processes controlling vertical coupling and …kklepeis/publications/klepeis et al....
TRANSCRIPT
UNCORRECTED PROOF
Processes controlling vertical coupling and decoupling between the upper
and lower crust of orogens: results from Fiordland, New Zealand
Keith A. Klepeisa,*, Geoffrey L. Clarkeb, George Gehrelsc, Jeff Vervoortc
aDepartment of Geology, University of Vermont, Burlington, VT, 05405-0122, USAbSchool of Geosciences, Division of Geology and Geophysics, University of Sydney, NSW 2006, Australia
cDepartment of Geosciences, University of Arizona, Tucson, AZ 85721, USA
Received 27 January 2002; received in revised form 15 July 2003; accepted 25 August 2003
Abstract
The pre-Cenozoic configuration of western New Zealand allows determination of the effects of magmatism and a changing lower crustal
rheology on the evolution of a Cretaceous orogen from upper to lower crustal levels (10–50 km). Beginning at ,126 Ma, a composite
batholith dominated by diorite was emplaced into the lower crust. During emplacement, deformation was partitioned into zones weakened by
magma and heat, leading to the development of two layer-parallel shear zones at the upper and lower contacts of the batholith. Transient
vertical decoupling of the crust above and below the batholith occurred from ,126 Ma until ,120 Ma as magma was emplaced into and
moved through a weak, thick lower crust. By ,116 Ma, however, much of the batholith had crystallized and the lowermost crust had cooled
from 750 8C , T , 850 8C to T ¼ 650–700 8C. Cooling was aided by the juxtaposition of pre-existing crust against hot new crust and by the
efficient extraction of partial melts out of the lower crust. Cooling together with dehydration of the lower crust and mafic compositions led to
the development of a strong, dry, lower crustal root by ,116 Ma. A strong lower crust resulted in high degrees of vertical coupling between
the upper and lower crust during contraction from ,116 to ,105 Ma even as magma continued to be emplaced into the mid-upper crust. A
narrow, focused orogenic style in the upper crust at this time reflected a highly viscous lower crust through which compressional stresses
were transferred vertically. The results imply that changes in plate boundary dynamics rather than the thermal weakening of thick lower crust
during convergence controlled the onset of regional extension at ,108–105 Ma.
q 2003 Published by Elsevier Ltd.
Keywords: Vertical coupling and decoupling; Magmatism; Orogen
1. Introduction
Studies of convergent margins worldwide have shown
that deformation patterns and the mechanical behavior of
continental crust vary according to crustal level and tectonic
setting (e.g. Sisson and Pavlis, 1993; Axen et al., 1998;
Klepeis and Crawford, 1999; Miller and Paterson, 2001;
Karlstrom and Williams, 2002; Teyssier et al., 2002).
Experimental data (Wilks and Carter, 1990; Rushmer, 1995;
Rutter and Neumann, 1995), numerical simulations (Harry
et al., 1995; Ellis et al., 1998; McKenzie et al., 2000), and
analytical models (Royden, 1996) indicate that lower crustal
strength and rheology especially affect how deformation is
partitioned vertically through the lithosphere during con-
vergence. These studies emphasize the critical role the
lower crust plays in linking the upper mantle with the upper
crust of orogens.
Despite this work, however, we still lack direct
information on the mechanisms by which deformation is
relayed vertically between different sections of the litho-
sphere, especially as physical and chemical conditions in the
lower crust change. Large, dipping shear zones that divide
the crust and upper mantle into different structural domains
have been observed or postulated in many orogenic belts
(Oldow et al., 1990; Harry et al., 1995; Willett, 1998;
McKenzie et al., 2000; Teyssier et al., 2002) but we do not
fully understand how deformation above, below, and within
these potentially transient features relate to one another or
affect orogenic evolution. This gap in knowledge arises
partly because orogens that allow direct observation of
processes at lower crustal levels and their relationship with
the upper crust are rare. In addition, the age and kinematic
0191-8141/$ - see front matter q 2003 Published by Elsevier Ltd.
doi:10.1016/j.jsg.2003.08.012
Journal of Structural Geology xx (0000) xxx–xxx
www.elsevier.com/locate/jsg
* Corresponding author. Tel.: þ1-802-656-0246; fax: þ1-802-656-0045.
E-mail address: [email protected] (K.A. Klepeis).
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significance of lower crustal fabrics identified in geophy-
sical studies (e.g. Warner, 1990; Mayer et al., 1997; Nemes
et al., 1997) commonly are difficult to confirm.
In this paper, we show how displacements were
transferred vertically from lower to upper crustal levels of
an ancient orogen by reconstructing pieces of a composite
crustal column now exposed in Fiordland and Westland
(Fig. 1). This approach is possible because of the exposure
of an Early Cretaceous mid–lower crustal section in
Fiordland (Fig. 1; 25–50 km paleodepths) and its originally
contiguous mid–upper crust in Westland (Fig. 1; 8–27 km
paleodepths). The Alpine Fault now separates rocks of the
Fiordland belt from those of similar Early Cretaceous and
older affinity in Westland. Excellent pre-Cenozoic markers,
including the western margin of the Median Batholith (Fig.
1, inset), indicate that ,460 km of offset have accumulated
along the Alpine Fault (Wellman, 1953; Molnar et al., 1999;
Sutherland et al., 2000). Once restored to their pre-Cenozoic
configuration, the Fiordland and Westland regions form
parts of the same orogenic belt (Fig. 1; Oliver, 1990;
Tulloch and Challis, 2000). Mid–late Cretaceous extension
exhumed much of the lower crustal parts of the belt in
Fiordland as parts of the upper plate (including Westland)
slid off to the SW and NE (Gibson et al., 1988; Tulloch and
Kimbrough, 1989; Gibson, 1990; Oliver, 1990). By
,90 Ma, the Fiordland rocks had cooled to ,400 8C and
were in the upper 10 km of the crust (Mattinson et al., 1986;
Nathan et al., 2000; Claypool et al., 2002). The results of
this differential exhumation and offset allowed us to
compare processes and events from ,126–90 Ma in
Fiordland with those that occurred during the same time
interval at upper crustal levels in Westland.
We present structural, metamorphic and geochronologic
data that reveal the evolution of shear zones that separate the
middle and lower crustal section into distinctive structural
domains. We compare these features to structural patterns
and events preserved in the mid–upper crust and describe
how strain was partitioned within the orogen during a
transition from lithospheric contraction to extension. The
data indicate that strong physical and kinematic links were
established between the different layers of the lithosphere
only a few (,3–4 Ma) million years after emplacement of a
major batholith in the lower crust. The results provide direct
physical evidence of transient vertical decoupling followed
by coupling between the upper and lower crust during the
period ,126–105 Ma. We discuss the controls on coupling
and decoupling processes and explain why the mechanical
behavior of the Fiordland–Westland orogen may differ
from other orogens that experienced larger degrees of partial
melting and pluton emplacement in the deep crust.
2. The Fiordland–Westland orogen
The Fiordland–Westland orogen (Fig. 1) records a
history of magmatism, metamorphism and deformation
that accompanied the development of an early Mesozoic arc
along the margin of Gondwana. A Western Belt (Fig. 1),
representing the ancient continental margin, contains
Paleozoic terranes that preserve a record of mostly
,380–300 Ma pluton emplacement, low- to high-grade
metamorphism, and convergence (Landis and Coombs,
1967; Bishop et al., 1985; Cooper and Tulloch, 1992; Muir
et al., 1996; Ireland and Gibson, 1998). An Eastern Belt
(Fig. 1) contains plutons and volcano-sedimentary terranes
that originally formed outboard of the margin during the
early Mesozoic (Mattinson et al., 1986; McCulloch et al.,
1987; Tulloch and Kimbrough, 2003). Between these two
provinces (Fig. 1) is a linear, N- and NE-trending belt of
early Mesozoic plutonic, volcanic and sedimentary rock
called the Median Tectonic Zone (Kimbrough et al., 1994;
Muir et al., 1994) or the Median Batholith (Mortimer,
1999a,b).
The Median Batholith contains several compositionally
distinctive plutonic suites. On the outboard (east) side of the
Gondwana margin, the Median Suite of Mortimer and
Tulloch (1996) and the Darran Suite (Fig. 2) of Muir et al.
(1998) were emplaced into a Permo–Triassic accretionary
complex (Brook Street terrane, Fig. 1) mostly during the
interval 170–128 Ma (Mortimer, 1999). These suites are
dominated by diorite although gabbro and smaller granite
plutons also are common. On the continent side of the
Median Batholith is a younger belt of ,126–105 Ma
plutonic rock that includes the Separation Point and Rahu
suites (Fig. 1; Bradshaw, 1990; Kimbrough et al., 1994;
Muir et al., 1994; Mortimer et al., 1999a; Tulloch and
Kimbrough, 2003). At mid–upper crustal levels, now
exposed in Westland and easternmost Fiordland, rocks of
these latter two suites are dominated by tonalitic, grano-
dioritic and granitic compositions. The lower crustal levels
of this belt, exposed in Fiordland, are represented by the
dioritic–monzodioritic Western Fiordland Orthogneiss
(WFO; Figs. 1 and 2). Gabbro also is common in the WFO.
The emplacement of plutons of the Separation Point
Suite into both Eastern and Western belts at ,126 Ma
indicate that these two provinces were together at that time
Fig. 1. Present configuration (top of inset) and Cretaceous reconstruction (bottom of inset and main diagram) of western New Zealand after Tulloch and Challis
(2000). Geologic relationships are from Wood (1972), Oliver and Coggon (1979), Bradshaw (1989), Daczko et al. (2002a) and Klepeis and Clarke (2003).
Abbreviations show key locations or features: MB—Median Batholith, SP—Separation Point, L—Largs Terrane; MS—Milford Sound; GS—George Sound;
CS—Caswell Sound; CHS—Charles Sound; DS—Doubtful Sound; LTA—Lake Te Anau; LM—Lake Manapouri. Metamorphic pressures from Fiordland
represent the peak of Early Cretaceous metamorphism at ,120 Ma and are from Bradshaw (1985, 1989a,b), Brown (1996), Klepeis et al. (1999), Clarke et al.
(2000) and Daczko et al. (2001a,b, 2002a,b). See text for discussion. Pressures from Westland show shallower early–mid-Cretaceous (125–105 Ma) pluton
emplacement depths (after Tulloch and Challis, 2000). Metamorphic and structural data from Fiordland show a south-tilted lower crustal section (b).
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(Williams and Harper, 1978; Mortimer et al., 1999a,b).
Hollis et al. (2003) obtained ages that suggest this
amalgamation occurred as early as ,136 Ma and certainly
by ,129 Ma. Tulloch and Kimbrough (2003) determined
that differences in composition and age reflect a configur-
ation where mantle-derived plutons of the outboard belt
were underthrust beneath Gondwana where they partially
melted at high pressures producing magma of the inboard
belt. Daczko et al. (2001a, 2002a) and Klepeis and Clarke
(2003) describe the contraction that accompanied this
amalgamation.
By ,108–105 Ma, regional extension affected parts of
Fiordland and Westland (Bradshaw, 1989; Tulloch and
Kimbrough, 1989; Gibson and Ireland, 1995; Spell et al.,
2000). Extensional metamorphic core complexes in the
Paparoa and Victoria ranges (Fig. 1) formed in the mid–
upper crust beginning at this time. Emplacement of the
,110 Ma Hohonu granitoids (Fig. 1; Waight et al., 1998)
Fig. 2. Structural map of Fiordland. Only main lithologic divisions are shown (see Turnbull (2000) and Klepeis and Clarke (2003) for details). Bold black lines
show boundaries of major shear zones. Structural measurements are from Bradshaw (1985, 1990), Blattner (1991), Klepeis et al. (1999), Daczko et al. (2002a),
Claypool et al. (2002), Turnbull (2000) and Klepeis and Clarke (2003). Foliation trajectories (thin black lines) show interpolation of structural trends. Plotted
U–Pb dates are from Hollis et al. (2003; white boxes) and this study (black boxes). Abbreviations show site localities: Mt. Daniel (MD), Mt. Edgar (ME),
Camp Oven Creek (CO), the Pembroke Valley (P), Mt. Ada (MA), Selwyn Creek (SC), Mt. Kepka (MK).
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and the youngest plutons of the Separation Point Suite at
,105 Ma (Tulloch and Kimbrough, 2003) may have
overlapped with the transition to extension. The Doubtful
Sound Shear Zone (Gibson et al., 1988; Gibson and Ireland,
1995) and the Anita Shear Zone (Hill, 1995; Klepeis et al.,
1999) record decompression and exhumation after ,108–
105 Ma (Fig. 1).
During the Cenozoic changes in relative motions among
the Pacific, Australian, and Antarctic plates led to the
development of the modern Pacific–Australian plate
boundary by ,25 Ma (Cooper et al., 1987; Sutherland,
1995; Lamarche et al., 1997). Approximately 70–75% of
current motion arising from the oblique convergence
between the Australian and Pacific plates is accommodated
by slip along the Alpine Fault (Sutherland et al., 2000;
Norris and Cooper, 2001). Claypool et al. (2002) review the
effects of late Cenozoic faulting and exhumation (#6 km)
on the structure of northern Fiordland.
2.1. Early Cretaceous crustal thickening and magmatism in
the mid–upper crust
In Westland, plutons of the ,126–105 Ma Separation
Point Suite (Fig. 1) record Early Cretaceous emplacement
depths of 8–27 km, with the greatest depths (17–27 km)
occurring on the western side of the Median Batholith
(Tulloch and Challis, 2000). The youngest plutons also
occur on this western side (Tulloch, 1979; Harrison and
McDougall, 1980; Kimbrough et al., 1994; Muir et al.,
1994). Early Cretaceous deformation was distributed across
a 50–75-km-wide zone in Westland. Plutons localized some
of this deformation. Ductile shear zones, including the
Wainui Shear Zone (Fig. 1), formed on the western side of
the Separation Point Suite, including in its amphibolite
facies contact aureole (Grindley, 1980; Bradshaw, 1993;
Tulloch and Challis, 2000). Paleozoic fault zones also
influenced Early Cretaceous deformation. A record of
multiple displacements in many fault zones (Cooper,
1979) probably reflects some Mesozoic reactivation.
