J. metamorphic Geol., 1999, 17, 271–286
Syn-tectonic pluton intrusion during contractional deformation:
microstructural and metamorphic evidence from the aureole of the
Acadian Victory Pluton, north-eastern Vermont, USA
K . A . H A N N U L A , J. S . L A C KE Y , E . MAT T O X , G . M C G R AT H , E. O N A SC H A N D J . WE R T H E IM
Geology Department, Middlebury College, Middlebury, VT 05753, USA
A B S T RA C T
Structures within the aureole of the Acadian (Devonian) Victory Pluton suggest that movement along
the Monroe Fault occurred during pluton intrusion. Pre- or syn-tectonic garnet textures, including discordant internal and external foliations, sigmoidal inclusion trails and deflection of the matrix foliation
around garnet, are found along the Monroe Fault. In sillimanite+K-feldspar grade rocks, the matrix
foliation, defined by biotite and fibrolitic sillimanite, wraps around garnet and sillimanite porphyroblasts.
Granite dykes and sills near the pluton contact are folded and boundinaged, and leucosome fills garnet
pressure shadows and shear bands. Away from the fault and pluton contact, microstructures indicate
much less deformation during metamorphism, suggesting that deformation was partitioned into the fault
or the pluton. Deformation ceased before crystallization of the main body of the Victory Pluton was
complete. It is possible that magmatism facilitated deformation along the Monroe Fault, and/or that the
magma was transported to the mid- to upper crust along the Monroe Fault. This study suggests that
Acadian deformation in north-eastern Vermont occurred as late as 390–370 Ma.
Key words: Acadian; fault-related emplacement; pluton; syn-tectonic emplacement; Vermont.
I N T R O D U C T IO N
Unfoliated cross-cutting plutons, nearly circular in
map view, have traditionally been considered posttectonic (Paterson et al., 1991). Such plutons have
commonly been used to place maximum ages on
ductile deformation, terrane accretion and other tectonism (e.g. Naylor, 1971; Zen, 1983). Recent work,
however, has focused increased attention on the
relationship between magmatism and deformation.
Melting can weaken rocks, enabling or localizing
deformation (Hollister & Crawford, 1986; Davidson
et al., 1994). Deformation can provide pathways
through which magma can migrate by increasing the
permeability of crystalline rocks (Brown & Solar,
1998). In extensional and strike-slip settings, deformation may provide a solution to the space problem
for magma intrusion (Hutton et al., 1990; Paterson
et al., 1991). Furthermore, it is becoming clear that
even apparently undeformed, cross-cutting plutons
may in fact have intruded during deformation
(Davidson et al., 1992; Rothstein et al., 1994; Davidson
et al., 1996; Brown & Solar, 1998). Many of these
bodies are elongate and concordant with the regional
foliation; however some, including the Bergell Pluton
(Davidson et al., 1996) and several western Maine
plutons (Brown & Solar, 1998) are circular in map
pattern.
The Devonian plutons of north-eastern Vermont
appear to be classic examples of post-tectonic granites
(Fig. 1). They are circular in map pattern, they appear
© Blackwell Science Inc., 0263-4929/97/$14.00
Journal of Metamorphic Geology, Volume 17, Number 3, 1999
to cut across regional foliation, fold and fault trends,
they have well-defined contact aureoles, and they have
magmatic foliations if they are foliated at all (Doll,
1951; Eric & Dennis, 1958; Johansson, 1963; Woodland,
1965). The Victory Pluton, in particular, appears to
cut across the Monroe Fault, which separates Silurian–
Devonian rocks of the Connecticut Valley trough from
rocks of the Bronson Hill belt (Fig. 1). Microstructures
along the fault within the aureole of the Victory
Pluton, however, suggest that the pluton intruded
during fault movement. The Victory Pluton may
therefore be an example of a mid- to upper-crustal
pluton which used an active ductile thrust fault as a
conduit for melt migration.
BACKGROUND
Regional geological setting
The Victory Pluton is one of a group of Devonian
plutons that straddle the boundary between two major
subdivisions of the New England Appalachians: the
Connecticut Valley trough and the Bronson Hill belt
(Fig. 1). The Connecticut Valley trough (Hatch, 1987)
is underlain by two major Silurian–Devonian stratigraphic units (Doll et al., 1961; Hatch, 1987). The
mixed pelitic and quartzose Gile Mountain Formation
is late Early Devonian in age (Hueber et al., 1990),
whereas the underlying calcareous Waits River
Formation is Silurian in age (Aleinikoff & Karabinos,
1990; Armstrong et al., 1997) (Fig. 1). Connecticut
271
272
K. A . HA NNULA ET A L .
Fig. 1. Simplified geological map of Vermont and New
Hampshire (from Doll et al., 1961 and Billings, 1955). The box
labelled ‘2’ indicates the area of Fig. 2. The Northeast
Kingdom batholith of Ayuso & Arth (1992) is indicated by
darker borders around New Hampshire series plutons. Inset:
simplified map of Acadian tectonic zones (after Bradley, 1983
and Rast & Skehan, 1993). CVT, Connecticut Valley trough;
BHB, Bronson Hill belt; CMB, Central Maine belt. The
location of the Connecticut Valley trough and Avalonia is from
Bradley (1983). The Bronson Hill belt includes the Piscataquis
Volcanic Arc of Bradley (1983); the Central Maine belt
includes the Merrimack trough and the Kearsarge–Central
Maine belt of Rast & Skehan (1993).
Valley trough rocks lie above and east of rocks
representing rift and post-rift sediments east of the
passive margin of Laurentia (Fig. 1). The margin of
Laurentia was partially subducted, and, along with
accretionary wedge sediments, imbricated along
numerous thrust faults (Stanley & Ratcliffe, 1985)
and metamorphosed to high-pressure greenschist or
epidote–amphibolite facies conditions (Laird & Albee,
1981) during the Ordovician Taconian orogeny.
To the east of the Connecticut Valley trough lie the
rocks of the Bronson Hill belt, which include the
Ordovician Oliverian gneisses, Albee Formation,
Ammonoosuc Volcanics and Partridge Formation, the
Silurian Clough Quartzite, the Silurian–Devonian
Fitch Formation, and the Devonian Littleton
Formation (Thompson et al., 1968; Rankin, 1996)
(Fig. 1). The rocks of the Bronson Hill belt are
separated from the rocks of the Connecticut Valley
trough by the Monroe Line (Hatch, 1988).
The nature of the contact between Connecticut
Valley trough and Bronson Hill belt rocks is controversial. Along the length of the belt, some workers
have argued that it is a fault (e.g. Johansson, 1963;
Hatch, 1987; Hatch, 1988; Rankin, 1996; Armstrong
et al., 1997), while others consider it an unconformity
(e.g. Doll et al., 1961; Thompson et al., 1997).
Discussion of the nature, significance and precise
location of the contact is complicated by disagreements
about the stratigraphy on both sides of the contact. In
southern and central Vermont, Thompson et al. (1997)
argued that rocks of the Connecticut Valley trough lie
unconformably on pre-Silurian rocks of the Bronson
Hill belt, and Thompson et al. (1993) argued that the
Devonian Littleton Formation lies unconformably
above the Gile Mountain Formation. Armstrong et al.
(1997), on the other hand, argued that some rocks
mapped as the Littleton Formation in eastern Vermont
belong to the Waits River Formation of the Connecticut
Valley trough, and that the contact between
Connecticut Valley and Bronson Hill rocks is controlled by two thrust faults: the pre-peak metamorphic
Skitchewaug Mountain fault and the post-peak metamorphic Westminster West fault.
In northern Vermont, there is general agreement
that the contact is a thrust fault, the Monroe Fault
(Hatch, 1987; Hatch, 1988; Moench et al., 1995;
Rankin, 1996), but disagreements about stratigraphic
relationships on both sides of the fault make any
estimate of total displacement along the fault difficult.
The origin and age of the rocks immediately east of
the fault are particularly controversial. Billings (1955)
and Rankin (1996) interpreted them to lie stratigraphically beneath the Ordovician Ammonoosuc
Volcanics, and mapped them as the Ordovician Albee
Formation. Moench (1992, 1996), however, argued
that they are correlative with Upper Ordovician to
Lower Devonian rocks found in western Maine, and
were emplaced west of the Bronson Hill belt along a
gravity slide as the Frontenac–Piermont allochthon.
