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Conference field trip, Sunday, June 14, 2015
An overview of the bedrock geology and tectonic
evolution of west-central, Vermont
Co-lead by Keith Klepeis & Laura Webb, Department of Geology, University of Vermont
with logistical support from UVM graduate students & the Vermont Geological Survey
Field trip itinerary
Time
Duration
8:00
Itinerary description
Location details (or comments):
Depart Stowe Mountain Lodge
Stowe Mountain Lodge, 7412 Mountain
Road, Stowe, VT 05672
1:00
Drive to UVM Geology Dept
0:15
Arrive at UVM Geology, rendevous with co-leader,
Keith Klepeis, and UVM Geology van shuttle drivers
Use restrooms (option: geology overview in lobby in
case of inclement weather)
9:00
9:15
Depart UVM Geology Dept
0:10
9:25
Drive to Lone Rock Point
Arrive at Lone Rock Point trail parking area across
from Episcopal Diocese of Vermont; porta potty
available at parking lot
0:20
Walk to Champlain Thrust outcrop
9:45
0:35
Field Trip Stop 1: Champlain Thrust
10:20
0:20
Walk back to bus
10:40
12:30
0:25
Depart Lone Rock Point trail parking area
Drive to Mount Philo State Park lower parking lot via
Spear Street (for views)
Arrive at Mount Philo State Park lower parking lot;
porta potty at parking lot
Walk or shuttle to Mount Philo overlook. Lunch at
overlook when you arrive. Men’s and women’s
restrooms with sinks available
Field Trip Stop 2: Champlain Thrust and Mt Philo
overlook
12:55
0:30
Walk or shuttle back to parking lot
0:40
11:20
1:10
13:25
13:40
Will need to shuttle coolers with lunches to
top of Mt Philo
13'9" underpass with this route, which was
chosen because approach from south has
covered bridge
Arrive at Charlotte Beach parking area
1375 Lake Rd, Charlotte, VT 05445
Walk to beach outcrop
0:30
Stop 3: Charlotte Beach Mesozoic dikes
0:25
Depart Charlotte Beach parking area
Drive to Hinesburg thrust stop on Place Rd via Pond
Rd
14:40
Arrive at Hinesburg thrust parking area
0:10
Walk to outcrop
0:45
Stop 4: Hinesburg thrust
15:35
Place Rd near intersection with Pond Rd,
Hinesburg; alternative bus parking in dirt lot
south of intersection on Pond Rd
Depart for Stowe Mountain Lodge
1:00
16:35
Mount Philo Parking Lot, State Park Road,
Charlotte, VT 05445
Drive to Charlotte Beach via Greenbush Rd to Lake Rd
(northern approach)
0:05
14:15
14:50
Episcopal Diocese of Vermont, 5 Rock Point
Road, Burlington, VT 05408
Depart for Charlotte Beach
0:15
13:45
180 Colchester Ave, University of Vermont,
Trinity Campus, Burlington, VT 05401
Drive to Stowe Mountain Lodge
Arrive back at Stowe Mountain Lodge
Stowe Mountain Lodge, 7412 Mountain
Road, Stowe, VT 05672
1
Overview
The geology of Vermont provides a record of the tectonic evolution of the eastern
margin of North America, including insight into the effects of multiple phases of collision,
accretion, and rifting. This field trip will visit several classic localities that serve as superb
educational outdoor laboratories and provide important clues on the complex history of
superposed orogenies in the northern Appalachians. Our first stop will be the Champlain
thrust at Lone Rock Point along scenic Lake Champlain (Fig. 1). From there we will visit
Mount Philo State Park, were we will enjoy lunch with panoramic views of the Champlain
Valley, Lake Champlain, and the Adirondack Mountains. Afternoon stops include Mesozoic
dikes at Charlotte beach and the Hinesburg thrust shear zone.
Figure 1. Location of field trip stops on Google Maps terrain map.
Please note that this field guide contains excerpts from New England Intercollegiate
Geological Conference (NEIGC) field trip guides by West et al. (2011), Kim et al. (2011), and
McHone and McHone (2012), as noted below (pdf versions of these field guides can be
provided upon request).
Introduction to the general geology of Vermont (Kim et al., 2011)
Vermont can be divided into several north-northeast trending bedrock belts of generally
similar age and tectonic affinity (Fig. 2). From west to east the belts are:
2
1) Champlain Valley: Cambrian–Ordovician carbonate and clastic sedimentary rocks
deposited on the eastern (present coordinates) continental margin of Laurentia (e.g., Stanley and
Ratcliffe, 1985). This continent was left behind after the Rodinian supercontinent rifted apart
during the Late Proterozoic and the intervening Iapetus Ocean formed between it and Gondwana
(e.g., van Staal et al., 1998). The margin was deformed and, in some parts, weakly
metamorphosed during the Ordovician Taconian Orogeny. The Devonian Acadian Orogeny
resulted in further deformation and metamorphism, particularly on the eastern side of the valley.
2) Taconic Allochthons: Late Proterozoic- Ordovician slices of clastic metasedimentary
rocks of oceanic and continental margin affinity that were thrust onto the Laurentian margin
(Champlain Valley Belt) by arc-continent collision during the Taconian Orogeny (e.g., Stanley
and Ratcliffe, 1985).
3) Green Mountain: Late Proterozoic–Cambrian rift- and transitional rift-related
metasedimentary and meta-igneous rocks that unconformably overlie Mesoproterozoic basement
rocks. These assemblages were deformed and metamorphosed during the Taconian Orogeny
(also during the Acadian Orogeny) (e.g., Thompson and Thompson, 2003).
4) Rowe-Hawley: Metamorphosed continental margin, oceanic, and suprasubduction zone
rocks of Late Proterozoic–Ordovician age that were assembled in the suture zone of the Taconian
Orogeny. These rocks also were deformed and metamorphosed during the Acadian Orogeny. Arc
components are part of a Shelburne Falls Arc that collided with the Laurentian margin, causing
the Taconian Orogeny (Karabinos et al., 1998).
5) Connecticut Valley: Silurian and Devonian metasedimentary and metaigneous rocks
deposited in a post-Taconian marginal basin. Tremblay and Pinet (2005) and Rankin et al. (2007)
suggested that this basin formed from lithospheric extension associated with post-Taconian
collisional delamination processes. These rocks were first deformed and metamorphosed during
the Acadian Orogeny.
6) Bronson Hill: Ordovician metaigneous and
metasedimentary rocks of magmatic arc affinity and
the underlying metasedimentary rocks on which the
arc was built (e.g., Stanley and Ratcliffe, 1985).
Recent studies show that this is a composite arc
terrane with juxtaposed components of Laurentian
and Ganderian/Gondwanan arc affinity (e.g.,
Moench and Aleinikoff, 2003; Aleinikoff et al.,
2007; Dorais et al., 2008; 2011). Accretion of the arc
terranes onto the composite Laurentia occurred
during the latest stage of the Taconian Orogeny and
Silurian Salinian Orogeny (van Staal et al., 2009).