Crustal thickening by magma addition in the Eastern Belt
is indicated by a peak in subduction-related magmatism at
,140 Ma (Tulloch and Kimbrough, 2003). A major thrust
fault on Stewart Island (Fig. 1) formed between the outboard
and inboard belts during ,125–105 Ma magmatism
(Allibone and Tulloch, 1997; Tulloch and Kimbrough,
2003). Southeast of Fiordland, the Early Cretaceous Largs
Terrane (L, Fig. 1; Williams, 1978; Mortimer et al., 1999b)
is folded and thrust over the Brook Street Terrane
(Bradshaw, 1989). In the Victoria Range (Fig. 1), argon
thermochronology (Spell et al., 2000) and metamorphic data
(Tulloch, 1979) suggest that thickening and burial of the
Western Belt to mid-crustal levels initiated by ,130–
120 Ma and occurred prior to ,105 Ma (Tulloch and
Challis, 2000; Tulloch and Kimbrough, 2003).
2.2. Early Cretaceous crustal thickening, magmatism and
partial melting in the mid–lower crust
In Fiordland, the WFO was emplaced into a mid–lower
crust composed of Paleozoic metasediment and orthogneiss,
and early Mesozoic intrusive rock (Hollis et al., 2003).
Published ages from the WFO range from ,126 to
,116 Ma (Mattinson et al., 1986; Muir et al., 1998). New
high-precision data reported by Tulloch and Kimbrough
(2003) suggest that the main dioritic phase of the WFO was
emplaced during the interval 126–124 ^ 2 Ma. Hollis et al.
(2004) reports zircon ion probe data from Mt. Daniel and
George Sound that suggest that the WFO crystallized at
121.8 ^ 1.7 and 120 ^ 2.6 Ma, respectively, at these
localities (Hollis et al., 2004). Part of the WFO near
Doubtful Sound may be as young as ,116 Ma (Hollis et al.,
2004). Tulloch and Kimbrough (2003) also report a
116.6 ^ 1.2 Ma age from Wet Jacket Arm south of
Doubtful Sound that supports this interpretation. Alterna-
tively, the young zircon may reflect recrystallization.
Bradshaw (1989a, 1990) and Bradshaw and Kimbrough
(1989) used metamorphic P–T paths to infer an up-pressure
metamorphic history for the WFO following its emplace-
ment. Clarke et al. (2000) confirmed an up-pressure loading
history from P , 8 to 14–16 kbar for rocks of the Arthur
River Complex on the basis of metamorphic mineral
assemblages. The Arthur River Complex (Fig. 2) is
composed of Paleozoic paragneiss and both Paleozoic and
Mesozoic orthogneiss that lie below the WFO (Tulloch et al.,
2000; Hollis et al., 2003).
Oliver (1990) and Brown (1996) inferred that the cause
of burial involved magma loading during and following
emplacement of the WFO. Muir et al. (1995, 1998)
suggested that part of the Median Batholith was thrust to
lower crustal depths beneath western Fiordland and melted
to produce the WFO. This interpretation is supported by
geochemical data reported by Tulloch and Kimbrough
(2003). Daczko et al. (2002a) showed that some tectonic
burial was accomplished by the stacking of thrust sheets
above the WFO. Klepeis and Clarke (2003) showed that
lower crustal thickening involved displacements on both
steep and gently dipping shear zones.
Granulite facies metamorphism in the lower crust
accompanied and followed WFO emplacement (Oliver,
1977; Bradshaw, 1985, 1989a,b; Gibson and Ireland, 1995;
Clarke et al., 2000; Daczko et al., 2001b). Thin meta-
morphic overgrowths on Paleozoic and Mesozoic zircon
yield an average age of ,120 Ma in northernmost Fiordland
(Tulloch et al., 2000; Hollis et al., 2003). This age is
interpreted to represent the peak of metamorphism at
temperatures of 750 8C , T , 850 8C. Pressures represent-
ing the peak of this metamorphism range from P ¼ 7–
9 kbar in the shallowest part of the section to P ¼ 11–
16 kbar at the deepest levels (data plotted in Fig. 1). Recent
work suggests that granulite facies metamorphism
accompanied the partial melting of metadiorite during and
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after WFO emplacement and was promoted by the
mobilization of water-poor melt in fractures (Clarke et al.,
2000; Daczko et al., 2001b). Partial melting was controlled
by the decomposition of hornblende ^ clinozoisite (Anti-
gnano, 2002; Klepeis et al., 2003). Antignano (2002) used
experiments to show that positive volume changes associ-
ated with the reaction of hornblende ^ clinozoisite to
produce melt were sufficient to fracture matrix feldspar
and quartz. Klepeis and Clarke (2003) outline relationships
that suggest the fractures formed by high fluid pore pressure
in melt pockets and propagation during the fluid-absent
melting of mafic lower crust.
Following peak metamorphism at the granulite facies,
the batholith and its lower crustal host rock hydrated and
cooled (Daczko et al., 2002b). Quantitative cation mapping
and thermodynamic modeling of kyanite- and paragonite-
bearing assemblage by Daczko et al. (2002c) suggested
isobaric cooling of the Arthur River Complex to conditions
of T ¼ 650–700 8C and P ¼ 11–15 kbar.
3. Structure of the middle and lower crustal sections
Continuous exposure along Fiordland’s waterways
combined with ridge sections above the fjords were used
to construct serial cross-sections from Milford Sound to
Caswell Sound (Figs. 1b, 3 and 4). This approach allowed
determination of the three-dimensional structure of northern
Fiordland (Figs. 2 and 5). The results show a tilted middle–
lower crustal section (,80 km horizontal distance) on the
western side of Fiordland that includes both the upper and
lower contacts of the WFO. The uppermost contact of this
batholith is well exposed at Caswell Sound; the lowermost
contact is well exposed at Mt. Daniel (Figs. 1b and 2). These
contacts dip variably to the S and W and are deformed by
shear zones of variable thickness and geometry. The shear
zones at these two localities formed at different paleodepths
and divide the crustal section into distinctive structural
domains.
The geometry of the titled section, with Caswell and
George sounds exposing the middle crust (25–30 km depth)
and areas north of Mt. Daniel exposing the lower crust (45–
50 km depth) is consistent with variations in Early
Cretaceous paleodepths determined using published ther-
mobarometric data (Fig. 1). Pressures reflecting the peak of
Early Cretaceous (,120 Ma) metamorphism at Caswell and
George sounds have a range of P ¼ 7–9 kbar and occur
within the WFO and its contact aureole (Daczko et al.,
2002a). Pressures reflecting peak metamorphic conditions
during and following emplacement of the WFO at Mt.
Daniel and Milford Sound are P ¼ 10–13 kbar (Bradshaw,
1985, 1989a,b) and P ¼ 12–16 kbar (Clarke et al., 2000),
respectively. South of Charles Sound the dip of the WFO-
country rock contact changes to the NE, subparallel to the
Doubtful Sound Shear Zone (Fig. 1b). Peak pressures in the
WFO at Charles Sound are recorded at P ¼ 8–10 kbar
(Bradshaw, 1985; Brown, 1996). South of Charles Sound
paleodepths increase. In the footwall of the Doubtful Sound
Shear Zone, pressures reflecting the peak of Early Cretac-
eous metamorphism are P ¼ 11–12.5 kbar (Gibson and
Ireland, 1995). The tilted depth section between Caswell
and Milford sounds thus lies in the hanging wall of the
Doubtful Sound Shear Zone (Fig. 1b).
The western boundary of the lower crustal section in
northern Fiordland coincides with the steep, upper amphi-
bolite facies Anita Shear Zone (Fig. 2). This shear zone cuts
all Early Cretaceous fabrics within the Arthur River
Complex and the WFO (Figs. 2 and 5) and separates the
high-grade rocks from Paleozoic rocks (including the
Greenland Group and Saint Anne Gneiss) to the west. The
western boundary of the lower crustal section coincides with
the eastern margin of a 10–15-km-wide shear zone named
the Indecision Creek Shear Zone by Klepeis and Clarke
(2003). This shear zone separates the granulites from
weakly metamorphosed ,150–130 Ma plutonic rock of
the Darran and Roxburgh suites to the east (Fig. 2). Most of
the Darran Suite is only weakly deformed except on its
western sides (see also Blattner and Graham, 2000). Highly
deformed rocks on this western side include the Selwyn
Creek Gneiss (SC, Fig. 2). Brittle faults deform the steep
margins of the Anita Shear Zone and the Darran Suite.
East of the Anita Shear Zone the lower crustal section
displays two structural domains (Figs. 2 and 5) that are
separated by a 4–5-km-wide transitional zone (Fig. 4a). The
western domain is composed of Paleozoic metasediment,
mafic dikes, and layered intrusions, including the ,129 Ma
Mt. Edgar Diorite (Fig. 2; Hollis et al., 2003). Igneous
layering and gneissic foliations in this domain dip
moderately to the S, W and SW (Fig. 2). The eastern
domain, collectively termed the Indecision Creek Complex
by Bradshaw (1990), is dominated by steep foliations of the
Indecision Creek Shear Zone (Figs. 2 and 4). This latter
domain contains gabbroic gneiss, dioritic gneiss and
deformed felsic dikes. In the transitional zone, the foliations
of the western domain are folded and transposed parallel to
the margins of the Indecision Creek Shear Zone. At George
Sound, a 4–10-km-wide zone of high strain, named the
George Sound Shear Zone by Klepeis and Clarke (2003),
displays foliations that are similar in geometry to those of
the Indecision Creek Shear Zone (Fig. 2). This shear zone
lies structurally below a mid-crustal fold–thrust belt at the
Caswell Sound (Fig. 2) that was first identified by Daczko
et al. (2002a).
4. Shear zone evolution in the middle and lower crust
4.1. The Caswell Sound fold–thrust belt
At Caswell Sound, garnet granulite and upper amphibo-
lite facies thrusts sole into a subhorizontal shear zone
located at and below the contact between the WFO and its
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UNCORRECTED PROOF
host rock (Fig. 3). Inside the WFO, the shear zone is defined
by a ,1-km-thick section of subhorizontal to gently dipping
upper amphibolite facies foliations (Fig. 3d and e). Country
rock is composed of calcsilicate gneiss, marble, and
metapsammitic schist that probably form part of the
Paleozoic–Triassic sequences of Gondwana (Bradshaw
and Kimbrough, 1991; Hollis et al., 2004). At the east end
of the sound, imbricated, W-dipping thrust splays formed
within the contact aureole of the WFO and locally cut across
it (Fig. 3). Tight to isoclinal, S-plunging folds between
thrusts deform a penetrative gneissic foliation (S1) in
metasedimentary country rock. To the west of the imbricate
series the fold geometry changes from tight and overturned
to open and upright, reflecting a decrease in strain intensity
above and away from the WFO margin and the basal shear
zone. A steep upper amphibolite facies foliation parallels
the axial planes of the folds (Fig. 3b). Farther west
(,12 km), the folds gradually tighten and overturn to the
east above an E-dipping thrust fault. This E-dipping thrust
separates the high-grade rocks of the thrust belt to the east
from the weakly deformed McKerr monzodiorite to the west
(Fig. 3).
Flattened clusters of coarse hornblende, clinozoisite and
garnet in a feldspar matrix define thrust plane foliations
inside the WFO. Inside the WFO and within 500 m of its
uppermost contact, feldspar in the thrust zones was
dynamically recrystallized along grain boundaries. In
contrast, greater than 500 m above the contact, feldspar
behaved in a brittle manner during deformation in thrust
zones although mylonitic textures also are common. Thrust
faults greater than 500 m and up to 2.5 km from the contact
with the WFO are defined by aligned chlorite, muscovite,
quartz, feldspar, clinozoisite and amphibole. Daczko et al.
(2002a) described these variations in mineral assemblage in
detail and showed that they reflect a temperature gradient of
T ¼ 700–800 8C within the (500 m thick) contact aureole
and T ¼ 550–600 8C outside of it. These strong links
among increasing metamorphic grade, the recrystallization
of feldspar, and proximity to the WFO suggests that the
thrusts were preferentially partitioned into crust that was
thermally softened by the emplacement of the batholith. The
subsolidus character of the thrust fabrics also indicates that
contraction outlasted emplacement and crystallization of the
WFO.
Structurally below the Caswell thrusts, the WFO is
composed of layered dioritic intrusions and folded rafts of
upper amphibolite facies country rock (Fig. 3). At George
Sound (Fig. 2) the rafts contain a folded layer-parallel
foliation (S1) that is cut by the dioritic intrusions and
metasedimentary rock is migmatitic within 500 m of the
Fig. 3. Composite NW–SE profile of region between Charles Sound and George Sound. See Fig. 2 for section locations. Light shaded regions in B–B0 are
metasedimentary (Paleozoic–Mesozoic) country rock. Shaded and black units in D–D0 are felsic and mafic dikes, respectively. Thin, solid, black lines are
lithologic layering, dashed lines are foliation trends, and bold, black lines are shear zones. Equal area stereoplots (a–h) show poles to foliations, lineations and
fold axes. All data were measured by the authors and Daczko et al. (2002a) except those in plot e, which are from Bradshaw (1985). Locations of samples
(995a, 9928) for U–Pb analyses are shown. Sample 995a is from the Mckerr monzodiorite.
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UNCORRECTED PROOF
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UNCORRECTED PROOF
WFO margin (Fig. 3). The preservation of this migmatite
supports our interpretation that thrust zones were preferen-
tially partitioned into a similar zone that was weakened by
melt and heat at Caswell Sound. Foliations inside the diorite
form two dominant structural trends. The first trend includes
highly variable magmatic flow foliations (SWFO) defined by
the planar alignment of clinopyroxene, hornblende, tabular
plagioclase and other minerals. The second trend includes
steeply to moderately dipping subsolidus foliations (e.g.
Fig. 3f and g) that are heterogeneously developed within the
batholith. These latter foliations are best developed in the
George Sound Shear Zone (Figs. 2 and 3f) where they cut
the older synmagmatic SWFO foliations. Tight, upright, S-
plunging folds of dikes with steep axial planar foliations
display geometries that are similar to the upright folds and
steep foliations of the Caswell fold–thrust belt (compare
Fig. 3b and f). The George Sound Shear Zone flattens up
section and merges with the subhorizontal shear zone
beneath the Caswell fold–thrust belt (Fig. 3).
