which is assumed to be of replacement origin here.

STOP 4. SE moat of Fort Crown Point
Approximately 3 meters of thickly laminated limestone and dolostone of the Valcour Formation are exposed in the southeast moat of the British Fort. The dominant facies here is alternating 0.5-2.0 cm thick laminae of limestone and dolostone, we term “ribbon rock.” The limestone ribbons are very fine-grained peloid grainstones or “calcisiltites” and appear blue-grey color on slightly weathered surfaces, and as indentations on more deeply weathered surfaces. The dolostone ribbons weather tan-brown in color, and consist of an interlocking mosaic of 20-300 micron dolomite crystals of replacement origin. Quartz silt grains are present in the dolostone ribbons, versus medium to fine-grained quartz sand in the limestone ribbons, suggesting the limestone horizons were slightly coarser-grained than the dolostone when deposited.
An erosional surface with 10 to 20 cm of relief is exposed near the base of the south wall of the moat (Figure 11). Similar erosional surfaces occur within this facies in other Chazy exposures in the Champlain Valley. We interpret these to represent micro-karstic solution surfaces on a tidal rock platform that developed during subaerial exposure of cemented limestone. Typically, the rock below the surface is mostly calcite limestone, suggesting that cementation and diagenetic stabilization of the limestone occurred prior to development of the erosional surfaces. Overlying rock horizons contain more dolomite. Maclurites shell hash can be found in pockets on the erosional surface. Dolomitized burrows cut across the limestone ribbons in some parts of the outcrop. Trough cross-strata filling low scour surfaces are also visible.
On the less weathered prominence on the SE corner of the moat, shallow scours containing brachiopod and gastropod debris can be seen. Intraclasts or pseudoclasts of limestone in dolostone are also present. Some “clasts” appear to be cored by dolomitized burrows.
We interpret the stratigraphy seen at this stop as representing a tidal flat to shallow subtidal facies. The alternating limestone/dolostone ribbon rock represents rhythms of slightly coarser-grained (limestone) and finer-grained (dolostone) sediment deposited on the lower reaches of a tidal flat, similar to the rhythmic bedding described by Reineck and Singh (1980) from the mud/sand tidal flats of the North Sea. These coarse-fine alternations might also reflect storm-related, ebb-surge deposition (Nelson, 1980). Early cementation of the slightly coarser-grained limestone ribbons made this lithology less susceptible to dolomitization, which affected the finer-grained muddy ribbons that became dolomitized. Variations in the intensity of burrowing is interpreted to reflect subtle differences in duration of subaerial exposure of the tidal flat and/or the extent of reworking by tidal currents. Limited in situ faunal diversity is to be expected in the tidal flat setting, where organisms are stressed by salinity fluctuations. The absence of mudcracks or evidence of evaporite minerals may indicate that only the lower portion of a humid climate tidal flat system is preserved here.

STOP 5. Parade Grounds near barracks
As we enter the parade ground from the southwest corner of the moat, note the array of carbonate rocks used in construction of the barracks walls. Chazy, Black River and (rarely) Trenton lithologies can be identified. Restoration of the barracks began in 1916 and in 1976 the New York State Division for Historic Preservation undertook protection and stabilization of the ruins.
The low rock pavement just north of the barracks is within the upper part of the “ribbon rock” unit seen at Stop 4. Immediately up-section, cross-stratified grainstone beds are exposed. Coarse-grained quartz and feldspar sand is easily seen on weathered surfaces. Trough cross-strata and herringbone co-sets of planar-tabular cross-strata are visible on the vertical surfaces. Large, angular clasts of slightly dolomitic grainstone and Maclurites are present on bedding surfaces. We interpret this facies as representing a current-dominated sand shoal environment, similar to that seen at Stop 2.

STOP 6. West Parade Grounds
Bechtel and Mehrtens (1993) suggested that the sandstone unit in the westernmost parade ground is the basal sandstone of the Black River Group, an interpretation which differs from that of Speyer and Selleck (1988), who suggested that this unit was part of the underlying Chazy Group. In thin section this sandstone is a quartz-rich sandstone, very poorly sorted, and containing fewer lithic fragments and phosphatic fragments than stratigraphically lower Chazy sandstones. Visible at the very easternmost portion of this ridge is a buff-colored dolostone bed containing pockets of quartz sand (burrow infills?). These basal sandstone and dolostone lithologies are very similar to those described by Walker (1972) in his description of the Pamelia Formation at its type locality in north-central New York. Alternatively, placement of the sandstone unit within the Chazy Group is consistent with the common presence of coarse-grained quartz and feldspar sand within the Chazy at the Crown Point Preserve, whereas the siliciclastic material in the Black River Group at Crown Point is mainly silt and clay. Whatever the stratigraphic placement of the sandstones and dolostones here, it marks an interval when sands were transported from a nearby source area to this possible peri-tidal setting. This interval was followed by marine reworking of the sand and deposition of fine-grained limestone of the basal Black River Group.

STOP 7. Northeastern moat
WARNING: poison ivy is common along the base of this exposure
There is approximately one meter of covered interval between Stop 6 in the parade grounds and Stop 7 in the northeastern moat. Along this wall are exposed several meters of the lower Black River Group (Lowville Formation, House Creek Member) which in Vermont is termed the Orwell Limestone. At the base of this exposure a series of stylolitized gastropod-bearing (Liospira) wackestone beds are overlain by thicker beds of Phytopsis-burrowed aphanitic mudstones. This sequence can be interpreted as a shallowing-up cycle (SUC) consisting of subtidal overlain by peritidal lagoonal muds. Examine the sharp contact of the aphanitic mudstone with the overlying wackestones, a contact which in thin section appears to be a firmground (Bechtel, 1993). These SUCs also comprise the base of the Black River Group at other localities in the Champlain Valley. There are three motifs of repetitious bedding that occur in the Black River Group and the cycles seen here at the base are the the thickest, occurring at a macroscopic scale, interpreted to represent 4th order (10,000 to 100,000 years) or smaller cycles. Examination of Phytopsis burrows in thin section (Figure 12) reveals that many are filled with graded (fining-up) geopedal silt, evidence of cementation in the meteoric vadose zone.
Continuing up section, several thick packstone beds are exposed. More detailed examination of these beds reveals that they consist of alternating one to six centimeter thick intraclast and oncolite-rich packstone horizons interpreted as tempestites, interbedded with fossiliferous wackestone/packstone horizons. The tempestites consist of graded and crudely imbricated intraclasts and skeletal fragments. Note the nature of the upper and lower contacts of these horizons. The second common bedding motif in the Black River Group are these tempestite horizons interbedded with in situ fossiliferous muds.
The uppermost third of the outcrop appears to be a massive bed of limestone, however closer examination also reveals small scale cycles of alternating wackestone/packstone and grainstone, the third motif of bedding in the Black River strata. These cycles are characterized by a base of thinly laminated or cross laminated grainstone horizons 0.5 to 1 cm thick, overlain by fossiliferous wackestones and packstones. In thin section the bases of the grainstones can be identified as firmgrounds, recognizable by the truncations of allochems and cements present in the underlying mud (Figure 13). The very top of this exposure (best seen at the next stop) exhibits a burrow mottled fabric with selected dolomitization of many burrows. Tetradium occurs in life position in these horizons.