One possible stratigraphic tie has been made across
the Monroe Fault in northern New Hampshire: the
Devonian Ironbound Mountain Formation of the
Connecticut Valley trough has been mapped in depositional contact with the controversial Albee Formation
(D. W. Rankin, pers. comm.) or the Silurian Smalls
Falls Formation of the western Maine sequence
(Moench et al., 1995).
There is evidence for both brittle, probably Mesozoic,
faulting and ductile, probably Acadian, faulting along
and near the Monroe Line in north-eastern Vermont
(Hatch, 1988; Rankin, 1996). Evidence for ductile
faulting includes deformed porphyroblasts, sheared-off
SYN-TECTO NIC PLUTON INT RUSION, VERM ONT
limbs of isoclinal folds, concentrations of quartz veins
(Hatch, 1988), and development of a high-strain fabric
in metamorphosed gabbro to tonalite dykes (Rankin,
1996). Placement of the pre-Devonian Albee Formation
on the Devonian Gile Mountain Formation along a
steeply east-dipping fabric suggests a thrust sense of
movement (Rankin, 1996).
273
north-eastern Vermont are generally similar, and
although they cannot be definitively shown to be the
same age in both places, the dominant S foliation
2
may be related to the movement of nappes in northeastern as well as south-eastern Vermont. Contact
metamorphic mineral growth has been traditionally
interpreted to post-date all deformation (Eric &
Dennis, 1958; Woodland, 1965).
Metamorphism and microstructures
The Acadian metamorphic history of north-eastern
Vermont has some similarities to the better-studied
metamorphic history of east-central and south-eastern
Vermont, but also has some important differences. In
south-eastern and south-central Vermont, peak
Acadian metamorphic P–T conditions were as high as
c. 10.5 kbar and c. 730 °C (Armstrong et al., 1992).
Silurian and Devonian metasediments experienced four
generations of deformation: S , usually bedding1
parallel; S , the dominant planar fabric; S , a millimetre
2
3
to centimetre scale spaced cleavage which intensifies
near shear zones; and S , a possibly Mesozoic cleavage
4
associated with kink bands (Armstrong et al., 1997).
Peak metamorphic mineral assemblages generally postdate S (Menard & Spear, 1994; Armstrong et al.,
2
1997). In some localities, garnet growth post-dates S
3
crenulations as well, but in others, garnet and staurolite
retrogression occurred during development of S
3
mylonites (Armstrong et al., 1997). Metamorphism in
south-eastern and south-central Vermont is generally
considered to be primarily the result of burial beneath
easterly derived nappes (Armstrong et al., 1992;
Menard & Spear, 1994).
In north-eastern Vermont, the highest metamorphic
conditions (sillimanite or sillimanite+K-feldspar
grade) are found in the aureoles of Devonian plutons.
The presence of andalusite and pseudomorphs of
andalusite in pelites and the composition of amphibole
in mafic rocks suggest that the rocks were metamorphosed under much lower pressure conditions than
were the rocks of south-eastern Vermont (Laird et al.,
1984). Outside the aureoles of the plutons, metamorphism reached biotite or chlorite grade (Doll et al.,
1961). Three phases of deformation have been identified
in both the Connecticut Valley trough and Bronson
Hill belts in north-eastern Vermont (Rankin, 1996),
although only the later two are easily distinguishable
microscopically (Doll, 1951; Dennis, 1956; Eric &
Dennis, 1958; Johansson, 1963; Woodland, 1965). The
earlier (S ) foliation, defined by the orientation of fine2
grained muscovite and, in a few places, biotite, is axial
planar to rarely observed isoclinal folds of bedding
(Woodland, 1965). In most places (on the limbs of
isoclinal folds), the early foliation and bedding are
subparallel and generally NE-striking and steeply
E-dipping (Woodland, 1965). The later foliation (S )
3
is a zonal crenulation cleavage that dips to the east,
less steeply than the early foliation (Woodland, 1965).
Descriptions of S and S from south-eastern and
2
3
Plutonism
Devonian plutons which intruded during or slightly
after the Acadian orogeny are found throughout
northern New England (Fig. 1). Billings (1956) named
the Devonian plutons the New Hampshire plutonic
series, and the name has been extended to refer to
similar plutons in the rest of New England (e.g. Page,
1968; Dallmeyer et al., 1982). The New Hampshire
plutonic series includes plutons that clearly pre-date
or are synchronous with Acadian deformation and
plutons that appear to post-date all deformation. In
New Hampshire, the Bethlehem Gneiss (c. 405 Ma,
U–Pb zircon, T.M. Harrison cited in Spear et al., 1990)
and the Kinsman Granite (403±3 Ma, U–Pb zircon,
Robinson & Tucker, 1996) are foliated and deformed
in nappes, whereas the younger Concord Granite is
considered post-tectonic (Billings, 1956). In central
and northern Maine, 409–404 Myr old plutons appear
to cut regional folds (Bradley et al., 1996), although
some are foliated and interpreted as late syn-kinematic
(Solar et al., 1998). The plutons of north-eastern
Vermont have previously been considered postkinematic (e.g. Naylor, 1971).
The New Hampshire series plutons of north-eastern
Vermont are also known as the Northeast Kingdom
batholith (Ayuso & Arth, 1992; Arth & Ayuso, 1997)
(Fig. 1). The plutons range in composition from quartz
diorite to muscovite-bearing granite and originated
from a variety of mantle- and crustal-derived melts
that underwent varying amounts of fractional crystallization (Arth & Ayuso, 1997). Rb–Sr whole-rock
isochron ages for plutons of the Northeast Kingdom
batholith range from 370±17 Myr old (Derby Pluton)
to 390±14 Myr old (Nulhegan Pluton) (Arth & Ayuso,
1997). Although the plutons appear mostly undeformed
and cut both foliations (Doll, 1951; Dennis, 1956;
Woodland, 1965), a few granitic dykes and sills are
boudinaged (Doll, 1951; Dennis, 1956) or folded (Doll,
1951), suggesting that the plutons did not post-date
all Acadian deformation. Xenoliths of country rock
are abundant near the contacts of the plutons (Ayuso
& Arth, 1992), and a zone of mixed granite and
country rock containing numerous xenoliths and
granitic dykes (the ‘granite–hornfels complex’ of
Woodland, 1965) up to two miles wide can be found
along the western contact and to the south-east of the
Victory Pluton (Fig. 2).
The Victory Pluton itself is very poorly exposed
274
K. A . HA NNULA ET A L .
Fig. 2. Simplified geological map of the Victory Pluton and its aureole showing stereo-net data (modified from Eric & Dennis,
1958, Johansson, 1963 and Woodland, 1965). S , dominant foliation; L , mineral elongation lineation. Inset: extent of cover in the
e
vicinity of the Victory Pluton. The grey patternDindicates areas of Quaternary
sedimentary cover (after Woodland, 1965).
(Fig. 2). Although the ‘granite–hornfels complex’
underlies many peaks in the area, including Kirby
Mountain, Burke Mountain and Umpire Mountain,
most of the pluton proper is found beneath the Victory
Bog. The mapped contact (Johansson, 1963; Woodland,
1965) between the pluton and its country rock follows
the break in slope at the base of hills surrounding the
bog in many places. The eastern lobe of the pluton,
which is mapped as cutting across the Monroe Fault,
is particularly poorly exposed. Although the presence
of sillimanite-grade country rock exposed above the
break in slope suggests that the pluton or the ‘granite–
hornfels complex’ is nearby, the exact shapes of the
Victory Pluton and its surrounding ‘granite–hornfels
complex’ are poorly known.
O B S E R VAT IO N S
Map relations and previous work suggest that the
Victory Pluton post-dates all deformation in northeastern Vermont. However, studies of microstructures
along the Monroe Fault and within the aureole of the
Victory Pluton, and of hand sample-scale structures in
the highest-grade rocks adjacent to the Victory Pluton,
imply that magmatism and deformation were coeval.
Structures
On both sides of the Monroe Fault, three foliations
are developed. The first foliation, S , is defined by
1
alignment of fine-grained white mica and in places,
chlorite. S has been isoclinally folded and is usually
1
only preserved in microlithons between S cleavage
2
domains (Fig. 3A). The spacing between S cleavage
2
domains is generally <1 mm; S is generally only
1
recognizable under the microscope, and in some places
S may be obliterated entirely. In micaceous layers, S
1
2
is wrinkled by a variably developed crenulation
cleavage, S (Fig. 3A). In general, S strikes north-east
3
2
and dips vertically to steeply south-east, while S
3
strikes north-east and dips moderately south-east
(Fig. 2). Along the Monroe Fault, only one foliation
can be recognized. It is not clear whether it is
equivalent to S or whether it is a younger foliation
3
developed only along the fault.