Figure 2. Bedrock Belts (modified from Doll et al., 1961). CV =
Champlain Valley; TA = Taconic Allochthons; GMB = Green
Mountain Belt; RHB = Rowe Hawley Belt; CVT = Connecticut
Valley Trough; BH = Bronson Hill. Box denotes region visited in
this field trip. B and S denote the approximate locations of
Burlington and Stowe Mountain Resort, respectively.
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Generalized geologic history of the region (West et al., 2011)
The field area for this trip encompasses the western part of the Green Mountain Belt
and the Champlain Valley geologic belts (Fig. 2). These belts represent the foreland and
hinterland of the Taconian Orogen of west-central Vermont, respectively (e.g., Stanley and
Wright, 1997). This region can be divided into three lithotectonic slices which are, from west to
east and from structurally lowest to highest: A) the parauthochthon, B) the upper plate of the
Champlain Thrust, and C) the upper plate of the Hinesburg Thrust (Figs. 3, 4, and 5). The
Champlain thrust fault forms the tectonic boundary between A and B, whereas the Hinesburg
thrust shear zone separates B and C. Although these slices were originally juxtaposed during the
Ordovician Taconian Orogeny, subsequent deformation occurred during the Acadian (Devonian)
and possibly later orogenies (e.g., Stanley and Sarkisian, 1972; Stanley, 1987).
Tectonostratigraphy of slices
The stratigraphy of the field area has been described in detail by Cady (1945), Doll et al.
(1961), Welby (1961), Dorsey et al. (1983), Gilespie (1975), Stanley (1980;1987), Stanley and
Sarkisian (1972), Stanley and Ratcliffe (1985), Stanley et al. (1987), Stanley and Wright (1997),
Mehrtens (1987; 1997), Landing et al. (2002), Thompson et al. (2003), Landing (2007), Kim et al.
(2007), and Gale et al. (2009). The legend in Figure 3 briefly summarizes the rock types found
within each of the units.
Figure 4 shows the tectonostratigraphy of each of the lithotectonic slices in the field area
from west to east. From west to east, it is immediately apparent that each slice cuts into
successively older rocks and, consequently, deeper structural levels. Below are descriptions of
the tectonic affinity and general rock types contained within each slice:
A) Parautochthon (a term used to describe a sequence of rocks that have only been
transported a short distance from their original site of deposition): Late Cambrian–Late
Ordovician packages of carbonate sedimentary rocks deposited on the Laurentian
continental margin that have both conformable and tectonic (normal faults) contacts
with adjacent Late Ordovician black shales. The shales are interpreted as flysch by
Rowley (1982) and Stanley and Ratcliffe (1985), which means they were initially derived
from the erosion of thrust slices that were driven westward during the Taconian
Orogeny and were subsequently overridden by these thrusts.
B) Upper plate of the Champlain Thrust: Early Cambrian–Middle Ordovician carbonate and
subordinate clastic sedimentary rocks that were deposited on the Laurentian
continental margin. This major thrust sheet of regional extent transports older rocks at
the southern end of this field trip (Middlebury area) than it does at the northern end.
C) Upper plate of the Hinesburg Thrust: Late Proterozoic rift clastic metasedimentary and
metaigneous rocks associated with the initial opening of the Iapetus Ocean, including
the Pinnacle (CZp) and Fairfield Pond (CZfp) formations. These rocks are overlain by
Iapetan drift stage clastic rocks (argillaceous quartzite and quartzite) of the Cheshire
formation (e.g., Stanley, 1980; Stanley and Ratcliffe, 1985).
4
1
4
3
2
5
Figure 3 (includes the map on the previous page and the legend above). Generalized
geologic map of western Vermont in the area between Middlebury and Burlington
(modified from Mehrtens, 1997). Underscored numbers show the locations of field trip
stops from West et al. (2011); numbers in boxes represent ESNM2015 field trip stops.
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Thrust faults
The parauthochthon is separated from the autochthon (a term used to describe rocks that
rocks that have not been significantly transported from their site of formation/deposition) in
New York by a major unnamed fault that lies beneath Lake Champlain. In the autochthon,
Mesoproterozic metamorphic rocks of the Adirondack Massif are unconformably overlain by
sedimentary rocks of the Beekmantown Group (Isachsen and Fisher, 1970). The interested
reader is referred to McLelland et al. (2011) and references therein for the details of the
geologic history of the eastern portion of the Adirondack massif in New York.
In the southern and central parts of the field area, the Champlain thrust fault juxtaposes the
basal dolomitic member of the Middle Cambrian Monkton Formation with the Late Ordovician
Stony Point Formation. At the northern part of the field area, the Champlain thrust cuts down
section about 650 meters (~2000 feet) into the Lower Cambrian Dunham Formation (well
exposed at Lone Rock Point just north of Burlington—our Stop 1) (Stanley, 1987). Between the
Burlington and the Quebec border, this thrust generally follows the base of the Dunham
Formation and then becomes the Rosenburg thrust fault in southern Quebec (e.g., Sejourne
and Malo, 2007). South of the field area, the Champlain thrust can be mapped continuously at
the base of the Monkton Formation to south of Snake Mountain near Middlebury, Vermont
(e.g., Coney et al., 1972; Stanley and Sarkisian, 1972, Stanley, 1987). Hayman and Kidd (2002a,
b), based on detailed geologic mapping south of Snake Mountain, have suggested that the
Champlain thrust system can be traced through the Champlain Valley of Vermont and into the
northern end of the Hudson River Valley in New York. Stanley (1987) suggested that total
displacement on the Champlain Thrust is 55–100 km (34–62 miles).
In the field area, metamorphosed Late Proterozoic–Early Cambrian rift clastic (e.g., Pinnacle
Fm.) to early drift stage passive margin sedimentary rocks (e.g., Cheshire Fm.) were driven
westward over very weakly metamorphosed sedimentary rocks of the upper plate of the
Champlain Thrust along the Hinesburg Thrust. Dorsey et al. (1983) proposed that this thrust
nucleated in an overturned fold/nappe that ultimately sheared out along its axial surface. North
and south of the field area, the Hinesburg Thrust appears to die out in large fold structures (M.
Gale, personal communication, 2011). For the southern extension of the Hinesburg Thrust, P.
Thompson (personal communication, 2011) proposed that it may actually root in Precambrian
basement in the northernmost basement massif. Stanley and Wright (1997) suggested a total
displacement of ~6.4 km (4 miles) on the Hinesburg thrust.
On the basis of fold structures in the upper plate of the Champlain thrust that are truncated
by the Hinesburg thrust, the Champlain Thrust is at least slightly older than the Hinesburg
Thrust (e.g., Doll et al., 1961; Gale et al., 2010). Stanley and Sarkisian (1972) and P. Thompson
(personal communication, 2011) suggested that the Champlain Thrust may have moved a
second time after formation of the Hinesburg Thrust, partly on the basis of its metamorphic
history (described briefly below). The detailed structural history of the Hinesburg Thrust has
been discussed by Gillespie (1975), Dorsey et al. (1983), Strehle and Stanley (1986), and is
further described in Kim et al. (2011). Details of the deformational history of the Champlain
Thrust are also found in Stanley and Sarkisian (1972) and Stanley (1987).
