Overview and Highlights

Through photos, field notes, and interpretations, this blog attempts to describe and contextualize the geologic formations of Western New England, as seen in outcrop on a series of five field trips.  Highlights are as follows:

• Precambrian rocks of the Grenville basement are the oldest units visited, and were seen primarily in field trip 5 and the first stop of field trip 1.
• During the Precambrian to Early Cambrian breakup of Rodinia and concurrent opening of the Iapetus Ocean, a series of rift sediments was deposited. These sediments are the protoliths for the various orogenically metamorphosed formations in field trips 1 and 2.  The specific lithologies of these rocks are indicative of the sea level in the locations at which they were deposited.
• The Hoosac Schist of field trip 2 as well as all the formations seen in field trip 3 are derived from passive margin slope-rise deposits, from deeper water than the synrift deposits.  The passive margin persisted from the Early Cambrian through the Middle Ordovician.
• The passive Laurentian margin ended with its Ordovician collision with the Shelburne Falls arc.  Outrcrops reflective of active tectonic processes such as subduction and volcanism were seen in field trip 4.  

  

  Figure 0.1: Tectonic Units of western Vermont and Massachusetts.  This figure will occur throughout the
  blog, with the relevant sequence highlighted.

References (entire blog):

Allen, J.S., W. A. Thomas, and D. Lavoie, 2010. The Laurentian margin of northeastern North America.  Geologic Society of America Memoirs, 206, 71-90.

Cawood, P. A., P. J. McCausland, and G. R. Dunning, 2001.  Opening Iapetus: constraints from the Laurentian margin in Newfoundland.  Geological Society of America Bulletin, 113, 443-453.

Courtillot, V., and G. E. Vink, 1983. How continents break up. Scientific American, v.? 43-49.

Karabinos, P., 2013.  Class and lab lecture/discussion. GEOS 401, Williams College, Fall 2013.

Karabinos, P., S. D. Sampson, J. C. Hepburn, and H. Stoll, 1998. Taconian orogeny in the New England Appalachians: collision between Laurentia and the Shelburne Falls arc. Geology, 26, 215-218.

Lister, G. S., M. A. Etheridge, and P. A. Symonds, 1986. Detachment faulting and the evolution of passive continental margins.  Geology, 14, 246-250.

Weaver, P. P. E., R. B. Wynn, N. H. Kenyon, and J. Evans, 2000. Continental margin sedimentation, with special reference to the north-east Atlantic margin. Sedimentology, 47, 239-256. 

Williams, H., and R. N. Hiscott, 1987. Definition of the Iapetus rift-drit transition in Western New Foundland. Geology, 15, 1044-1047.

Withjack, M. O., R. W. Schlishe, and P. E. Olsen, 1998. Diachronous rifting, driting, and inversion on the pasive margin of central eastern North America: an analog for other passive margins, AAPG Bulletin, 82, 817-835.

Bennington, Vermont: Rift to Shelf Deposits (Field Trip 1)

Figure 1.1:Rift Clastics seen at these outcrops include Cheshire quartzite, Monkton quartzite,
Winooski dolomite, and Bascom marble.

Stop 1: Mount Holly Complex

Figure 1.2 Banded gneiss at this roadside outcrop contains
mafic and felsic bands, some of which are truncated.  The
greenish color is primarily due to the alteration of feldspar.
Arrow points to bedding truncation.  Pencil photographed
for scale.

This outcrop consists primarily of gneiss with bands of pegmatite.  The pegmatite is close to granite in composition.  Smaller bands of quartz, biotite, and epidote can also be seen.  In some areas, banding is truncated.  These truncations, as well as foliations in the gneiss can be seen in the photo at right.
     From argon dating, we know that this outcrop is roughly 1.1 billion years old, significant in that it tells us that the Taconic orogeny did not generate enough heat in this area to reset the argon clock.  The sediments were originally deposited in shallow water   The strike and dip are 220, 66.  In the photo at right, foliations in the gneiss can be seen.

Figure 1. 3: An anticline in this outcrop is its most distinctive structural feature;
line drawn in is parallel to bedding.

Stop 2: Cheshire quartzite

Figure 1.4: From closer up, cross-bedding can faintly be
observed in the fairly homogenous quartzite.  The cross-
bedding, asymptotically truncated, can be used as a "way up"
structure.  Here, up is up. 