STOP 8. East-west ridge across the service road
A black chert layer near the top of Stop 7 provides the correlation to Stop 8, the outcrop across the service road. The limestone beds on this ridge commonly consist of alternating wackestone/packstone and planar to cross laminated grainstone beds, as seen at Stop 7, however bedding plane exposures permit identification of many fossils in these, the most faunally diverse beds in the Black River Group. Specimens of gastropods (Liospira, Lophospira, Hormotoma), Lambeophyllum, Tetradium, stromatoporoids, the bivalve Cyrtodonta, the brachiopod Stromphomena sp. and cephalopods are recognizable. This ridge exposure is most notable for its bedding plane exposures of Tetradium and Lambeophyllum. It is interpreted as representing a wave baffle margin lithofacies similar to that described by Walker (1972) at the Black River type section.
A prominent reentrant in the exposure along the path to the north marks a volcanic ash horizon, or K-bentonite (Figure 14). This interval is currently under study, but likely represents the “M/H” bentonite (Brett and Cornell, 2000) that falls within the M4A (upper Lowville Formation) sequence. This prominent ash layer is traceable from the Roaring Brook section in the Black River Valley through outcrop and core in the Mohawk Valley and normal-fault bounded Wells outlier within the south-central Adirondack Massif.
A third cycle motif present in this interval, laminated grainstone overlain by bioturbated wackestone, is interpreted to represent smaller scale 4th order cycles, which could be the result of facies mosaicing and/or small scale base level changes.
There are at least two distinct types of chert occurrences in the Black River Group. One, the infilling of horizontal burrows, is fabric selective. Chert also occurs less frequently as broad bedding plane parallel sheets. The uppermost chert horizon on this ridge, traceable down to the shoreline, is of this latter variety. Clearly, there was a significant source of silica available for chert formation, perhaps a combination of silica derived from sponges (Tetradium?) and bentonite alteration. In thin section the chert cross cuts all previous cements, including late fracture-filling calcite, and it is therefore the youngest form of diagenesis present in these rocks.

STOP 9. Quarry to the north
WARNING: be extremely careful around the quarry as the thick algal scum in the quarry water obscures where the grass ends and the quarry wall drops off.
The older weathered south walls of the quarry show, by color differentiation, two cycles. There are 3 to 4 cm thick beds of planar laminated skeletal and peloidal hash overlain by burrowed wackestones overlain by intraclast-rich horizons. These cycles represent our third motif of Black River bedding. Interbedded with these cycles are tempestite couplets of mudstone/wackestone and fossil hash layers in which brachiopod-rich layers are abundant. The abundant quarried blocks lying about provide the opportunity to look for cycles, and from these, topping directions.

STOP 10. Shoreline of eastern Bulwaga Bay
The uppermost horizons in the quarry can be traced down to the shoreline to the north where the uppermost Black River strata can be seen (Chaumont Formation). Shoreline bedding planes exhibit horizontal traces of Chondrites and opercula of Maclurites. There is a thin covered interval of approximately one meter to the basal beds of the Glens Falls Limestone of the Trenton Group. The Glens Falls is characterized by thin beds of nodular to wavy bedded wackestones, mudstones and rare grainstones. Bedding planes along the shoreline contain mostly Chondrites and Helmenthopsis burrows, however as one moves up section, recognizable fragments of Cryptolithus, Isotelus, orthid brachiopods, Stictopora, and Prasopora simulatrix can be found; the latter is important because it permits the correlation of the lower Glens Falls here in the Champlain Valley to the lower Denley Limestone at the Trenton type section in central New York (Mehrtens and Barnett, 1979).
MacLean (1986) interpreted the lithofacies of the basal Glens Falls visible here to represent sedimentation in a shallow subtidal environment periodically influenced by storm activity. In thin section the nodular and wavy bedded wackestones appear thoroughly bioturbated, a process which would influence and enhance subsequent differential compaction. Grainstone beds exhibit more planar bases with basal skeletal fragment lags or finely crushed debris of brachiopod, trilobite and crinoidal material and capped by carbonate mud. MacLean interpreted these as tempestite deposits. Moving further up section (in horizons not seen at Crown Point) the Glens Falls records progressively deepening conditions from the subtidal, storm influenced conditions seen here to bioclastic turbidites separated by authochthonous shale horizons. For those familiar with the lower Trenton Group localities in central New York State (ex., Rathbun Brook, City Brook, Inghams Mills), the paucity of fossiliferous bedding planes here at Crown Point is noteworthy. The overall fine grain size and ichnofauna suggest that bathymetry increased significantly and rapidly from the Black River into the Glens Falls, a transition that might reflect not only rising sea level but base level changes as well. The sedimentologic and faunal transitions from the Glens Falls to the overlying Cumberland Head Argillite and the Stony Point Shale are much more gradational than that of the Black River/Glens Falls contact.