4.2. The Mount Daniel Shear Zone
A few hundred meters above the lowermost contact of
the WFO exposed at Mt. Daniel, the main phase of the WFO
is a coarse-grained diorite that contains pods of gabbro. A
hornblende cumulate layer occurs at the base of the diorite
(Fig. 6). These rocks display primary igneous layering and,
locally, a coarse-grained magmatic flow foliation (SWFO)
defined by aligned hornblende, clinozoisite, and plagioclase
Fig. 4. Cross-sections across the Mt. Daniel region (a) and north of Milford Sound (b). See Fig. 2 for locations. Equal area stereoplots show poles to foliations,
lineations and fold axes. Symbols the same as in Fig. 3. Plots i, ii, and iv in (a) include data from Bradshaw (1985). P and M are samples of orthogneiss dated by
Tulloch et al. (2000). PB3 in (b) is post-tectonic dike discussed in the text. # represent granulite facies fracture arrays.
Fig. 5. Composite block diagram of region between Caswell Sound (south end) and Milford Sound (north end) constructed using structural relationships shown
in Figs. 1a, 2, 3 and 4. Equal area stereoplots show poles to foliations, mineral lineations and fold axes. Plots (a)–(d) represent the western domain; plots (e)–
(h) represent the eastern domain. Symbols the same as in Fig. 3. Plot (c) is from Bradshaw (1985). WFO is the Western Fiordland Orthogneiss.
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UNCORRECTED PROOF
that parallel a moderately (40–508) SW-dipping basal
contact (Fig. 4a, part iii).
Below the cumulate layer is a banded igneous complex
(Fig. 6) that preserves evidence of both suprasolidus and
subsolidus deformation. Sheeted tonalite intrusions in this
zone display undulate, diffuse contacts with slightly more
mafic tonalitic bodies that reflect injection into an
incompletely crystallized host (see also Fig. 5 of Klepeis
and Clarke, 2003). Discordant mafic dikes with sharp,
straight contacts cut some tonalite sheets and, in turn, are cut
by veins that originate from the surrounding tonalite host.
These mutually crosscutting relationships indicate the
simultaneous emplacement of tonalitic and more mafic
sheets.
The central and lower parts of the banded igneous
complex preserve relationships indicating the accumulation
of high strains while the rocks were partially molten.
Migmatitic tonalite and trondhjemite sheets are complexly
interfolded and stretched. Recumbent folds of dikes and
igneous layering are common. Interfolded, transposed
sheets are cut by less deformed sheets, indicating that
deformation coincided with the periodic emplacement of the
sheeted intrusions (see also Fig. 5d of Klepeis and Clarke,
2003). Tightly folded pegmatites display axial planes that
parallel the margins of tonalite layers. However, despite the
evidence of the high strains required to produce these tight
folds, many of the pegmatite dikes are not foliated. Coarse
biotite in the dikes forms radial or misaligned patterns and
plagioclase exhibits a clean, interlocking igneous texture
with little evidence of subsolidus recrystallization. These
features define a thin melt-enhanced shear zone at the base
of the WFO (Fig. 6).
Beneath the basal shear zone, metagabbroic–metadioritic
orthogneiss forms part of the Arthur River Complex (Fig. 6).
Greater than 100 m below the contact zone at Mt. Daniel,
the dominant rock type is a biotite-rich dioritic orthogneiss.
This latter unit is compositionally similar to the ,129 Ma
Mt. Edgar Diorite (Fig. 2). Between the biotite orthogneiss
and the basal shear zone is a 30–50-m-thick zone of
granulite facies, garnet-rich metagabbro. The metagabbro
contains little plagioclase and no biotite. These observations
suggest that the metagabbro represents a depleted part of the
Arthur River Complex that resulted from partial melting of a
biotite-rich dioritic host similar to that which occurs below
the metagabbro (see also Daczko et al., 2002b).
In addition to evidence of suprasolidus deformation, a
heterogeneous subsolidus overprint also occurs within the
basal shear zone. Thin (,10 m wide) upper amphibolite
facies shear zones locally cut the margins of some folded
intrusions and preserve evidence that plagioclase grain sizes
were reduced during dynamic recrystallization. The shear
zones parallel the axial planes of recumbent, SE-plunging
folds that are geometrically similar to the melt-enhanced
folds we described earlier. However, the former contain a
weak axial planar foliation (S2) defined by flattened
plagioclase and aligned biotite and hornblende (Fig. 4a,
parts ii and iv). The heterogeneous development of this
subsolidus foliation and a lack of transposition during
folding resulted in the preservation of the migmatitic
features at the base of the WFO. In addition, the subsolidus
folds deform the entire lower contact of the basal shear zone
and form part of a set that also occurs at the base of the Mt.
Edgar Diorite (Fig. 4a). We refer to these folds as F2
structures because they deform primary igneous layering
and older gneissic foliations (S1). The presence of S2 and the
subsolidus shear zones indicate that F2 folding either post-
dated or outlasted crystallization of the melt-enhanced shear
zone. We suggest that they are related to crustal thickening
Fig. 6. Vertical profile of the Mt. Daniel Shear Zone. Bold lines are upper amphibolite facies shear zones discussed in the text. WFO is Western Fiordland
Orthogneiss; ARC is Arthur River Complex. # represent granulite facies fracture arrays.
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UNCORRECTED PROOF
during and slightly after emplacement and burial of the
WFO.
Another set of folds occurs in a narrow zone (,20 m
thick) between parallel minor shear zones in the Arthur
River Complex (Fig. 6). These folds are disharmonic,
mostly upright and plunge gently to moderately to the S
(Fig. 4a, part v). A spaced crenulation cleavage approxi-
mately parallels the axial planes of these folds. This same
style of folding occurs in the transitional zone between the
western and eastern domains (Figs. 2 and 4a). On the basis
of crosscutting relationships, we refer to these upright folds
as F3 structures (Fig. 4a, part v).
Finally, above and below the basal shear zone are arrays
of discordant veins and fractures filled with leucosome that
cut all ductile fabrics in the basal shear zone (# symbols in
Figs. 4 and 6). Narrow (6–7 cm) dehydration zones
containing garnet- and clinopyroxene-bearing assemblages
surround garnet-bearing leucosome in some fractures.
Inside these zones, hornblende–clinozoisite-bearing assem-
blages that define foliation in these rocks are replaced by
garnet, clinopyroxene, and rutile. Daczko et al. (2002b)
reported symplectic intergrowths of clinopyroxene and
kyanite and also of clinozoisite, quartz, kyanite and plagio-
clase that partially replace the older hornblende and
clinozoisite assemblage. These assemblages and reaction
textures record dehydration of the WFO and its host rocks at
the garnet granulite facies following cooling of the WFO
and its basal shear zone at Mt. Daniel (Daczko et al., 2002b).
None of the fractures or leucosome appears folded.
4.3. The Indecision Creek and George Sound Shear Zones
The Indecision Creek and George Sound Shear Zones
(Figs. 2–4) display vertical to steeply dipping, upper
amphibolite facies foliations (SSZ) that strike to the N, NE
and NNE and dip variably to the NW and SE. These steep
foliations parallel the axial planes of tight, S-plunging folds.
Steeply to gently plunging hornblende and plagioclase
mineral lineations occur on foliation planes (Figs. 2–4).
Rock fabrics are locally mylonitic and the folds and
intrusive contacts are transposed in the shear zone.
South of Milford Sound, the transitional domain (Fig. 4a)
preserves crosscutting relationships among shear zone
fabrics and other fabrics and fold sets of northern Fiordland.
In this zone, all igneous layering and gneissic foliations of
the western domain (S1, SWFO, S2) and F2 folds are tightly
folded into south-plunging F3 folds (Figs. 2 and 4a). From
W to E across this zone, F3 fold axial planes steepen to near
vertical, interlimb angles decrease, and the F3 folds
gradually are transposed parallel to the steep SSZ foliations
(Fig. 4a, parts v, vi and vii). The dominant SSZ foliation
parallels the axial planes of these tight F3 folds. These
changes define a 3–4-km-wide positive strain gradient that
increases from W to E into the central part of the shear zone.
North of Milford Sound, in the Pembroke Valley (P, Fig.
2), steep fabrics of the Indecision Creek Shear Zone envelop
a large lens of dioritic and gabbroic gneiss (Fig. 4b). This
locality preserves features that record the progressive
development of mineral assemblages and fabrics in the
Indecision Creek Shear Zone. One of the most spectacular
features is a lattice-like array of 3–5-cm-wide dehydration
zones that surround steep, orthogonal sets of leucosome-
filled fractures in dioritic and gabbroic orthogneiss. The
dehydration zones contain coronas of garnet and clinopyr-
oxene mantling hornblende. These reactions zones, like
those found at Mt. Daniel and elsewhere, record dehydration
at the garnet granulite facies. Descriptions of the petrology,
P–Tconditions (P ¼ 13–16 kbar, T . 750 8C), and origin
of these features are provided by Blattner (1976), Oliver
(1977), Bradshaw (1989a,b), Clarke et al. (2000) and
Daczko et al. (2001b).
The dehydration zones and vein sets form markers (Fig.
7a) that record the evolution of two sets of superposed shear
zones (Daczko et al., 2001a). The first set includes pairs of
thin, 1–3-m-wide sinistral and dextral shear zones (Fig. 7b).
The sinistral set is dominant and displays a steep mylonitic
foliation that strikes to the E and NE. The dextral set is
subordinate in size and abundance to the sinistral set. This
latter set dips gently to moderately to the SW. Both shear
zone sets contain gently plunging hornblende and clinozoi-
site mineral lineations. Superimposed on the sinistral and
dextral shear zone pairs is a younger set of vertically
stacked, layer-parallel shear zones that dip gently to the SE
(Fig. 4b, part i and Fig. 7c). Each of these shear zones
contains a 7–10-m-thick central zone where asymmetric
pods of coarse-grained gneiss is surrounded by thin (,1 m
thick) mylonitic to ultramylonitic shear bands. The asym-
metric pods form imbricated, antiformal stacks. The thin
shear bands locally dip steeply to the NW and SE and swing
into parallelism with the gently SE-dipping shear zones
located above and below them. Hornblende mineral
lineations on foliation planes plunge to the SE (Fig. 4b,
part i). The vertical spacing between parallel shear zones is
,50–100 m. The exact thickness of the stack is unknown.
East and west of the Pembroke Valley, the delicate
dehydration zones and the layer-parallel thrusts are mostly
transposed and recrystallized by the Indecision Creek Shear
Zone (Fig. 7d). This shear zones cuts the lower (eastern
side) contact of the WFO southeast of Mount Daniel (Figs.
2b, 4a and 5) and records retrogression of granulite facies
mineral assemblages to the upper amphibolite facies.
Southeast of the Indecision Creek Shear Zone, the George
Sound Shear Zone cuts across the central part of the WFO
(Figs. 2 and 3).
5. Kinematic relationships
The kinematics of deformation that occurred while the
WFO batholith was partially molten are recorded best in the
melt-enhanced basal shear zone exposed at Mt. Daniel.
Oblique foliations and the truncation of sheeted intrusions
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UNCORRECTED PROOF
by successive sheets indicate top-to-the-E and -NE, thrust-
style displacements parallel to W- and SW-plunging
mineral lineations (Fig. 4a, part iii). The plunges of the
lineations indicate a sinistral component to the deformation.
The subsolidus shear zones that parallel the axial planes of
F2 folds give identical oblique-thrust senses of shear.
Below the WFO, kinematic indicators in the steeply
dipping shear zone pairs of the Pembroke Valley (Fig. 7b)
include oblique foliations, micro-faulted garnet, and asym-
metric tails on feldspar porphyroclasts. These shear zones
record mostly NE–SW stretching parallel to the arc with a
component of sinistral displacement. Daczko et al. (2001a)
showed that they also record subhorizontal shortening at
high angles to the trend of the Median Batholith. The
Pembroke Thrust Zone (Fig. 7c) contains asymmetric
hornblende and clinozoisite fish, minor shear bands, and
asymmetric tails on feldspar clasts that record a top-to-the-
NW sense of shear. The style of displaced, asymmetric pods
that are stacked on top of one another also reflects a
component of layer-perpendicular thickening. Together, the
steep shear zone pairs and Pembroke Thrust Zone record
subhorizontal (layer-parallel) shortening normal to the
batholith, sinistral arc-parallel displacements, and vertical
(layer-perpendicular) thickening. This result is consistent
with the oblique-thrust style displacements recorded in the
melt-enhanced shear zone at Mt. Daniel, suggesting that this
style of deformation began during emplacement of the
WFO.
Above the WFO, the Caswell fold–thrust belt also
records arc-normal contraction and crustal thickening
following crystallization of the WFO. Lineation trends are
similar to those that characterized those in the basal shear
zone at Mt. Daniel (Fig. 5a and d). The spread of lineation
plunges on foliation planes (Fig. 3c) also suggests that the
imbricated thrusts record a component of sinistral displace-
ment. The conjugate style of W-dipping thrusts with an
E-dipping back thrust (Fig. 3) indicates compression
directions at high angles to the trend of the arc.
Below the Caswell Sound fold–thrust belt, the Indeci-
sion Creek Shear Zone records shortening at high angles to
the arc leading to the development of steep foliation planes.
This shortening is best illustrated by the progressive change
in fold geometry, including tightness and the steepening of
fold axial surfaces, from W to E across the transition zone.
Outcrop-scale sense of shear indicators, including hornble-
nde and clinozoisite fish, asymmetric tails of biotite and
hornblende around garnet porphyroblasts, asymmetric
boudinage, and minor shear zones, mostly occur in areas
of low–intermediate strain at the eastern and western edges
of the shear zone. The sense of shear in these areas is
dominantly sinistral parallel to gently and moderately
plunging mineral lineations.
In the central part of the Indecision Creek Shear Zone,
changes in orientation of hornblende lineations with
increasing strain provide additional kinematic information.