7
Figure 4. The tectonostratigraphy of each of the lithotectonic slices in the field area
from west to east. Refer to Kim et al. (2011) for tectonostratigraphic details.
8
Brief tectonic history
1) Neoproterozoic: The Rodinian supercontinent composed of Mesoproterozoic crust
(remnants of which are exposed in the Adirondack and Green Mountains) breaks apart
and rift basins form between the newly separated continents of Laurentia and
Gondwana (e.g., van Staal et al., 1998).
2) Neoproterozoic–Early Cambrian: These rift basins filled with immature clastic
sedimentary rocks and mafic igneous rocks (Coish, 2010) as the incipient Iapetus Ocean
widens (represented by the Pinnacle and Fairfield Pond formations).
3) Early Cambrian–Late Ordovician: Largely continuous carbonate dominated deposition
occurs on the Laurentian passive margin from the Early Cambrian (Cheshire Formation)
through the Middle Ordovician (Bridport Formation) as the region transitions into the
drift stage (e.g., Stanley and Ratcliffe, 1985).
4) Late Cambrian–Middle Ordovician: The Iapetus Ocean begins to close, first along an
east-dipping subduction zone. Magmatism associated with this subduction builds a
volcanic arc (Shelburne Falls arc of Karabinos et al., 1998) within the Iapetus (initially
well offshore of the Laurentian margin).
5) Middle–Late Ordovician: The Shelburne Falls arc collides with the Laurentian continental
margin. The earliest stages of this “collision” resulted in normal faulting that affected
the thicknesses of the Late Ordovician deposits (e.g. Jacobi, 1980; Mehrtens and Selleck,
2002). Continued convergence and collision caused regional deformation (folding, thrust
faulting, etc.) in the foreland region. Of the large scale structures, the Champlain thrust
fault likely formed first, followed by the Hinesburg thrust (e.g., Stanley and Ratcliffe,
1985).
6) Late Ordovician–Devonian Further folding and faulting of the Taconian foreland region
(e.g., Stanley and Sarkisian, 1972).
Regional trends (Kim et al., 2011)
From the edge of Lake Champlain eastward across the Champlain and Hinesburg thrusts,
several regional trends are evident (Fig. 5). Nearest the lake, the rocks record mostly brittle
deformation that includes blind normal faults. Farther east, in the hanging wall of the
Hinesburg Thrust, the rocks record mostly ductile deformation, including superposed folds sets,
transposed cleavages, and ductile shear bands. In between, the outcrops exhibit an interesting
interplay between ductile and brittle styles of deformation. This interplay has generated a
spectacular variety of mesoscopic (outcrop scale) structures. These include many different
types of sense of shear indicators that provide a wealth of information on the slip history of two
thrusts, as well as the several phases of deformation that predate and postdate thrust faulting.
In addition to changes in the overall style of deformation, the variety of structures
preserved along the transect collectively record a first-order increase in finite strain toward the
east, with local maxima occurring within a few hundred meters of both the Champlain and
Hinesburg thrusts. In the footwall of the Champlain Thrust, F1 folds of bedding planes (S0)
tighten as their axial planes rotate from steep and moderately east-dipping to shallowly eastdipping (Figs. 5b, 5c). The styles and mechanisms of these folds also change from localized faultbend folds several kilometers below the thrust (Fig. 5b), to penetrative fold trains that formed
9
by a combination of interlayer slip and ductile flow near the thrust (Fig. 5c). The appearance of
two cleavages reflects this increase in finite strain. These include an early penetrative slaty
cleavage (S1) that formed during F1 folding and a second localized pressure solution cleavage
(S2) that marks the presence of intraformational thrusts (Fig. 5b). A similar increase in strain
occurs near the Hinesburg Thrust. At Stop 8, part of a faulted anticline lies structurally below
the Hinesburg Thrust. Here, isoclinal intrafolial folds of bedding (S0), stretched pebbles and
disarticulated compositional layers reflect a generally high magnitude of finite strain. Stop 4,
where the Hinesburg Thrust is exposed at Mechanicsville (Fig. 5e), records even higher strains
in mylonitic rock of the Cambrian Cheshire Formation.
Figure 5. Simplified diagram showing the regional structural trends from west to east.
Another interesting regional trend is the influence of rock type on the style and partitioning
of deformation within the section. In general, deformation associated with the emplacement of
the two major thrust sheets is expressed differently in competent units than it is in the weaker
shales. For example, in the footwall of the Champlain Thrust variations in the thickness and
abundance of competent limestone layers have produced distinctive fold styles. In the shale
units, ductile flow during contraction resulted in recumbent isoclinal folds that became rootless
at high strains. In contrast, thick competent limestone layers deformed mostly by interlayer
slip, resulting in large inclined folds, preserve numerous en echelon vein sets. A similar pattern
exists at the regional scale where most of the deformation that accompanied the formation of
the Champlain Thrust is partitioned into the weak Stony Point Shale units in the footwall. In this
latter locality, the deformation is widely distributed. In contrast, deformation in the thick,
10
competent quartzite layers of the Monkton Formation in the hanging wall tends to be more
localized and mostly involves interlayer slip (Fig. 5d).
This influence of lithology and rheological contrasts on structural style also has resulted in
many different types of kinematic indicators throughout the section. At Stop 4, competent
metapsammite layers located above the Hinesburg Thrust (Fig. 5e) preserve asymmetric vein
sets and folds that record a top-to-the-northwest sense of shear. In the weaker pelitic layers it
is recorded mostly by shear band cleavages. Although these structures generally show similar
top-to-the-west and -northwest senses of motion, the wide variety of types reflect different
starting materials. These and many other examples illustrate one of the basic principles of
interpreting the great variety of structures observed along this transect: differences in the
strength and rheology of the rock units as they deformed can explain much of the great variety
of structures observed in the Champlain Valley and in the lithotectonic slices to the east.
Since brittle structures, with the exception of normal faults, are not portrayed on Figure 5,
we will give a brief summary of the characteristics of the dominant fracture sets. Fractures that
have strikes orthogonal to the dominant planar fabrics (E-W to NE-SW) and steep dips are
common throughout the field area. Since Cretaceous dikes intruded along many of these
fractures, we know that these fractures are at least Cretaceous in age. Some fracture sets have
north-south strikes with moderate-steep dips and can sometimes be associated with fracture
cleavages associated with late generation folding (Fig. 5C). NW-SE trending steep fractures are
also common, but are of uncertain origin.
Metamorphism
Stanley and Wright (1997) summarized that the Taconian foreland rocks of the
parautochthon and upper plate of the Champlain Thrust are “essentially unmetamorphosed”
(p. B1-1) with temperatures of ~200°C and pressures corresponding to depths of ~2.5 km.
Stanley and Sarkisian (1972) reported prograde chlorite in fractures in the Monkton Formation
in the Upper Plate of the Champlain Thrust, and used this occurrence to suggest that this thrust
underwent multiple episodes of motion.