The direction of truncation of the cross-bedding demonstrates that the rocks are currently right side up.  The quartzite at this stop is fairly homogeneous, probably having formed from very pure quartz sand from when the formation was part of a beach.  An alternative explanation for the homogeneity of this quartzite is that during diagenesis fluid might have weathered out other minerals such as feldspar; however, this would require a substantial amount of fluid.  Like in stop 1, the features of this outcrop are indicative of a wet climate, and show that chemical weathering is dominant.  Out west, in contrast, where the climate is drier, physical weathering is dominant.

Between the Mount Holly Complex and the Cheshire quartzite, we bypassed the Dalton Formation, also part of the rift phase of the breakup of Rodinia.   

Stop 3: Monkton quartzite thrust over Winooski Dolomite

Figure 1.5: The contact between the older Monkton

Quartzite and younger Winooski Dolomite is thrust fault.

In the Monkton, the rusty looking rock is the quartzite,

 while interbedding of phyllite and marble accounts for the other colors.

At this stop, an acid test confirmed the presence of dolomite.  Quartz veins pervade throughout the dolomite matrix. Bedding in the dolomite of the lower formation is continuous.The Monkton is composed of phyllite, quartzite, and dolomitic marble.  Truncations in both formations indicate that the contact between them is a fault.

     Though the Monkton quartzite is older than the Winooski dolomite, they are both from Cambrian sedimentary deposits.  Both were deposited in a shallow marine setting.

Stop 4: Scolithus tubes

Figure 1.6: The dark, vertical linear features are the trace
fossils from burrowing.  Pen photographed for scale.

 Due to bioturbation, this outcrop of Cheshire Quartzite has not retained sedimentary structures.  The fact that bioturbation didn't destroy its own evidence completely indicates that sedimentation was rapid.

Stop 5: Bascom Marble

Layers of folded grey calcite marble alternate with darker beds of marble.  Bedding thickness ranges from about 5-40 cm.  The white calcite veins parallel to bedding are not part of the depositional sedimentary history; they precipitated from fluid along bedding planes later on.

     Slip is going toward the core of the fold, where cleavage is at the highest angle to bedding.  

     The calcite beds are of Ordovician age, and were likely deformed during the Taconic orogeny.

Figure 1.7: This syncline includes layers of clay-rich dark calcite marble as well as lighter-colored calcite marble. The folding caused stress to overpower the strength of the rock, inducing a fault.

Additional Tectonic Context

• The unconformity between Mt. Holly Gneiss basement and rift sediments suggest that the Laurentian margin at the breakup of Rodinia was a lower-plate margin (Lister 1986).  The uplift and subsidence of the lower plate resulted in high rates of deposition of the adjacent trough.  

Figure 1.8: Adapted from Lister et. al., 1986.  The Laurentian margin was the lower plate margin associated with the detachment faulting when Rodinia broke up.  When the crust bowed down due to removal of asthenosphere below, a basin formed, where the Cheshire quartzite,  Dalton quartzite, Monkton quartzite, and Winooski dolomite were then deposited. 

• Courtillot and Vink (1983) discuss how parts of rifted continents can be matched back together by examining the pattern of magnetic anomalies associated with rifting, which are interrupted by the edges of continents.  This does not work, however, if the rift has propagated.  Propagation occurs because the.rifting crust is not uniformly weak; the strongest sections remain attached longest, creating locked zones.  As volcanic material intrudes through the weak zones of crust, it eventually stretches and breaks apart the locked zones.  
• During rifting, clastic deposition of clastic volcanic rocks occurred at the Laurentian margin, and were overlain by carbonate shelf deposits once the margin was passive (Allen et. al., 2010).  Sandstones that were metamorphosed into the Cheshire quartzite make up part of the rift zone.  In Western Vermont and Massachusetts, these sandstones overlie the Grenville basement directly because teh Pinnacle Formation, which can be observed further north, pinched out completely.  The Pinnacle Formation is composed of synrift volcanic clastics.  Since the Cheshire Formation lies between volcanic clastic deposits/basement and carbonates, it marks the transition from rift deposits to passive margin deposits.  
• Allen et. al. (2010) further divides the carbonate deposits into a shallow to deep water succession, ending in the Winooski dolomite.  That the shelf carbonates interfinger with slope deposits suggest rapid subsidence during rifting, as Lister et. al. (1986) suggest, due to the instability of the lower-plate margin during and immediately following rifting.  