From W to E across the transition zone, the lineations
change from gently and moderately S-plunging to near
vertical and steeply plunging (compare Fig. 4a, parts vi and
vii). The migration of these mineral lineations toward the
dip line of the steep shear zone indicates stretching parallel
Fig. 7. Block diagrams showing the sequence of deformation recorded in structures at Pembroke Valley (modified from Fig. 3 of Clarke et al. (2000) and Fig. 13
of Daczko et al. (2001a)). (a) Garnet granulite facies fracture arrays cut S1. (b) Steep, sinistral–dextral shear zone pairs deform fracture arrays. (c) Gently SE-
dipping thrust zones (see also Fig. 4b) cut sinistral shear zones. (d) Steep upper amphibolite facies foliation (SSZ) envelops shear zones at Pembroke Valley.
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UNCORRECTED PROOF
to this direction (Lin et al., 1998; Jiang and Williams, 1998).
The reference frame provided by the gneissic layering of the
western domain and well-defined boundaries indicate that
the shear zone was thickening vertically during contraction.
The kinematic evolution of the George Sound Shear
Zone has not been studied in detail due to its remote locality.
However, the increase in fold tightness and rotation of
hornblende lineations to down dip with increasing strain
suggest that it also records arc-normal contraction, near
vertical stretching. Minor shear zones show both dextral and
sinistral displacements. Given that these styles are similar to
those of the Indecision Creek Shear Zone (Fig. 5) we
suggest that the kinematic evolution of the two shear zones
is similar also.
6. The ages of lower crustal deformation, magmatism,
and metamorphism
Published dates and four new age determinations were
used to estimate the age of high-grade fabrics and intrusions.
At Caswell Sound, U–Pb spot analyses on single zircons
from the McKerr monzodiorite (sample 995a, Figs. 2 and 3)
and from a dioritic dike within the zone of imbricated
thrusts (sample 9928, Figs. 2 and 3) allowed us to place a
lower limit on the age of deformation within eastern and
western parts of the fold–thrust belt. Twenty analyses were
conducted on the cores of zircon from sample 995a using a
beam diameter of 50 microns (analytical procedures are
described in Appendix A). Nineteen grains yielded analyses
that are apparently of the same 206Pb/238U age, and one
additional grain is discordant due to inheritance (Fig. 8a).
The weighted mean of these analyses yields an interpreted
crystallization age of 116.8 ^ 3.7 Ma at the 2s level (Fig.
8a and b). Twenty-two analyses were conducted on the
cores of zircon grains from sample 9928 using a slot
diameter of 50 microns. Twenty grains yield analyses that
are apparently of the same age, and two additional analyses
are apparently discordant due to a slight amount of
inheritance. The final 206Pb/238U age is 118.7 ^ 3.8 Ma
(2s level; Fig. 8c and d). These data suggest that the
Caswell Sound fold–thrust belt evolved during and after the
interval 122.5–113 Ma.
Two samples of syntectonic dikes from within the
Indecision Creek Shear Zone (Ada2 and 0221K, Fig. 2)
provided an approximate lower age limit of deformation in
the shear zone. Forty-five analyses were conducted on
zircon cores from sample Ada2 using a laser beam diameter
of 25 microns. This dike from Mt. Ada (MA, Fig. 2) cuts the
steep foliation (SSZ) of the shear zone and also is folded
within it. These analyses yield two clusters of ages (Fig. 8e
and f). The older age of 204.0 ^ 6.1 Ma from zircon cores is
interpreted to record igneous crystallization. The rim ages of
115.7 ^ 3.8 Ma are interpreted to record the growth of
metamorphic zircon. The spread of the rim ages indicate
that shear zone deformation continued through the interval
119.5–112 Ma.
Fifty analyses were conducted on zircon grains from
sample 0221K using a beam diameter of 25 microns. This
dike (from near Mt. Kepka, Fig. 2) also cut steep foliation
planes (SSZ) in the Indecision Creek Shear Zone and is
folded within it. Most analyses were conducted on core
areas of the zircon grains, with a smaller number of analyses
on the rims (tips) of the grains. The rim analyses generally
yield ages that are younger than the core ages. The
occurrence of two distinct clusters of ages (Fig. 8g and h)
suggests that the grains record two phases of zircon growth:
an older phase at 365.3 ^ 11.4 Ma that reflects crystal-
lization of the dike, and a younger phase at 129.5 ^ 4.2 Ma
that reflects the growth of metamorphic zircon. The young
rim ages obtained from the 0221K and Ada2 samples are in
agreement with crosscutting relationships indicating that
deformation in the Indecision Creek Shear Zone outlasted
emplacement of the WFO.
These new ages are compatible with other published ages
(Fig. 8i). Tulloch et al. (2000) identified Paleozoic
(355 ^ 10 Ma) oscillatory-zoned cores and Early Cretac-
eous (134 ^ 2 Ma) sector-zoned cores from the Arthur
River Complex. Both these core types displayed thin low-U
rims that yield an average age of ,120 Ma but some with
ages as young as ,105 Ma (Tulloch et al., 2000). Similar
rim ages (Fig. 2) have been obtained from the deformed
western margin of the Darran Suite (Selwyn Gneiss of
Hollis et al., 2003). The Jurassic core age of sample Ada2 is
compatible with similar ages obtained from the Darran Suite
(Muir et al., 1998; Blattner and Graham, 2000). The rim age
of sample 0221K is similar to ,136–129 Ma crystallization
ages from intrusive rocks in the Arthur River Complex (Fig.
2; Hollis et al., 2003). Metamorphism leading to zircon
growth also could have accompanied emplacement of
intrusive rocks prior to the WFO (see also Tulloch et al.,
2000).
Zircon ages of ,82 Ma from a post-tectonic dike (PB3;
Fig 4b) indicate that ductile deformation in northernmost
Fiordland terminated in the Late Cretaceous (Hollis et al.,
2003). This age of a few million years younger than K–Ar
ages on hornblende (Nathan et al., 2000) and U–Pb dates on
apatite (Mattinson et al., 1986) indicate that the Arthur
River Complex had cooled to T ¼ 300–400 8C by ,90 Ma.
Near Doubtful Sound, Gibson and Ireland (1995) dated
thermal conditions of T . 800 8C at 107.5 ^ 2.8 Ma from a
sample of the WFO deformed by the Doubtful Sound Shear
Zone (sample D in Fig. 8i). This age could reflect the
recrystallization of zircon in the shear zone. However, the
chemistry and age of the zircon suggested to Gibson and
Ireland (1995) that it represents a new generation of zircon
growth during metamorphism. The age is consistent with
regional geologic relationships indicating that extension
began by ,108–105 Ma (Tulloch and Kimbrough, 1989,
2003; Gibson and Ireland, 1995; Spell et al., 2000). K–Ar
cooling ages of ,93 and ,77 Ma on amphibole and biotite,
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UNCORRECTED PROOF
Fig. 8. U–Pb isotopic data from zircon collected using an inductively coupled plasma mass spectrometer (ICPMS). Plots (a), (c), (e) and (g) are concordia plots (ellipses shown with dashed lines were not used to
calculate mean ages). An explanation is provided in Appendix A. Plots (b), (d), (f) and (h) show error analyses and distribution of analyses (solid lines) used to calculate mean ages and errors (2s level). Locations
of samples are shown in Fig. 2. Part (i) shows a comparison with published ages from Mattinson et al. (1986), Muir et al. (1998), Tulloch et al. (2000) and Hollis et al. (2003). Ages representing early phases of the
Median Batholith are from both the Darran Suite and the Arthur River Complex (A). Samples P and M are from Tulloch et al. (2000) located in Fig. 4b. Post-tectonic dike is sample PB 3 reported by Hollis et al.
(2003) and shown in Fig. 4b. Sample D is a zircon age from the extensional Doubtful Sound Shear Zone (Gibson and Ireland, 1995).
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UNCORRECTED PROOF
respectively, also support a late Cretaceous age for the
Doubtful Sound Shear Zone (Gibson et al., 1988).
7. Correlation of structures within Northern Fiordland
Crosscutting relationships, geochronology, similarities in
style and metamorphic grade, and the results of physically
tracing structures above and below the WFO allowed us to
correlate fabrics between Milford and Caswell Sounds. We
use these correlations to reconstruct the sequential evolution
of the section (Table 1, Figs. 9 and 10).
One especially useful marker unit is the WFO. The
regional extent of this batholith and its ,126–120 Ma age
allowed us to divide structures into groups that predated,
accompanied, and post-dated its emplacement. Structures
that predate emplacement occur in Paleozoic and early
Mesozoic host gneiss located above and below the WFO or
as xenoliths within it. Primary igneous layering and the
gneissic foliations (S1) in the Arthur River Complex and
Darran Suite are included in this group (Table 1). These
structures are locally cut by the WFO and mostly occur in
the western domain (Figs. 2 and 4a).
The second group of structures includes all magmatic
foliations (SWFO) that formed within the WFO during its
emplacement, including the Mt. Daniel Shear Zone (Table
1). These structures all exhibit evidence of deformation
while the batholith was still partially molten. The links
among increasing metamorphic grade, feldspar recrystalli-
zation, and proximity to the WFO also suggest that
deformation in the Caswell Sound fold–thrust belt began
during this stage. We include in this group the recumbent F2
folds and axial planar foliations (S2) that formed at the lower
contact of the WFO.
In northernmost Fiordland the melt-induced fracture
arrays and dehydration zones record the peak of granulite
facies metamorphism in the lower crust during or immedi-
ately after WFO emplacement. The development of these
arrays across large areas of the section (Pembroke Valley,
Mt. Daniel, George Sound, Doubtful Sound) suggests that
similar processes controlled their development. However,
these features may exhibit slightly different ages across the
section. Metamorphic rims on zircon suggest that in the
Milford Sound region this metamorphism mostly occurred
between ,123 and ,116 Ma with a clustering at ,120 Ma
(Tulloch et al., 2000; Hollis et al., 2003; this study). Near
Doubtful Sound a zircon age of 107.5 ^ 2.8 Ma may
indicate that granulite facies metamorphism there is
younger (Gibson and Ireland, 1995). Despite this age
range, we correlate these distinctive garnet granulite
reaction zones (GRZ, Table 1). The recrystallization of
these features in the Indecision Creek and George Sound
shear zones indicates that they formed prior to the shear
zones north of Caswell Sound.
The third group of fabrics includes the sinistral–dextral
shear zone pairs (Fig. 7b), the Pembroke thrust zone (Fig.
7c), and the steep fabrics of the Indecision and George
Sound Shear Zones (Fig. 7d). On the basis of structural
relationships exposed in the Pembroke Valley, Clarke et al.
(2000) and Daczko et al. (2001a) referred to the shear zone
Table 1
Time–space correlation of structures (dashed lines) from SW (left) to NE (right) within northern Fiordland. GRZ is garnet granulite reactions zones, S is
foliation where subscript refers to rock unit and generation: WFO is Western Fiordland orthogneiss
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UNCORRECTED PROOF
pairs and the ductile thrust fabrics as D3 and D4 structures,
respectively. However, the occurrence of these features in
the transitional zone of the Indecision Creek Shear Zone and
evidence of arc-normal contraction during each phase of
deformation suggest that they all reflect slightly different
stages of the same contractional event. We also include in
this group the tight F3 folds. This correlation is consistent
with crosscutting relationships indicating that all minor
shear zones and the F3 folds in the transitional zone deform
S1, S2, F2, SWFO and GRZ structures. These relationships
and the rim ages obtained from sample Ada2 (Figs. 6g and
8e) indicate that this deformation occurred after ,120 Ma
and outlasted the emplacement of a ,116 Ma dike.
South of George Sound, the George Sound Shear Zone
merges with the subhorizontal shear zone exposed at
Caswell and Charles Sounds (Figs. 3 and 5). On the basis
of crosscutting relationships with respect to the WFO both
the Caswell fold–thrust belt and the George Sound Shear
Zone developed following crystallization of the WFO and
before the onset of regional extension (Table 1). Hollis et al.
(2004) reports metamorphic rims on zircon from George
Sound (an average age of ,120 Ma with ages spread across
the interval ,138–106 Ma) that support this interpretation.
Together, the combined isotopic data from Tulloch et al.
(2000), Hollis et al. (2004) and from samples 995a, 9928
and Ada2 indicate that the Caswell fold–thrust belt and the
Fig. 9. (a) Interpretive cross-section showing the setting of the Fiordland–Westland orogen during the period ,126–105 Ma (modified from Klepeis et al.,
2003). (b) Block diagram showing correlations of structures in Fiordland and Westland. Diagram was constructed using pre-Cenozoic configuration of the
orogen and the relative structural position of the Fiordland cross-sections shown in Figs. 1b, 3, 4 and 5 (localities listed on right side of diagram). Section
representative of Westland is from the north shore of Westland (separation point). Paleodepths were calculated from data shown in Fig. 1. Symbols (#)
represent garnet granulite facies fracture arrays.
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UNCORRECTED PROOF
Indecision Creek and George Sound Shear Zones all
evolved during the interval ,122.5–105 Ma.
The last group of structures includes the Anita Shear
Zone (ASZ, Figs. 5 and 9b; Table 2) and the Doubtful Sound
Shear Zone (Table 1). Both shear zones record the
exhumation (Fig. 9b) of the lower crust after ,108–
105 Ma (Mattinson et al., 1986; Gibson et al., 1988; Nathan
et al., 2000; Claypool et al., 2002). The Anita Shear Zone
also preserves fabrics (including ASZ1 at P ¼ 12 kbar and
ASZ2 at P ¼ 8 kbar, Fig. 10) that record Late Cretaceous–
Cenozoic decompression (Klepeis et al., 1999).
8. Interpretation of vertical coupling and decoupling
within the crustal column
The results of our analyses show that deformation
accompanying the ,126–120 Ma emplacement of the
WFO was localized within and at its upper and lower
contacts (Figs. 1b, 9b and 10c). In contrast, deformation a
few kilometers below the WFO during this period was weak
to nonexistent (Table 1). The two layer-parallel shear zones
at Mt. Daniel and Caswell Sound both separate areas of
melt-enhanced deformation inside the WFO (SWFO) from
areas outside it where older structures (S1 and igneous
layering) and mineral assemblages are well preserved. In
addition, structural relationships above and below the Mt.
