Batch-wise assembly and zoning of a tilted calc-alkaline batholith:
Field relations, timing, and compositional variation
N. Coint1, C.G. Barnes1, A.S. Yoshinobu1, K.R. Chamberlain2, and M.A. Barnes1
1
Department of Geosciences, Texas Tech University, Lubbock, Texas 79409-1053, USA
Department of Geology and Geophysics, University of Wyoming, Laramie, Wyoming 82071-3006, USA
2
ABSTRACT
INTRODUCTION
The Wooley Creek batholith is a tilted,
zoned, calc-alkaline plutonic complex in the
Klamath Mountains, northern California,
USA. It consists of three main compositionaltemporal zones. The lower zone consists of
gabbro through tonalite. Textural heterogeneities on the scale of tens to hundreds of
meters combined with bulk-rock data suggest that it was assembled from numerous
magma batches that did not interact extensively with one another despite the lack of
sharp contacts and identical ages of two
lower zone samples (U-Pb [zircon] chemical
abrasion–isotope dilution–thermal ionization mass spectrometry ages of 158.99 ± 0.17
and 159.22 ± 0.10 Ma). The upper zone is
slightly younger, with 3 samples yielding ages
from 158.25 ± 0.46 to 158.21 ± 0.17 Ma, and
is upwardly zoned from tonalite to granite.
This zoning can be explained by crystalliquid separation and is related to upward
increases in the proportions of interstitial
K-feldspar and quartz. Porphyritic dacitic
to rhyodacitic roof dikes have compositions
coincident with evolved samples of the upper
zone. These data indicate that the upper
zone was an eruptible mush that crystallized
from a nearly homogeneous parental magma
that evolved primarily by upward percolation of interstitial melt. The central zone
is a recharge area with variably disrupted
synplutonic dikes and swarms of mafic
enclaves. Central zone ages (159.01 ± 0.20 to
158.30 ± 0.16 Ma) are similar to both lower
and upper zones crystallization ages. In the
main part of the Wooley Creek batholith, age
data constrain magmatism to a short period
of time (<1.3 m.y.). However, age data cannot
be used to identify distinct magma chambers
within the batholith; such information must
be extracted from a combination of field
observations and the chemical compositions
of the rocks and their constituent minerals.
Suprasubduction zone magmatism is responsible for the generation of large volumes of
magma expressed as batholiths and volcanic
arcs. Because arcs develop episodically, a great
deal of recent research has focused on the tempo
of intrusive and volcanic activity, on the temporal and petrological relationships between plutonic and volcanic magmas, and on the conditions necessary for large volumes of magma
to exist in subvolcanic magma chambers. A
number of workers have used high-precision
U-Pb (zircon) ages to show that some large
batholiths are assembled over several million
years (e.g., Coleman et al., 2004; Glazner et al.,
2004; Matzel et al., 2006; Grunder et al., 2008;
Schaltegger et al., 2009). In the case of large
zoned intrusions such as the Tuolumne batholith (California, USA), these results have been
interpreted to indicate that large intrusions result
from emplacement and amalgamation of many
distinct magma batches that were not able to
interact extensively with each other (Coleman
et al., 2004; Glazner et al., 2004; Bartley et al.,
2008). Nevertheless, detailed studies of magmatic fabrics and zoning patterns in parts of the
Tuolumne batholith indicate that distinct intrusive contacts do not always separate batches of
different ages (e.g., Memeti et al., 2010).
Models in which individual magma batches
crystallize completely, preventing extensive
interaction with one another, are contradicted
by the presence of voluminous ignimbrites that
are erupted during single events (e.g. Bacon
and Druitt, 1988; Bachmann et al., 2002;
Hildreth, 2004; Christiansen, 2005; Lipman,
2007). These eruptions require storage of large
volumes of magma in the crust, at least for
short periods of geologic time. Discrepancies
between models for development of ignimbrite magma systems and models for magma
batches unable to interact with one another
have led some to suggest that large-volume
ignimbrites are unrelated to typical batholith-
scale plutons (Mills and Coleman, 2010; Tappa
et al., 2011).
It is clear that there is a need for field and
laboratory tests to determine whether magma
batches emplaced incrementally remain as compositionally isolated entities or whether collections of magma batches are able to interact and
perhaps homogenize in large crustal magma
chambers (e.g., Tepley et al., 2000; Ohba et
al., 2007; Turnbull et al., 2010; Ruprecht et al.,
2012). The first logical test involves high-precision dating, commonly U-Pb dating of zircon.
In cases where emplacement ages span a range
of time longer than the time in which a magma
batch should solidify, incremental emplacement with little to no communication between
the batches is supported. However, it is possible
that emplacement of batches unable to interact
with one another occurred in such a short period
of time that even high-precision dating cannot
distinguish batches; in such cases, field, petrographic, and geochemical data must be used. For
example, batch-wise emplacement with little or
no communication between batches might be
expected to leave mineral and geochemical evidence of isolated magma units unrelated by in
situ differentiation processes such as fractional
crystallization, magma mixing, or assimilation.
We apply such tests to the Wooley Creek batholith (WCb), a large (326 km2), tilted plutonic
complex in the Klamath Mountain province
of northern California (USA; Fig. 1). Detailed
mapping and petrographic data are combined
with high-precision U-Pb dating of zircon and
bulk-rock geochemical data to characterize zoning of the batholith, determine the physical and
temporal relationships between magma zones,
and develop a model for batholith assembly.
GEOLOGIC SETTING
Klamath Mountains
The Klamath Mountains geologic province,
northwestern California and southwestern
Geosphere; December 2013; v. 9; no. 6; p. 1729–1746; doi:10.1130/GES00930.1; 9 figures; 1 table; 1 supplemental table.
Received 12 March 2013 ♦ Revision received 17 July 2013 ♦ Accepted 29 August 2013 ♦ Published online 23 October 2013
For permission to copy, contact editing@geosociety.org
© 2013 Geological Society of America
1729
Coint et al.