Florida, Massachusetts: Rift to Slope Deposits (Field Trip 2)

Figure 2.1: Though the Stanford Granite Gneiss is basement rock, the conglomerate and meta-arkose seen
in this field trip are derived from slope rise deposits.

Stop 1: Weakly Foliated Stanford Granite Gneiss

Figure 2.2: Weakly foliated "Rapakivi"gneiss includes

squashed feldspars surrounded by thin rims of plagioclase.

Pen photographed for scale.

The granite contains plagioclase, K-feldspars, bluish quartz, garnets, epidote and clusters of biotite.  The granite is termed "Rapakivi" granite, meaning that there are large phenocrysts of K-feldspar surrounded by thin rinds of plagioclase.  Quartz veins, not weakly foliated like the surrounding rock, must have been deposited fairly late.  Some of the feldspars are squished into eye-shapes, so texturally the rock can be classified as augen gneiss.  Strike and dip are approximately 320, 14.
     A gently plunging anticline (not pictured)  exhibits the least deformation at its core.




Stop 2: Strongly Foliated Stanford Granite Gneiss

Figure 2.3: Strongly foliated Stanford granite gneiss, in

which quartz grains are  prominent and feldspars are 

flattened significantly.

Compared to the first outcrop, the gneiss here is more strongly foliated with fewer grain size variations.  Quartz is more dominant, feldspars are more elongated, and folding here is strong enough to obliterate the Rapakivi texture.  Overall, it appears that stresses here, at the boundary of the Stanford granite gneiss, were stronger than at the formation's center.  Strike and dip are about 260, 16.
     In both of the first two stops, the granite protolith of the gneiss is Precambrian.






Stop 3: Dalton Formation

Figure 2.4: Dalton conglomerate, in which many lithic clasts

are flattened parallel to bedding almost to the extent of being 

reabsorbed by the matrix.  Quartz grains, however, are largely

undeformed.  Pen is photographed for scale.

The conglomerate here includes grains of various sizes and levels of deformation including large grains flattened almost to the point of being reabsorbed by the matrix. Quartz, unlike other minerals, is not stretched, indicating the sediment did not reach an extremely high temperature during metamorphism.  Flattening occurs in a direction almost parallel to bedding, which is still visible.  The sediment was deposited roughly at the end of the Precambrian.  Strike and dip are 285, 5.

Figure 2. 5: Note that in this conglomerate, there is more

deformation in the bed below the drawn line than in that

above it.  This suggests either different lithologies or

different stresses.

Stop 4: Conglomerate with Magnetite
This conglomerate is similar to that from stop 3, and is of the same age, but is more deformed and includes octahedral magnetite grains.  ~40 cm thick bedding is still visible.  The outcrop is cut through by a fault, through which quartz precipitates.  Strike and dip are 280, 12.


   







Stop 5: Meta-arkose/ Conglomerate
Arkose and conglomerate are interbeded.  Bedding is thicker than at either stops 3 or 4.  Grain size is smaller than at the previous stop.  Strike and dip are 305, 16.  The outcrop is of late Precambrian age.

Stop 6: Age equivalent of Cheshire Quartzite
This  rock is a garnetiferous schist.  Compared to the other outcrops of this field trip, this rock was formed in deeper water, though was still part of the continental shelf.  Strike an dip are roughly 260, 17. Additional Tectonic Context

• The stratigraphic rift to drift transition lies between a less stable sedimentary layer of synrift basin rocks and a more uniform and stable layer of sediment (Williams and Hiscott, 1987).  Synrift deposits can be recognized by various characteristics, including substantial and rapid thickness changes, boundary by listric faults defining basin highs, and the presence of large quantities of volcanic rocks.  
• The creation of the Iapetus Ocean, at about 550 Ma, marks the rift-drift transition (Williams and Hiscott, 1987), though precision in dating might be limited by the presence of lock-zones (see previous post).
• According to Withjack et. al. (1998), the transition from rift to drift was diachronous (began earlier in southern North America) and was marked by a switched from northwest-southeast extension to shortening at the same time as sills and dikes intruded the continental crust.  Normal faulting at this time was replaced by reverse faulting because the maximum stress before rifting and the minimum stress after rifting were subvertical.   