B5: AN OVERVIEW OF THE EARLY PALEOZOIC STRATIGRAPHY OF THE CHAMPLAIN VALLEY OF NEW YORK STATE
James C. Dawson
Earth & Environmental Science
State University of New York,
Plattsburgh, NY, 12901
TRIP ABSTRACT
This, Sunday only, trip provides participants with an opportunity to visit some of the well-exposed, classic Early/Medial Cambrian to Medial Ordovician clastic and carbonate shelf sequences and the Medial Ordovician foreland basin carbonates and calcareous shales. Stops include the Potsdam Group (Lower/Middle and Upper Cambrian), the Theresa Formation and the Beekmantown Group (Lower Ordovician), the Chazy Group (Middle Ordovician) and the Trenton Group (Middle Ordovician). The trip essentially repeats the author’s 2002 Joint New England Intercollegiate Geological Conference/New York Geological Association trip (Dawson, 2002) and is ideal for high school teachers and undergraduates seeking an introduction to these rocks.
The field trip begins at the Comfort Inn, near Exit 37, I-87 ‘the Northway’, Plattsburgh, NY at 8:30 AM and will end about 4:00 PM at the Comfort Inn, Plattsburgh, NY. This trip visits many of the localities that are commonly visited in the undergraduate classes at SUNY Plattsburgh. The field trip description is largely based on the work of Fisher (1968) and Isachsen, et al. (2000) and is not based on any significant research by this author.

B5: AN OVERVIEW OF THE EARLY PALEOZOIC STRATIGRAPHY OF THE CHAMPLAIN VALLEY OF NEW YORK STATE

Note the 'old base' stone buildings
On your right you will note three stone buildings constructed by the U. S. Army before the area became Plattsburgh Air Force Base (PAFB). The oldest building, the Old Stone Barracks, is the one furthest from the road. Construction on this started in 1839, but due to construction and funding issues, the building was not occupied until 1843. The two younger replica barracks date from the 1930s. All three buildings are constructed of local Keeseville Sandstone.

STOP 1. Montreal Member of the Glens Falls Limestone and Dike
Montreal Member of the Glens Falls Limestone of the Trenton Group and Lamprophyre (Monchiquite Dike) at the former PAFB Marina.
From the parking area walk down the obvious roadway to the lake shore where the former PAFB marina and warf can be seen on Marina Point. To reach the outcrop, hike, approximately 700 or so paces north along the cobble beach shore of Lake Champlain. The Pleistocene sediments forming the bluff between the railroad track and Lake Champlain along the way include a contact between earlier Lake Vermont varved and later Champlain Sea sediments. The Montreal Member and a cross cutting Lampophyre Dike are exposed to varying degrees depending on the water level of Lake Champlain. The Montreal Member is a medium-bedded limestone with some shale partings that weather to medium dark gray. Thin laminations within the medium-thickness beds have been interpreted as turbidity current deposits. Crab Island (the smaller island that can be seen in Lake Champlain two miles/3.2 km southeast of this stop) takes its name from the trilobite, Isotelus, that is commonly found in exposures of the Montreal Member on the island. Fisher (1968) mapped the Lamprophyre Dike as a monchiquite dike, containing phenocrysts of titanaugite with some biotite and barkevikite (a monoclinic amphibole), phenocrysts of olivine and titanomagnetite. After visiting the outcrop return to the vehicles.

Parking for Stop 1A

STOP 1A. (Alternate) Montreal Member
After parking, walk north along the paved pedestrian walkway for 0.3 miles/0.5 km to a pedestrian bridge on your right/east. Cross over the Canadian Pacific mainline railroad tracks on the pedestrian bridge and follow the path down to the lake shore. This will bring you to Sailor’s Point, the former PAFB beach and picnic area. The Montreal Limestone is exposed at the south end of the cobble beach. From the outcrop you can see Crab Island to the southeast and Marina Point to the south.

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

City Hall and Macdonough monument
Continue north on City Hall Place past the Plattsburgh City Hall on your left and the Macdonough (American naval commander during the Battle of Plattsburgh) Monument on your right.

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

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

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

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

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

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

STOP 6. Valcour Limestone of the Chazy Group at Sheldon Lane.
(THIS IS NOT A HAMMER STOP.)
This is the first of two opportunities that we will have to see the bioherms in the Valcour Limestone.
See the Stop 10 description for the details of what to see at this stop.

Road to old Quarry
Just past a faded red barn on the right/east there is an amateurish ‘stonehenge’ and an unpaved road. The road leads to the abandoned International Lime and Stone quarry where the Pamelia Dolostone, the Lowville Limestone and the Isle La Motte Limestone formations of the Black River Group (Fisher, 1968 – Figure 25) were once exposed. A few years ago much of the remainder of this outcrop and the quarry floor today consists of Valcour Limestone. Some of the Black River Group remains on the north side of the quarry on private property; but, this exposure is only accessible when the water in the quarry is frozen.

STOP 7. Crown Point Limestone of the Chazy Group at the Northway in Chazy
Maclurites magnus Le Sueur, Death Assemblage
(THIS IS NOT A HAMMER STOP.)
To reach the outcrop cross the open area to the north and enter the woods on a vague trail behind and to the right of the obvious transmission pole. The vague trail heads north a few yards in a white cedar forest and then bends left/west toward the Northway. The outcrop is close to the Northway, a short distance north of the Fiske Road Bridge and is illustrated in Fisher, 1968 (Figure 20). The Crown Point Limestone is mainly a medium to dark gray calcilutite. This locality exhibits an unusually large assemblage of the planispiral gastropod Maclurites magnus along with trilobite and brachiopod fragments. Be sure to note the insipient karst development.

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

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

STOP 9. Lacolle Melange at Coopersville.
The small outcrop is located adjacent to the south side of Route 9B and is figured in Fisher, 1968 (Figure 29). The breccia clasts include a wide range of sizes and the lithology of the clasts is predominately carbonate fragments, likely of Black River Group and Trenton Group origin. Also refer to the Stop 8 description.