Legend
123°20’00”W
Jurassic Wooley
Creek batholith &
Slinkard pluton
Ji
Condrey Mt.
Schist
~800 MPa
Slinkard pluton
Western Klamath
Terrane (Jurassic)
Late Jurassic
plutons
Western
Klamath terrane
gr/gd gb/di/
qd
lt
au
Or
le
an
s Thrus
t
F
c. 152 Ma
Orleans thrust
Western/eastern
wHt
eHt Hayfork terranes
RCt
gd
Galice Formation/
GFm
RCt Rattlesnake Creek
RCt
terrane
RCt
gb/di/
qd
ton/qd
Condrey Mt. Schist
gb/di/qd
Ji
RCt
wHt
eHt
650 MPa,
700-800 °C
RCt
Ji
500 MPa
reek
C
y
le
Woo tholith
ba
41°30’30” N
RCt
RCt
Explanation
Orleans thrust fault (teeth on HW)
eHt
Thrust fault (teeth on HW)
Extensional fault (box on HW)
wHt
300 MPa Contact metamorphic pressure
determination
Fig. 2
Igneous zonal boundary
wHt
Field area
350
MPa
GFm
North
Klamath
Mts.
300 MPa
CA
0
5 km
W
Figure 1. Wooley Creek batholith and Slinkard pluton map with host terrane boundaries
(modified from Barnes, 1983; Barnes et al., 1990). HW—hanging wall; gr—granite; gd—
granodiorite; gb—gabbro; di—diorite; qd—quartz diorite; ton—tonalite.
Oregon, consists of a sequence of tectonostratigraphic terranes accreted to North America. The
province is subdivided into four belts, each of
which consists of a number of terranes. The
belts are separated by east-dipping shear zones
with reverse sense of motion. In general, the
overall age of the belts becomes younger to
the west, structurally downward (see Irwin,
1960; Davis, 1968; Saleeby et al., 1982; Snoke
and Barnes, 2006). Most of the component terranes of the province have oceanic affinities
(ophiolites, mid-oceanic ridge basalt–related
mafic assemblages, subduction mélanges, arc
sequences, flysch deposits) and many formed in
suprasubduction zone settings.
1730
The WCb and Slinkard pluton (Fig. 1) were
emplaced into three lithologically, chemically,
and isotopically distinct tectonostratigraphic
terranes of the western Paleozoic and Triassic
belt (Fig. 1; Hotz, 1971; Irwin, 1994; Snoke
and Barnes, 2006). The structurally lowest
host unit is the Triassic–Jurassic Rattlesnake
Creek terrane, an ophiolitic mélange that contains metaserpentinite, metagabbro, metabasite, hemipelagic metasedimentary rocks, and
marble (Donato et al., 1982) metamorphosed
ca. 168 Ma (Garlick et al., 2009). The Middle
Jurassic western Hayfork terrane overlies the
Rattlesnake Creek terrane and consists of an
arc-related clastic, volcaniclastic, and hemi-
Geosphere, December 2013
pelagic sequence (Fig. 1; Donato et al., 1982,
1996; Wright and Fahan, 1988). The volcanogenic rocks of the western Hayfork terrane
range from basaltic to andesitic (Barnes et al.,
1995). The structurally highest host-rock unit
is the Triassic eastern Hayfork terrane, which
is a chert-argillite mélange and broken formation with local olistostromal units (e.g., Wright,
1982; Donato et al., 1982; Ernst et al., 2008).
Mélange blocks include metasandstone, marble, and serpentinized peridotite. Contractional
deformation beginning ca. 152 Ma associated
with the Nevadan orogeny placed these terranes
along with the WCb in thrust contact above
low-grade metasedimentary rocks of the Galice
Formation and the Condrey Mountain Schist
(Fig. 1; Irwin, 1972; Jachens et al., 1986). The
Condrey Mountain Schist is exposed in a domal
window northeast of the WCb (Fig. 1). The metamorphic grade of hanging-wall rocks above
the Condrey Mountain dome increases toward
the dome from greenschist (at ~300 MPa) to
granulite (at ~800 MPa) facies (Donato, 1987,
1989; Petersen, 1982; Lieberman and Rice,
1986; Garlick et al., 2009), indicating that doming caused radial tilting of the hanging wall to
the southwest, away from the dome (current
geographical coordinates). Effects of this tilting were observed in the contact metamorphic
assemblages in the WCb aureole, where mineral
assemblages give pressure estimates of ~300 ±
150 MPa along the southern contact, farthest
from the dome, to ~650 ± 150 MPa along the
northeastern contact, nearest the dome (Barnes
et al., 1986b). These pressure differences show
that regional tilting of ~15° and subsequent erosion have exposed ~9 km of structural relief
from north to south across the WCb.
The intrusive system consists of two plutons
(Fig. 1), the WCb and the Slinkard pluton, plus
basaltic, andesitic, and dacitic dikes that crop out
along the southwestern margin of the WCb. The
Slinkard pluton, which crops out northeast of the
WCb (Fig. 1), is structurally, compositionally,
and temporally linked to the WCb, and therefore extends the exposed vertical extent of the
system to at least 12 km (Barnes et al., 1986b,
1990). The original thermal ionization U-Pb
(multicrystal zircon) dating of the WCb and
Slinkard pluton yielded an age of ca. 161 Ma +4,
−2 Ma (Late Jurassic; Barnes et al., 1986a). In
this study we focus on the WCb, because of the
availability of high-precision U-Pb (zircon ages)
and the fact that many samples of the Slinkard
pluton have undergone deuteric alteration.
BATHOLITH ZONATION
The WCb is broadly zoned from more mafic
rocks in the structurally deeper northern and
Map of igneous
lithologies
Samples with red
outline have modal
analysis
Ji
WCB
1408
Diorite/gabbro
Quartz diorite
Tonalite
Quartz monzodiorite
Granodiorite
Quartz monzonite
Granite
Geochronology sample
Lower zone
Ji
Ten Bear
Mtn.