Mafic dykes are present in the ‘Albee Formation’
SYN-TECTO NIC PLUTON INT RUSION, VERM ONT
(a)
(b)
275
coarse hornblende is broken and bent, and fine-grained
recrystallized hornblende is parallel to the fault-related
foliation (Fig. 3C).
Mineral lineations and pressure shadows are rare in
the study area. However, along the Monroe Fault,
lineations, defined by pressure shadows around garnet
porphyroblasts and by alignment of fine-grained
hornblende, plunge steeply down-dip (Fig. 2). Within
the sillimanite+K-feldspar zone east of the fault,
lineations are defined by pressure shadows around
garnet porphyroblasts and by aligned fibrolitic sillimanite. Within the staurolite zone west of the Monroe
Fault, biotite porphyroblasts are aligned to form a
weak subhorizontal mineral lineation. Many of these
biotite porphyroblasts have quartz-filled pressure
shadows visible in sections cut parallel to the mineral
lineation and perpendicular to the foliation.
Metamorphism
Metamorphic grade generally increases toward the
Victory Pluton, as would be expected in a contact
aureole. Mineral assemblages and preliminary thermobarometry can be used to determine the approximate depth
at which the Victory Pluton intruded. They suggest
somewhat different conditions on either side of the
Monroe Fault. An appendix containing tables of mineral
compositions on both sides of the fault is available from
the Journal of Metamorphic Geology www sites (URLs
are given on the journal cover).
West of Monroe Fault
(c)
Fig. 3. Photomicrographs. (a) Sample of Gile Mountain
Formation from the hinge zone of an F2 fold, illustrating three
foliations (S –S ). The thin section is cut perpendicular to the
3
intersection 1between
S and S . Lines indicate the orientation
2
of each foliation. The field
of 3view is 1.8 mm wide. (b)
Undeformed amphibolite dyke from the Albee Formation. The
scale bar is 1 mm. (c) Deformed dyke in the Albee Formation.
This sample shows one large hornblende crystal (sigma-shaped
clast in the centre of the photomicrograph) that has been bent
and truncated by the fault-related foliation, defined by finegrained aligned hornblende. The scale bar is 1 mm.
just east of the Monroe Fault. At distances greater
than 20 m from the fault these dykes contain mediumto coarse-grained, randomly oriented hornblende
(Fig. 3B). Along the fault, however, the grain size of
the hornblende is drastically reduced (Onasch, 1997),
The metamorphic grade west of the Monroe Fault
ranges from biotite grade far from the pluton to
sillimanite grade near the pluton contact (Fig. 2). In
the garnet zone west of the Monroe Fault, the typical
mineral assemblage is biotite+chlorite+quartz+
graphite±garnet±muscovite±plagioclase±ilmenite±
apatite. Thermobarometry using the internally consistent dataset of Berman (1988) as formulated by Gordon
(1992) and Gordon et al. (1994) yields T =499 °C, P=
4.7 kbar for a sample within the Gile Mountain
Formation (Fig. 4, P–T conditions B), and T =507 °C,
P=4.8 kbar for a sample of Meetinghouse Slate
adjacent to the Monroe Fault (Fig. 4, P–T conditions
A).
In the staurolite zone west of the Monroe Fault, the
typical mineral assemblages are staurolite+garnet+
biotite+muscovite+quartz±plagioclase,
garnet+
biotite + muscovite + chlorite + quartz ± plagioclase,
and staurolite+biotite+muscovite+chlorite+quartz±
plagioclase. These mineral assemblages are generally
consistent with pressures similar to those calculated
for the garnet zone. Staurolite is commonly replaced
by decussate muscovite+chlorite, and coarse chlorite
porphyroblasts in the matrix crosscut S and S
2
3
and appear to have grown during retrograde
metamorphism.
276
K. A . HA NNULA ET A L .
(a)
Fig. 4. Petrogenetic grid for high-temperature pelites after
Spear et al. (in press), showing estimates of metamorphic
conditions at various locations. Ovals indicate the uncertainty
for P–T estimates based on thermobarometry. (a) Garnet zone,
Meetinghouse Slate in Monroe Fault zone. (b) Garnet zone,
Gile Mountain Formation. (c) Sillimanite zone, Gile Mountain
Formation. (d) Garnet zone, Albee Formation. (e)
Sillimanite+K-feldspar zone, Miles Mountain migmatite
complex. Numbered reactions: (2) St+Ms=
Grt+Bt+As+H O; (3) Ms+Bt+H O=St+Grt+As+L; (4)
2
2
Ms+Ab=As+Kfs+H
O; (5) St+Ms=Grt+Bt+As+L;
(6)
2 O.
Bt+Grt=Opx+Crd+H
2
In the sillimanite zone west of the Monroe Fault,
the typical mineral assemblage is fibrolitic sillimanite+
biotite+muscovite+quartz±plagioclase±staurolite±
ilmenite, with several samples containing pseudomorphs of andalusite. In some samples from near the
sillimanite-in isograd, the andalusite pseudomorphs
are composed of decussate muscovite and tiny, subhedral staurolite grains (Fig. 5A), suggesting the retrograde
reaction
And+Bt+Qtz+H O=St+Chl+Ms
2
This reaction has a moderate slope in the andalusite
and sillimanite stability fields (Fig. 4), and could be
the result of either an increase in pressure or a decrease
in temperature. In other samples, andalusite pseudomorphs are composed of coarse sillimanite. Although
the sillimanite-zone mineral assemblage is unsuited for
thermobarometry, the assemblage is consistent with
temperatures between c. 575 °C and c. 650 °C and
pressures between c. 3.5 and 7 kbar based on the
petrogenetic grid of Spear et al. (in press) (Fig. 4, P–T
conditions C).
East of the Monroe Fault
Metamorphic grade east of the Monroe Fault ranges
from chlorite grade near the New Hampshire border
to sillimanite+K-feldspar grade at Miles Mountain,
south-east of the Victory Pluton (Fig. 2). In the garnet
zone, mineral assemblages are similar to those east of
(b)
Fig. 5. (a) Photomicrograph of muscovite+staurolite
pseudomorph of andalusite. The small, high-relief, subhedral
grains surrounded by muscovite are staurolite. The scale bar is
1 mm. (b) Leucosome fills strain shadows around garnet
porphyroblasts in the Miles Mountain migmatite complex. The
coin is 1.8 cm in diameter.
the fault, although graphite is not present and mineral
proportions are different. Unlike garnet west of the
fault, however, garnet east of the fault is at least
partially retrograded to chlorite in nearly all samples.
One unretrograded sample yields T =514 °C, P=
5.3 kbar (Fig. 4, P–T conditions D) using the internally
consistent database of Berman (1988) as formulated
by Gordon (1992) and Gordon et al. (1994). In a
narrow sillimanite+muscovite zone exposed in stream
gullies along the north-eastern boundary of the pluton,
samples contain sillimanite (fibrolite)+biotite+
muscovite+quartz+plagioclase±garnet±staurolite.
To the south-east of the pluton is a >5 km wide
migmatite complex (Fig. 2), in which about half the
outcrop consists of granite and half the outcrop
consists
of
migmatite
containing
garnet+
sillimanite+biotite+K-feldspar+plagioclase+quartz.
The contacts of the migmatite complex with lowergrade rocks are covered, and the exact location of the
SYN-TECTO NIC PLUTON INT RUSION, VERM ONT
staurolite, sillimanite and sillimanite+K-feldspar isograds surrounding the migmatite complex are uncertain.
Both fibrolitic and coarse varieties of sillimanite are
present in the migmatite. In many samples, retrograde
muscovite replaces sillimanite and chlorite replaces
biotite. Textures in one sample suggest that
melt+garnet were produced (Fig. 5B) by a reaction
such as
Bt+Sil+Qtz=Grt+Kfs+melt
( Vielzeuf & Holloway, 1988)
or
Bt+Sil+Qtz+H O=Grt+melt
2
( Vielzeuf & Holloway, 1988)
Peak temperatures are clearly higher south-east of
the pluton than anywhere else in the immediate
area. The presence of sillimanite+K-feldspar and
garnet+sillimanite+biotite assemblages limits the
temperature to at least 650 °C (Fig. 4). Internally
consistent thermobarometry (using the database of
Berman (1988) and the formulation of Gordon (1992)
and Gordon et al. (1994)) on the assemblage
+ biotite + sillimanite + plagioclase + quartz,
using core compositions of 1 cm diameter garnet
crystals and core composition of matrix plagioclase,
yielded T =680 °C, P=5.8 kbar. These conditions are
broadly consistent with those predicted by the mineral
assemblages present in the migmatite complex (Fig. 4,
P–T conditions E), although they may not reflect peak
metamorphic conditions due to re-equilibration
during cooling.