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

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

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

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

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

STOP 13b.
Continue west on the Gadway Road and park at the next significant open area on the right/north. At the north end of the open area near a red posted sign there is a 4 meter long (10 cm wide) Protichnites track. Yochelson and Fedonkin (1993) illustrate several similar examples of this trace fossil in the Keeseville Sandstone of northern New York. Additional trace fossils can be seen near the Adirondack Nature Conservancy sign. Recent literature suggests that the trace fossil Protichnites may be the track of a eurypterid-like stem arthropod (Seilacher, 2007; Hagadorn and Seilacher, 2009) or by euthycarcinoid arthropods travelling in pairs (Collette and Hagadorn, 2010; Collette, et. al, 2012).

STOP 13c.
Park at the turn around loop. Walk about 100 paces beyond the turn-around loop and pass an arrow sign. On the left/south side of the road you can locate additional Protichnites tracks. One track is 1.4 meters long (11 cm wide) and some partial tracks cross over one another.

Bridge across Great Chazy River
Continue south on the Devil Den Road as it crosses the Great Chazy River. Just upstream from the bridge to the right/west is the McGregor Powerhouse, in Spanish architectural style, and the LaSalle Dam that were constructed by William H. Miner in 1923. Excellent exposures of the Keeseville Sandstone can be found on both sides of the river between the road bridge and the powerhouse.

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

Route 190, ‘the Military Turnpike’.
You will soon see one of the local wind farms set back on both sides of the road. These units feed their power to the nearby Power Authority of New York’s Duley, Altona, NY, substation.

Pass the historic Robinson Tavern on the left/east.
On the left/east side of the road you pass the overgrown ruins of Lewis Sage Robinson’s Tavern, erected in 1837. The opposite/west side of the road is the site of the Original Log Tavern erected in 1810. This 1810 tavern was visited by President James Madison in 1817.

STOP 15. Ausable Arkose of the Potsdam ss on the Military Turnpike.
Examples of the Altona Member of the Potsdam Sandstone can be found in and along the ditch from near the ‘Truck 0.5 Mile’ sign south to the near the bottom of the hill. The best exposures are toward the top of the hill. Grenville meta-anorthosite can be found just across the small drainage at the bottom of the hill.
The prominent hard reddish rock is a hematitic, micaceous quartzose dolostone. When highway crews clear the ditch it is also possible to see interbedded red shales; but, these tend to get grassed over in between ditch clearings.
About 2 miles/3.2 km east of here, SUNY Plattsburgh has a series of research hydrology wells on the William H. Miner Agricultural Research Institute’s Altona Flat Rock. One of these wells penetrated 235 feet of Altona Dolostone/Shale before entering the underlying Grenville meta-anorthosites and basalt dikes.

Pass the Rand Hill dikes on both sides of the road
Pass the Murtagh Hill Road on the right/west. Just south of the Murtagh Hill Road on both sides of Route 190 there are good exposures of Grenville meta-anorthosites that have been intruded by a series of diabase dikes, also of Grenville age, and locally known as the Rand Hill Dikes.

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

B6: ICE RETREAT AND READVANCE ACROSS THE GREEN MOUNTAIN FOOTHILLS: BOLTON AND JERICHO, VERMONT
Stephen F. Wright
Burlington, Vermont 05405
swright@uvm.edu
George E. Springston
Department of Earth and Environmental Sciences
Norwich University
Northfield, Vermont 05663
John G. Van Hoesen
Department of Environmental Studies
Green Mountain College
Poultney, Vermont 05764
This field trip visits several key sites in the towns of Bolton and Jericho Vermont that shed light on the glacial, ice-contact, and lacustrine environments that existed during a relatively short period of time during the retreat of the Laurentide ice sheet across this area. This area lies on the western slope of the Green Mountains, an area dominated by steep slopes and narrow stream valleys that are cut by the broad, low-gradient Winooski River valley (Fig. 1). Stop 1 is a section displaying evidence of at least one cycle of glacial retreat followed by readvance and eventual retreat. Stop 2 visits a working gravel pit where the internal structures of a large delta formed where glacial meltwater flowing south from Bolton Notch entered Glacial Lake Mansfield which occupied the Winooski River valley. Farther north in Bolton Notch Stop 3 visits a series of nested meltwater channels and other ice-contact landforms. The trip finishes at a large landslide section along the Lee River (Stop 4) which exposes the full transition from quiet water lake sediments through deltaic sediments where the Lee River delta prograded into the Coveville Stage of Glacial Lake Vermont. The location of all the field trip stops are shown in Figure 2 and more detailed maps accompany the road log. Many of the field stops described in this guide are located on private land. Permission from landowners, noted in the following guide, must be secured before visiting these sites.