Central zone
Pigeon
Roost
wms
WCB5109
Cuddihy
Lakes
basin
Lower Zone
At the outcrop scale, the lower zone shows
varying degrees of heterogeneity. In some locations, for example in the Pigeon Roost area
(Fig. 2), rock types vary on the scale of meters
to tens of meters, whereas in other parts local
modal layering is observed (Fig. 3A).
The lower zone consists of gabbro through
tonalite. Lower zone samples are generally
characterized by the assemblage augite + orthopyroxene + biotite ± Ca-amphibole (hereafter
hornblende) (Figs. 4A, 4B). Samples that lack
pyroxene contain clusters of hornblende or
more commonly actinolitic amphibole rimmed
by hornblende. These clusters are interpreted
to be relict pyroxene surrounded by magmatic
(nonperitectic) hornblende. As with the variation in rock type, there is little or no systematic
large-scale spatial variation of the mafic mineral
assemblage; however, texture varies at scales
of tens to hundreds of meters. Oxide minerals
(magnetite and ilmenite) are in low abundance
in most samples.
Variably shaped masses of pyroxenite, melagabbro, and gabbro are locally abundant in the
Pigeon Roost region (location in Fig. 2; Figs.
3B, 3C). Grain size in such bodies is typically
coarser than in the more common quartz diorite–tonalite. Contact relationships between
quartz diorite–tonalite host and these coarser
bodies are variable but are consistent with the
coarse-grained pyroxenites-melagabbro-gabbro
bodies being intrusive into and synmagmatic
with their hosts. Evidence for this relationship
consists of (1) coarse pyroxene grains from the
pyroxenite-melagabbro as xenocrysts in the surrounding rocks, (2) the presence of melagabbro
dikes into the host diorite (Fig. 3B) with local
123°20’00”W
WCB
4809,
4909,
Z5
WCB
10510
WCB
7809
41°30’30” N
eastern parts of the pluton to the most felsic
rocks in the structurally shallower southwestern area (Barnes, 1983). This zoning is shown
in detail in Figure 2, which illustrates the variation in rock type and in varietal minerals across
the batholith. On the basis of rock type and mineral assemblage, lower and upper zones can be
mapped. These zones are locally separated by
a 2–3-km-wide region that contains numerous
synplutonic mafic dikes, microgranitoid enclaves
and enclave swarms, and a variety of host-rock
xenoliths in a sheeted tonalitic to quartz dioritic
matrix. Locally intense hypersolidus deformation is observed. We refer to this zone as the
central zone, and discuss its relationship to the
upper and lower zones in the following. Along
the western, southwestern, and southern margins, discontinuous mafic selvages of pyroxenebearing quartz diorite and gabbro crop out; these
zones reach 1 km in width (Fig. 2).
Biotite hornblende
(Pyroxene) biotite hornblende
Biotite hornblende pyroxene
Two-pyroxene (+/-biotite)
Assembly of a tilted batholith
WCB
7909
Medicine
Mtn.
Upper zone
sms: southern mafic
selvage
wms: western mafic
selvage
MMB
377
Gradational
compositional
boundary
Late granite
WCB
2308
region of abundant
“Roof zone” dikes
sms
Ji
WCB
2408
0
N
5 km
Figure 2. Simplified geologic map of the modal distribution of minerals and
rock type for the Wooley Creek batholith (sample prefix WCB) and dikes
and internal gradational compositional boundaries (dashed lines). Sample
numbers of rocks dated by chemical abrasion–isotope dilution–thermal ionization mass spectrometry are indicated. Semitransparent gray ellipses and
polygons represent regions noted in the text. Legend as in Figure 1.
Geosphere, December 2013
1731
Coint et al.
A
B
two pyroxene diorite
pyroxene-rich dike
C
D
Pyroxene-rich block
Roof
dike
biotite
pyroxene-rich rock
biotite hornblendite
Granodiorite
Quartz-rich vein
E
quartz diorite
Salmon river
F
WCB-4909
MME
mgr
qtz-dio
MME
WCB-4809
fgr
qtz-dio
MME
syn-plutonic dike
fgr
qtz-dio
mgr
qtz-dio
Figure 3. Outcrop photographs. (A) Modal layering defined by the variable proportions of hornblende and pyroxene compared to
plagioclase in the lower zone. (B) Pyroxene-rich melagabbro cutting across two-pyroxene diorite in the lower zone. (C) Pyroxene-rich
blocks surrounded by a reaction rim of hornblendite, brecciated by quartz-rich veins in the lower zone. (D) Upper zone granodiorite
cut by one of the roof dikes. (E) Folded synplutonic dikes in quartz diorite, central zone. (F) Alternating sheets of mafic magmatic
enclave (MME) free fine-grained quartz diorite (fgr qtz-dio) and MME-rich medium-grained (mgr) quartz diorite in the central zone.
1732
Geosphere, December 2013
Assembly of a tilted batholith
B
A
Px
Px
Px
Hbl
Px
Hbl
Pl
Bt
Px
Pl
Hbl
Bt
Pl
2 mm
Bt
C
Pl
2 mm
D
Bt
Bt
Hbl
Bt
Pl
Pl
Qz
Pl
Qx
Hbl
Hbl
K-spar
Hbl
Pl
Qz
Hbl
Pl
Qz
Hbl
Pl
Pl
Pl
1 mm
2 mm
E
F
Chl
Qz
Hbl
Cpx
Plag
Hbl
Cpx
Qz
Qz
Cpx
Qz
1 mm
Hbl
2 mm
Figure 4. Photomicrographs of the main rock types found in the Wooley Creek batholith in crossed-polarized light. Abbreviations: Px—
pyroxene; Pl—plagioclase; Hbl—hornblende; Bt—biotite; Qz—quartz; K-spar—potassium feldspar; Cpx—clinopyroxene. (A) Biotite
hornblende pyroxene diorite from the lower zone. (B) Hornblende biotite tonalite displaying a subsolidus fabric, defined by the alignment of
biotite, overprinting the magmatic one defined by alignment of plagioclase and hornblende. (C) Biotite hornblende granite from the upper
zone. Note the broken zoned plagioclase phenocrysts in a poikilitic K-feldspar. (D) Biotite hornblende quartz diorite from the central zone.