Relative timing of metamorphism and deformation
The relative timing of porphyroblast growth and
deformation varies spatially throughout the study area.
Garnet zone
Along the Monroe Fault, garnet appears to have
grown before or during development of the dominant
foliation. We have found garnet porphyroblasts along
the Monroe Fault in three places. South of the Victory
Pluton (Fig. 6, location 7A), the Meetinghouse Slate
member of the Gile Mountain Formation along the
fault displays a single intense foliation. Garnet porphyroblasts contain slightly curved inclusion trails which
are truncated at a high angle to the dominant foliation,
indicating that garnet growth preceded development
of the dominant foliation (Fig. 7A). In Buzzell Gap,
north of the Victory Pluton (Fig. 6, location BG),
garnet porphyroblasts near the fault zone have an
unusually flattened shape parallel to the dominant
foliation. The porphyroblasts contain inclusion trails
parallel to the dominant foliation, but the dominant
foliation tightens significantly around the garnet porphyroblasts, suggesting that they grew during develop-
277
ment of the foliation, but were not rotated by simple
shear during their growth. Along Granby Stream, in
the aureole of the Maidstone Pluton (Fig. 6, location
7B), garnet crystals contain curved inclusion trails, the
outer ends of which parallel the matrix foliation,
suggesting that garnet grew during deformation
(Fig. 7B).
Garnet crystals from samples more than 10 m from
the Monroe Fault reveal one of two possible growth
patterns. Two samples contain garnet which completely
overgrows both S and S (Fig. 6, location 7C, Fig. 7C).
2
3
The majority of the garnet porphyroblasts in the
garnet zone appear to have grown after the development of S , but before the development of S . This
2
3
growth history is best expressed by garnet crystals that
contain straight inclusion trails while the external
foliation is crenulated around them (Fig. 7D). Some
garnet has been retrograded to chlorite. S is generally
3
deflected around the pseudomorphs without deforming
them (Fig. 7D), indicating that the pseudomorphing
reaction occurred before S formed.
3
Staurolite zone
There is little pattern to the timing of staurolite growth
in samples cut perpendicular to the intersection
lineation L . In eight samples north-west of the
2×3
pluton, staurolite porphyroblasts clearly overgrow all
foliations. In many samples, S crenulations tighten
3
around staurolite porphyroblasts. In some samples,
staurolite contains an internal foliation that is discontinuous with S , the dominant foliation in the matrix
2
(Fig. 8A).
Most garnet crystals in the staurolite zone do not
contain inclusions, so it is more difficult to determine
their timing compared to deformation. S crenulations
3
are deflected about the majority of garnet porphyroblasts, and S is deflected about garnet in one sample.
2
In many samples, none of the foliations are deflected
about garnet, suggesting that the garnet grew after all
deformation ceased.
Many staurolite porphyroblasts have been replaced
by pseudomorphs of muscovite and chlorite. These
pseudomorphs display the same range of timing
relations as do the staurolite porphyroblasts. In some
samples, both the pseudomorphs and the original
staurolite appear to post-date all deformation. In other
samples, S or S is deflected around decussate,
2
3
perfectly staurolite-shaped aggregates of muscovite and
chlorite, suggesting that some deformation occurred
after staurolite growth, but before the pseudomorphing
reaction occurred.
Biotite textures
Biotite also forms porphyroblasts in the garnet and
staurolite zones. Along the Monroe Fault, biotite is
found parallel to the dominant foliation. Away from
the fault, however, biotite forms porphyroblasts that
278
K. A . HA NNULA ET A L .
Fig. 6. Simplified geological map of the
study area showing the timing of mineral
growth with respect to deformation. Circles
show the location of garnet samples,
triangles show the location of staurolite
samples, and squares show the location of
sillimanite samples. Black symbols indicate
post-tectonic porphyroblasts, grey symbols
indicate that S crenulations post-date
3 and open symbols indicate
porphyroblasts,
that S is deflected around porphyroblasts.
2
Stars indicate
syn-tectonic garnet, which is
only found along the Monroe Fault. Black
squares indicate unoriented fibrolite, grey
squares indicate lineated fibrolite, and open
squares indicate lineated coarse sillimanite.
Numbers indicate the location of samples
photographed in Figs 7 & 8. BG, Buzzell
Gap. Eric & Dennis (1958), Johansson
(1963), and Woodland (1965).
do not lie parallel to the dominant foliation, although
in many samples they have a preferred orientation
aligned subparallel to L
and define a mineral
2×3
lineation on S surfaces. Biotite porphyroblasts reveal
2
a variety of textures similar to staurolite. In some
samples, biotite porphyroblasts overgrow all foliations.
In other samples, biotite porphyroblasts overgrow a
straight foliation, while the external foliation is crenulated by S (Fig. 8B). In still other samples, the biotite
3
porphyroblasts contain straight inclusion trails at an
angle to the dominant foliation. These samples are
typically those containing a gentle and widely spaced
crenulation cleavage. It is possible that all the biotite
porphyroblasts post-date development of S and that
2
some pre-date S , while others post-date both S
3
2
and S .
3
Microstructures parallel biotite lineation
In staurolite-zone samples away from the Monroe
Fault which are cut parallel to the biotite lineation
and perpendicular to S , microstructures show more
2
consistent evidence for deformation during or after
mineral growth than in sections cut perpendicular to
the intersection lineation. Only one foliation is typically
visible, because the biotite lineation is subparallel to
the L
intersection lineation. The foliation is
2×3
deflected about garnet porphyroblasts, many of which
contain inclusion trails parallel to the matrix foliation
(Fig. 8C). The foliation also wraps around staurolite
pseudomorphs, and the muscovite and chlorite which
replace staurolite are aligned with the foliation. Biotite
porphyroblasts are elongate parallel to the foliation
and have quartz pressure shadows around them
(Fig. 8C). Their cleavage is generally subparallel to the
section (parallel to the biotite lineation and perpendicu-
lar to S , or NE-striking and gently NW-dipping).
2
The pressure shadows around biotite are generally
symmetric except where two porphyroblasts impinge
on one another. In sections of the same sample cut
perpendicular to L , garnet appears to post-date the
2×3
S foliation but pre-date the S crenulation cleavage,
2
3
and biotite is generally oriented with cleavage planes
at a high angle to S (Fig. 8D), suggesting that the
2
deformation observed in sections parallel to the
lineation occurred during D .
3
Sillimanite and sillimanite+K-feldspar zones
Both coarse and fibrolitic sillimanite are found on
both sides of the Monroe Fault. Rocks in the sillimanite
zone do not typically preserve evidence of the three
foliations observed at lower grades. Some samples are
compositionally layered with little or no mineral
alignment. Other samples, particularly those within
the migmatite complex east of the Monroe Fault, have
a single well-developed foliation. It is unclear whether
this foliation is S or S . Some of the well-foliated
2
3
samples are lineated as well, with lineations defined by
pressure shadows on garnet or by aligned sillimanite
grains.
West of the fault, coarse sillimanite is found in
randomly oriented pseudomorphs after andalusite.
Fibrolitic sillimanite forms epitaxial growths on biotite.
In some samples, the fibrolite is randomly oriented,
while in others it is weakly aligned with the dominant
foliation (Fig. 6).
East of the fault, sillimanite is found in the Miles
Mountain migmatite complex co-existing with
K-feldspar as well as in a narrow sillimanite+
muscovite zone bordering the eastern edge of the
Victory Pluton (Fig. 6). Within the sillimanite+
SYN-TECTO NIC PLUTON INT RUSION, VERM ONT
279
(a)
(b)
(c)
(d)
Fig. 7. Photomicrographs to illustrate garnet microstructures. Sample locations are shown in Fig. 6. (a) Syn-tectonic garnet along
the Monroe Fault south of the Victory Pluton. The scale bar is 1 mm. (b) Syn-tectonic garnet from the aureole of the Maidstone
Pluton. The scale bar is 1 mm. (c) Post-tectonic garnet from the Gile Mountain Formation. The scale bar is 1 mm. (d) Post-S ,
pre-S garnet from the Albee Formation. The garnet has been partly replaced by post-tectonic chlorite, and contains straight 2
3
inclusion
trails. In the matrix, the dominant foliation (S ) has been crenulated (S ). The scale bar is 0.5 mm.