Daniel Shear Zone at the base of the WFO are discordant.
This discordance and evidence that the structures above and
below the basal shear zone formed at different times and
under different physical conditions suggest that the crust
above and below the lower contact of the WFO was
decoupled during emplacement of the batholith.
As the batholith cooled and crystallized, the role of the
layer-parallel shear zones began to change. The develop-
ment of melt-induced fracture arrays and dehydration zones
(Fig. 10c) that cut the lower contact of the WFO mark the
abandonment of the melt-enhanced shear zone at Mt.
Fig. 10. Pressure–temperature–time path (a) for the lower crust exposed between Caswell and Milford sounds (modified from Daczko et al., 2002c). P–T data
are compiled from Bradshaw (1985, 1989a,b), Gibson and Ireland (1995), Klepeis et al. (1999), Clarke et al. (2000) and Daczko et al. (2001a,b, 2002a,b,c).
Shaded regions incorporate errors for specific metamorphic mineral assemblages discussed by Daczko et al. (2002c). Block diagrams ((b)–(e)) illustrate four
stages in the tectonic evolution of the Fiordland–Westland belt. Patterns are the same as those in Fig. 1a. Parts (c)–(e) are modified from Klepeis et al. (2003).
Abbreviations are: ASZ—Anita Shear Zone, ARC—Arthur River Complex, WFO—Western Fiordland Orthogneiss, ME—Mt. Edgar, SC—Selwyn Creek,
C—Caswell Sound, G—George Sound, MD—Mt. Daniel, DSSZ—Doubtful Sound Shear Zone, P—Paparoa Range, V—Victoria Range.
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1824
1825
1826
1827
1828
1829
1830
1831
1832
1833
1834
1835
1836
1837
1838
1839
1840
1841
1842
1843
1844
1845
1846
1847
1848
1849
1850
1851
1852
1853
1854
1855
1856
1857
1858
1859
1860
1861
1862
1863
1864
1865
1866
1867
1868
1869
1870
1871
1872
1873
1874
1875
1876
1877
1878
1879
1880
1881
1882
1883
1884
1885
1886
1887
1888
1889
1890
1891
1892
1893
1894
1895
1896
1897
1898
1899
1900
1901
1902
1903
1904
UNCORRECTED PROOF
Table 2
U–Pb isotopic data
Sample U
(ppm)
206Pb/204Pb
ratio
206Pb/238U ^ ratio
(%)
207Pb/235U ^ ratio
(%)
206Pb/207P ^ ratio
(%)
Error
correction
206Pb/238U ^ age
(Ma)
207Pb/235U ^ age
(Ma)
206Pb/207Pb ^ age
(Ma)
207Pb correction ^ age
(Ma)
28-1c 40 NA 0.0205 ^ 0.5 0.1308 ^ 5.8 21.579 ^ 5.8 0.09 130.6 ^ 0.7 124.8 ^ 6.8 15 ^ 70 131.0 ^ 0.8
28-2c 33 NA 0.0195 ^ 0.7 0.1491 ^ 4.7 18.022 ^ 4.6 0.16 124.4 ^ 0.9 141.1 ^ 6.1 432 ^ 51 123.3 ^ 1.0
28-3c 42 NA 0.0188 ^ 0.6 0.1113 ^ 6.1 23.238 ^ 6.1 0.09 119.8 ^ 0.7 107.1 ^ 6.2 2 166 ^ 76 120.6 ^ 0.8
28-4c 39 NA 0.0186 ^ 0.8 0.1282 ^ 4.5 19.962 ^ 4.4 0.17 118.6 ^ 0.9 122.5 ^ 5.2 199 ^ 51 118.3 ^ 1.0
28-5c 61 NA 0.0193 ^ 1.1 0.1293 ^ 3.9 20.555 ^ 3.7 0.29 123.1 ^ 1.3 123.4 ^ 4.5 131 ^ 43 123.0 ^ 1.4
28-6c 65 NA 0.0180 ^ 0.7 0.1324 ^ 5.0 18.723 ^ 5.0 0.14 114.9 ^ 0.8 126.3 ^ 5.9 346 ^ 56 114.2 ^ 0.9
28-7c 88 NA 0.0194 ^ 0.5 0.1615 ^ 1.6 16.558 ^ 1.5 0.30 123.8 ^ 0.6 152.0 ^ 2.2 618 ^ 16 122.0 ^ 0.6
28-8c 99 NA 0.0190 ^ 0.4 0.1491 ^ 1.6 17.582 ^ 1.5 0.26 121.4 ^ 0.5 141.1 ^ 2.1 487 ^ 17 120.1 ^ 0.5
28-9c 134 NA 0.0188 ^ 0.4 0.1415 ^ 1.4 18.321 ^ 1.3 0.31 120.1 ^ 0.5 134.4 ^ 1.8 395 ^ 15 119.1 ^ 0.5
28-10c 141 NA 0.0184 ^ 0.6 0.1387 ^ 1.3 18.293 ^ 1.2 0.44 117.6 ^ 0.7 131.9 ^ 1.6 399 ^ 13 116.7 ^ 0.7
28-11c 207 NA 0.0180 ^ 0.7 0.1252 ^ 1.6 19.831 ^ 1.4 0.44 115.1 ^ 0.8 119.8 ^ 1.8 215 ^ 16 114.8 ^ 0.8
28-12c 127 NA 0.0191 ^ 0.4 0.1546 ^ 1.7 17.014 ^ 1.6 0.26 121.9 ^ 0.5 146.0 ^ 2.3 559 ^ 18 120.3 ^ 0.5
28-13c 144 NA 0.0188 ^ 0.4 0.1343 ^ 1.6 19.262 ^ 1.6 0.25 119.8 ^ 0.5 127.9 ^ 1.9 282 ^ 18 119.3 ^ 0.5
28-14c 77 NA 0.0209 ^ 1.2 0.1565 ^ 3.0 18.402 ^ 2.8 0.40 133.2 ^ 1.6 147.6 ^ 4.1 385 ^ 31 132.3 ^ 1.6
28-15c 155 NA 0.0180 ^ 0.4 0.1224 ^ 2.4 20.320 ^ 2.4 0.17 115.3 ^ 0.5 117.3 ^ 2.7 158 ^ 28 115.2 ^ 0.5
28-16c 47 NA 0.0176 ^ 1.2 0.0924 ^ 6.0 26.258 ^ 5.9 0.21 112.4 ^ 1.4 89.7 ^ 5.1 2 479 ^ 77 113.8 ^ 1.4
28-17c 80 NA 0.0182 ^ 0.7 0.1121 ^ 3.4 22.338 ^ 3.4 0.20 116.0 ^ 0.8 107.9 ^ 3.5 2 68 ^ 41 116.5 ^ 0.8
28-18c 57 NA 0.0192 ^ 0.9 0.1467 ^ 3.6 18.025 ^ 3.5 0.25 122.4 ^ 1.1 139.0 ^ 4.7 432 ^ 39 121.4 ^ 1.1
28-19c 24 NA 0.0185 ^ 0.6 0.1087 ^ 2.0 23.464 ^ 1.9 0.32 118.1 ^ 0.7 104.8 ^ 2.0 2 190 ^ 24 119.0 ^ 0.8
28-20c 40 NA 0.0186 ^ 0.5 0.1159 ^ 7.5 22.120 ^ 7.4 0.07 118.8 ^ 0.6 111.4 ^ 7.8 2 44 ^ 90 119.3 ^ 0.8
28-21c 36 NA 0.0187 ^ 0.5 0.1155 ^ 8.1 22.364 ^ 8.1 0.06 119.7 ^ 0.6 111.0 ^ 8.5 2 71 ^ 98 120.2 ^ 0.8
28-22c 68 NA 0.0184 ^ 0.4 0.1414 ^ 3.8 17.914 ^ 3.8 0.11 117.3 ^ 0.5 134.2 ^ 4.7 445 ^ 42 116.2 ^ 0.6
5a1c 24 NA 0.0190 ^ 1.1 0.0940 ^ 13.8 27.826 ^ 13.7 0.08 121.2 ^ 1.3 91.2 ^ 12.0 2 636 ^ 188 123.1 ^ 1.5
5a2c 24 NA 0.0185 ^ 1.0 0.1592 ^ 6.9 16.040 ^ 6.8 0.15 118.3 ^ 1.2 150.0 ^ 9.5 686 ^ 73 116.2 ^ 1.3
5a3c 23 NA 0.0248 ^ 0.5 0.1744 ^ 14.4 19.614 ^ 14.4 0.03 158.0 ^ 0.7 163.3 ^ 21.5 240 ^ 166 157.7 ^ 1.6
5a4c 31 NA 0.0187 ^ 0.8 0.1084 ^ 7.6 23.806 ^ 7.6 0.11 119.5 ^ 1.0 104.5 ^ 7.5 2 226 ^ 95 120.5 ^ 1.1
5a5c 27 NA 0.0180 ^ 0.9 0.0856 ^ 17.8 28.890 ^ 17.8 0.05 114.7 ^ 1.1 83.4 ^ 14.2 2 740 ^ 249 116.6 ^ 1.4
5a6c 27 NA 0.0176 ^ 1.0 0.0858 ^ 11.4 28.296 ^ 11.3 0.08 112.5 ^ 1.1 83.6 ^ 9.1 2 682 ^ 156 114.4 ^ 1.2
5a7c 28 NA 0.0179 ^ 0.6 0.1245 ^ 7.9 19.820 ^ 7.9 0.08 114.3 ^ 0.7 119.1 ^ 8.9 216 ^ 92 114.0 ^ 0.9
5a8c 31 NA 0.0174 ^ 1.0 0.1216 ^ 6.8 19.771 ^ 6.7 0.14 111.4 ^ 1.1 116.5 ^ 7.4 222 ^ 77 111.1 ^ 1.2
5a9c 39 NA 0.0178 ^ 1.0 0.1153 ^ 7.3 21.276 ^ 7.2 0.14 113.7 ^ 1.1 110.8 ^ 7.6 49 ^ 86 113.9 ^ 1.2
5a10c 27 NA 0.0182 ^ 0.9 0.1258 ^ 10.7 19.889 ^ 10.6 0.09 115.9 ^ 1.0 120.3 ^ 12.0 208 ^ 123 115.7 ^ 1.3
5a11c 35 NA 0.0180 ^ 0.5 0.0627 ^ 13.6 39.460 ^ 13.6 0.03 114.7 ^ 0.5 61.8 ^ 8.1 2 1711 ^ 237 118.0 ^ 0.7
5a12c 39 NA 0.0181 ^ 0.6 0.0992 ^ 12.9 25.174 ^ 12.9 0.04 115.7 ^ 0.7 96.0 ^ 11.8 2 369 ^ 168 116.9 ^ 1.0
5a13c 22 NA 0.0180 ^ 0.8 0.0975 ^ 9.5 25.394 ^ 9.4 0.08 114.8 ^ 0.9 94.5 ^ 8.5 2 391 ^ 123 116.0 ^ 1.1
5a14c 48 NA 0.0178 ^ 0.8 0.1025 ^ 5.4 23.944 ^ 5.4 0.15 113.8 ^ 0.9 99.1 ^ 5.1 2 241 ^ 68 114.7 ^ 1.0
5a15c 24 NA 0.0186 ^ 0.8 0.1403 ^ 6.8 18.239 ^ 6.8 0.11 118.5 ^ 0.9 133.3 ^ 8.5 405 ^ 76 117.6 ^ 1.0
5a16c 42 NA 0.0176 ^ 0.9 0.0996 ^ 6.6 24.380 ^ 6.6 0.13 112.6 ^ 1.0 96.4 ^ 6.1 2 287 ^ 84 113.6 ^ 1.0
5a17c 20 NA 0.0186 ^ 0.9 0.1028 ^ 12.1 24.929 ^ 12.0 0.07 118.7 ^ 1.0 99.4 ^ 11.4 2 344 ^ 155 120.0 ^ 1.3
5a18c 32 NA 0.0185 ^ 0.9 0.1570 ^ 9.4 16.198 ^ 9.4 0.09 117.8 ^ 1.0 148.1 ^ 12.9 665 ^ 101 115.9 ^ 1.3
5a19c 36 NA 0.0194 ^ 1.1 0.1317 ^ 9.3 20.285 ^ 9.2 0.11 123.8 ^ 1.3 125.7 ^ 10.9 162 ^ 108 123.6 ^ 1.5
5a20c 25 NA 0.0188 ^ 0.9 0.1471 ^ 8.4 17.584 ^ 8.4 0.11 119.8 ^ 1.1 139.3 ^ 10.9 486 ^ 92 118.5 ^ 1.3
221K-1t 267 18051 0.0206 ^ 0.7 0.1309 ^ 2.8 21.669 ^ 2.7 0.25 131.3 ^ 0.9 124.9 ^ 3.7 5 ^ 33 NA
221K-2t 351 29570 0.0200 ^ 0.5 0.1267 ^ 1.9 21.757 ^ 1.8 0.26 127.6 ^ 0.6 121.2 ^ 2.4 2 4 ^ 22 NA
SG
14
56
—3
0/9
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SF
OR
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—8
24
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br,ed
,cor
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.K
lepeis
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yxx
(00
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xxx1
8
AR
TIC
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SS
19
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19
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19
07
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19
10
19
11
19
12
19
13
19
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19
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19
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19
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19
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19
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19
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19
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19
42
19
43
19
44
19
45
19
46
19
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19
48
19
49
19
50
19
51
19
52
19
53
19
54
19
55
19
56
19
57
19
58
19
59
19
60
19
61
19
62
19
63
19
64
19
65
19
66
19
67
19
68
19
69
19
70
19
71
19
72
19
73
19
74
19
75
19
76
19
77
19
78
19
79
19
80
19
81
19
82
19
83
19
84
19
85
19
86
19
87
19
88
19
89
19
90
19
91
19
92
19
93
19
94
19
95
19
96
19
97
19
98
19
99
20
00
20
01
20
02
20
03
20
04
20
05
20
06
20
07
20
08
20
09
20
10
20
11
20
12
20
13
20
14
20
15
20
16
UNCORRECTED PROOF
Table 2 (continued)
Sample U
(ppm)
206Pb/204Pb
ratio
206Pb/238U ^ ratio
(%)
207Pb/235U ^ ratio
(%)
206Pb/207P ^ ratio
(%)
Error
correction
206Pb/238U ^ age
(Ma)
207Pb/235U ^ age
(Ma)
206Pb/207Pb ^ age
(Ma)
207Pb correction ^ age
(Ma)
221K-3t 428 36509 0.0198 ^ 0.6 0.1297 ^ 1.6 21.040 ^ 1.4 0.41 126.3 ^ 0.8 123.8 ^ 2.0 76 ^ 17 NA
221K-4c 262 43226 0.0364 ^ 1.4 0.2498 ^ 1.9 20.096 ^ 1.4 0.71 230.5 ^ 3.2 226.4 ^ 4.9 184 ^ 16 NA
221K-5c 40 4497 0.0577 ^ 0.9 0.4339 ^ 4.0 18.347 ^ 3.9 0.23 361.8 ^ 3.3 365.9 ^ 17.3 392 ^ 43 NA