(E) Porphyritic two-pyroxene andesite roof dike. (F) Porphyritic rhyodacite roof dike. Note the presence of euhedral quartz phenocrysts.
Geosphere, December 2013
1733
Coint et al.
back-veining and/or disruption of pyroxenitemelagabbro by the surrounding rocks, and
(3) rare enclaves of the host quartz diorite–
tonalite enclosed in pyroxenites-melagabbro;
in such instances a 1–3-cm-thick rind of hornblende separates the two rock types. Compositional banding that consists of centimeter to
tens of meters variations in the proportions of
mafic minerals is also present in the lower zone
(Fig. 3A). Elsewhere, for example at Medicine
Mountain (Fig. 2), large areas are underlain
by quartz diorite–tonalite with scant variations
in color index and rare bodies of melagabbropyroxenite. Many parts of the lower zone are
cut by planar, commonly net-veined mafic
dikes. These dikes typically are parallel walled
and fine grained and range from gabbroic to
quartz dioritic, with hornblende ± biotite as the
mafic phases.
Upper Zone
The upper zone ranges from medium- to
coarse-grained tonalite to granite and appears
fairly homogeneous at the scale of the outcrop
(Fig. 3D). Figure 2 shows that this compositional zonation, while variable, is from more
mafic to more felsic rocks toward the westsouthwest, structurally upward. The texture of
nearly all upper zone samples is hypidiomorphic granular, with sparse centimeter-scale
hornblende and plagioclase phenocrysts. The
groundmass is composed of a seriate distribution of hornblende, plagioclase, biotite, quartz,
and K-feldspar from ~5 mm to ~0.2 mm in
diameter. Hornblende and plagioclase are
euhedral to subhedral and are weakly oriented
in samples with magmatic foliation. Biotite is
euhedral to anhedral and in most samples is
not oriented in the foliation plane. Quartz varies from interstitial in tonalitic rocks to euhedral
and/or subhedral in granite. Euhedral quartz is
a common inclusion in poikilitic K-feldspar in
the granites, whereas K-feldspar is everywhere
interstitial to poikilitic (Fig. 4C). Broken plagioclase crystals form inclusions in K-feldspar
in granodioritic and granitic samples (Fig. 4C).
In summary, the mineral assemblages and overall textures of upper zone rocks are essentially
identical from one sample to another; samples
differ only in mineral proportions and quartz
habits. Accessory phases in the upper zone are
apatite, zircon, allanite, and epidote.
Felsic dikes occur sparsely in the upper zone,
and in the southern part of the pluton a small,
medium-grained granitic intrusion with associated pegmatites, referred to as the late granite,
cuts the upper zone rocks (Fig. 2). Mafic dikes
are also uncommon; however, enclave swarms
and accumulations of mafic enclaves (pillows)
1734
are common, particularly in the structurally lowest part of the zone (see following). It is common that the color index of tonalitic and granodioritic rocks near swarms of mafic enclaves is
higher than the color index of rocks distal from
enclave swarms.
Central Zone
The central zone (Fig. 2) represents the most
heterogeneous part of the batholith in terms
of lithology, mineralogy, structure, and intrusive relationships. In areas of excellent exposure, such as the Cuddihy Lakes basin (Fig.
2), the central zone is underlain by sheets of
medium- to coarse-grained quartz diorite and
tonalite (Fig. 3F). Some sheets contain abundant mafic magmatic enclaves and others are
essentially enclave free (Fig. 3F). Central zone
samples contain hornblende and plagioclase
with the same habit as in the upper zone (Fig.
4D). Biotite varies from euhedral to interstitial and quartz and K-feldspar are interstitial.
Most samples lack pyroxene but some contain
clusters of amphibole similar to those from
the lower zone that are interpreted to be relict
pyroxene. Accessory minerals are apatite, zircon, and scant allanite.
The central zone is also characterized by
abundant synplutonic dikes (Fig. 3E). These
mafic-intermediate dikes are similar to those in
the lower zone, fine to medium grained and consisting of hornblende + plagioclase ± biotite and
accessory minerals. Net veining of these dikes is
pervasive and with increasing structural height,
ductile to brittle deformation and disruption of
dikes becomes common (see following) (Fig.
3E; Barnes et al., 1986a, 1990).
Appinitic dikes and masses crop out mainly
in the central zone, but are also sparsely present
in the lower and upper zones. In nearly all cases,
these dikes are back veined by the host. The
appinites contain euhedral, centimeter-scale
blocky hornblende set in an interstitial to poikilitic groundmass of calcic plagioclase ± quartz
± K-feldspar. It is common for euhedral, millimeter-scale pyroxene crystals to be enclosed in
poikilitic plagioclase.
BATHOLITH–HOST ROCK CONTACTS
Contact relationships were described in
Barnes (1983) and are reviewed here, along
with new data and descriptions from the southwestern contact. The contact between the lower
zone and host rocks along the northeastern
contact is best described as a contact zone in
which plutonic and metamorphic rocks are
interleaved over a 30–100 m distance. This part
of the contact currently dips shallowly to the
Geosphere, December 2013
north, whereas the eastern and western contacts
between the lower zone and the host rocks are
steep to vertical. Along the eastern contact of the
lower zone, foliations in the batholith and host
rock are subparallel to the contact, but along the
northwestern margin of the pluton such foliations are discordant to the contact (Fig. 5A).
In this area, the northwestern contact is locally
defined by a younger, brittle, high-angle fault
that overprints and deforms the contact. However, south of the fault, the foliation remains discordant to the contact (Fig. 2).