Figure 2.6: Adapted from Withjack et. al. (1998)  Extensional forces on either side of the rift acted on the lithosphere (left panel), allowing asthenosphere to upwell from below.  The asthenosphere intruded through the weakened crust (right panel), forming a volcanic wedge.  These intrusions exerted force away from the ridge at a rate faster than extension, which necessitated compensational forces as a sort of rebound effect. 

Charlemont, Massachusetts: Oceanic Realm Crust and Deposits (Field Trip 3)

Figure 3.1:During this field trip, we moved from the slope-rise deposits of the Hoosac schist to the deep
oceanic realm Rowe schist.

Stop 1: Hoosac Schist

Figure 3.2: Different colors in the schist indicate graphite

(dark grey), muscovite and albite (lighter grey), and quartz

veins (rusty color).  Hammer photographed for scale.

The green color of the schist at this outcrop is from the associated chlorite and epidote.  The pelitic schist is of Ordovician age, and also contains muscovite, garnet, albite crystals and quartz in veins and as part of the matrix.  Graphitic inclusions include graphite cores derived from biogenic material with  with white rims where feldspar forms after the graphite.
     Foliation is fairly strong; some quartz veins are isoclinally folded with parallel limbs, indicating that they were intruded before deformation. Folding of foliation indicates that the rock underwent two deformations; the original pelitic sediment was metamorphosed in the Taconic orgeny, and folding occurred during the Acadian orogeny.  Strike and dip are 345, 25.

Stop 2: Rowe Schist

Figure 3.3: This schist is lustrous and silver due to the presence

of chloritoid.  The most notable deformation features are the

lineations.  Pencil photographed for scale. 

Compared to the Hoosac schist at stop 1, the Rowe schist is finer-grained, more lustrous, and more finely layered although schistocity is similar.  Composition is more uniform, though there are still many quartz veins.  There is an abundance of chloritoid, a mark of aluminum rich rocks.  There is no albite in this rock, which is lower in both magnesium and iron than the Hoosac schist, but higher in sodium and Aluminum.  Instead of muscovite, there is paragonite, which has a similar crystal structure, but is the sodium end-member.  Like at the first stop, pelitic sediment was deposited here and underwent metamorphism during the Taconic orogeny.  At this outcrop, there is no evidence for a later deformation.
     Strike and dip are 001, 82, and the trend and plunge of crenulation lineation is 140, 50.  




Stop 3: Rowe Schist continued on Whitcomb Hill Road

Figure 3.4: Color banding in this schist is resultant of
quartzite-rich and chlorite-rich layers.  Pen photographed
for scale. 

This schist has a greenish tint similar to at the previous stop, but is less shiny, probably because of a smaller percentage of paragonite and other mica. Quartzite is interbedded with the chlorite, which calls into question whether the Rowe schist is really part of the oceanic realm because there shouldn't be much quartzite far away from land.  The presence of quartz "pebbles" can be explained if they are not actually pebbles, but boudinaged quartz veins.
     Strike and dip of schistocity are 345, 66, and trend and plunge of lineation is 130, 48.






Stop 4: Carbonaceous Rowe Schist

Figure 3.5: The contact between the two forms of Rowe
Schist is distinct in color and consistency differences. The
typical Rowe schist is similar to that in Figure 3.4, while
the carbonaceous schist falls apart, is darker, and is not
lustrous.

The rusty, platy "cruddy black schist" of the Rowe Formation forms a contact here with the paragonite-dominated Rowe schist.  The cruddy black schist is porous, weak, and covered in vegetation.  Although it is not mica-rich, it breaks parallel to schistocity.  The sedimentation rate must have been high, because the fact that quartz was weathered out is indicative of reducing conditions, which means the material must have been buried before it got oxidized.

Stop 5: Ultramafics at the Reed Brook Preserve

Figure 3.6: This rock is harder than the others we visited,
because it has a mantle source.  Note the hint of dark green
ultramafic in the right-hand foreground.  Hammer is photo-
graphed for scale. 

To the southeast of the Rowe formation, small blocks of ultramafic material indicate mantle rock.  The rock is too fine-grained for olivine to be present.  The general belief is that it is part of a suture zone in the Iapetus Ocean, but it might actually be from near the Laurentian continental margin.

Figure 3.7: Mafic schist of the Moretown formation is 

intruded by large veins of quartz.