2
3
K-feldspar zone, sillimanite further than 3 km from the
Monroe Fault is randomly oriented (Fig. 6). Closer
than 3 km to the Monroe Fault, fibrolitic sillimanite is
generally aligned in the dominant foliation and
deflected around garnet porphyroblasts (Fig. 8E), and
is frequently lineated. Fibrolitic sillimanite also wraps
around coarse sillimanite porphyroblasts, which are
frequently oriented at high angles to the fibrolite
lineation.
Structures within the ‘granite–hornfels complex’ and
migmatite complex
West of the fault, the country rock in the ‘granite–
hornfels complex’ is typically fine-grained except for
coarse sillimanite pseudomorphs after andalusite, and
the dominant foliation is defined by lithological
layering, with some weak alignment of biotite, muscov-
ite and sillimanite. Leucosome-filled tension gashes are
found within the country rock on Kirby Mountain
(Fig. 2), perpendicular to a weak mineral lineation
defined by andalusite pseudomorphs; the leucosome is
interpreted as indicating the former presence of melt.
Sills of granite intrude parallel to the foliation (Fig. 9A).
Some of these sills are isoclinally folded (Fig. 9B), and
some have a weak foliation defined by plagioclase
phenocrysts. Within the plutonic parts of the ‘granite–
hornfels complex’, there are numerous xenoliths.
Within the migmatite complex, the country rock is
partially migmatized, well-foliated, and contains
sillimanite+garnet+biotite+K-feldspar assemblages.
Some leucogranite dykes and sills are parallel to the
foliation, and are folded along with the country rock
foliation, whereas others crosscut all metamorphic
structures (Fig. 9C, D). Other granite sills are boudinaged (Fig. 9E). In other places, leucosome fills shear
280
K. A . HA NNULA ET A L .
(a)
(b)
(d)
(c)
(e)
Fig. 8. Photomicrographs to illustrate the relationship of foliations and mineral growth. (a) Pre-S staurolite. The scale bar is
2
0.5 mm. (b) Post-S , pre-S biotite. The scale bar is 0.5 mm. (c) Section cut parallel to mineral elongation
lineation and
perpendicular to S2. Note 3the strain shadows around biotite porphyroblasts. The curved line indicates deflection of S around
3
garnet. The garnet 2in the centre of the photograph is 0.5 mm in diameter. (d) Section cut perpendicular to mineral elongation
lineation and S . The large garnet porphyroblast is 0.5 mm in diameter. (e) Fibrolite, aligned in the foliation, wraps around a garnet
porphyroblast. 2The scale bar is 5 mm.
bands, garnet pressure shadows and boudin necks
(Fig. 5B); again the leucosome is interpreted as indicating the former presence of melt. Near the contact with
the pluton, the foliated country rock is brecciated and
intruded by a small volume of granite (Fig. 9F). Most
granite bodies within the migmatite complex have a
weak magmatic foliation; where a contact with the
country rock is exposed, the foliation is parallel to
the contact.
D I S C U SS I O N
Regional correlation of structures
It can be difficult to correlate foliations, outcrop-scale
structures and microstructures, particularly in poorly
exposed areas in which foliation orientations only
differ by 10–30°, as S and S do near the Victory
2
3
Pluton. Correlating the ages of microstructures in
SYN-TECTO NIC PLUTON INT RUSION, VERM ONT
281
(a)
(b)
(c)
(d)
(e)
(f)
Fig. 9. Photographs to show relationships in the field. (a) Western border zone of the Victory Pluton on Kirby Mountain. Granite
preferentially intruded parallel to the foliation in the country rock. In the upper part of the photograph, leaves and roots partially
cover country rock. The hammer head is 12 cm long. (b) Crosscutting leucogranite dykes are ptygmatically folded in the ‘granite–
hornfels complex’ in the Kirby Mountain area, west of the Monroe Fault. Axial planes of the folded dykes are parallel to the
dominant foliation. The hammer handle is 4 cm wide. (c) Tightly folded granitic segregation, Miles Mountain. The dominant
foliation is axial planar to the folds. Note that garnet porphyroblasts are flattened and aligned in the foliation. The pencil is
approximately 8 mm in diameter. (d) Crosscutting granite dyke, Miles Mountain. The dominant foliation is defined by biotite and
sillimanite, and is folded by the chevron folds visible in photograph. The hammer is approximately 30 cm long. (e) Boudinaged
leucogranite sill, Miles Mountain. The pencil is approximately 14 cm long. (f ) Intrusion breccia, Miles Mountain. Both lightcoloured foliated parts of the outcrop and angular dark fragments are xenoliths. The hammer is approximately 20 cm long.
rocks that have experienced different temperatures or
different amounts of strain can be particularly difficult,
because high-temperature recrystallization or highstrain deformation can obliterate evidence of earlier
deformations. Furthermore, competent rock types such
as granite, amphibolite, quartzose metasediments and
coarsely crystalline metamorphic rocks tend to be
resistant under low-strain deformation, and are unlikely
282
K. A . HA NNULA ET A L .
to develop the style of crenulations found in finegrained phyllites and schists. However, in a contact
aureole, peak metamorphic mineral growth may
provide a time marker to determine which structures
pre-date, post-date, or are nearly synchronous with
pluton intrusion.
At staurolite and garnet grades west of the fault,
deformation during pluton intrusion appears restricted
to development of the weak S crenulation cleavage
3
and associated subhorizontal biotite lineation. S pre1
dates peak metamorphism. In most samples, S clearly
2
pre-dates garnet, staurolite and biotite porphyroblast
growth; in the remaining ambiguous samples, it is
possible that S was crenulated broadly after porphyro2
blast growth, causing a discordance between the matrix
foliation and the internal foliation (e.g. Fig. 8A) but
not developing a clear S cleavage. S pre-dates
3
3
porphyroblast growth in some samples and post-dates
porphyroblast growth in others, in some cases over a
small area. This type of variation in textures could be
the result of multiple phases of mineral growth, local
variation in the timing of development of the crenulation cleavage, or a combination of both factors, and
we consider S to be regionally approximately the
3
same age as peak metamorphism. The weak subhorizontal mineral elongation lineation defined by biotite
with quartz pressure shadows developed at the same
time as S . Retrograde minerals, such as chlorite
3
porphyroblasts, chlorite replacing garnet, and
chlorite+muscovite replacing staurolite, post-date S
3
in most samples.
In the sillimanite zone west of the Monroe Fault,
there is a single dominant foliation present defined
primarily by compositional layering. Coarse sillimanite
generally lies parallel to compositional layering, but is
not lineated, and fibrolite is randomly oriented in most
samples and only weakly oriented in others. Xenoliths
in the ‘granite–hornfels complex’ west of the fault have
a similar foliation defined by compositional layering.
The dominant foliation in sillimanite-zone rocks west
of the pluton pre-dates peak metamorphism and pluton
intrusion, and is probably equivalent to S in lower2
grade rocks. The weak alignment of fibrolitic sillimanite
correlates with S in lower-grade rocks.
3
In the garnet zone east of the Monroe Fault, some
biotite porphyroblasts pre-date S or D folding of S ,
2
3
2
while other biotite clearly post-dates S and pre-dates
2
S . Garnet growth occurred after S , but before S .
3
2
3
Retrogression of garnet to chlorite occurred after S
3
formed. Because it is possible that some biotite
porphyroblasts grew during regional low-grade metamorphism that pre-dated the contact metamorphism,
we interpret the timing of deformation and metamorphism to be similar to that in rocks west of the fault:
S developed before contact metamorphism and pluton
2
intrusion, and S developed in the short interval
3
between peak metamorphism and cooling.
Within the migmatite complex east of the Monroe
Fault, a single foliation is found. It has been folded on
the outcrop scale (Fig. 9D) and probably at larger
scales based on variation in its orientation (Fig. 2).
The alignment of biotite and sillimanite in the foliation
and the deflection of sillimanite around garnet porphyroblasts suggest that the deformation recorded by the
dominant foliation in the migmatite occurred at high
temperatures. Leucosome-filled strain shadows around
garnet porphyroblasts (Fig. 5B) suggest that melt was
present and able to migrate into low-strain areas
during deformation. The presence of leucosome along
axial planes and hinge zones of folds, within shear
bands, in boudin necks and in numerous sills parallel
to the country rock foliation is also consistent with
the presence and migration of magma during deformation (Amato et al., 1994; Collins & Sawyer, 1996;
Davidson et al., 1996; Brown & Rushmer, 1997; Brown
& Solar, 1998. Isoclinally folded and boudinaged sills
(Fig. 9B. D. E) suggest that deformation continued
after thin magma bodies crystallized (Rothstein et al.,
1994). However, the presence of undeformed dykes
(Fig. 9D) and the absence of sub-solidus fabrics in the
main body of the Victory Pluton suggest that deformation ended before the magma in the main body
crystallized below the rheological critical melt percentage (Arzi, 1978). Deformation in the migmatite zone,
including foliation development and folding of that
foliation, occurred during peak metamorphism and
partial melting.