221K-6c 44 15551 0.0560 ^ 1.0 0.4324 ^ 4.7 17.856 ^ 4.6 0.22 351.2 ^ 3.7 364.8 ^ 20.4 453 ^ 51 NA
221K-7c 15 7790 0.0569 ^ 1.2 0.3726 ^ 16.6 21.039 ^ 16.6 0.07 356.4 ^ 4.4 321.6 ^ 60.9 76 ^ 197 NA
221K-8c 90 38799 0.0580 ^ 0.8 0.4120 ^ 3.5 19.416 ^ 3.4 0.23 363.6 ^ 3.0 350.3 ^ 14.5 263 ^ 39 NA
221K-9t 118 12730 0.0202 ^ 1.5 0.1271 ^ 6.8 21.893 ^ 6.6 0.23 128.8 ^ 2.0 121.5 ^ 8.7 2 19 ^ 80 NA
221K-10t 56 6095 0.0261 ^ 1.3 0.2049 ^ 9.5 17.538 ^ 9.5 0.14 165.9 ^ 2.2 189.3 ^ 19.6 492 ^ 104 NA
221K-11c 94 18096 0.0502 ^ 1.2 0.3895 ^ 7.1 17.782 ^ 7.0 0.17 315.9 ^ 3.9 334.0 ^ 27.7 462 ^ 77 NA
221K-12c 101 16611 0.0570 ^ 1.0 0.4402 ^ 2.0 17.864 ^ 1.8 0.49 357.5 ^ 3.6 370.4 ^ 9.0 452 ^ 19 NA
221K-13c 99 26804 0.0583 ^ 1.0 0.4178 ^ 3.1 19.225 ^ 3.0 0.32 365.0 ^ 3.8 354.5 ^ 13.1 286 ^ 34 NA
221K-14c 74 22844 0.0575 ^ 1.1 0.4156 ^ 2.2 19.089 ^ 1.9 0.51 360.6 ^ 4.2 352.9 ^ 9.3 302 ^ 22 NA
221K-15t 67 11424 0.0205 ^ 1.8 0.1253 ^ 11.0 22.504 ^ 10.8 0.16 130.5 ^ 2.4 119.8 ^ 13.9 2 86 ^ 133 NA
221K-16c 152 29614 0.0513 ^ 2.9 0.3736 ^ 4.9 18.927 ^ 4.0 0.58 322.4 ^ 9.4 322.3 ^ 18.4 322 ^ 45 NA
221K-17c 95 23260 0.0496 ^ 1.1 0.3565 ^ 3.6 19.191 ^ 3.4 0.31 312.2 ^ 3.5 309.6 ^ 12.8 290 ^ 39 NA
221K-18c 100 18092 0.0535 ^ 3.2 0.3782 ^ 7.7 19.504 ^ 7.0 0.42 336.0 ^ 11.1 325.7 ^ 29.1 253 ^ 80 NA
221K-19c 491 23329 0.0574 ^ 0.9 0.4329 ^ 2.3 18.274 ^ 2.1 0.40 360.1 ^ 3.4 365.8 ^ 10.3 401 ^ 24.0 NA
221K-20c 196 9298 0.0567 ^ 0.6 0.4178 ^ 2.6 18.723 ^ 2.5 0.22 356.3 ^ 2.0 354.9 ^ 11.7 346 ^ 28.0 NA
221K-21c 121 4240 0.0422 ^ 0.6 0.2875 ^ 3.1 20.220 ^ 3.0 0.19 266.2 ^ 1.5 257.2 ^ 9.2 169 ^ 35.0 NA
221K-22c 827 8487 0.0190 ^ 3.0 0.1263 ^ 3.7 20.737 ^ 2.1 0.82 121.8 ^ 3.7 121.4 ^ 5.2 110 ^ 25.0 NA
221K-23t 964 8645 0.0206 ^ 0.3 0.1402 ^ 2.1 20.257 ^ 2.1 0.16 131.5 ^ 0.4 133.7 ^ 3.6 165 ^ 24.0 NA
221K-24c 370 19432 0.0501 ^ 1.0 0.3744 ^ 2.5 18.451 ^ 2.3 0.40 315.3 ^ 3.1 323.2 ^ 2.7 379 ^ 25.0 NA
221K-25c 548 33147 0.0594 ^ 0.9 0.4508 ^ 2.2 18.153 ^ 2.0 0.41 372.1 ^ 3.5 378.3 ^ 10.3 416 ^ 23.0 NA
221K-26c 882 4216 0.0862 ^ 1.7 3.1545 ^ 4.6 3.770 ^ 4.2 0.38 533.8 ^ 9.6 1446.0 ^ 136.8 3278 ^ 33.0 NA
221K-27t 166 1708 0.0591 ^ 0.9 0.6847 ^ 5.1 11.890 ^ 5.0 0.18 370.0 ^ 3.6 530.6 ^ 35.1 1295 ^ 49.0 NA
221K-28t 480 22242 0.0448 ^ 1.6 0.3304 ^ 2.6 18.677 ^ 2.1 0.59 282.2 ^ 4.5 290.1 ^ 9.1 352 ^ 24.0 NA
221K-29c 1150 30148 0.0480 ^ 2.2 0.3685 ^ 3.0 17.952 ^ 2.0 0.74 302.2 ^ 6.8 319.8 ^ 11.9 441 ^ 22.0 NA
221K-30c 829 29551 0.0594 ^ 0.5 0.4506 ^ 2.1 18.183 ^ 2.0 0.22 372.4 ^ 1.8 378.4 ^ 9.0 412 ^ 22.0 NA
221K-31c 108 7676 0.0589 ^ 0.9 0.4254 ^ 2.6 19.100 ^ 2.5 0.33 369.0 ^ 3.3 360.5 ^ 11.1 301 ^ 28.0 NA
221K-32c 179 10102 0.0580 ^ 0.4 0.4291 ^ 2.4 18.619 ^ 2.4 0.16 363.5 ^ 1.4 363.1 ^ 10.7 359 ^ 27.0 NA
221K-33c 819 33551 0.0596 ^ 0.8 0.4537 ^ 2.2 18.100 ^ 2.0 0.35 373.1 ^ 2.9 380.8 ^ 10.4 422 ^ 23.0 NA
221K-34t 143 13634 0.0184 ^ 4.5 0.1306 ^ 11.9 19.389 ^ 11.0 0.38 117.7 ^ 5.3 125.9 ^ 16.3 267 ^ 126.0 NA
221K-35t 68 2536 0.0560 ^ 0.5 0.4308 ^ 4.6 17.934 ^ 4.6 0.10 351.8 ^ 1.7 364.2 ^ 20.2 443 ^ 51.0 NA
221K-36c 136 4135 0.0574 ^ 0.7 0.4827 ^ 2.8 16.388 ^ 2.7 0.24 360.2 ^ 2.5 400.0 ^ 14.8 640 ^ 29.0 NA
221K-37t 278 24793 0.0561 ^ 0.7 0.4215 ^ 2.3 18.361 ^ 2.2 0.32 352.4 ^ 2.6 357.4 ^ 10.1 390 ^ 24.0 NA
221K-38t 234 8850 0.0233 ^ 0.9 0.1544 ^ 3.6 20.767 ^ 3.5 0.25 148.9 ^ 1.3 146.6 ^ 6.2 107 ^ 41.0 NA
221K-39c 265 22663 0.0606 ^ 0.3 0.4561 ^ 2.2 18.310 ^ 2.1 0.16 379.0 ^ 1.3 382.3 ^ 10.7 396 ^ 24.0 NA
221K-40t 155 4240 0.0537 ^ 1.1 0.3853 ^ 3.0 19.218 ^ 2.8 0.36 337.3 ^ 3.8 331.7 ^ 12.6 287 ^ 32.0 NA
221K-41t 51 16902 0.0586 ^ 1.1 0.4362 ^ 3.7 18.535 ^ 3.6 0.30 367.1 ^ 4.2 368.1 ^ 16.6 369 ^ 40.0 NA
221K-42t 49 4837 0.0584 ^ 1.1 0.4686 ^ 5.4 17.185 ^ 5.3 0.21 366.0 ^ 4.2 0.3 ^ 26.3 537 ^ 58.0 NA
221K-43t 42 2299 0.0595 ^ 1.1 0.4142 ^ 5.2 19.809 ^ 5.1 0.21 373.3 ^ 4.3 352.7 ^ 22.0 217 ^ 59.0 NA
221K-44t 192 6125 0.0203 ^ 2.3 0.1206 ^ 4.0 23.251 ^ 3.3 0.57 130.6 ^ 3.0 116.5 ^ 5.5 2 167 ^ 41.0 NA
221K-45t 187 45340 0.0576 ^ 0.5 0.4275 ^ 2.3 18.578 ^ 2.3 0.23 361.3 ^ 2.0 361.0 ^ 10.1 364 ^ 26.0 NA
221K-46t 35 4453 0.0588 ^ 1.7 0.4522 ^ 8.9 17.926 ^ 8.8 0.19 368.5 ^ 6.3 379.2 ^ 40.3 444 ^ 98.0 NA
221K-47c 59 10614 0.0573 ^ 0.9 0.4391 ^ 4.5 17.975 ^ 4.5 0.20 359.1 ^ 3.0 370.5 ^ 20.6 438 ^ 49.0 NA
(continued on next page)
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lepeis
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tructu
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yxx
(00
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21
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21
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UNCORRECTED PROOF
Table 2 (continued)
Sample U
(ppm)
206Pb/204Pb
ratio
206Pb/238U ^ ratio
(%)
207Pb/235U ^ ratio
(%)
206Pb/207P ^ ratio
(%)
Error
correction
206Pb/238U ^ age
(Ma)
207Pb/235U ^ age
(Ma)
206Pb/207Pb ^ age
(Ma)
207Pb correction ^ age
(Ma)
221K-48c 67 9742 0.0580 ^ 1.0 0.4275 ^ 4.6 18.699 ^ 4.4 0.23 363.8 ^ 4.0 361.1 ^ 20.4 349 ^ 50.0 NA
221K-49c 90 18019 0.0591 ^ 1.0 0.4423 ^ 4.0 18.413 ^ 3.9 0.25 370.4 ^ 4.0 372.8 ^ 18.9 384 ^ 44.0 NA
221K-50c 26 2991 0.0601 ^ 1.1 0.4130 ^ 4.4 20.052 ^ 4.3 0.24 376.4 ^ 4.0 351.1 ^ 18.0 189 ^ 50.0 NA
ADA2-1c 314 16959 0.0323 ^ 0.9 0.2236 ^ 1.6 19.907 ^ 1.3 0.54 204.8 ^ 1.8 204.9 ^ 3.6 206 ^ 15 NA
ADA2-2c 729 3605 0.0319 ^ 0.5 0.2051 ^ 1.0 21.437 ^ 0.9 0.49 202.3 ^ 1.0 189.4 ^ 2.0 31 ^ 10 NA
ADA2-3c 125 6567 0.0315 ^ 1.0 0.2113 ^ 4.2 20.563 ^ 4.0 0.25 200.0 ^ 2.1 194.6 ^ 8.9 130 ^ 47 NA
ADA2-4c 128 4641 0.0320 ^ 0.4 0.1987 ^ 3.2 22.200 ^ 3.2 0.12 203.0 ^ 0.8 184.0 ^ 6.5 2 53 ^ 39 NA
ADA2-5c 53 17623 0.0302 ^ 0.6 0.1870 ^ 7.4 22.282 ^ 7.4 0.07 191.9 ^ 1.1 174.0 ^ 14.0 2 62 ^ 90 NA
ADA2-6c 70 2532 0.0319 ^ 0.6 0.2188 ^ 4.8 20.080 ^ 4.7 0.13 202.2 ^ 1.3 200.9 ^ 10.6 186 ^ 55 NA
ADA2-7c 30 15567 0.0323 ^ 1.5 0.2033 ^ 16.4 21.890 ^ 16.3 0.09 204.8 ^ 3.1 187.9 ^ 33.3 2 19 ^ 197 NA
ADA2-8c 81 2624 0.0318 ^ 0.6 0.1967 ^ 5.6 22.305 ^ 5.6 0.11 201.9 ^ 1.3 182.3 ^ 11.1 2 65 ^ 68 NA
ADA2-9c 135 212500 0.0316 ^ 2.0 0.2363 ^ 3.8 18.429 ^ 3.2 0.53 200.4 ^ 4.1 215.4 ^ 9.0 382 ^ 36 NA
ADA2-10c 59 176999 0.0317 ^ 0.9 0.2102 ^ 12.0 20.796 ^ 11.9 0.08 201.2 ^ 1.9 193.7 ^ 25.2 103 ^ 141 NA
ADA2-11c 54 6500 0.0271 ^ 1.1 0.1906 ^ 11.4 19.602 ^ 11.3 0.10 172.3 ^ 2.0 177.1 ^ 21.8 242 ^ 131 NA
ADA2-12c 94 3222 0.0322 ^ 0.5 0.2014 ^ 6.7 22.062 ^ 6.7 0.07 204.4 ^ 1.0 186.3 ^ 13.7 2 38 ^ 82 NA
ADA2-13c 48 1570 0.0258 ^ 0.7 0.1465 ^ 14.3 24.243 ^ 14.3 0.05 164.0 ^ 1.1 138.8 ^ 21.0 2 272 ^ 181 NA
ADA2-14c 43 13125 0.0236 ^ 1.4 0.1503 ^ 18.9 21.688 ^ 18.9 0.07 150.6 ^ 2.1 142.2 ^ 28.4 3 ^ 227 NA
ADA2-15c 162 32272 0.0325 ^ 1.0 0.2277 ^ 2.9 19.675 ^ 2.7 0.34 206.2 ^ 2.0 208.3 ^ 6.6 233 ^ 31 NA
ADA2-16c 132 12743 0.0327 ^ 0.6 0.2176 ^ 3.7 20.728 ^ 3.6 0.16 207.5 ^ 1.2 199.9 ^ 8.1 111 ^ 43 NA
ADA2-17c 90 18709 0.0328 ^ 0.5 0.2286 ^ 5.4 19.775 ^ 5.4 0.09 208.0 ^ 1.1 209.0 ^ 12.5 221 ^ 62 NA
ADA2-18c 125 3783 0.0322 ^ 0.8 0.2325 ^ 3.9 19.085 ^ 3.8 0.20 204.2 ^ 1.6 212.2 ^ 9.1 303 ^ 43 NA
ADA2-19c 63 3726 0.0325 ^ 1.1 0.2299 ^ 6.5 19.478 ^ 6.4 0.18 206.0 ^ 2.4 210.1 ^ 14.9 256 ^ 73 NA
ADA2-20c 38 13947 0.0279 ^ 1.1 0.1673 ^ 13.1 22.960 ^ 13.1 0.08 177.1 ^ 2.0 157.0 ^ 22.1 2 136 ^ 162 NA
ADA2-21c 44 8748 0.0240 ^ 0.7 0.1499 ^ 12.3 22.026 ^ 12.3 0.06 152.6 ^ 1.1 141.9 ^ 18.5 2 34 ^ 149 NA
ADA2-22c 99 26510 0.0332 ^ 0.6 0.2388 ^ 4.4 19.175 ^ 4.4 0.14 210.6 ^ 1.3 217.5 ^ 10.7 292 ^ 50 NA
ADA2-23c 162 37708 0.0316 ^ 0.6 0.2342 ^ 2.6 18.621 ^ 2.6 0.23 200.8 ^ 1.2 213.7 ^ 6.3 359 ^ 29 NA
ADA2-24c 330 48018 0.0329 ^ 0.8 0.2430 ^ 1.6 18.657 ^ 1.4 0.48 208.5 ^ 1.6 220.8 ^ 4.0 354 ^ 16 NA
ADA2-25c 39 7083 0.0328 ^ 1.7 0.2370 ^ 12.2 19.084 ^ 12.1 0.14 208.1 ^ 3.5 216.0 ^ 28.9 303 ^ 137 NA
ADA2-26t 145 2506 0.0177 ^ 6.2 0.1220 ^ 9.3 19.989 ^ 7.0 0.66 113.0 ^ 7.0 116.8 ^ 11.5 196 ^ 82 NA
ADA2-27t 43 1987 0.0252 ^ 1.3 0.1605 ^ 14.2 21.628 ^ 14.1 0.09 160.2 ^ 2.0 151.1 ^ 22.8 10 ^ 170 NA
ADA2-28t 44 5893 0.0187 ^ 1.0 0.1249 ^ 4.4 20.597 ^ 4.3 0.22 119.1 ^ 1.2 119.5 ^ 5.6 126 ^ 51 NA
ADA2-29t 107 4837 0.0196 ^ 0.5 0.1382 ^ 1.1 19.529 ^ 1.0 0.42 124.9 ^ 0.6 131.4 ^ 1.6 250 ^ 12 NA
ADA2-30t 77 1348 0.0180 ^ 1.1 0.1040 ^ 2.3 23.825 ^ 2.1 0.45 114.8 ^ 1.2 100.4 ^ 2.5 2 228 ^ 26 NA
ADA2-31t 75 20322 0.0175 ^ 1.1 0.1416 ^ 6.4 17.042 ^ 6.3 0.17 111.8 ^ 1.2 134.4 ^ 9.1 555 ^ 69 NA
ADA2-32t 205 1111 0.0175 ^ 0.9 0.0982 ^ 1.9 24.544 ^ 1.6 0.49 111.7 ^ 1.0 95.1 ^ 1.9 2 304 ^ 21 NA
ADA2-33t 107 1503 0.0190 ^ 1.3 0.1139 ^ 1.7 23.000 ^ 1.2 0.74 121.4 ^ 1.6 109.6 ^ 2.0 2 140 ^ 14 NA
ADA2-34t 20 33533 0.0196 ^ 0.7 0.1456 ^ 2.2 18.590 ^ 2.1 0.32 125.3 ^ 0.9 138.0 ^ 3.2 362 ^ 23 NA
ADA2-35t 79 7649 0.0187 ^ 4.9 0.1440 ^ 13.3 17.874 ^ 12.3 0.37 119.2 ^ 5.9 136.6 ^ 19.2 450 ^ 137 NA
ADA2-36t 363 4293 0.0181 ^ 1.2 0.1192 ^ 6.7 20.934 ^ 6.6 0.18 115.6 ^ 1.4 114.3 ^ 8.1 88 ^ 78 NA
ADA2-37t 27 1387 0.0188 ^ 2.4 0.1489 ^ 6.6 17.411 ^ 6.1 0.37 120.1 ^ 2.9 140.9 ^ 9.9 508 ^ 67 NA