On the basis of field and petrographic observations presented by Strehle and Stanley
(1986), Stanley et al. (1987), and this volume (Stop 4), the metamorphic rocks from the
westernmost Taconian hinterland (upper plate of the Hinesburg Thrust), reached
chlorite/sericite grade. In the field area, there is a pronounced metamorphic contrast between
the rocks above and below the Hinesburg Thrust.
Geochronology
The maximum age of motion (first episode) on the Champlain and Hinesburg thrusts in the
field area are constrained by the youngest stratigraphic ages of rocks located below these
faults. In the case of the Hinesburg Thrust, the maximum age is Middle Ordovician (Bascom
Formation, Oba) whereas for the Champlain Thrust it is Late Ordovician (Stony Point Shale,
Osp). Rosenberg et al. (2011) used the K/Ar method to obtain ages of illite from the fault zone
of the Champlain Thrust at Lone Rock Point in Burlington. The ages they obtained range from
Carboniferous (~325 Ma) to Late Jurassic (~153 Ma). These authors speculated that post11
Taconian illite growth may reflect fluid flow associated with the Alleghenian Orogeny and the
Jurassic–Cretaceous unroofing of the Adirondacks and New England (e.g., Roden-Tice, 2000;
Roden-Tice et al., 2009).
Mesozoic igneous rocks (McHone and McHone, 2012)
Mesozoic dikes and small intrusions can be found in Vermont. There are about 1100 dikes
on the Mesozoic intrusion map by McHone (1984; Fig. 6) but many more exist, and new ones
are frequently exposed by construction projects. Field work for the new Vermont bedrock map
(Ratcliffe and others, 2011) added a hundred or more dikes, and this excellent map is one of the
few that show dikes on a state-wide scale (but you have to look close!). Many of the dikes must
be co-magmatic and connected, that is, mafic magmas have probably split and branched from
larger into smaller dikes as they rose from mantle sources. These were fluids of relatively low
viscosity due to high volatile contents, so they moved quickly with mainly laminar flow
McHone, 1978). There is little to no evidence that the small dikes reached the surface, which
was probably a few km above the present levels in Early Cretaceous times, but it is possible that
some created small volcanoes. A big volcano must have existed over the current Ascutney
Mountain plutonic complex, which contains large inclusions interpreted as volcanic products
that settled into the upper magma chamber (Daly, 1903). The same is true for some large
Mesozoic plutons in New Hampshire (Creasy and Eby, 1993).
Dike rocks of the different generations can often be categorized through careful
examination of hand samples. Most of the dikes are fine-grained but holocrystalline as well as
porphyritic, and many have obvious igneous structures such as flow bands. Thin sections are
always very useful, and in some cases they are indispensable for characterizing the petrography
and classification. It is important to improve our knowledge of the types and distribution of
Mesozoic intrusions because there appear to be definite boundaries to these igneous provinces
in New England (Fig. 6; McHone and Butler, 1984; McHone and Sundeen, 1995). The igneous
province boundaries could be due to tectonic controls such as terrane boundaries and major
faults, or specific mantle melt zones may be featured. We will discuss the physical and
geographic distinctions of these rocks during this trip.
Even before the determination of Mesozoic ages for plutons of the White Mountain magma
series (principally by Foland and others, 1971 and Foland and Faul, 1977), many geologists
lumped all post-metamorphic igneous rocks of northern New England as White Mountain
Magma Series, commonly labeled WMMS. Because most of the WMMS dikes, stocks, and
batholiths display alkalic igneous characteristics (Ti-rich clinopyroxene and alkali amphibole;
abundant alkali feldspar; high K, Na and Th contents), it has seemed logical to relate them to
differentiation or fractionation of intra-plate mantle or crustal melts that have some common
genesis (Foland and others, 1988; Creasy and Eby, 1993). Although some examples show
hydrothermal alteration and igneous fabrics, the Mesozoic igneous rocks lack metamorphic
foliations and S-type granites that distinguish the common Paleozoic plutons in New
Hampshire.
Additional work (McHone and Butler, 1984) reinforced the age divisions of Foland and Faul
(1977) into groups with Middle Triassic (220-235 Ma), Early to Middle Jurassic (175-195 Ma)
and Early Cretaceous (100-130 Ma) ages, with a few important exceptions (Belknap Mountain is
12
probably close to 160 Ma). We use names and acronyms of Coastal New England (CNE), White
Mountain Magma Series (WMMS), and New England-Quebec (NEQ) for the provinces, from
older to younger. Additional and separate tholeiitic magmatism around the Triassic-Jurassic
boundary (201 Ma) filled Early Mesozoic rift basins of eastern North America
Figure 6. Mesozoic igneous province of New England; figure from McHone and McHone (2012) NEIGC field trip
guide and modified after McHone and Butler (1984). Province abbreviations: CNE = Coastal New England; WMMS =
White Mountain Magma Series; NEQ = New England–Quebec
13
with thick flood basalts derived from volcanic fissures, which are now shown by very large
feeder dikes (McHone, 1996). The largest of these giant dikes is the Higganum-HoldenChristmas Cove dike (Fig. 6). The younger Bridgeport-Pelham Dike trends toward west-central
New Hampshire and may be present near Claremont (Filip, 2010), but more confirmation is
needed. Apparent boundaries for Mesozoic igneous provinces in northern New England are
shown in Figure 6.
A major problem is that Early Jurassic WMMS intrusions include many alkalic diabase and
lamprophyre dikes that are similar to members of the Cretaceous NEQ and the Triassic CNE
provinces (described below). To date, however, all such look-alikes are found only in New
Hampshire and western Maine (McHone and Trygstad, 1982; McHone, 1992; McHone and
Sundeen, 1995), a geographic association that makes it easy to call them members of the White
Mountain Magma Series. McHone and Butler (1984) proposed that the alkalic Early Jurassic
plutons and dikes of the WMMS are a cohesive province only in central and northern New
Hampshire (into northeastern-most Vermont), and western Maine (at least to the Rattlesnake
Mountain pluton). We do not know of any proven Triassic or Early Jurassic dikes in Vermont,
only Early Cretaceous intrusions of the New England-Quebec igneous province.
Although magmas of all the Mesozoic divisions overlap in New Hampshire, the Early
Cretaceous intrusions range across a much wider province that includes the Monteregian Hills
of southern Quebec, dense swarms of lamprophyre dikes in western Vermont and eastern New
York, and scattered dikes through New Hampshire and southern Maine (McHone, 1984;
McHone and Sundeen, 1995). Based on similar ages and petrological characteristics, the Early
Cretaceous intrusions were regrouped as the New England-Quebec (NEQ) igneous province by
McHone and Butler (1984).
Local dike types include alkali diabase, monchiquite, camptonite, spessartite, and bostonite.
All of these igneous types are presumed to be related through some upper mantle to upper
crustal paths of fractional melting, differentiation, contamination, and crystallization, but it is
unlikely that they were at one time all co-magmatic. However, the alkali syenite and gabbro
members of the Ascutney Mountain complex are derived from the same magmatic sources as
the dikes (also see Eby, 1985; Schneiderman, 1991).