Along the Monroe Fault itself, two garnet-bearing
samples have textures indicating deformation during
or after garnet-grade metamorphism. These samples
appear to be within the contact metamorphic garnet
zone of the Victory and Maidstone Plutons. Both of
them lie within 5 km of a pluton contact, a shorter
distance from the pluton than some samples in which
garnet overgrew both S and S . If the garnet along
2
3
the Monroe Fault grew during contact metamorphism,
some movement along the fault must have taken place
during pluton intrusion. This is the primary evidence
for pluton intrusion during movement along the
Monroe Fault.
In summary, the regional dominant foliation, S ,
2
developed before pluton intrusion. Pluton intrusion
was accompanied by a variety of different deformational styles: along the Monroe Fault, localized
high-strain deformation occurred; in the Miles
Mountain migmatite complex, penetrative ductile
deformation forming and deforming a single dominant
foliation accompanied high-temperature metamorphism; and in other areas throughout the aureole on
both sides of the fault, a weak crenulation cleavage
(S ) striking north-east and dipping moderately south3
east developed.
Strain partitioning and variation in deformational style
Both deformational style and the amount and orientation of strain during pluton intrusion vary from the
Monroe Fault to the remainder of the Victory Pluton’s
SYN-TECTO NIC PLUTON INT RUSION, VERM ONT
aureole. Along the fault, high strain occurred with a
down-dip finite stretching direction, whereas further
from the fault strain was lower and the finite stretching
direction was subhorizontal. Evidence for relatively
high strain along the fault includes grain-size reduction,
the pervasive evidence of deformation during or after
porphyroblast growth, and the single strong foliation
that reorients minerals even in competent rock types
such as the amphibolite on the east side of the fault.
The down-dip mineral lineation along the fault suggests
that the long axis of the finite strain ellipsoid is steeply
dipping. Away from the fault, lower strain during
pluton intrusion is suggested by the preservation of
early S and S fabrics, and by the widely spaced,
1
2
weakly developed nature of the S crenulation cleavage.
3
A weak subhorizontal mineral lineation and subhorizontal strain shadows around biotite and garnet
porphyroblasts provide evidence for a weak subhorizontal finite stretching direction during pluton
intrusion.
Intrusion-related structures suggest that both ductile
and brittle deformation occurred during pluton
intrusion. In the migmatite zone east of the fault,
ductile deformation is implied by folded granitic dykes,
foliated high-grade minerals, the deflection of the
dominant foliation around garnet and sillimanite
porphyroblasts, and the presence of leucosome-filled
shear bands and garnet pressure shadows. Brittle
deformation, however, is indicated by a diatremelike breccia (Fig. 9F). West of the pluton, ductile
deformation is suggested by isoclinally folded sills in
sillimanite-grade rocks, whereas leucosome-filled tension gashes at both garnet and sillimanite grade
suggest brittle deformation. Based on metamorphic
pressures in its aureole (4.1–5.8 kbar), the Victory
Pluton intruded the mid-crust, at a depth between 14
and 20 km. The shallower depths correspond to the
approximate location of the brittle–ductile transition
for a typical continental geothermal gradient of
20–25 °C km−1. Intrusion of the Victory Pluton heated
its country rocks well above typical brittle–ductile
transition temperatures, thermally weakening them at
low strain rates and facilitating ductile deformation.
Magma intrusion could also locally decrease effective
stresses, however, causing higher strain rate brittle
deformation.
Amount of syn-intrusion movement on the Monroe Fault
Microstructures are useful for determining the relative
age of metamorphism and deformation, but cannot be
used to determine total displacement along a fault.
Because regional stratigraphic relations are controversial (Hatch, 1987; Moench, 1992; Moench et al.,
1995; Moench, 1996; Rankin, 1996; Armstrong et al.,
1997; Thompson et al., 1997), it is difficult to determine
the regional significance of the Monroe Fault. For
instance, if rocks correlative to the Gile Mountain
Formation lie unconformably above the sequence of
283
rocks immediately east of the Monroe Fault (Moench
et al., 1995; D.W. Rankin, personal communication),
the total displacement on the Monroe Fault in northeastern Vermont might be relatively small. However, if
the rocks east of the fault are correlative with rocks
from western Maine (Moench, 1992, 1996) and if Gile
Mountain-equivalent rocks do not lie unconformably
above them, total displacement could be regionally
very significant. Furthermore, it is likely that movement
along the Monroe Fault during intrusion of the Victory
Pluton only accounts for a fraction of the total
movement along the fault.
Metamorphic P–T estimates across the Monroe
Fault provide one way to measure displacement
following peak metamorphism. Peak pressure estimates
are slightly higher on the east side of the fault than
they are on the west side. West of the Monroe Fault,
peak metamorphic pressures are 4.7–4.8 kbar (Fig. 4,
points A & B), equivalent to approximately 16 km
depth at crustal densities of 3.0 g cm−3. East of the
Monroe Fault, peak metamorphic pressures are
5.3–5.8 kbar (Fig. 4, point D), equivalent to depths of
18–20 km. Errors for pressure estimates are at least
±1 kbar however, and post-peak metamorphic
offset along the fault of 0–4 km could be consistent
with the calculated pressures. The maximum peak
metamorphic temperature conditions are also different
across the fault: west of the fault, the highest grade
rocks contain sillimanite+muscovite+biotite assemblages (c. 575–650 °C, Fig. 4, point C), whereas east
of the fault, the highest grade rocks contain
sillimanite+K-feldspar (>650 °C, Fig. 4, point E). This
suggests that at least some movement along the fault
occurred after peak metamorphism.
Reaction textures west of the fault provide evidence
of fault movement during metamorphism. Although
andalusite is not found in the study area, two types of
andalusite pseudomorphs are present west of the
Monroe Fault. Pseudomorphs of sillimanite after
andalusite are consistent with either increasing
temperature or increasing pressure during metamorphism. The second type of pseudomorph,
muscovite+staurolite after andalusite (Fig. 5A), suggests that the retrograde reaction
And+Bt+Qtz+H O=St+Chl+Ms
2
has taken place. This reaction has a moderate slope in
the andalusite stability field (Fig. 4, Spear et al., in
press), and could occur during either increasing
pressure or decreasing temperature. Peak metamorphic
conditions (4.1–4.7 kbar) are slightly higher than the
maximum pressure at which andalusite+biotite
(+quartz+muscovite+H O) are stable (c. 3.8 kbar,
2
below the intersection of reaction (1) and andalusite=
sillimanite) (Fig. 4). The presence of andalusite pseudomorphs and both pseudomorphing reactions could be
explained by an increase in pressure following heating
to andalusite-grade conditions. This P–T history would
284
K. A . HA NNULA ET A L .
be expected for rocks buried in the lower plate of a
thrust fault during intrusion of a granitic magma.
Metamorphic conditions across the fault and metamorphic textures west of the fault suggest that
movement of <1–4 km along the steeply dipping fault
could have occurred during the intrusion of the Victory
Pluton. Map relations suggest that the largest amounts
of displacement are unlikely to have occurred, because
neither the pluton nor the isograds surrounding it
appear offset by the fault. This argument is fairly weak,
however, because the outcrop along the fault is
extremely poor (Fig. 2). The mapped eastern lobe of
the pluton is entirely unexposed, and the covered
region could be underlain by a migmatite complex
similar to that exposed on Miles Mountain rather
than by the pluton itself. The broad northward curve
in the garnet and staurolite isograds south of the
migmatite complex (Fig. 2) is similarly poorly exposed,
and it is possible that those isograds are in fact offset.
The lack of intracrystalline deformation in the
pluton provides more constraint on the possible
amount of displacement during pluton intrusion.