B6: ICE RETREAT AND READVANCE ACROSS THE GREEN MOUNTAIN FOOTHILLS: BOLTON AND JERICHO, VERMONT

Vermont Meeting Point

STOP 1. Preston Brook Landslide
Stop 1 is a large landslide along the east side of Preston Brook in Honey Hollow, a steep, narrow N–S valley typical of the Green Mountains (Fig. 5). There are two ways of accessing this field site. If the water level in Preston Brook is low enough, the bottom of the slide can be accessed by driving 0.5 mile up Honey Hollow road (see above road log) and parking at a small pull-off where a red iron gate blocks an overgrown logging road. The base of the slide is approximately 200 m SE of the iron gate. There is no trail. When Preston Brook is too high to cross, the top of the slide can be accessed by hiking ∼1 km up the Catamount Trail to where a gently sloping alluvial terrace separates the trail from the top of the landslide.
Wright first observed this section in 2008 and as of this writing the exposures are still quite good. Please anticipate that they may not remain so in the future. Continuing slope movements change which parts of the section are visible at different times. Parts of the section are very steep and quite wet. Caution and sturdy/waterproof footwear are recommended.
The landslide exposes an approximately 30 m high section containing lacustrine sediments at three stratigraphic levels separated by both diamict and coarser-grained ice-proximal sediments (Fig. 6). The description that follows divides the section into “Lower,””Middle,” and “Upper” parts which are separated by covered intervals. Lacustrine sediments exposed in the “Lower” and “Middle” parts of the section are deformed. One objective of this stop is to ascertain whether structures in these deformed sediments formed by (1) a readvance of the ice sheet or (2) whether they formed during one or more landslides (debris flows) in this narrow, steep-sided mountain valley.
Lower Section
The lowest part of the section (stream-level) consists of fine-grained lacustrine sediments that are folded, faulted, and over-consolidated (Fig. 6). These deformed sediments are overlain by a boulder/cobble/pebble-rich diamict(?). This in turn is overlain by a boulder/cobble gravel that generally fines upwards into a sequence of sediments consisting largely of silt and very fine sand, with irregular lenses of coarser sand, pebbles, and cobbles (Fig. 6). Bedding is sometimes coherent, but also occurs as a slumgullion of lenses and irregular blebs or isolated fragments of silt/clay or rocks surrounded by finer grained sediments. The finer sediments are very compact and appear to be over-consolidated. Near the top of the lower section an approximately 10 cm thick bed of laminated silt and very fine sand has been both folded and apparently injected (dike-like) into the overlying sediments (Fig. 7). To the right (south) of the fold hinge bedding is continuous at least a short ways down section. To the left (north) of the fold hinge bedding has been pealed/delaminated away from the underlying sediment and highly deformed sediment occupies the space above the now overturned limb of the fold. Figure 6: Description and interpretation of the Preston Brook landslide section that extends upwards from the brook to an abandoned fluvial terrace (see Fig. 3 for location). The section exposes three separate units of lacustrine sediments separated by both diamict and glaciofluvial sediments. The lower two units are deformed as a consequence of emplacement of the diamict as till during a glacial readvance or as till remobilized during an underwater landslide. The uppermost part of the section consists of alluvium deposited when Preston Brook flowed across the sediment-filled valley into Glacial Lake Vermont.
Middle Section
The middle part of the section is largely covered and water-saturated, but small exposures consist of silt and clay layers (varves) that have been deformed into meter-scale folds (Fig. 6). The present limited exposure makes it impossible to know whether (1) the entire section is deformed or (2) whether the deformation is intraformational (i.e. the folds occur only within one stratigraphic horizon and are overlain by undeformed sediments).
Upper Section
The upper part of the section consists of diamict (till?/landslide debris?) overlain by rhythmically bedded (varved) very fine sand, silt, and clay (Fig. 6). While some layers contain intraformational slumps, particularly at the base, this section is undeformed. The lacustrine sediments are overlain by a thin veneer of alluvium which extends across a gently north-sloping terrace (between the landslide headwall and the Catamount trail) and marks the top of the section. An old alluvium-filled stream channel cuts across the uppermost section of lacustrine sediments (Fig. 6).
Interpretation
The Preston Brook valley (Honey Hollow), similar to other valleys in the Winooski River basin, was partially flooded by Glacial Lakes Winooski, Mansfield, and Vermont (upper Coveville stage) as the ice sheet retreated NNW (down) the Winooski River valley (Fig. 4). The top of the landslide scarp is at an elevation of ∼600 feet (∼185 m) which is at or below the elevation of all these lakes (Fig. 6). It is unclear which of the lacustrine sediment units exposed in this section were deposited in which of these lakes or within a lake that may have preceded the formation of Glacial Lake Winooski. In the interpretation that follows, we have suggested a sequence of events that is consistent with the sediments exposed in this section and our understanding of the glacial history in the Winooski River valley, but other interpretations are also possible.
The fine-grained lacustrine sediments at the bottom of the section indicate that at the time these sediments were deposited, the ice sheet had retreated down the Winooski River valley west of the mouth of Preston Brook in order to flood the valley with lake water. We suggest that this lake formed in the Winooski River valley before the Middlesex Readvance that Larsen (1999) referred to as Glacial Lake Merwin. The tilted, faulted, folded, and over-consolidated nature of these sediments, in conjunction with the overlying diamict, suggest that they were overridden during a readvance of the ice sheet into the Preston Brook valley, the Middlesex Readvance. This readvance occurred during an approximately 100-year long cold period between 13,900 and 14,000 yr b2k and is coeval with the Littleton Readvance in northwestern New Hampshire (Ridge et al., 2012, Thompson, 1999). During this readvance the ice sheet reclaimed most of the Winooski River drainage basin as evidenced by sections of deformed sediments in many of the river’s tributary valleys (Larsen 1999, 2001; Larsen et al. 2003; Wright 1999, 2002, 2015).
Glacial Lake Winooski began to form at the end of the readvance when the ice once again began retreating WNW down the river valley. The coarse boulder/cobble gravels that overlie the lower diamict suggest they were deposited near the mouth of an esker tunnel as the ice sheet retreated across this area. The fining up sequence of fluvial sediments that overlie the coarse gravels were most likely deposited in a subaqueous fan that became progressively finer grained as the ice sheet retreated west of the Preston Brook valley. Continued retreat of the ice established a quiet-water arm of Glacial Lake Winooski in the valley when the varved silt/clay sediments in the middle section were deposited. However, by the time the ice retreated far enough down the Winooski River valley to expose the Preston Brook valley only 2 km of additional ice retreat allowed Glacial Lake Winooski to catastrophically drain through the Huntington River and Hollow Brook valleys across two different thresholds to form Glacial Lake Mansfield (Fig. 4). Elsewhere in the Winooski River valley this initial lake elevation drop of ∼80 m is stratigraphically instantaneous and clearly recorded by an abrupt increase in grain size (Larsen et al., 2003). Very little of the middle part of the section is exposed, but we suspect that these fine grained, ice-distal, lacustrine sediments were deposited in both Glacial Lakes Winooski and Mansfield and, if better exposed, would show the transition between the two lakes.
The large folded layers of silt/clay visible in the exposed part of the Middle Section may have formed by slumping or they may have formed from shear stresses imparted when the upper diamict was emplaced. One interpretation is that the diamict formed from a debris flow (consisting of remobilized till mixed with lacustrine sediments) originating on the steep valley side that flowed into the lake-flooded valley. After entering the lake the debris flow triggered further slumping/deformation in the already-deposited sediments now exposed in the lower and middle parts of the section (Fig. 7). This slumping may have been triggered by the catastrophic drainage event associated with the transition from Glacial Lake Winooski to Glacial Lake Mansfield.
If the upper diamict is interpreted as a debris flow deposit, then the lacustrine sediments deposited on top of the diamict are most likely Glacial Lake Mansfield sediments as they were deposited in relatively quiet water and the elevation of Glacial Lake Vermont (Coveville Stage) in the valley is only slightly higher than the elevation of these sediments (i.e. when Glacial Lake Vermont occupied the valley this area would have been close to its shoreline).
Another interpretation is that the diamict is a readvance till and deformation in the underlying section was produced by movement of the overlying ice sheet. If this diamict was produced during the Middlesex Readvance, then all the underlying sediments are older than Glacial Lake Winooski and the thin section of lacustrine sediments lying above the diamict were deposited in Glacial Lake Winooski. It is also possible that this diamict records a more recent readvance of the ice sheet, although there isn’t any indication in the ice core records of another cold period capable of producing a readvance during the waning stage of Glacial Lake Winooski (∼13,700 yr b2k). Unlike the lower diamict, the upper diamict is overlain by sediments deposited in quiet water. If this upper diamict is a till deposited during a readvance, subglacial conduits did not funnel high-energy sediments in the valley as the ice was retreating.
The uppermost lacustrine sediments are cut by a large channel filled with a poorly sorted pebble/cobble/boulder gravel and are overlain by a thin layer of similar sediments interpreted to be historical Preston Brook alluvium. The abandoned stream terrace at the top of the section grades to the elevation of the Coveville stage of Glacial Lake Vermont (Fig. 6). This indicates that the valley was locally filled with at least 100 feet (30 m) of sediment when Preston Brook began incising its present channel when the elevation of Glacial Lake Vermont dropped from the Coveville to Fort Ann stage.