The western, southwestern, and southeastern contacts are primarily between rocks of the
upper zone and host rocks. This contact is generally steep and sharp, but is locally gradational
over <10 m of interleaved plutonic rock and
host rocks. The discontinuous mafic selvages
along the western contact (Fig. 2) are underlain
by pyroxene-bearing mafic and/or intermediate rocks (mafic selvages; Barnes, 1983) that
separate upper zone rocks from host rocks of
the western Hayfork terrane. Contacts between
the selvages and host rocks are sharp, but the
internal contacts between the selvages and the
upper zone rocks are gradational (Barnes, 1983;
Coint, 2012). Other such contact zones may
exist because the color index of upper zone
rocks is locally higher adjacent to the upper
contact than it is 100 m inside the pluton. However, the quality of exposures does not permit
mapping of such zones. The southeastern contact exposed in the Salmon River (Fig. 6) dips
moderately to the west.
The southernmost contact with the host rocks
is a plexus of mafic to intermediate rocks that
intrude chert-argillite breccia and calc-silicate
rocks of the eastern Hayfork terrane (Fig. 6).
This region was mapped as a mafic selvage similar to that observed on the western margin (e.g.,
Barnes et al., 1986a); however, the southern
contact zone is distinct from the western contact
zones in at least three ways. First, the southern
selvage consists in part of coarse- to fine-grained
gabbro with variable crystal size and texture.
Some gabbro samples contain sparse plagioclase and pyroxene phenocrysts with relict
olivine, others have granoblastic textures, and
still others contain abundant glomerocrysts of
subhedral to euhedral orthopyroxene and augite
with interstitial oxide minerals set in a matrix
of euhedral, aligned plagioclase in which the
foliation wraps around the glomerocrysts. Second, the contact between the southern selvage
and the upper zone granodiorite is sharp, but
lobate. Within the gabbroic rocks the proportion
of plagioclase phenocrysts decreases away from
the contact with granodiorite (Fig. 7B), suggesting that these crystals were entrained from
the granodiorite. In some places, mafic lobes
Assembly of a tilted batholith
65
123°20’00”W
68
37
17
79 82
Ji
75 59
42
60
76
72
75
74 80 65
85
83
71
82 57
74
50 75
65
85
77 78 80 75
40
67
72
84
42
62
32
83
52
48
65
60
43
50
72 83
63
80
60
71 47
58
63
80
80
89
46
29 73 60 83 72 60
74
25
80
58
33 49
82 57 62
81
58
85
40
75 71
45
55
49
56
80
60
35
85 59
53
60
54
75
50
40
79
33
54 39
68
52
42
46
57 40
73
24
71
45
50
3870 68 49
84
38
54
58
58
58
27 17
65 72 78 58
58
69
88
46
80
58 36 n = 13
75
58
75
41
74
26
32
49
38
35
46 58
62
87
Upper
WCb
48
80
45
inclined, vertical
magmatic
foliation
6975
51
n = 28
49
55
74
50 25
85
62
52
37
32
58
60
20 72
83 64 85
70 78
zones of
increased
abundance
of xenoliths
n = 179
80
70 46
59 58
45
60
26
55
N
45
83
74
C
N
30
38 86
65
Ji
72
49
85
N
14
80
82
80
59 66
Ji
35
41°30’30” N
85
57
16
22
34 1418 n = 114
50
14
70 75
44 51
80 35
20
60
70
42 14
40
83
123°20’00”W
20
37
Ji
D
38
26
37
74
55 53
25
Lower
WCb
26
35
58
N
45
58
35
B
34
41°30’30” N
A
56
58 76 77
52
50
0
inclined, vertical
metamorphic
foliation
40
North
0
5 km
5 km
Ji
Figure 5. (A) Geologic map of the Wooley Creek batholith (WCb) displaying host-rock structures and internal magmatic foliations. (B) Poles
to foliations (top) and lineations (bottom) from the lower part of the Wooley Creek batholith (WCb). Kamb contours, contour interval =
1σ. (C) Upper WCb. (D) Trend line map of magmatic foliations. Also shown are regions where metamorphic xenoliths are conspicuous.
Patterns and legend from Figures 1 and 2.
detached from the gabbro form gabbro enclaves
within the granodiorite (Fig. 7B). Therefore,
the southernmost selvage is interpreted to be
comagmatic with granodiorite of the upper zone.
A xenolith found in the granodiorite, preserving
the contact between the gabbro and the eastern
Hayfork terrane (Fig. 7A), indicates that part of
the southern selvage was emplaced before the
granodiorite. Third, the southern selvage is cut
by late andesite and granodioritic dikes; the lat-
ter are rich in magmatic enclaves. These dikes
are generally orthogonal to one other, with the
andesitic dikes oriented approximately eastwest and the granodioritic dikes approximately
north-south (Fig. 7A).