Andalusite pseudomorph textures west of the fault and
garnet inclusion textures along the fault suggest that
fault movement occurred after heating caused by
pluton intrusion began, and the lack of intracrystalline
deformation in the pluton suggests that deformation
ended before significant cooling occurred. These observations constrain the fault movement to the period of
time between pluton intrusion and crystallization. A
granite pluton 10 km in diameter intruding at 15 km
depth (or into 375 °C country rock at a geothermal
gradient of 25 °C km−1) should crystallize within one
million years of intrusion. One to four km of displacement in one million years implies an average slip rate
of 1–4 mm year−1, somewhat less than the modern
slip rate on thrust faults such as the Oak Ridge Fault
in the Ventura basin, southern California (5 mm year−1,
Yeats, 1988). The amount of fault slip suggested by
differences in metamorphic pressures is therefore
reasonable.
C O N C L U S IO N S A N D I MP L IC AT ION S
Some of the deformation along the Monroe Fault
occurred during intrusion of the Victory Pluton.
Movement of <1–4 km occurred between initial
intrusion of the pluton and crystallization below the
rheological critical melt fraction. Away from the fault,
deformation during intrusion was lower-strain and
involved subhorizontal stretching rather than subvertical stretching. The regional dominant foliation in
north-eastern Vermont (S ) pre-dates the intrusion of
2
the Victory Pluton.
It is not possible to infer emplacement or transport
mechanisms for a pluton, or its relationship to regional
or local deformation, simply from the map pattern of
the pluton. Furthermore, syn-kinematic plutons can
appear undeformed in hand-sample or thin sections as
long as deformation ceased before the pluton crystallized below the rheological critical melt fraction. It
is necessary to integrate observations from the pluton
margins and aureole at a variety of scales and combine
those observations with map-scale observations to
determine the true relationship between a pluton and
regional deformation.
This study suggests that either the magma that fed
the Victory Pluton made use of the Monroe Fault to
travel through the crust, or the presence of melt
weakened the crust and focused deformation along the
Monroe Fault. Perhaps both processes occurred simultaneously: magma took advantage of a pre-existing
shear zone to travel through the crust, and the presence
of melt weakened the fault, making movement possible
under lower differential stresses. This feedback does
not appear to have continued after the Victory Pluton
was entirely crystallized. Because retrograde minerals
in the aureole (chlorite after garnet, muscovite after
sillimanite, chlorite and muscovite after staurolite)
appear nearly entirely post-kinematic, deformation was
not partitioned into the country rocks as crystallization
strengthened the pluton.
Although the Victory Pluton appears to be intimately linked to the Monroe Fault, the other plutons
of the Northeast Kingdom batholith do not straddle
candidates for major faults. However, boudinaged and
folded dykes and sills are found around other plutons
in north-eastern Vermont (Doll, 1951; Dennis, 1956),
suggesting that local deformation occurred around
them as well. Given the lack of detailed internal
stratigraphy in the Waits River and Gile Mountain
Formations, of sufficient outcrop surrounding many of
the plutons, and of recent detailed mapping in much
of north-eastern Vermont, there may be ductile faults
that have not yet been recognized.
Although the dominant regional foliation in northeastern Vermont formed before intrusion of the Victory
Pluton, some contractional deformation occurred
during intrusion. This suggests that Acadian deformation continued until 390–370 Ma, if the Victory
Pluton is similar in age to other plutons of the Northeast Kingdom batholith.
ACKNOWLEDGEMENTS
Support for this research was provided by grants from
VT-EPSCoR. We thank D. Rankin for introducing us
to the Monroe Fault. This paper was improved by
discussions with T. Rushmer and by reviews from T.
Armstrong and C. Davidson.
R E F ER E N C E S
Aleinikoff, J. N. & Karabinos, P., 1990. Zircon U–Pb data for
the Moretown and Barnard Volcanic Members of the
Missisquoi Formation and a dike cutting the Standing Pond
Volcanics, southeastern Vermont. In: Summary Results of the
Glens Falls CUSMAP Project, New York, Vermont, and New
Hampshire. U. S. Geological Survey Bulletin, 1887, D1–D10.
SYN-TECTO NIC PLUTON INT RUSION, VERM ONT
Amato, J. M., Wright, J. E., Gans, P. B. & Miller, E. L., 1994.
Magmatically induced metamorphism and deformation in the
Kigluaik gneiss dome, Seward Peninsula, Alaska. T ectonics,
13, 515–527.
Armstrong, T. R., Tracy, R. J. & Hames, W. E., 1992. Contrasting
styles of Taconian, Eastern Acadian, and Western Acadian
metamorphism, central and western New England. Journal of
Metamorphic Geology, 10, 415–426.
Armstrong, T. R., Walsh, G. J. & Spear, F. S., 1997. A transect
across the Connecticut Valley sequence in east-central
Vermont. In: Guidebook to Field T rips in Vermont and Adjacent
New Hampshire and New York (eds Grover, T. W., Mango, H.
N. & Hasehohr, E. J.). New England Intercollegiate Geological
Conference, 89, A6-1–A6-56.
Arth, J. G. & Ayuso, R. A., 1997. The Northeast Kingdom
batholith, Vermont: geochronology and Nd, O, Pb, and Sr
isotopic constraints on the origin of Acadian granitic rocks.
In: T he Nature of Magmatism in the Appalachian Orogen (eds
Sinha, A. K., Whalen, J. B. & Hogan, J. P.) Geological Society
of America Memoir, 191, 1–18.
Arzi, A. A., 1978. Critical phenomena in the rheology of partially
melted rocks. T ectonophysics, 44, 173–184.
Ayuso, R. A. & Arth, J. G., 1992. The Northeast Kingdom
batholith, Vermont: magmatic evolution and geochemical
constraints on the origin of Acadian granitic rocks.
Contributions to Mineralogy and Petrology, 111, 1–23.
Berman, R. G., 1988. Internally-consistent thermodynamic data
for minerals in the system Na O–K O–CaO–MgO–
2 . 2 Journal
FeO–Fe O –Al O –SiO –TiO –H O–CO
of
2 29,
3 445–522.
2 3
2
2 2
2
Petrology,
Billings, M. P. (compiler), 1955. Geologic Map of New Hampshire.
Scale 15250 000. New Hampshire State Planning and
Development Commission, Concord, NH.
Billings, M. P., 1956. T he Geology of New Hampshire: Part II.
Bedrock Geology. New Hampshire State Planning and
Development Commission, Concord, NH.
Bradley, D. C., 1983. Tectonics of the Acadian orogeny in New
England and adjacent Canada. Journal of Geology, 91,
381–400.
Bradley, D., Tucker, R. & Lux, D., 1996. Early Emsian position
of the Acadian orogenic front in Maine. Geological Society of
America Abstracts with Programs, 28, 500.
Brown, M. & Rushmer, T., 1997. The role of deformation in the
movement of granitic melt: views from the laboratory and the
field. In: Deformation-Enhanced Fluid T ransport in the Earth’s
Crust and Mantle (ed. Holness, M. B.). Chapman & Hall,
London.
Brown, M. & Solar, G. S., 1998. Shear-zone systems and melts:
feedback relations and self-organization in orogenic belts.
Journal of Structural Geology, 20, 1365–1393.
Collins, W. J. & Sawyer, E. W., 1996. Pervasive granitoid magma
transfer through the lower-middle crust during non-coaxial
compressional deformation. Journal of Metamorphic Geology,
14, 565–579.
Dallmeyer, R. D., Vanbreemen, O. & Whitney, J. A., 1982.
Rb–Sr whole-rock and 40Ar/39Ar mineral ages of the Hartland
stock, south-central Maine: a post-Acadian representative of
the New Hampshire plutonic series. American Journal of
Science, 282, 79–93.
Davidson, C., Hollister, L. S. & Schmid, S. M., 1992. Role of
melt in the formation of a deep-crustal compressive shear
zone: the Maclaren Glacier metamorphic belt, south central
Alaska. T ectonics, 11, 348–359.
Davidson, C., Schmid, S. M. & Hollister, L. S., 1994. Role of
melt during deformation in the deep crust. T erra Nova,
6, 133–142.
Davidson, C., Rosenberg, C. & Schmid, S. M., 1996. Synmagmatic
folding of the base of the Bergell Pluton, Central Alps.
T ectonophysics, 265, 213–238.
Dennis, J. G., 1956. The geology of the Lyndonville area,
Vermont. Vermont Geological Survey Bulletin, 8.
Doll, C. G., 1951. Geology of the Memphremagog quadrangle
and the southeastern portion of the Irasburg quadrangle,
Vermont. Vermont Geological Survey Bulletin, 3.
285
Doll, C. G., Cady, W. M., Thompson, J. B., Billings, M. P.