Alkali diabase is relatively plagioclase rich although the feldspars may be altered to clay or
sericite. Augite and minor biotite in the groundmass often look oxidized, and some contain
small but completely altered olivine crystals. Most alkali diabase dikes are fine-grained and nonporphyritic, but some show prominent plagioclase phenocrysts. The diabase is essentially
altered dolerite in dikes of hypabyssal alkali basalt or sub-alkaline tholeiite. Monchiquite is a
very mafic, granular, analcite-bearing, olivine or biotite-bearing, augite-rich alkali basalt similar
to nephelinite, often also with calcite in separate grains or clumps (or replacing mafic minerals),
phlogopitic mica, and kaersutitic hornblende. Feldspars are poorly developed or lacking.
Monchiquite is commonly dark gray in color and relatively dense, but phenocrysts of mafic
minerals are visible. Camptonite can look much like monchiquite, except that olivine is rare or
absent, kaersutite is common to abundant, and plagioclase is more abundant than analcite.
Phenocrysts are also mafic (augite and/or kaersutite), only very rarely felsic. Camptonite dikes
usually have a brownish to gray range of colors, lighter than monchiquite. Camptonite is the
hypabyssal equivalent of basanite.
14
Spessartite dikes lack olivine and analcite, but plagioclase (intermediate Ca) is well
developed and present as phenocrysts as well as intergrown with augite in the groundmass.
Phenocrysts or megacrysts of kaersutite serve to distinguish spessartite from alkali diabase
dikes common in eastern New England. Spessartite often shows a distinctly greenish or purplish
cast as well as gray colors. Bostonite is a name that in a strict sense applies only to felsic
(anorthoclase-rich) dikes that have a "felty" clumped-grain feldspar texture, which is not always
present. Trachyte, although used for volcanic rocks as well, is a better general term. Minor
minerals include oxidized biotite, quartz, and clay products. Some examples show well-formed
alkali feldspar and/or quartz phenocrysts. Trachyte dikes may be iron-stained, but they are
generally light brown to cream-colored on fresh surfaces. Felsite dikes associated with the
Ascutney Mountain plutonic complex (Balk and Krieger, 1936) may be bostonite.
Although small, lamprophyre dikes of the New England-Quebec province are numerous and
widespread, with a consistent range of compositions across the 400+ km width of the province
(McHone, 1978). The dikes both predate and postdate the larger plutonic complexes (although
perhaps not by much), while their distribution suggests that the alkalic lamprophyres are not
offshoots from those plutons. Major and trace element compositions, and isotopic ratios of the
lamprophyres are similar to world-wide alkali-olivine basalts, basanites and nephelinites
(McHone, 1978; Eby, 1985), although lamprophyres have higher concentrations of H2O+ (1–3
%) and CO2 (commonly 2–4 %). As per the definitions of Rock (1977), alkali lamprophyres
generally have phenocrysts of mafic minerals but few or none of feldspar.
Work by Hodgson (1968), McHone (1978), Eby (1985) and others shows that the regional
lamprophyric magmas cannot be derived from a common parent through differentiation or
assimilation, but must instead reflect different initial compositions followed by crystal and
chemical fractionation. Eby (1985) has argued that isotopic characteristics of the dike rocks
show their origin to be partial melts of a heterogeneous spinel lherzolite mantle.
Mantle xenoliths in the North Hartland dike and other localities in New England and Quebec
are spinel-bearing types that represent mantle rocks above the source melts, with a few highpressure cumulates from ponded magmas (Raeside and Helmstaedt, 1982). Thus, the xenoliths
indicate a direct mantle origin for lamprophyres in common with other alkalic basalts. Melting
of the mantle to form such magmas may depend upon metasomatic events that add K, Ti,
water, and other components (Windom and Boettcher, 1980; Eby, 1985).
Mesozoic uplift and fault tectonism in New England have been shown to extend into
Cretaceous and possibly later times by apatite and zircon fission track studies (Doherty and
Lyons, 1980; Roden-Tice and others, 2012). Regional bedrock formations, metamorphism, and
orogenic events may be old, but the mountains, valleys, and other landforms we see today are
relatively young. As Yngvar Isachsen explained to us [sic; i.e., the McHones] (pers. comm. 1973;
Isachsen, 1981), erosion rates are too fast for high-relief mountains to be very old, and
relatively high elevations of the Appalachians and Adirondacks are due to Late Mesozoic and
Tertiary uplift (Roden-Tice and others, 2009).
Thus some major high-angle faults in the region have significant offset dating from Middle
Jurassic times, with activity continuing to affect NEQ dikes as well. Several mafic dikes in this
zone are cut by high-angle faults with offsets of a few meters or less, although these may only
be sympathetic movements along particularly weak foliation planes, which tend to be steeply
15
dipping and sub-parallel to the major faults. The Early Cretaceous NEQ intrusions are aligned
with NW-SE dike trends in Quebec and possibly northwestern Vermont (McHone and Shake,
1992), but large WMMS plutons trend more N-S in New Hampshire (Fig. 1). Dike trends across
eastern Vermont through New Hampshire and western Maine vary, but most commonly they
are NE-SW (McHone, 1984). We see little evidence that dikes intrude faults, but they do follow
local joints and/or cleavage foliations in many outcrops.
----------------------------------------Stop 1: The Champlain Thrust at Lone Rock Point. Episcopal Diocese of Vermont, 5 Rock Point
Road, Burlington, VT 05408. Pull into and park in a small parking lot on the right across
from a low brick building labeled “Diocesan Offices.” The GPS coordinates of this parking
lot are 44.496668°, -73.240245°. Description is from West et al. (2011).
The Episcopal Diocese maintains a series of hiking trails in this area that are open to the
public. However, during normal weekly office hours, visitors should check in at the Diocesan
Office and obtain permission to use the trails (this is never a problem). The outcrops of interest
are about a 15 minute walk from the parking lot. From the parking area walk along the paved
road towards the west (in the direction you were driving towards) and you will soon cross over
a small bridge over the bike trail. Immediately past this bridge, bear to the left following signs
for “Trails.” This road eventually becomes unpaved and eventually turns into a footpath just
beyond a farmhouse. Continue on the footpath into the woods and within 50 meters bear to
the left on a trail marked by a sign for “OverThrust.” This trail will lead into a small grassy field,
and then back into the woods where it will soon descend steeply down to the shores of Lake
Champlain. Upon reaching the shore, walk north (to the right) about 20 meters and the feature
of interest will be very apparent! This outcrop is located at 44.490956°, -73.248892°.
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. Similar to what
we will observe at Mechanicsville (Stop 4) and Mt. Philo (Stop 2), the “older-on-top-ofyounger” 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—based on the fact that the Lower Cambrian Dunham
Formation is found above the Middle Ordovician Iberville Formation (Fig. 7). Estimates of fault
displacement (the amount of offset accommodated by the fault) are ~80 km (Stanley, 1987),
which is significantly higher than the calculated displacement on the Hinesburg Thrust.