The contact aureole of the WCb is variable in
width and degree of emplacement-related metamorphism and deformation. At the map scale
(e.g., Fig. 2), host-rock structures are weakly
to strongly discordant except for regions within
Geosphere, December 2013
the contact aureole where significant ductile
deformation and emplacement-related dynamothermal metamorphism occurred. Preexisting
foliations in the host terranes along the margins
of the intrusion are locally deflected, subparallel
to the contact within a contact aureole from tens
of meters to ~1 km wide. In Barnes (1983), it
was noted that intense deformation and isoclinal
folding occurred within 50 m of the intrusion
along the northwestern and northeastern mar-
1735
81
60
80
52
70
4
4
81
45
76 64
WCb
37
80
56
80
45
North
74
1
60
57
2
27
82
granodiorite
~ 10 m
pyroxene
gabbro
64 85
60 74
7
1 80
pyroxenite
km
80
34
32
80 64
WCb
76
45
57
84
75
45
on
Gra
81
63
30
diorite gabbro
Salmon River Road
granodiorite
23
76
45
75
45
74000mN
48
55
24
WHt-ms
Explanation
WCb, upper zone (outcrops (dots) and stream
exposures (lines) by lithology)
granite (gr), granodiorite, tonalite
diorite to gabbro
gabbro
45
86
nt
ek
Cre
75
55
54
81
88 72
WHt-mv
r
Ri ve 74
75
80 85
74000mN
77
~5m
pyroxene
gabbro
B
64 85
70
55
74
63
lm
78
75
45
76
23
15
Sa
n
Salmo
Figure 7
62
72
80 8
3
85
45
75
65
84 80 55
57
WHt-mv
50
77
78
45
81 76
77
45
20
r
47
Grant Creek
20
76
70
65
65
40
79
85
gabbro
60
15
diorite
45
78
18
68
R i ve
79
0
45
WHt
76
WCb
57
24
C
45
69
68000mE
52
26
59
gr
62
74
45
45
4
4
67
55
55
54
57
80
78
45
67
66
4
A
45
4
65
64
4
4
63
Coint et al.
57
24
24 64
73
57
bedding,
inclined,
vertical
pyroxenite
metamorphic
foliation inclined,
vertical
magmatic
foliation,
lineation
57
dike inclined,
vertical
Host rocks
57 24
45
contact, approximated
Calcareous rocks of the Eastern Hayfork terrane
Metasedimentary rocks of the Western Hayfork terrane (WHT-ms)
thrust fault
Metavolcanic rocks of the Western Hayfork terrane (WHT-mv)
Figure 6. (A) Detailed geologic map of the southern contact of the Wooley Creek batholith (WCb) along the
Salmon River. (B) Outcrop map of xenoliths and screens of calcareous rocks of the eastern Hayfork terrane in diorite and gabbro directly west of Grant Creek. (C) Outcrop map of intrusive relationships between different phases
of the upper WCb as exposed along the north shore of the Salmon River directly north and east of the confluence
with Grant Creek. See text for details. Topography is from Orleans Mountain and Forks of Salmon (1:24,000 U.S.
Geological Survey quadrangle sheets); datum is mean sea level geodetic datum NAD 27 UTM (North American
Datum, Universal Transverse Mercator) grid zone 10, contour interval 50 ft (~15.2 m).
1736
Geosphere, December 2013
Assembly of a tilted batholith
A
Explanation
WCb, upper zone
Fig. B
Salmon
80
Rive
granodiorite
granite, granodiorite, tonalite
diorite to gabbro
gabbro
Calcareous rocks of the Eastern
Hayfork terrane
r
76
75 Fig. D
40 m
B
81
80 magmatic
foliation
C
Host rocks
76 inclined
dike
15 cm
D
Granodiorite
Gabbro
15 cm
Figure 7. (A) Outcrop sketch map of the eastern contact zone of the Wooley Creek batholith (WCb) along the Salmon River depicting
intrusive relationships between upper zone granite, granodiorite, tonalite, and diorite to gabbro. (B) Photograph of granodiorite
intruded by gabbro with crosscutting dioritic, fine-crystalline dike. (C) Sketch of photo in B illustrating enclaves of granodiorite of
the upper zone included in gabbro. (D) Photograph of intricate, interfingering contact between upper zone granodiorite and andesitic
dike interpreted to indicate comagmatic intrusion.
gins. However, in many localities preexisting
foliations and other structures are clearly discordant with the intrusive contact.
ROOF-ZONE DIKES
The southwestern contact of the batholith
is the highest structural level exposed and is
referred to here as the roof zone (Fig. 2; Barnes
et al., 1986b). The intrusive contact is steeply
dipping and does not restore to a subhorizontal roof; however, we retain the usage so as
to avoid confusion with the earlier papers of
Barnes et al. (1986a, 1986b). The roof zone
is characterized by porphyritic dikes (to 3 m
wide) (Fig. 3D) that range from basaltic to
rhyodacitic in composition (Figs. 4E, 4F;
Barnes et al., 1986a, 1990). A few of the mafic
dikes are hornblende pyroxene microgabbro.
However, most mafic dikes show evidence
for static (contact) metamorphism or contain
elongate to acicular Ca-amphibole in a matrix
of epidotized plagioclase, suggestive of crystallization from an H2O-saturated melt. Intermediate-composition dikes are predominantly
andesitic with phenocrysts of plagioclase,
augite, and enstatite (Fig. 4E). These dikes
intrude rocks of the upper zone within 2 km
of the roof zone as well as the host rocks structurally above the batholith (Fig. 2). A smaller
number of intermediate dikes are pyroxenehornblende andesite.
Dacitic to rhyodacitic roof dikes contain phenocrysts of plagioclase + hornblende + quartz
± biotite ± rare augite (Fig. 4F). Dike groundmass textures vary from granophyric through
hypidiomorphic granular, and it is common to
find decimeter-scale rounded mafic magmatic
enclaves in these dikes. Some dacitic dikes
intrude andesitic dikes, some form the center
of composite andesite + dacite dikes, and others
are cut by andesitic dikes (Barnes et al., 1986a).
isolated, 10-cm-scale bodies. In contrast, both
isolated enclaves and enclave swarms are common in the central zone and the lower part of
the upper zone. The enclave swarms may consist of rounded to tabular enclaves oriented
parallel to magmatic foliation or of collections
of rounded enclaves similar to so-called pillow
swarms (Barnes, 1983; cf. Wiebe et al., 2002;
Wiebe and Collins, 1998); in the former case,
enclaves tend to vary in texture and phenocryst
proportions, whereas in the latter case they
tend to be texturally similar to one another.
MAGMATIC ENCLAVES
XENOLITHS
Magmatic enclaves are widespread in the
WCb, but their proportions vary. In general,
these enclaves have a higher color index
than their host rocks, but a few have equivalent or lower color index and differ from the
host mainly in texture. The mafic magmatic
enclaves vary from fine- to medium-grained
and equigranular masses to porphyritic, with
phenocrysts of plagioclase and hornblende,
and rarely biotite. Groundmass phases are
hornblende, plagioclase, and biotite ± quartz
± K-feldspar, with accessory apatite and scant
Fe-Ti oxides ± zircon ± allanite.