Stop 6: Moretown Formation
The mafic schist with slaty cleavage was originally a basalt or diabase mafic intrusive, containing a lot of alternating biotite and quartz-rich layers, feldspars, amphibole, quartz veins, and some pyrite, chalcopyrite, epidote, and actinolite.   The Moretown is supposedly from the Shelburne Falls forearc region, but zircon dating data has not confirmed this.





Additional Tectonic Context

• The presence of the Moretown Formation, broken off from Gondwana accreted onto the postrift Laurentian Margin, is explained by Cawood et. al. (2001).  They suggest that rifting at the Iapetus margin has multiple stages, where microcontinents broke off separately after the main rift.  They also disagree with Williams and Hiscott (1987), who claim that the rift-drift transition is marked by a change from siliciclastic to carbonate deposits.  Instead, they point to siliciclastic and volcanic rocks bounded by faults as the rift transition zone.  They consider the oxidation of these rocks to be an indicator that time passed between their accumulation and further siliciclastic deposition.  

Shelburne Falls, Massachusetts: Arc Domain Rocks (Field Trip 4)

Figure 4.1: The outcrops of this field trip are related to the collision of the Shelburne Falls arc with Laurentia. 

Stop 1: Moretown Formation

Figure 4.2: Darker and lighter layers are the more mafic and
felsic materials, respectively.  A precipitated quartz vein can be
seen running from top-left to bottom-right of photo, with an
arrowpointing to it.  Hammer is photographed for scale. 

The Moretown Formation was originally part of Gondwana.  Detrital zircons have given the age of the earliest sediments in this formation as 514 Ma.  The oldest intrusion is dated at 504 Ma.  Older models had suggested that pieces of Gondwana had collided with Laurentia during the Devonian; the ages of Moretown zircons suggest that this fragment of Gondwana would have actually arrived during the Ordovician.  
     In general, this outcrop consists of highly foliated gneiss with alternating quartz- and mica-rich bands intruded by mafic dike and quartz veins.  Mineralogy also includes reddish garnets ranging from <.5 mm to about 7 mm in size, green epidote, and actinolite in the mafic intrusion.  The thinly laminated micaceous layers are composed mostly of muscovite.  Along the top of the outcrop, actinolite and weathered mafic textures were particularly noticeable.  Boudinage, where more ductile layers are deformed preferentially around more robust layers, suggests at least a local extension of fold limbs.
     Strike and dip at this outcrop are approximately 045, 49. The trend and plunge of lineation is about 100, 42.

Stop 2: Hallockville Pond Gneiss

Figure 4.3: Crenulated gneissic texture serves as evidence
for two deformations.

From zircon dating, the age of this outcrop, composed of intrusive rock forming part of the Shelburne Falls arc, has been determined to be 475 Ma.  The mica-rich gneiss is intruded by numerous quartz veins.  Evidence in the fabric suggests that it has been subjected to at least two deformations, possibly as part of the Taconic and Acadian orogenies.  Evidence consists of aplite dikes that have been folded and redeformed, as well as crenulated gneissic texture.







Stop 3: Hawley Volcanics

Figure 4.4: Possible deformed pillow lavas.  Field
notebook is photographed for scale.

Dating of detrital zircons show a clearly Laurentian origin for this outcrop, with a date of approximately 475 Ma for the mafic rock, intruded by a more felsic sill. The mafic rock consists of quartz, feldspar, and amphibole, with small amounts of garnet and epidote that likely formed through hydrothermal alteration of basalts in a marine environment by percolating seawater.
     In pillow lava; the cusp of the pillow should be pointing down with the bulbous part pointing up.  Some structures at this outcrop appear similar to pillows.  However, if pillow lavas, they must be highly deformed as the cusp and top are not readily distinguishable.



Stop 4: Shelburne Falls Dome: Collinsville Formation

Figure 4.5: Mafic dike intruded into tonalite

and intruded by quartz veins. 

This outcrop is composed primarily of foliated tonalite, rich in plagioclase and quartz and deficient in K-feldspar, a composition typical of volcanic arc rocks.  The tonalite is intruded by mafic dikes, which would have required a magma temperature of 1200 degrres C, compared with roughly 700 degrees C for the felsic material.  Formation required partial cooling and the beginning of crystallization of the tonalite followed by mafic magma intrusion.  Strike and dip are about 100, 10.  Trend and plunge of lineation on the approximately isoclinal folds are 275, 6.   Mica is strongly aligned with the fold axis.