An additional 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 dictate 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
16
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. 7a). Deformation
is not restricted to slip along the fault plane, and can be dived into two domains as described by
Figure 7: Sketches showing the general geometry of features observed at the Champlain Thrust fault
at Lone Rock Point. (A) Block diagram of the Champlain Thrust fault with the resistant Dunham
Formation (massive dolostones) above the weak Iberville Formation (calcareous shales). The faultzone is divided into two domains consisting of the inner fault-zone and the outer fault-zone. The inset
sketch depicts the relationships between folded and faulted quartz-rich layers (dark gray), clay rich
shale material (light gray), alignment trajectories of clay minerals (dashed lines), and the geometries
of calcite veins (white). (B) Block diagrams of two mechanisms of structural slickenline formation
(modified from Means, 1987). The left diagram shows the formation of fault corrugations while the
right diagram illustrates gouging caused by resistant material within the upper plate. (C) Formation of
mineral slickenlines by the infilling of voids caused by steps within the fault surface.
17
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.
Mullions. 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 7 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 7b).
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 7a).
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 7c).
Cleavage. Cleavage planes are formed from the preferential alignment of mineral grains due
to flattening, which is accommodate 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.
Secondary 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
18
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.
----------------------------------------Stop 2: The Champlain Thrust at Mt. Philo and regional overlook (with lunch). Mount Philo
Parking Lot, State Park Road, Charlotte, VT 05445. The GPS coordinates of this parking
lot are 44.278557°, -73.222327°. Description is from West et al. (2011) and Kim et al.
(2011). See Figure 8 for a park map.
This stop is designed to give a geomorphic overview of the four major tectonic zones in
west-central Vermont and eastern New York. From west to east, and from structurally lowest to
highest, these zones include: 1) the autochthon (rocks that have not been significantly
transported from their site of formation/deposition), 2) the parautochthon (rocks that have only
been transported a short distance), 3) the upper plate of the Champlain Thrust, and 4) the
upper plate of the Hinesburg Thrust.
At the famous Lone Rock Point exposure of the Champlain Thrust in Burlington (e.g.,
Stanley, 1987), the fault juxtaposes the Lower Cambrian Dunham Dolostone with the Iberville
Shale; this suggests the presence of an along-strike ramp that climbs ~610 m up section
between Burlington and Mt. Philo. The stratigraphic throw of the Champlain Thrust is ~2743 m
at Lone Rock Point and ~1830 m at Mt. Philo (Stanley, 1987). Total displacement along the
Champlain Thrust ranges from 55–100 km (34–62 miles; Stanley, 1987; Rowley, 1982).
The autochthonous rocks on the west side of Lake Champlain include Middle Proterozoic
metamorphic rocks of the Adirondack massif that are uncomformably overlain by early
Paleozoic sedimentary rocks of the Beekmantown Group (NY State Geologic Map; Isachsen and
Fisher, 1970). The autochthon is structurally overlain by 1) Late Cambrian–Middle Ordovician
sedimentary rocks of the parautochthon, 2) very weakly metamorphosed sedimentary rocks of
the upper plate of the Champlain Thrust; and 3) low grade (chlorite-sericite) rift clastic
metasedimentary rocks of the Upper Plate of the Hinesburg Thrust.
19
Faults of inferred Mesozoic age are exposed along the cliff faces of Mt. Philo, and may be
related to the formation of a graben between the Adirondacks and the Green Mountains.
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.
Figure 8. Map of Mt. Philo State Park. From the lower parking lot, field participants have the option to take a van
shuttle to the top of the mountain, or to walk the auto road and/or trails. The Middle Ordovician Stony Point shale
formation comprises the lower plate of the Champlain Thrust here; outcrops are exposed in roadcuts exposed just
uphill of the auto road intersection. The Devil’s Chair Trail is recommended for the most extensive outcrops of the
lower dolomotic sandstone member Monkton Formation (upper plate of the Champlain Thrust) and brittle faults;
please note that this trail is steep, narrow, and rocky in places. Exposures of the upper member of the Monkton
Formation, a ferruginous quartzite, are exposed along the old carriage road and overlook ridgeline.
At the top of Mt. Philo, you are standing on the upper member of the Middle Cambrian
Monkton Formation, which is a deep red-brown-colored ferruginous quartzite. The lower
dolomitic sandstone member of the Monkton Formation is exposed just below these red
quartzites along the cliffs below (excellent exposures can be seen on Devils Chair trail). Both of
these members of the Monkton Formation are within the Upper Plate (hanging wall) of the
Champlain thrust fault. The low-lying area between Mt. Philo and Lake Champlain is underlain
by rocks of the Parauthochthon structurally beneath the Champlain thrust fault. At Mt. Philo,
the fault thrusts Middle Cambrian rocks of the Monkton Formation over Middle Ordovician
rocks of the Stony Point Formation (older rocks on top of younger rocks violates the principal
of superposition). From the overlook, the thrust surface is approximately 50 m (160 ft) below
your feet. The Stony Point Formation is not very well exposed at Mt. Philo, but small overgrown
road cuts of calcareous shale can be found about 100 m downhill from where the House Rock
Trail crosses the Auto Road.
20
The Champlain thrust fault is a regional scale north-south trending structure in western
Vermont extending from well south of Middlebury to north of the Canadian border (over 300
km in length: see Stanley, 1987 for details). On a clear day from the top of Mt. Philo, one can
easily see the geomorphic expression of the Champlain thrust fault as the steep western slope
breaks on Shellhouse, Buck, and Snake Mountains to the south, and on Pease Mountain to the
north. At each of these localities, erosionally-resistant older rocks of the Monkton Formation
have been thrust towards the west over less-resistant younger carbonate/shale rocks.
----------------------------------------Stop 3: Charlotte Beach Mesozoic dikes. 1375 Lake Road, Charlotte, VT 05445. The GPS
coordinates of the parking lot are 44.333960°, -73.281574°. Description is from West et
al. (2011).
Low outcrops along the edge of the lake expose dark gray graphitic limestones with
interbedded shale of the Late Ordovician Stony Point Formation. These rocks have been
intruded by a series of tan-colored, fine-grained, syenite dikes of probable Cretaceous age. The
Stony Point Formation lies structurally beneath the Champlain thrust fault and is part of the
parautochthon (a thrust sheet that has only been transported a short distance), which is the
lowest structural level in the field area. At this location, bedding in the sedimentary rocks
strikes approximately north-south and dips steeply to the east.
The steeply dipping dikes (a total of seven with a combined thickness of about 2 meters)
exploited the east-west trending fractures in the country rocks that are common throughout
the area. McHone and McHone (1999) refer to the dike rocks as “bostonites”—a name that has
largely disappeared from modern petrologic literature, but refers to fine-grained hypabyssal
(shallow depth) intrusive rocks rich in alkali feldspar. The dikes are likely correlative with the
Barber Hill stock (located about 3.5 km southeast of here) which has been dated at 111 ± 2 Ma
(K/Ar biotite age; Armstrong and Stump, 1971). A whole rock Rb-Sr isochron age of 125 ± 5 Ma
on seven trachyte dikes from the Burlington area was reported by McHone and Corneille
(1980), and probably provides an upper limit on the age of these dikes.