(compilers), 1961. Centennial Geologic Map of Vermont. Scale
15250 000. Vermont Geological Survey, Montpelier, VT.
Eric, J. H. & Dennis, J. G., 1958. Geology of the ConcordWaterford area, Vermont. Vermont Geological Survey
Bulletin, 11.
Gordon, T. M., 1992. Generalized thermobarometry: solution of
the inverse chemical problem using data for individual species.
Geochimica et Cosmochimica Acta, 56, 1793–1800.
Gordon, T. M., Aranovich, L. Y. & Fed’kin, V. V., 1994.
Exploratory data analysis in thermobarometry: and example
from the Kisseynew Sedimentary Gneiss Belt, Manitoba,
Canada. American Mineralogist, 79, 973–982.
Hatch, N. L., 1987. Lithofacies, stratigraphy, and structure in
the rocks of the Connecticut Valley Trough, eastern Vermont.
In: Guidebook for Field T rips in Vermont (ed. Westerman, D.
S.), New England Intercollegiate Geological Conference, 2,
192–212.
Hatch, N. L., 1988. New evidence for faulting along the ‘Monroe
Line’, eastern Vermont and westernmost New Hampshire.
American Journal of Science, 288, 1–18.
Hollister, L. S. & Crawford, M. L., 1986. Melt-enhanced
deformation: a major tectonic process. Geology, 14, 558–561.
Hueber, F. M., Bothner, W. A., Hatch, N. L. & Aleinikoff, J.,
1990. Devonian plants from southern Quebec and northern
New Hampshire and the age of the Connecticut Valley
Trough. American Journal of Science, 290, 360–395.
Hutton, D. H. W., Dempster, T. J., Brown, P. E. & Becker, S.
D., 1990. A new mechanism of granite emplacement: intrusion
in active extensional shear zones. Nature, 343, 452–455.
Johansson, W. I., 1963. Geology of the Lunenburg–Brunswick–
Guildhall area, Vermont. Vermont Geological Survey Bulletin,
22.
Laird, J. & Albee, A. L., 1981. High-pressure metamorphism in
mafic schist from northern Vermont. American Journal of
Science, 281, 97–126.
Laird, J., Lanphere, M. & Albee, A., 1984. Distribution of
Ordovician and Devonian metamorphism in mafic and pelitic
schists from northern Vermont. American Journal of Science,
284, 376–413.
Menard, T. & Spear, F. S., 1994. Metamorphic P–T paths from
calcic pelitic schists from the Strafford Dome, Vermont, USA.
Journal of Metamorphic Geology, 12, 811–826.
Moench, R. H., 1992. The ‘Piermont allochthon’ in the Littleton–
Moosilauke area of west central New Hampshire—Reply.
Geological Society of America Bulletin, 103, 1541–1545.
Moench, R. H., 1996. Stratigraphic basis for the Piermont–
Frontenac allochthon, Bath to Piermont, New Hampshire,
and Bradford, Vermont. In: Guidebook to Field T rips in
Northern New Hampshire and Adjacent Regions of Maine and
Vermont (ed. van Baalen, M. R.), New England Intercollegiate
Geological Conference, 88, 133–153.
Moench, R. H., Boone, G. M., Bothner, W. A., Boudette, E. L.,
Hatch, N. L., Hussey, A. M. I. I. & Marvinney, R. G., 1995.
Geologic map of the Sherbrooke–Lewiston area, Maine, New
Hampshire, and Vermont, United States, and Quebec, Canada.
US Geological Survey Miscellaneous Investigations Map,
I-1898-D. Scale: 15250 000 Washington, DC.
Naylor, R. S., 1971. Acadian orogeny: an abrupt and brief event.
Science, 172, 558–560.
Onasch, E., 1997. Analysis of structures and metamorphism across
the Monroe L ine northeast of the V ictory Pluton, Granby,
Vermont. Unpublished Senior Thesis, Middlebury College,
Middlebury, Vermont, USA.
Page, L. R., 1968. Devonian plutonic rocks in New England. In:
Studies in Appalachian Geology: Northern and Maritime (eds
Zen, E., White, W. S., Hadley, J. B. & Thompson, J. B.).
Interscience Publishers, New York.
Paterson, S. R., Vernon, R. H. & Fowler, T. K. Jr, 1991. Aureole
Tectonics. In: Contact Metamorphism (ed. Kerrick, D. M.).
Mineralogical Society of America, Reviews in Mineralogy,
26, 673–722.
286
K. A . HA NNULA ET A L .
Rankin, D. W., 1996. Bedrock geology of the Littleton,
Vermont–New Hampshire 15-minute quadrangle northwest
of the Ammonoosuc Fault. In: Guidebook to Field T rips in
Northern New Hampshire and Adjacent Regions of Maine and
Vermont (ed. van Baalen, M. R.). New England Intercollegiate
Geological Conference, 88, 5–39.
Rast, N. & Skehan, J. W., 1993. Mid-Paleozoic orogenesis in
the North Atlantic: the Acadian orogeny. In: T he Acadian
Orogeny: Recent Studies in New England, Maritime Canada,
and the Autochthonous Foreland (eds Roy, D. C. & Skehan,
J. W.). Geological Society of America Special Paper, 275, 1–25.
Robinson, P. & Tucker, R. D., 1996. The ‘Acadian’ in central
New England: new problems based on U–Pb ages of igneous
zircon and metamorphic monazite and sphene. Geological
Society of America Abstracts with Programs, 28, 94.
Rothstein, D. A., Karlstrom, K. E., Hoisch, T. D. & Morrison,
J., 1994. Synkinematic contact metamorphism: implications
for the timing of pluton emplacement and regional deformation
in the Scanlon shear zone, south-eastern California. Journal
of Metamorphic Geology, 12, 709–721.
Solar, G. S., Pressley, R. A., Brown, M. & Tucker, R. D., 1998.
Granite ascent in convergent orogenic belts: testing a model.
Geology, 26, 711–714.
Spear, F. S., Hickmott, D. D. & Selverstone, J., 1990.
Metamorphic consequences of thrust emplacement, Fall
Mountain, New Hampshire. Geological Society of America
Bulletin, 102, 1344–1360.
Spear, F. S., Kohn, M. J. & Cheney, J. T., in press. P–T paths
from anatectic pelites. Contributions to Mineralogy and
Petrology.
Stanley, R. S. & Ratcliffe, N. M., 1985. Tectonic synthesis of the
Taconian orogeny in western New England. Geological Society
of America Bulletin, 96, 1227–1250.
Thompson, J. B., Robinson, P., Clifford, T. N. & Trask, N. J. J.,
1968. Nappes and gneiss domes in west-central New England.
In: Studies in Appalachian Geology: Northern and Maritime
(eds Zen, E., White, W. S., Hadley, J. B. & Thompson, J. B.),
pp. 203–218. Interscience Publishers, New York.
Thompson, J. B., Rosenfeld, J. L. & Chamberlain, C. P., 1993.
Sequence and correlation of tectonic and metamorphic events
in southeastern Vermont. In: Field T rip Guidebook for the
Northeastern United States: 1993 Boston GSA (eds Cheney,
J. & Hepburn, J. C.), University of Massachusetts Department
of Geology and Geography Contributions, 67, B-1–B-26.
Thompson, J. B., Rosenfeld, J. L., Hepburn, J. C. & Trzcienski,
W. E., 1997. How does New Hampshire connect to Vermont?
In: Guidebook to Field T rips in Vermont and Adjacent New
Hampshire and New York (eds Grover, T. W., Mango, H. N.
& Hasehohr, E. J.). New England Intercollegiate Geological
Conference, 89, C1-1–C1-17.
Vielzeuf, D. & Holloway, J. R., 1988. Experimental determination
of the fluid-absent melting relations in the pelitic system:
consequences for crustal differentiation. Contributions to
Mineralogy and Petrology, 98, 257–276.
Woodland, B. G., 1965. The geology of the Burke quadrangle,
Vermont. Vermont Geological Survey Bulletin, 28.
Yeats, R. S., 1988. Late Quaternary slip rate on the Oak Ridge
fault, Transverse Ranges, California: implications for seismic
risk. Journal of Geophysical Research, 93, 12137–12149.
Zen, E., 1983. Exotic terranes in the New England
Appalachians—limits, candidates, and ages: a speculative
essay. In: Contributions to the T ectonics and Geophysics of
Mountain Chains (eds Hatcher, R. D., Williams, H. & Zietz,
I.). Geological Society of America Memoir, 158, 55–81.
Received 27 April 1998; revision accepted 8 October 1998.