STOP 2. Bolton Notch Delta
The active gravel pit on the west side of the road is excavating deltaic sediments deposited by a north-flowing stream originating in Bolton Notch. This is a large delta and many of the foreset and topset beds are composed of very coarse gravels with some very large (m-scale) lag(?) boulders (Fig. 8). The size of the drainage basin is small which suggests that the high-energy stream that built the delta was fed directly by melting ice retreating up the notch and/or by drainage out a glacial lake that formed on the north side of the notch (Fig. 9). No deltas occur (or are not preserved) in other stream valleys farther east suggesting that the Bolton Notch delta formed quickly during a time span too short for significant non-glacial deltas to form in adjacent stream valleys.
Well-developed foreset and topset beds are usually exposed in the pit. These are overlain by a thin (∼2 m) sheet of medium to fine sand. Duck Brook has eroded a deep channel through this delta isolating an erosional remnant of this deltaic terrace on the west side of the brook. The elevation of the topset/foreset contact, measured using a Trimble GPS, is ∼233 m (± 1 m). With ∼4 m of isostatic uplift (based on an almost due north uplift slope of 0.7 m/km (Rayburn, 2004), this corresponds with the higher 229 m Gillette Pond outlet of Glacial Lake Mansfield (Larsen, 1987, Fig. 7).
The geometry of the ice margin at this time suggests that the Gillette Pond outlet was tenuous at best (Fig. 9).The medium to fine sand overlying the coarse topset beds of the delta was most likely deposited when the lake deepened for a short period of time, drowning the delta (i.e. the delta moved farther up the Bolton Notch valley and only fine-grained bottomset sediments here). A minor readvance of the ice sheet could easily have closed the Gillette Pond outlet allowing the lake elevation to rise for a limited period of time (Fig. 9). Similarly, a relatively small 1–2 km retreat of the ice margin uncovered a lower elevation outlet (Hollow Brook) which allowed the lake elevation to drop ∼25 m.