In the lower zone, mafic magmatic enclaves
are generally oblate ellipsoids and occur as
Xenoliths are widespread in parts of the
WCb but vary significantly in size, rock type,
and abundance (Fig. 5D). In the lower zone,
xenoliths are common within 500 m of the
contact but are sparse in the rest of the zone.
They consist primarily of amphibolite, biotiterich schist, migmatitic quartzofeldspathic and
calc-silicate gneiss, plus scant metaperidotite
and rare garnet metagabbro.
A discontinuous zone across the center of
the pluton that broadly overlaps with the central
zone (e.g., Fig. 5D) is particularly rich in xenoliths, which range in size from centimeter scale
to at least 50 × 10 m. The largest elongate blocks
are oriented subparallel to the contact. Some
Geosphere, December 2013
1737
Coint et al.
xenoliths are enclosed by gabbro, whereas others are enclosed by tonalite, indicating that the
blocks were isolated during emplacement of
multiple magma batches. Rock types include
migmatitic quartzofeldspathic and calc-silicate
gneiss, metaquartzite and/or metachert, and rare
skarn. Xenoliths in the upper zone occur almost
exclusively within 200–500 m of the westernsouthwestern contact. In the Ten Bear Mountain area (Figs. 2 and 5B), swarms of migmatitic quartzofeldspathic and calc-silicate gneiss
xenoliths crop out in zones as wide as 500 m.
Along the southwestern contact in exposures
in the Salmon River, xenoliths from the eastern
Hayfork terrane are abundant, range from centimeter to meter scale, and form a ghost stratigraphy within the southern selvage (Fig. 6). These
xenoliths consist of calc-silicate, metachert, and
meta-argillitic rock types, and some are migmatitic. However, most xenoliths in the southern
selvage also preserve preemplacement metamorphic fabrics, in contrast to the gneissic xenoliths
that crop out further from intrusive contacts.
INTERNAL DEFORMATION AND
FABRIC DEVELOPMENT
Structures within the batholith can be divided
into three domains based on orientation, degree
of development, and inferred conditions of formation. These domains correspond to the lithologic zonation described here. Primary structures within the lower zone include variably
developed magmatic foliations defined by the
alignment of plagioclase laths (Fig. 4A), elongate pyroxene prisms, biotite [001] faces, and
sparse ellipsoidal magmatic enclaves. Within
the lower zone near the northern contact this
fabric is locally overprinted by recrystallized
biotite and to a lesser extent hornblende (Fig.
4B). Plagioclase displays undulose extinction
and minor subgrain development. However,
in some samples interstitial quartz crystals are
undeformed, indicating that quartz grew in the
absence of significant strain before final solidification of the magma. Tonalitic rocks along the
northern (structurally lowest) contact contain a
crystal-plastic texture defined by recrystallized
biotite and quartz aligned with, and overprinting, the magmatic foliation (Fig. 4B). The foliation is subparallel to the northern contact and
varies in intensity along strike (Fig. 5). Away
from host-rock contacts, magmatic foliations
within the lower zone are generally north striking and steeply dipping, although a weak great
circle with a shallowly north plunging B-axis
may be defined (Fig. 5B). However, regional
fold patters have not been observed in the host
rocks; therefore, it is unlikely that the entire
batholith–host rock system was folded while
1738
still partially molten. Magmatic lineations are
difficult to discern, but where observed in the
field are within the foliation plane and plunge
shallowly north-south (Fig. 5B).
Foliation in the upper zone is primarily magmatic and is defined by the weak to moderate
alignment of hornblende and plagioclase. In
contrast to the lower zone, euhedral to subhedral biotite is randomly oriented and does not
define the foliation. Despite an overall subhedral shape, quartz crystals display evidence of
subgrain formation and minor recrystallization
in some samples. Poles to foliations within the
upper zone are widely distributed but generally
define a great circle with a shallowly southwest
plunging B axis (Fig. 5C). Magmatic lineations
within the upper zone plunge shallowly with a
variety of trends (Fig. 5C).
Fabric development in the central zone is
dominantly magmatic and includes foliations,
lineations, folds, and boudinage in mafic to intermediate dikes. Compared to the lower and upper
zones, foliations are generally more strongly
developed, are north-striking and steeply dipping, and are axial planar to several generations of open to isoclinally folded, synplutonic
dikes (e.g., fig. 4B in Barnes et al., 1986a).
Magmatic enclaves occur as ellipsoids and as
angular blocks aligned in the foliation plane.
Mafic dikes display various states of disruption,
pinch-swell features, and boudinage (Fig. 3E).
Late crosscutting basaltic to basaltic-andesite
dikes are gently folded about the magmatic
foliation and contain an axial-planar foliation
defined by hornblende and plagioclase phenocrysts. The axial-planar foliation is parallel to
the fabric in the host quartz-diorite to tonalite.
Other finely crystalline synplutonic dikes show
evidence for early brecciation followed by folding and boudinage.
Figure 5B displays an interpretive magmatic
foliation trend line map, based on structural
measurements throughout the batholith (Fig.
5A). Foliations in the lower zone are generally
north trending, whereas in the upper zone foliations define overlapping onion-skin patterns.
From outcrop to map scale, magmatic foliations
are slightly to strongly discordant to internal,
gradational lithologic contacts, and therefore
must postdate lithologic zonation in the area
where discordance is observed.