----------------------------------------Stop 4: Hinesburg Thurst (note the Rolfe Stanley commemorative plaque). Place Rd near
intersection with Pond Rd, Hinesburg, Vermont. The GPS coordinates of outcrop are
44.352129°, -73.107476°. Description is from West et al. (2011) and Kim et al (2011).
Background. Mechanicsville is an important locality in western Vermont because it provides
a superb three-dimensional exposure of the Hinesburg thrust “fault”, which—as we will
describe below—is better referred to as a “shear zone.” This feature is one of several large
thrust faults exposed in the Champlain Valley that formed during the Ordovician Taconic
Orogeny (e.g., Stanley and Ratcliffe, 1985). The generally north-south trending Hinesburg thrust
21
juxtaposes moderately metamorphosed rocks of the Green Mountain belt (i.e., Fairfield Pond
and Pinnacle formations) to the west against the very weakly metamorphosed early Paleozoic
rocks of the Champlain Valley belt to the east.
Interpreting the Mechanicsville exposures. One of the first questions one might ask about
the Mechanicsville outcrop is “How do we know it is a thrust fault?” There are several ways to
address this question. First, by definition, thrust faults place older rocks on top of younger
rocks. At Mechanicsville, an irregular, undulating surface separates fractured and sheared
dolomite of the Ordovician Bascom Formation below from fine-grained schist and quartzite of
the Neoproterozoic to Early Cambrian Fairfield Pond Formation above (Fig. 9). This “older-ontop-of-younger” stratigraphic relationship across this surface is the first suggestion of thrust
faulting (basically we have 550 million year old rocks sitting on top of 480 million year old
rocks). Second, the sense of motion involved in reordering the stratigraphy in this way implies
that the schist that now forms the upper plate was once located at greater depths than rocks in
the lower plate. In support of this claim, the schist in the upper plate displays greenschist-facies
metamorphic mineral assemblages indicative of temperatures ≥350°C (Strehle and Stanley,
1986; Stanley et al., 1987). In general, temperatures increase with depth in thrust belts,
implying that rocks of the Fairfield Pond Formation were transported from deeper levels in the
crust than the colder, virtually unmetamorphosed rocks of the Bascom Formation. This
interpretation is consistent with the interpretation of Stanley and Wright (1997), who
suggested a total displacement of ~6.4 km on the Hinesburg thrust.
Whereas inferring the sense of motion on fault surfaces using stratigraphic and
metamorphic relationships is useful, it is even better to be able to identify structures that more
directly record the history of motion on the fault itself. This is another reason why the
Mechanicsville exposures are so valuable: they preserve a superb array of these kinds of
structures (Fig. 10). In general the category of structures that record how fault blocks once
moved is called sense-of-shear or kinematic indicators. The great variety of these indicators at
Mechanicsville allows us to reconstruct the history of motion on the fault.
Upon closer inspection of this outcrop, one may notice that, in addition to the differences in
color and rock type, the rocks above and below the fault appear to have responded differently
to deformation. The layers of schist above the fault display a very fine grain size and preserve
an abundance of folded layers. Some of these layers have been folded so tightly that they have
rotated onto their sides and have limbs that parallel one another (Fig. 9, bottom). In other
places, the tight folds are themselves folded again in large trains that lean to the northwest
(toward the left side of the outcrop). In the quartzite layers, white quartz-filled fractures (veins)
form S-shapes that taper to narrow points at their tips (the old ones tend to be S-shaped, the
young ones tend to be straighter). The quartz in the S-shaped veins display a milky-white
appearance where it is difficult to see individual crystals, indicating that they have recrystallized
into new smaller grains. These features are typical of regions where rock as deformed by
ductile flow rather than breaking.
22
Figure 9: Block diagram (top) and sketch of cliff face (bottom) of the
outcrops at Mechanicsville. Note the distinction between the Hinesburg
Thrust and the Hinesburg Thrust Shear Zone.
Ductile deformation involves changes in a rock’s shape by flowing much like plastic does
when it is heated over a stove. As they flow, the minerals become stretched, attenuated, and
thinned to the extreme. The overall grain size of the rock is reduced as the minerals deform and
recrystallize by crystal-plastic processes. This type of fine-grained crystalline rock, which is
called a mylonite (after the Greek word ‘mylos’, meaning to mill down), are diagnostic of zones
of concentrated ductile flow called shear zones. Shear zones are simply ductile fault zones that
tend to form in areas, such as at depth, where temperatures are in excess of 250-350°C or in
zones where deformation is slow enough to avoid breaking the rock. At Mechanicsville the
mylonitic schist located above the fault that separates the Fairfield Pond and Bascom
Formations define the Hinesburg thrust shear zone. This zone of ductile flow may represent an
initial phase of deformation that occurred when the rocks of the Fairfield Pond Formation were
23
still deeply buried and flowing ductilely. Later, as displacement on the thrust accumulated and
mylonitic rocks were uplifted to shallower depths, the deformation became brittle and was
Figure 10: The exposures at Mechanicsville (Stop 4) preserve a great variety of structures that record the history of
motion in the Hinesburg Thrust Shear Zone. These include: (a) Rotated, sigma-type clasts. These structures form
because the clasts, which can be of any size, are more rigid than the surrounding ductilely flowing matrix. Because
they are stronger, the clasts help shield the matrix on its flanks from high strains during flow. These shielded areas,
or pressure shadows, form a wedge-shaped area around the clast that becomes warped as the object rotates. The
name sigma clast comes from the resulting asymmetric shape, which resembles the Greek letter σ. (b) Shear bands.
Shear bands are thin (typically less than 1 mm) bands of very high shear strains that mimic the geometry of the
larger shear zone. Between parallel shear bands, foliation planes are deflected toward parallelism with the bands.
The sense of rotation of the foliation as it approaches the shear band yields the sense of shear. (c) Sheared quartz
veins. These veins initially formed in relatively strong quartzite layers as a result of high fluid pressure, which caused
fracturing. After they formed, the fluid pressure was reduced and the rock began to flow again. Where the tips of
some of these veins penetrate into the schist on either side of the quartzite they are asymmetric. This asymmetry
results because the schist flows more easily than the quartzite. The sense of rotation of the vein tips as they
approach the schist yields the sense of shear. (d) Asymmetric folds. Asymmetric folds can yield a sense of motion if
you can find areas where the folds record different intensities. The intensity of folds can be measured qualitatively
by their tightness or the degree of parallelism of their limbs. To use this type of indicator, find a fold whose limbs
are close to parallel (i.e., a tight to isoclinal fold). Draw a line that connects the point of maximum curvature in the
hinge region of each successive layer that is folded (called the axial trace). Now find a fold whose limbs are not
parallel (an open fold) and find its trace. The sense of rotation from trace of the open fold to the trace of the tight
fold yields the sense of shear.
superimposed on the older ductile structures. An alternative explanation suggests that the
younger, brittle faulting formed during the Devonian Acadian orogeny and is unrelated to the
older (Taconic) ductile shear zone.