STOP 3. Bolton Notch
Bolton Notch is a north-south U-shaped valley bounded by near vertical bedrock cliffs on either side of a valley floor filled with an unknown thickness of talus, lacustrine, and fluvial sediments. A poorly drained part of the valley floor marks the drainage divide at an elevation of 1180–1200 ft (∼365 m) asl (Fig. 10). The objective of this stop is to observe and offer interpretations of (1) a series of terraces and channels that likely formed during ice retreat, (2) a group of kettles occurring immediately south of the gravel pit, and (3) ice-contact and lacustrine sediments exposed in the gravel pit. All of these features occur at or immediately north of the drainage divide. A detailed map of many of the small-scale landforms occurring near the drainage divide of the notch was made at an original scale of 1:4,000 using a hand-held GPS device (Fig. 11).
Striation Data
Striation measurements in the area surrounding the notch indicate a very consistent NW to SE direction of ice flow (Figs. 3, 10). Few measurements have been made on the steep valley sides; however some striations on the lower slopes show deflection to a NNW to SSE direction suggesting that ice flow rotated parallel to the valley as the ice sheet thinned, although the length of time the ice flowed in this direction was probably relatively short. Therefore, despite the distinctive “U” shape of the valley, this form may be more strongly linked to differential weathering and erosion of a relatively weak lithology underlying the valley rather than erosion by moving ice.
Terraces
A nested group of terraces occur along the eastern side of Bolton Notch, near the drainage divide. They are all aligned NE–SW, oblique to the N–S orientation of the valley, and all slope gently to the southwest (Fig. 11). The southernmost terraces are the highest and they progressively step down to lower elevations farther north. The lower-elevation terraces cross-cut adjacent higher-elevation terraces. The elevation difference between adjacent terraces typically ranges between 2–5 m. The most northerly of the terraces leads into a distinctive channel (now flooded) that occupies most of the area at the drainage divide. Farther south this pond drains and forms the modern stream channel that cross-cuts all of the higher terraces (Fig. 11). Most of these terraces are floored with till, but some also contain poorly rounded cobbles and boulders, particularly notable where these materials have been incorporated into stone walls.
These nested terraces are interpreted to be abandoned glacial meltwater channels that formed from water discharging from a tunnel beneath the glacier. This water flowed between the retreating ice front and the steep valley side or older, higher terraces. From south to north, these different terraces, at successively lower elevations, mark the position of the retreating glacier, perhaps at yearly intervals (Fig. 11). It was this glacial meltwater, flowing south through the notch and towards the Winooski River valley, that supplied most of the sediment that accumulated in the Bolton delta (see Stop 2).
Kettles
Several kettles occur in the bottom of the valley south of the gravel pit and two of the larger ones are shown on the detailed map (Fig. 11). Several of the abandoned meltwater channels abruptly drop into the pond suggesting that the pond too is a kettle or network of kettles and streams emanating from the glacier flowed directly across stagnant ice. The implication is that as the retreating ice thinned across the drainage divide, it thinned enough so it could no longer flow and became stagnant and disconnected from the thicker, still active ice front north of the divide.
Gravel Pit Geology
An old gravel pit lies along the north side of the road, at the northern end of Bolton Notch (Figs. 10, 11). The pit exposes sediments deposited in both an esker and a small glacial lake that formed as soon as the ice front retreated north of the drainage divide. The trace of the esker is shown on Figure 11. While many of the pit faces are slumped, an amazing vertical walled “canyon” has formed between two levels of the pit and offers excellent exposures of sedimentary structures (Fig. 12). Material eroded from this canyon has formed a large alluvial fan immediately below the northern end of the canyon.
Much of the esker has been quarried away, but many large rounded boulders occur on the floor of the pit at its southern end and very coarse sediment with additional boulders occurs along the pit’s slumped face immediately below and northeast of the road. Many of the boulders deposited in the esker tunnel are high-grade metamorphic rocks (Grenville-age gneisses) that are distinctly erratic to northwestern Vermont and were transported here from the northern Adirondack Mountains or the Laurentian Mountains north of Montréal.
The bulk of the material currently exposed in the pit consists of medium to very fine sand and silt with occasional, relatively thin beds of coarse sand and pebble gravel. Dropstones are common. Cross-bedding and climbing ripples, frequently in beds >1 m in thickness, occur in many beds and all indicate that currents were dominantly flowing to the south, towards the notch. Structureless beds of medium to fine sand imply that these beds are slumped. Coherent blocks of bedded sediment have also slid off the steep side of the esker into deeper water (Fig. 13). There is a general fining-up pattern where the thickest beds of coarse sediment occur in the deepest parts of the pit and the finest interlayered fine sand/silt horizons occur at the uppermost part of the pit. Soft sediment deformation by slumping and/or differential loading is extensive in some beds.
The fine lacustrine sediments exposed here indicate that a small glacial lake existed here once the ice sheet retreated north of the drainage divide (Figs. 10, 11). The lowest elevation channel, leading to and through the pond, marks the final outlet to this lake at an elevation of ∼1,190 ft/363 m (Fig 11). This lake was likely small and short-lived as less than 1 km of further ice retreat to the northwest would have uncovered a much lower outlet to the Winooski River valley via the Indian Brook drainage (outlet elevation 920 ft/280 m).
A ground penetrating radar (GPR) survey shows that in the central part of the pit an esker is buried by a sequence of lacustrine sediments that outcrop in most of the pit (Fig. 14). Another GPR survey line emphasizes the continuity of these beds (Fig. 15). The 100 GHz antenna shows coherent reflectors to a “depth” of 300 ns (Fig. 14) whereas penetration with the 200 GHz antenna does not extend much below 150 ns (Fig. 15). However, the 200 GHz antenna shows sedimentary structures in much more detail.
The large volumes of sand, frequently deposited as meter-thick beds of climbing ripples from south-flowing currents, implies that sediments in this glacial lake originated from a subglacial drainage system, i.e. meltwater and sediments discharging through a tunnel, which formed a subaqueous fan that fined upwards as the ice front receded to the north. Given the small size of the lake basin and the high sedimentation rates at the mouth of subglacial tunnel, the lake probably rapidly filled with sediment. At the northern end of the canyon a large >15 m wide) channel filled with coarse sand and pebble/cobble gravel cuts across the finer lacustrine sand. This is interpreted to be a stream channel extending from the receding glacier to the drainage divide that was eroded through the recently deposited lacustrine sediment that largely filled the lake basin.

STOP 4. Lee River Delta (2 hours)
A large landslide has been recurrently active along an outside meander bend of the Lee River for at least the last 15 years revealing an almost 100 ft (30 m) high section (Fig. 16). From year to year exposures vary, but key parts of the section can be excavated where slumped. The section offers the opportunity to see the almost complete transition between relatively deep water lacustrine sediments upwards through a classic Gilbert-style delta which was deposited in the higher, Coveville stage, of Glacial Lake Vermont. This site lies on private property. Please contact the McLoughlin family (Plains Road, Jericho) for permission to visit the site.
The lowest part of the section is frequently slumped, but consists of rhythmically bedded silt and clay (varves) with occasional dropstones. Bedding is approximately horizontal. Farther up the section the summer silt layers frequently contain beds of medium to fine sand, some of which contain well-preserved asymmetric ripples. Ripples and cross-beds here and farther up the section consistently indicate current flow was to the northwest at the time these sediments were deposited, parallel to the modern course of the Lee River.
The proportion of sand in the section rapidly increases upsection. Winter clay layers are usually present and indicate that the annual accumulation of sand often exceeded 1 m/year. Asymmetric ripples, climbing ripples, and graded beds are common in these still horizontally bedded sediments (Fig. 17). Soft sediment deformation resulting from differential loading has produced folds and flame structures in some layers. Calcite-cemented concretions are also common. These are the bottomset beds of a delta encroaching from the east, fed by the Lee River.
Higher in the section the first beds containing coarse sand and pebble gravel appear bedding is inclined to the west northwest. The average grain size continues to increase in these foreset beds. The very top of the section consists of horizontal beds of coarse pebble/cobble gravel, the topset beds of the delta. Poor exposure usually makes it difficult to ascertain whether the transition between the foreset and topset beds is smooth or truncated. The topset/foreset contact lies ∼2 m below the terrace at the top of the landslide. The elevation of that terrace has not been surveyed, but lies between 620–640 ft (189–195 m) asl based on the topographic map and pins the elevation of the Coveville Stage of Glacial Lake Vermont in this part of Vermont (Fig. 18).