ADA2-38t 20 1362 0.0183 ^ 3.7 0.1107 ^ 7.1 22.817 ^ 6.0 0.53 117.0 ^ 4.4 106.6 ^ 7.9 2 120 ^ 74 NA
ADA2-39t 37 3061 0.0182 ^ 2.0 0.1087 ^ 6.3 23.083 ^ 6.0 0.32 116.3 ^ 2.4 104.8 ^ 7.0 2 149 ^ 74 NA
ADA2-40t 143 18329 0.0180 ^ 3.7 0.1280 ^ 11.8 19.412 ^ 11.2 0.31 115.2 ^ 4.3 122.3 ^ 15.2 264 ^ 129 NA
ADA2-41t 644 15254 0.0184 ^ 0.9 0.0870 ^ 3.6 29.202 ^ 3.5 0.24 117.7 ^ 1.0 84.7 ^ 3.2 2 770 ^ 49 NA
ADA2-42t 30 5509 0.0189 ^ 0.8 0.1384 ^ 1.6 18.837 ^ 1.4 0.50 120.8 ^ 1.0 131.6 ^ 2.3 333 ^ 16 NA
SG
14
56
—3
0/9
/20
03
—1
4:5
8—
SF
OR
ST
ER
—8
24
96
—M
OD
EL
5–
br,ed
,cor
K.A
.K
lepeis
eta
l./
Jou
rna
lo
fS
tructu
ral
Geo
log
yxx
(00
00
)xxx–
xxx2
0
AR
TIC
LE
INP
RE
SS
21
29
21
30
21
31
21
32
21
33
21
34
21
35
21
36
21
37
21
38
21
39
21
40
21
41
21
42
21
43
21
44
21
45
21
46
21
47
21
48
21
49
21
50
21
51
21
52
21
53
21
54
21
55
21
56
21
57
21
58
21
59
21
60
21
61
21
62
21
63
21
64
21
65
21
66
21
67
21
68
21
69
21
70
21
71
21
72
21
73
21
74
21
75
21
76
21
77
21
78
21
79
21
80
21
81
21
82
21
83
21
84
21
85
21
86
21
87
21
88
21
89
21
90
21
91
21
92
21
93
21
94
21
95
21
96
21
97
21
98
21
99
22
00
22
01
22
02
22
03
22
04
22
05
22
06
22
07
22
08
22
09
22
10
22
11
22
12
22
13
22
14
22
15
22
16
22
17
22
18
22
19
22
20
22
21
22
22
22
23
22
24
22
25
22
26
22
27
22
28
22
29
22
30
22
31
22
32
22
33
22
34
22
35
22
36
22
37
22
38
22
39
22
40
UNCORRECTED PROOF
Daniel. In addition, the steep upper amphibolite facies
foliations (SSZ) of the Indecision Creek Shear Zone cut
across the lower (and eastern) contact of the WFO east of
Mt. Daniel (Figs. 9b and 10d). These relationships indicate
that once the WFO crystallized deformation was no longer
concentrated in a narrow melt-enhanced zone at the base of
the batholith. By ,116 Ma (and possibly earlier) this zone
had been replaced by the steep shear zones up to 15 km
wide. This represents a progressive widening of the zone of
deformation below the batholith to include the entire lower
crustal section between Milford and Caswell Sounds.
Following dehydration of the lower crust and abandon-
ment of the Mt. Daniel Shear Zone physical and kinematic
links developed above and below the WFO. The Caswell
fold–thrust belt continued to evolve in the middle crust
following crystallization of the batholith and records arc-
normal contraction and vertical thickening with a sinistral
component of arc-parallel displacement. The George Sound
and Indecision Creek Shear Zones also record subhorizontal
arc-normal contraction and vertical (layer-perpendicular)
thickening with components of sinistral arc-parallel dis-
placement. The simultaneous or near simultaneous devel-
opment of these contractional structures confirms indicate
that the middle and lower crust were coupled kinematically
above and below the batholith.
The style of contractional deformation affecting the
upper crust prior to ,105 Ma resembles the style of
deformation in the Caswell Thrust Belt (Fig. 9b). Narrow
thrust zones and steeply dipping foliations with down-dip
lineations were localized within the contact aureoles of
plutons on the western side of the Median Batholith in
Westland and on the eastern side east of Fiordland
(Bradshaw, 1989a; Tulloch and Challis, 2000). On Stewart
Island, dipping thrust faults formed inside the batholith
during ,125–105 Ma magmatism (Allibone and Tulloch,
1997; Tulloch and Kimbrough, 2003). These structural
patterns where deformation was concentrated in a narrow
(50–75 km wide) zone inside and on both sides of the
Median Batholith are remarkably similar to those preserved
in Fiordland where thrusts deform both pluton and country
rock in the contact aureoles of the WFO (Figs. 9a and 10d).
These relationships provide evidence that the upper, middle
and lower crusts were coupled kinematically during ,116–
105 Ma contraction as plutons were emplaced into the upper
crust.
9. Discussion: controls on vertical coupling and
decoupling in the crust
9.1. Changing strength and rheology of the lower crust
Structural and metamorphic relationships within north-
ern Fiordland suggest that decoupling of the middle and
lower crust above and below the WFO occurred as large
volumes of magma (.10 km thickness) were emplaced intoTab
le2
(co
nti
nued
)
Sam
ple
U (pp
m)
206P
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(%)
207P
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(%)
206P
b/2
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rati
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(%)
Err
or
corr
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on
206P
b/2
38U^
age
(Ma)
207P
b/2
35U^
age
(Ma)
206P
b/2
07P
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age
(Ma)
207P
bco
rrec
tion^
age
(Ma)
AD
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-43
t2
92
47
142
0.0
17
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0.7
0.1
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2.3
19
.51
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2.2
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11
13
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0.8
12
0.1
^2
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25
NA
AD
A2
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t2
21
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947
0.0
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1.6
0.1
19
8^
5.1
20
.73
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4.9
0.3
01
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1.8
11
4.9
^6
.21
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58
NA
AD
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t5
26
59
50
.01
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1.2
0.1
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2.8
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2.5
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21
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1.3
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4.9
^3
.73
19^
28
NA
Anal
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of
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info
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.U
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and
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Pb
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ake
rad
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ic206P
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and
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d206P
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SG 1456—30/9/2003—14:58—SFORSTER—82496— MODEL 5 – br,ed,cor
K.A. Klepeis et al. / Journal of Structural Geology xx (0000) xxx–xxx 21
ARTICLE IN PRESS
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UNCORRECTED PROOF
and moved through a lower crust weakened by melt and
heat. Two of the dominant processes controlling decoupling
during this period were the rates of emplacement and
crystallization of the WFO and the duration of the thermal
pulse that accompanied the magmatism. Rapid rates of
emplacement and cooling are implied by available geo-
chronologic and metamorphic data (Fig. 10a). Crystal-
lization ages (Tulloch and Kimbrough, 2003; Hollis et al.,
2004) suggest that most of the WFO was emplaced into the
lower crust within the period ,126–120 Ma. By ,120 Ma,
garnet granulite facies metamorphism at temperatures of
750 8C , T , 850 8C affected large regions of the batholith
and its lower crustal host (Clarke et al., 2000; Tulloch et al.,
2000; Hollis et al., 2003). By ,116 Ma kyanite- and
paragonite-bearing assemblages replaced the older garnet-,
clinopyroxene-bearing granulite facies assemblages indi-
cating that the lower crust below the batholith had cooled to
temperatures of T ¼ 650–700 8C (Daczko et al., 2002c).
The age of ,116 Ma is an average of the metamorphic rims
dates obtained from the Caswell Sound fold–thrust belt and
the Indecision Creek Shear Zone (samples 995a and Ada2,
Fig. 8).
Simple one-dimensional models of conductive heat loss
using standard solutions to the heat flow equation (Spear,
1993, p. 43) support the interpretation of rapid cooling of the
WFO by ,116 Ma. Magma compositions and the liquidus
temperature of basalt suggest that initial intrusions tem-
peratures were close to T ¼ 1200 8C following the estimates
of Petford and Gallagher (2001). Advective removal of heat
via fluid and melt transfer and the thrusting of pre-existing
arc crust above and below the batholith during emplacement
increases the rate of cooling. Burial and crustal thickening
tend to slow the rate of cooling. However, metamorphic and
geochronologic data from northern Fiordland indicate that
most of the burial of the Arthur River Complex occurred
prior to ,120 Ma granulite facies metamorphism (Fig.