Some of the most spectacular sense-of-shear-indicators occur in ductile shear zones such as
the Hinesburg Thrust. An interesting observational exercise one can complete at this outcrop is
to determine the sense of motion in the Hinesburg Thrust. There are several steps to doing this.
24
First, one must look for certain kinds of small-scale structures that record increments of
deformation in the shear zone. These types of structures tend to be curved and are asymmetric.
The asymmetry and curvature results because the structures rotate in the flowing matrix as
they form. Because they do not form instantaneously, but take some finite period of time to
grow, the oldest parts of these structures will have rotated slightly more than the youngest
parts, resulting in these types of structures shown in Figure 10. This figure provides schematic
illustrations of four types of asymmetric sense-of-shear-indicators observed at Mechanicsville
and how to interpret their geometry. The sense of rotation determined by the asymmetry
yields the sense of shear in the shear zone.
Second, one must look for these asymmetric structures on the right surface. The correct
surface to look on is one oriented perpendicular to the shear zone foliation and parallel to its
mineral lineation. A shear zone foliation forms as minerals are flattened and align into planes
during ductile flow. In addition, as the minerals are stretched, some crystals or groups of
crystals are smeared out into stretching lineations on these shear planes. See if you can find the
aligned chlorite and sericite on the foliation surfaces at Mechanicsville (Hint: look carefully at
the undersides of foliation surfaces). These latter structures are important because they
indicate the direction of flow. The orientations of mineral lineations in this shear zone indicate
that the direction of motion was either to the NW or to the SE. If the upper plate moved down
to the southeast, the shear zone is a ductile normal fault. If it were up to the northwest it is a
thrust fault. Only the asymmetric structures shown in Figure 10 reveal which interpretation is
correct.
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 9, 11a). 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 (Fig. 9). The V3 vein set is only weakly deformed and less recrystallized than the V2 veins,
and is mostly symmetric to slightly asymmetric. A black, quartz-poor pressure solution selvage
surrounds the veined quartzite layers (Fig. 11a) 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 crystalplastic 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 (Fig. 9).
The fanning of the V2 vein sets around fold hinges is a good indicator that they are folded (Figs.
11c, 11d). The tightest folds are recumbent and tend to occur nearest the Hinesburg Thrust
25
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 (Figs. 9, 11d). 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 (Fig. 11d). 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 shear sense given by the F2 folds and the asymmetric geometry of the superposed V2
and V3 veins sets offer important clues about how the veins formed. The concentration of the
veins within the quartzite layers suggests that these layers were more competent than the
surrounding schists. Assuming that the V2 and V3 vein sets reflect similar stresses, a comparison
of the younger vein set with the older one yields a top-to-the-northwest rotation sense
identical to that indicate by the F2 folds (Fig. 11b). 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 (Fig. 11a). 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 (Fig. 11b). 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 continues to open during shearing. 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 and the latter record finite strains. 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 (Fig. 11). 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 (Fig. 11c). 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.
26
Figure 11. Cartoon showing the preferred model of progressive formation of F2
fold and veins structures in the mylonitic hanging wall of the Hinesburg Thrust
at Mechanicsville. This model requires the veins to finish opening prior to the
onset of noncoaxial shear.
Stage 4 (Fig. 11d). As noncoaxial shear continues the F2 folds continue to rotate and tighten,
recording a progressive increase in finite strain during ductile thrusting. The S2 crenulation
cleavage rotates to the northwest along with the folds. The V2 veins exhibit a characteristic
27
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. This fold geometry mimics a kilometer scale dome
and basin interference pattern formed by north-plunging (F3) and east-plunging (F4) folds across
the field trip area (e.g., Five Tree Hill and Oak Knoll localities). These orthogonal fold sets are
among the youngest ductile structures preserved at Stop 4. On the thrust surface itself two
orthogonal crenulation lineations (L3, L4) mark the presence of the corrugation folds. 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.
Figure 12. Block diagram showing the relative geometry of orthogonal fold sets that deform the
Hinesburg Thrust surface and the mylonites at Mechanicsville, producing a dome and basin pattern.
The two fold sets (F3 and F4) are associated with two steeply dipping orthogonal cleavages.
28
BONUS: Glacial fun facts (courtesy of Stephen Wright, UVM Geology)
The ice sheet, when its terminus was at
its farthest extent (Cape Cod/Gulf of
Maine; Fig. 13) was between 2.5 and 3
km thick across northern Vermont, i.e.
all the New England mountains were
under ice.
The large-scale motion of the ice sheet
was from NNW to SSE when the ice was
thick enough to cover the mountains.
Glacial striations along the crest of the
mountains are all in this orientation.
Ice flow was topographically controlled
when the ice sheet thinned, e.g. in the
Champlain valley and Route 100 valley
(Stowe) ice flow shifted to ~N-S.
Vermont was deglaciated between
~15,500 (southeast corner) and ~13,200
(northwest corner) years ago. The
mountainous area between Stowe and
Burlington between ~13.8 and 13.5 ka.
(Like yesterday compared to those
rocks!)The retreating ice front dammed
glacial lakes in north- and west-flowing
Figure 13. Major glacial lakes in New England (from drainages (e.g. Champlain valley,
Rittenour et al., 2000).
Winooski Basin, Lamoille Basin)
Glacial Lake Winooski is the biggest and highest level lake in the Winooski and Lamoille valleys (Stowe
was underwater). Outlet (low point in the bathtub) is the drainage divide ~5 km south of
Williamstown (Rt 14 valley south of Barre) at 915 ft (229 m). Isostatic rebound has tilted the land
surface upwards to the NNW (mimicking the regional ice flow) such that the elevation of Glacial
Lake Winooski in the Stowe area was ~1,100 ft (if participants drive up Rt 108 from Stowe towards
Smuggler’s Notch, they’ll climb up on one of the deltas that formed in this lake just before reaching
the base of the toll road).
There was at least one “catastrophic” drainage event when the ice retreated far enough west, down the
Winooski River valley to uncover a lower outlet.
Glacial Lake Vermont in the at the latitude of Burlington was at ~620 ft at its high stage (UVM under
~300 ft of water) and about 100 ft lower at its low stage.
A higher-elevation glacial lake in the Lake Ontario basin catastrophically flooded into Glacial Lake
Vermont when the ice sheet retreated north of the Adirondacks.
These merged lakes (the lower level stage of Glacial Lake Vermont) expanded into Québec as the ice
sheet retreated northwards
Glacial Lake Vermont catastrophically flooded into the North Atlantic when the ice sheet retreated north
of the St Lawrence River valley ~12 ka.
The isostatically-depressed land surface was flooded with ocean water forming the Champlain Sea.
Isostatic uplift and stabilizing eustatic sea level (most of the ice sheet had melted) cut off access to the
North Atlantic initiated the modern fresh water Lake Champlain ~8 ka.
29
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