B7: BOREHOLE GEOPHYSICAL DEMONSTRATION AT ALTONA FLAT ROCK
Edwin A. Romanowicz
SUNY Plattsburgh
Geologic Setting
Altona Flat Rock (Flat Rock) is a 2,800 hectare field site located near Altona, New York. It is one of several sandstone pavements in the region that make up a discontinuous belt extending southeast into the Champlain valley from Covey Hill (at the New York and Québec border). The sandstone pavements were created by the denudation of overlying material during a catastrophic breakout of Lake Iroquois. Cambrian Potsdam Sandstone is exposed at the site. An exposed normal fault cuts through the site. Cold Spring Brook flows along the fault trace. The fault scarp is exposed clearly between the reservoir and Chasm Lake (Map 1 and Map 2). The downthrown block is to the north east of the fault.
The lack of surficial deposits makes the site an ideal location for studying fracture-flow hydrogeology. More than 25 bedrock wells from 12 to 140 meters depth have been installed at the site. Elevations and UTM coordinates for each well were measured using a survey-grade GPS. The wells are cased about one meter into bedrock. Below the casing, the wells are open to bedrock.
Potsdam Sandstone has two members; Keeseville and underlying Ausable. The Keeseville is a cream color, thinly bedded quartzitic sandstone (Williams et al 2010). The Ausable is a pink to gray coarse to medium arkosic sandstone (Williams et al 2010). The Altona Formation underlies the Ausable member of the Potsdam (Landing et al 2007). The Altona is argillaceous with hematite, arkosic sandstone, shale and dolomite (Landing et al 2007, Williams et al 2010).
Well 102
Well 102 is the deepest well at the site (Map 3). It is 140 m deep, fully penetrating the Cambrian units and extending into the Precambrian basement metamorphic and igneous rock. It has been used extensively for class demonstrations and undergraduate research. Several fractures in well 102 have been identified. Many of these fractures intersect wells nearby (Wells 500A, 500B and 103). Over the past several years, we have studied changes in chemistry of water in well 102 (Klein et al., 2013; Dorsey et al., 2014 and Mesuda et al., 2015). These changes in water chemistry occur due to changes in the relative contribution of groundwater from different sources flowing through different fractures. Well 102 is dominated by an upward flow of water originating in the Altona. Superimposed on this upward flow is a shallow flow system that is sensitive to surface events (precipitation and snow melt). In a 10 meter length of the borehole these two flow systems merge, resulting in rapid changes of water chemistry as conditions change.
Wells 500A and B are each 40 m deep and well 103 is 24 m deep. Despite sharing common fractures with well 102, the changes in borehole water chemistry have not been observed in the other wells. This suggests that an important driving mechanism for the changes observed in well 102 is the occurrence of upward flow from deeper in the well. The other wells are too shallow to intersect the upward flow system.
Borehole Geophysics Demonstration
This field trip will consist of a demonstration of borehole geophysical techniques at well 102. During the demonstrations we will log gamma, caliper, fluid temperature and resistivity and create an acoustical image of the borehole. We will also measure the vertical velocity of water flow within the borehole using a heat-pulse flow meter. An example of the well log is shown in figure 1.

B7: BOREHOLE GEOPHYSICAL DEMONSTRATION AT ALTONA FLAT ROCK

Meeting Place

STOP 1. Well with geophysical instruments
We will use a Mt. Sopris® Borehole Geophysical system. The system includes a 500-m electric winch with a Matrix® computer interface. Every probe will be zero-referenced to the top of the well. Thus, the top of the well is 0 m depth. The depth of the probe is measured from its bottom. The processing software takes into account the actual location of the sensor(s), so the depths reported in the data are those of the sensor. As the probe is lowered (4-5 m·min-1), the probe will automatically take measurements at a pre-programmed depth intervals.
The fluid temperature and resistivity probe measures the temperature and electrical resistance of the fluid in the well casing. Abrupt changes in fluid temperature or resistivity may indicate fractures or lithological units that are contributing water to the well. In deeper wells (>150 m) it is possible to observe the thermal gradient of the crust in the well.
The gamma probe counts gamma rays emitted naturally from the bedrock as a part of gamma decay. This probe is particularly useful in separating shales and carbonates from other sedimentary rocks. The major sources of gamma particles in rocks are potassium, thorium, uranium and radium. The energies associated with these different gamma sources differ. Some probes are able to differentiate between the sources. The probe we are using is not able to differentiate. The gamma probe uses a scintillation crystal. The crystal lights up when exposed to gamma rays. A photomultiplier counts the number of times the scintillation crystal lights up. This number is converted to counts per second (cps).
Gamma logs are not a quantitative tool. The number of counts is not as important as the differences in counts at different depths in the borehole. Shales are often high emitters of gamma particles. The source of the gamma particles is K in clays. Carbonates are typically low gamma emitters. Quartz-rich sandstone would have a very low gamma count, while arkosic sandstone will be high.
Caliper log shows the diameter of the borehole. Three prongs are extended from the caliper probe at the bottom of the well. The probe is then raised to the surface, allowing the prongs to drag along the walls of the borehole. The probe is sensitive enough to detect fractures intersecting the borehole.
The acoustical borehole imager scans the interior of the borehole using RADAR. The return time and reflected energy are recorded. The probe is also equipped with an accelerometer and magnetometer so that one can determine its orientation in three-dimension space. The resulting image from the probe is the wall of the borehole splayed open. Magnetic south is referenced along the center axis of the image. Magnetic north is referenced along the two edges. Fractures, foliations or changes in lithology often result in changes in return energy and travel times. The changes will appear in the image as changes colors. Planar features that intersect the well at perpendicular angles will be shown as bands of alternating colors perpendicular to the long axis of the figure. Features intersecting the borehole at other angles will appear as a sine wave, with a distinct peak and trough. Since the orientation of the probe is known, it is possible to determine the orientation of the strike and dip of planar features intersecting the borehole. By comparing the acoustical borehole image to the caliper log, one can distinguish between filled fractures, open fractures or foliation.
The heat-pulse flow meter measures vertical fluid flow in the borehole. Two thermistors, one above and one below a heating element monitor ambient fluid temperatures. Baffles direct water flow through the probe, past the sensors and heating element. From the surface, we trigger the heating element to heat for an instant. If there is vertical flow, one of the two thermistors will indicate an increase in ambient temperature. The time between firing the heating element and observing the peak increase in ambient temperature is used to calculate the water velocity or discharge.