"Glacial" potholes are present, and contain quartz veins where other minerals have been eroded away preferentially.  Pegmatites indicate areas where unusual elements that did not mix well with the rest of the magma crystallized into larger grained rock.

Figure 4.6: Pegmatite vein in tonalite

Additional Tectonic Context

• According to Karabinos et. al. (2013), the Taconian orogeny occurred when the Shelburne Falls arc, at 485 to 470 Ma, formed over an east-dipping subduction zone and collided with Laurentia soon after.  The polarity of the subduction zone then switched, ending the period of convergence.  The Bronson hill arc then formed above the new, west-dipping subduction zone.
• The Collinsville Formation represents the core of the volcanic arc, while the Hallockville Pond gneiss is a pluton that originated from subduction.  The Hawley volcanics are erupted extrusives from the volcanic arc. 

Monterey, Massachusetts: Berkshire Massif (Field Trip 5)

Figure 5.1: In this field trip, outcrops are of gneiss from the Grenville basement, though they contain felsic/granitic sills.  Some of the sills intruded in a rifting environment; others are dated as being almost as old as the basement rock.

Stop 1A: Becket Quarry Granite

Figure 5.2: Biotite gneiss overlies granite; contact between
the two is shown.  A band of pegmatite precipitated within
the gneiss.

A granitic sill, dated at 432 Ma, intrudes basement 1150 Ma.  On the north end of the quarry are basement xenoliths of biotite-rich gneiss.  On that same side of the quarry is a contact between granite overlain by biotite gneiss.  The gneiss is crosscut by a wide band of pegmatite, which is the same age as the granite.

Stop 1B: Becket Quarry sills

Figure 5.3: A pegmatite vein lies at the contact between
granite sill and Hoosac schist.

At this outcrop, two granite sills intrude through Proterozoic Hoosac schist containing plagioclase, quartz, and micas; pegmatite can also be seen.  The schist is too high-grade to show contact metamorphism, and its contact with the granite shows no indication of faulting.  The granite is interpreted to be about 434 Ma.  Grain size and composition for the two rocks are similar; the difference is mainly in the degree of foliation.  The strike and dip of the granite sill are about 340, 25.
        These sills may have been intruded during basement rifting; it is difficult for such bodies to intrude in a compressional zone.




Stop 2: Felsic sill in Tyringham gneiss
The two rock types here are a fine-grained felsic sill of tonalite and a coarser grained gneiss of granodioritic composition.  This Tyringham gneiss contains eye-shaped feldspars called augen, shaped in response to deformation.  Dark material, most likely magnetite, is contained within the augen.

Figure 5.4: Augen texture in the Tyringham gneiss. Pencil
is photographed for scale.  

     Felsic sills intruded the Tyringham gneiss. Zircon dating indicates that the Tyringham gneiss was crystallized at about 1179 Ma, and the felsic sill intruded at 1004 Ma.  Zircon dating, and the fact that grain structure is roughly east-west aligned compared with a typical post- Precambrian north-south alignment, suggests that the Taconic orogeny did not affect this area.  Oscillatory zoning of the zircons might be a result of pulses of magma, each with a slightly different composition and temperature.
     Contact metamorphism can most likely not be seen here because small intrusions like these cool relatively fast.  Strike and dip are about 275, 27.  Trend and plunge of lineation are 285, 28.

Stop 3: Tyringham Cobble

Figure 5.5: The mylonitic texture of this rock indicates that
metamorphism took place at high temperatures.  Pen is
photographed for scale.

Walking up the east slope of Tyringham Cobble, marble and then mylonitic gneiss is encountered. Argon dating in this locality and nearby monazite dating of around 400 Ma suggests that recrystallization occurred as a result of emplacement of the Berkshire Massif.

Figure 5.5: Biotite gneiss at the top of Tyringham Cobble.

At the top of Tyringham Cobble, biotite-rich gneiss looks similar to the basement rock.  It is intruded by a sill of granodioritic composition.  The sill hasn't been dated yet, but is probably similar in age to the sills from the previous stops, and so probably intruded at around 1000 Ma, meaning the thrust is not related to the Taconic orogeny.