10c). Calculations using metamorphic data to constrain
ambient temperatures at ,120 Ma (T ¼ 750–800 8C in
crust below the batholith and T ¼ 550–600 8C in crust
above it) suggest that the thermal pulse accompanying WFO
emplacement decayed to temperatures of T ¼ 650–700 8C
within ,3–4 million years. These results agree well with
the mineral reactions that record cooling of the lower crust
below the batholith (Daczko et al., 2002c). The well-known
dependence of lower crustal strength and rheology on
temperature, fluid activity and melt fraction imply that the
lower crust experienced rapid changes in strength (effective
viscosity) during the ,120–105 Ma period.
In addition to rates of cooling and magma emplacement,
deformation patterns also were influenced by rheological
contrasts and mechanical anisotropies created by the layered
architecture of the lower crust. The influence of these
features are revealed by the formation of shear zones in the
contact aureoles of the batholith where steep temperature
gradients and melt created strength contrasts and their
parallelism with compositional layering (S1) in host rock.
The importance of inherited anisotropies has been observed
in other orogens also (Klepeis and Crawford, 1999; Miller
and Paterson, 2001).
Unlike the Mt. Daniel Shear Zone, the Caswell fold–
thrust belt at the top of the WFO continued to develop as
steep shear zones formed in the lowermost crust (Table 1).
This heterogeneity appears to reflect differences in lithology
at these two locations. As the WFO cooled, the strength
contrast between the dioritic rock of the WFO and
metasedimentary host rock at Caswell Sound continued to
focus deformation. However, at Mt. Daniel similar bulk
compositions between the WFO and its host rock meant that
once the batholith had cooled to ambient temperatures, the
contrast in strength across this zone was small. This
interpretation explains differences in the thickness and
duration of the shear zones. The relationships illustrate that
crustal decoupling and the formation of subhorizontal or
dipping shear zones in the deep crust strongly reflects local
rheological contrasts and temperature gradients (see also
Karlstrom and Williams, 2002).
The period of orogenesis characterized by high degrees
of vertical coupling (Fig. 10d) coincided with the develop-
ment of a strong, dry, mafic lower crustal root within
northernmost Fiordland. A strengthening lower crust is
implied in part by cooling. However, other processes also
greatly influenced lower crustal strength. Piston cylinder
experiments performed on unmelted samples of dioritic
gneiss from the Pembroke Valley (Antignano, 2002)
indicate that melt fractions remained low (#10 vol%) at
all temperatures up to T ¼ 975 8C. Fluid absent melting of
the gneiss was controlled by the decomposition of
hornblende ^ clinozoisite to produce garnet þ melt and
resulted in low melt volumes (Antignano, 2002; Klepeis
et al., 2003). Klepeis et al. (2003) pointed out that although
partial melting occurred in large parts of the lower crustal
section (Fig. 10c), the total volumes of melt probably
remained low. Low melt volumes would have helped the
lower crust remain strong even as it partially melted. In
addition, the formation of melt-induced fractures and vein
networks by ,120 Ma aided melt extraction and escape
from the lower crust (Daczko et al., 2001b). Klepeis and
Clarke (2003) showed that ductile deformation in shear
zones also aided melt escape. Davidson et al. (1994),
Roering et al. (1995) and Rushmer (1995) report similar
mechanisms of melt-enhanced embrittlement in the deep
crust. These processes helped prevent the wholesale
weakening of the lower crust following batholith
emplacement.
In summary, the relationships we describe suggest that
the development of a strong lower crustal root was aided by
the following processes: (1) rapid cooling and crystal-
lization of the WFO as pre-existing arc crust was thrust
above and below the batholith; (2) the dehydration of the
lower crust at granulite facies conditions; (3) the inherently
high strength of mafic lower crustal rocks even as they
underwent mineral reactions involving partial melting; (4)
SG 1456—30/9/2003—14:58—SFORSTER—82496— MODEL 5 – br,ed,cor
K.A. Klepeis et al. / Journal of Structural Geology xx (0000) xxx–xxx22
ARTICLE IN PRESS
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2391
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2455
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2464
UNCORRECTED PROOF
low volumes of partial melt in the lower crust controlled by
the decomposition of hornblende and clinozoisite; and (5)
the rapid extraction of partial melt via a regionally extensive
networks of fractures, veins and ductile shear zones.
The interplay among these processes illustrates the time-
dependent nature of lower crustal rheology. This variability
is consistent with data derived from rock deformation
experiments that indicate melt volume, temperature, the
mobilization of melt, and crustal composition cause
variations in rock strength (Dell’Angelo and Tullis, 1988;
Rushmer, 1995; Huerta et al., 1996; Rutter, 1997). The close
relationships we observed among magmatism, elevated
temperatures, melt-assisted strain localization, and vertical
decoupling during the ,126–120 Ma period also are
consistent with numerical and analytical models indicating
that crust–mantle decoupling is characteristic of orogens
with a weak lower crust. A weak, low viscosity zone in the
lower crust causes stresses arising from mantle convergence
preferentially to be transmitted horizontally rather than
vertically through the crust (Royden, 1996; Ellis et al., 1998;
Willett, 1998). The disparate strain paths we observed
between different parts of the crustal section in Fiordland
during WFO emplacement appear to reflect this process.
Decoupling of the crust above and below the WFO is
consistent with experimental data that show stress cannot be
transmitted across zones containing more than ,30–45%
volume melt (Rutter and Neumann, 1995).
The Fiordland example also illustrates spatially (both
vertical and horizontal) variability of lower crustal strength
profiles. The development of a relatively cool (T ¼ 650–
700 8C) lower crust in northernmost Fiordland by ,116 Ma
contrasts with the elevated temperatures (T . 800 8C) at
,107.5 Ma recorded in granulite facies mineral assem-
blages ,100 km to the south in Doubtful Sound (Gibson
and Ireland, 1995). Geochronologic and geochemical data
also indicate that Na-rich magmas that reflect lower crustal
sources continued to be emplaced into the middle and upper
crust until ,105 Ma (Tulloch and Kimbrough, 2003). This
magmatic activity at shallow crustal levels is compatible
with data that suggest dikes and deformation ductile shear
zones helped extract melt from the lower crust during the
,116–105 Ma period (Klepeis and Clarke, 2003; Klepeis
et al., 2003). High degrees of vertical coupling thus occurred
even as melt moved out of the lower crust and was emplaced
into the upper crust until ,105 Ma. Together these
observations imply that the effects of magmatism and
partial melting on strength profiles and rheological tran-
sitions in the lower crust are much more transient and
spatially heterogeneous than previously believed.
Finally, the features in the upper and lower crusts of the
Fiordland–Westland orogen suggest that a highly viscous
(strong, dry, mafic) lower crust influenced the overall
structural style and partitioning of deformation vertically
within the orogen. The narrow (50–75 km wide) focused
style of contraction in the upper crust matches that predicted
for orogens characterized by a highly viscous lower crust,
where stresses derived from convergence and flow in the
upper mantle preferentially are transmitted vertically
through the crust (Royden, 1996; Ellis et al., 1998). In
these types of orogens, the lower crust is not sufficiently
weak to allow channelized lower crustal flow and crust–
mantle decoupling during convergence.
9.2. Conditions influencing late-orogenic extension
The Fiordland–Westland orogen records the onset of
extension (Fig. 10d) and the cessation of contraction by
,108–105 Ma (Tulloch and Kimbrough, 1989; Gibson and
Ireland, 1995; Spell et al., 2000). This pattern invites
comparison with other orogens that also exhibit evidence of
late orogenic extension following a period of crustal
thickening. The mechanisms that can contribute to this
process are variable, including thermal relaxation of
thickened lower crust (Vanderhaeghe and Teyssier, 1997,
2001), changes in convergence rate (England and House-
man, 1989), and the convective removal of a lithospheric
root (Molnar et al., 1993). Below, we evaluate the possible
roles of these mechanisms for the Fiordland–Westland
orogen.
In vertically coupled orogens characterized by a highly
viscous lower crust, contraction and extension can be
balanced between different layers of the crust (e.g. Burg
et al., 1984; Burchfield et al., 1992; Burg and Ford, 1997).
Focused extension in the middle and upper crusts can occur
simultaneously with shortening in the lower crust as a result
of return flow away from the zone of mantle convergence
(Royden, 1996). However, the timing of normal faulting in
both the upper and lower crusts in the Fiordland–Westland
eliminates this as a possible mechanism for this orogen. An
alternative mechanism of late orogenic extension is a
thermal weakening of previously thickened lower crust. In
the Canadian Cordillera and the Himalaya–Tibet region, the
onset of late orogenic extension has been linked to the
thermal weakening of the middle or lower crust and an
increased role of buoyancy forces due to partial melting
(Royden, 1996; Vanderhaeghe and Teyssier, 1997; Ellis
et al., 1998; Jamieson et al., 1998; Willett, 1998). This
mechanism explains crosscutting relationships and differ-
ences in timing between contractional structures and normal
faulting. However, we suggest that because we have
evidence for a viscous, cooling, mafic lower crust and
strong vertical coupling prior to the onset of extension (Fig.
10), we also can eliminate this as a possible mechanism of
late orogenic extension. The relationships we observe in
Fiordland suggest that late orogenic extension was caused
by changes in horizontal forces derived from changing plate
boundary dynamics (e.g. England and Houseman, 1989)
rather than a changing lower crustal rheology or an increase
in buoyancy forces arising from a thermal weakening of the
lower crust during convergence.
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10. Conclusions
A viscous, dry, mafic lower crust resulted in strong
vertical coupling between the upper and lower crusts during
Early Cretaceous contraction and orogenesis in western
New Zealand. Vertical coupling lasted from ,116–105 Ma
and may have begun by ,120 Ma. Transient vertical
decoupling of the crust coincided with the emplacement of
.10 km (thickness) of mafic–intermediate magma into the
lower crust during the interval ,126–120 Ma. The
development of a strong lower crustal root by ,116 Ma
reflected the following conditions and processes: (1) cooling
and crystallization of magma as pre-existing arc crust was
thrust above and below the WFO batholith; (2) extensive
dehydration of the lower crust at granulite facies conditions;
(3) the inherently high strength of mafic lower crustal rocks
as they underwent mineral reactions involving partial
melting; (4) low volumes of partial melt in the lower crust
controlled by the decomposition of hornblende ^
clinozoisite in mafic crust; and (5) the rapid extraction of
partial melt via melt-enhanced fracture networks, veins and
ductile shear zones. Rheological contrasts and a viscous
lower crust controlled the partitioning of deformation
vertically within the orogen and produced a narrow (50–
75 km wide), focused structural style in the upper crust.
High degrees of vertical coupling occurred even as partial
melts and magma moved out of the lower crust and were
emplaced into the middle and upper crusts. The results
imply that the onset of late orogenic extension at ,108–
105 Ma was caused by a change in plate boundary dynamics
rather than lateral flow of a weak lower crust during
convergence. The Fiordland exposures provide an important
natural example of the time-dependent, heterogeneous
nature of lower crustal strength and rheology.
11. Uncited References
Karlstrom and Williams, 1995. Landis et al., 1999.
Oliver, 1980. White and Jiang, 1994.
Acknowledgements
Funding to support this work was provided by National
Science Foundation funding to KAK (EAR-0087323);
Australian Research Council funding to KAK and GLC
(ARC-A10009053); and funding from the Geological
Society of America. We thank A. Tulloch for a helpful
review; and N. Mortimer, I. Turnbull, T. Rushmer and W.C.
Collins for discussions and assistance. We thank the
Department of Land Conservation in Te Anau for
permission to visit and sample localities in the Fiordland
National Park. J. Hollis and T. Ireland provided assistance in
compiling geochronologic data from northern Fiordland.
J. Stevenson, A. Papadakis, A. Claypool, S. Marcotte and
W.C. Simonson provided assistance in the field.
Appendix A. Analytical procedures for U–Pb isotopic
analyses
Zircons were analyzed at the University of Arizona with
a Micromass Isoprobe multicollector ICPMS equipped with
9 faraday collectors, an axial Daly detector, and 4 ion-
counting channels. The Isoprobe is equipped with a DUV
193 laser ablation system from New Wave Research. The
laser is a Compex 102 ArF Excimer laser, manufactured by
Lamda Physik, with an emission wavelength of 193 nm. The
analyses were conducted on 25–50 micron spots with an
output energy of ,32 mJ and a repetition rate of 8 Hz. Each
analysis consisted of one 30-s integration on the back-
grounds (on peaks with no laser firing) and twenty 1-s
integrations on peaks with the laser firing. The depth of each
ablation pit is ,20 microns. The collector configuration
allows simultaneous measurement of 204Pb in a secondary
electron multiplier while 206Pb, 207Pb, 208Pb, 232Th, and238U are measured with Faraday detectors. All analyses
were conducted in static mode.
Correction for common Pb for most samples was
performed by measuring 206Pb/204Pb, with the composition
of common Pb from Stacey and Kramers (1975) and
uncertainties of 1.0 for 206Pb/204Pb and 0.3 for 207Pb/204Pb.
For two of the samples analyzed (995a and 9928), 204Pb was
not measured reliably, and common Pb was accounted for
with a 207Pb correction (Ludwig, 2001).
Fractionation of 206Pb/238U and 206Pb/207Pb during
ablation was monitored by analyzing fragments of a large
concordant zircon crystal that has a known (ID-TIMS) age
of 564 ^ 4 Ma (2s; G.E. Gehrels, unpublished data).
Typically this reference zircon was analyzed once for every
four unknowns. The uncertainty arising from this calibration
correction contributed ,2% systematic error to the206Pb/238U and 206Pb/207Pb ages (2s level). The uncertainty
from decay constants and common Pb composition yielded
an additional ,1% error to each analysis.
The reported ages are based primarily on 206Pb/238U
ratios because the errors of the 207Pb/235U and 206Pb/207Pb
ratios are significantly greater (Fig. 8; Table 2). This is due
in large part to the low intensity (commonly ,1 mv) of the207Pb signal from these young grains. For each of the five
samples (Fig. 8), the age data are plotted on a Pb/U
concordia diagram, and the critical 206Pb/238U ages are
shown on a separate age plot. The final age calculations are
based on the weighted mean of the cluster of 206Pb/238U
ages, with the error expressed both as the uncertainty of this
mean and as the error of the age. The age error is based on
the quadratic sum of the weighted mean error and the
systematic error. Both are expressed at the 2s level.
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