GEOCHEMICAL DATA
Geochronology
Multicrystal Isotope Dilution (ID)–Thermal
Ionization Mass Spectrometry Analysis
The original mapping and geochronologic
studies of the WCb and Slinkard pluton indi-
Geosphere, December 2013
cated that contacts between the zones of the
pluton were gradational and that all magmatic
units were coeval within the +4, −2 m.y. uncertainties of multicrystal thermal ionization mass
spectrometry (TIMS) dating of zircon (Barnes,
1983; Barnes et al., 1986a). Our recent work has
attempted to test these conclusions on the basis
of further detailed mapping and single-crystal
chemical abrasion (CA) TIMS dating of zircon.
Field work from 1984 to 2012 resulted in
multiple traverses across boundaries between
the lower and central zones and between the
central and upper zones, as well as boundaries
between western mafic selvages and the upper
zone. It has not been possible to identify mappable contacts, or even contact zones, in these
areas. Instead, the boundaries have been identified in terms of mineral assemblages and textures. Moreover, field evidence for magma mixing is widespread in the central zone (Barnes et
al., 1986a). In addition, the discussion herein
(and see Fig. 2C) shows that the trajectories
of magmatic foliation crosscut zone boundaries without deflection. All of these field data
indicate that the lower, central, and upper zones
of the batholith were in a magmatic state at the
same time, albeit with varying magma viscosities and crystal proportions.
New Single-Crystal ID-TIMS Data
Three samples from the lower zone, two from
the central zone, and five from the upper zone
were dated by CA-TIMS at the University of
Wyoming (see Table 1).
The analytical method is in Table 1. Crystals
were selected to avoid xenocrysts and inherited
grains; therefore, the ages obtained are considered crystallization ages. Antecrysts as defined
in Miller et al. (2007) were excluded from the
calculations as they represent inherited zircon
crystals that grew earlier within the same igneous system.
All but one of the ages reported here result
from the average of three to five individual single-crystal ages (Table 1), whereas the age of
the late-stage granite represents a single dated
zircon grain.
Two samples from the lower zone were dated:
a two-pyroxene diorite and a biotite hornblende
tonalite. Zircon grains in these samples are
large (300–600 µm) euhedral crystals that are
broadly zoned in cathodoluminescence (CL)
images. The ages are identical within the analytical uncertainty, 159.22 ± 0.10 Ma and 158.99
± 0.17 Ma.
In the upper zone, dated samples were collected from structurally lower to higher levels
and are representative of the compositional
range, from biotite hornblende tonalite to biotite hornblende granite. In all samples from the
Geosphere, December 2013
two-pyroxene diorite
20.0
373.6
25.9
58.8
11.3
143.2
16.6
681.4
WCB 1408
sA
sD
sF
sG
10.2
1.6
3.6
17.1
biotite hornblende tonalite
6.8
228.7
6.0
22.0
125.7
3.2
11.2
81.9
2.1
16.2
45.4
1.2
19.4
72.4
1.9
WCB 5109
sA
sB
sC*
sD*
sF
204
41
41
283
41
71
23
19
36
biotite hornblende quartz diorite
6.3
88.4
2.3
14
2.7
166.2
4.2
11
1.9
247.6
6.1
12
WCB 4909
sA
sC
sD
Lower zone
biotite hornblende tonalite
6.3
122.8
3.8
1.9
152.2
4.0
4.5
80.0
2.1
3.1
149.6
3.8
1.5
231.2
5.9
24
8
9
12
9
34
32
31
36
48
WCB Z5
sA
sE*
sF
sI
sJ
Central zone
biotite hornblende tonalite
8.6
158.4
4.0
4.7
267.0
6.7
2.5
494.6
12.4
4.3
329.5
8.4
6.9
270.4
6.9
30
25
13
WCB 7809
sC
sD
sG
sH
sI*
173
biotite hornblende granite
7.6
159.2
4.0
5.1
191.0
4.9
2.3
231.6
5.6
11.0
(ppm) Sample
Pb
(pg)
WCB 7909
sA
sB
sC
426.7
U
(ppm)
biotite hornblende granodiorite
2.4
306.0
8.0
19
5.1
3172.6
81.2
411
9.1
101.7
2.5
23
15.8
Weight
(µg)
WCB 2308
sB
sD
sE
Upper zone
MMB 377 sA
Late granite
Sample
17.5
2.8
1.6
2.2
2.6
0.6
0.9
1.1
2.0
1.3
0.4
1.0
7.1
0.7
0.4
0.7
1.8
0.8
0.5
0.5
3.4
1.8
1.8
1.5
0.7
0.7
10.8
0.7
3.8
Pbc
(pg)
10.9
14.1
26.3
127.6
15.6
112.0
25.2
18.0
18.0
10.9
28.9
11.4
2.8
11.5
23.2
17.2
4.7
42.4
62.7
62.8
10.6
26.1
16.7
16.4
17.1
26.5
37.4
32.5
45.1
Pb*
Pbc
0.44
0.46
0.41
0.36
0.53
0.47
0.41
0.51
0.48
0.51
0.41
0.34
0.51
0.48
0.47
0.43
0.39
0.41
0.40
0.40
0.43
0.39
0.39
0.44
0.28
0.55
0.40
0.38
0.55
Th
U
697
890
1669
8142
970
6939
1596
1117
1128
683
1834
748
192
719
1436
1091
316
2693
3967
3971
678
1662
1070
1040
1130
1624
2369
2074
2747
Pb
204
Pb
206
0.02443
Pb
238
U
(rad.)
206
(0.76)
(% error)
Pb
U
(rad.)
0.1656
235
207
Corrected atomic ratios
(0.9)
(% error)
Pb
Pb
(rad.)
0.0491
206
207
159.22 ± 0.10 Ma 95% confidence (MSWD 0.1), 4 points
0.14
0.02500
(0.08)
0.1704
(0.53)
0.0494
0.15
0.02500
(0.19)
0.1706
(0.93)
0.0495
0.13
0.02501
(0.14)
0.1714
(0.94)
0.0497
0.11
0.02502
(0.21)
0.1701
(0.30)
0.0493
158.99 ± 0.17 Ma 95% confidence (MSWD 0.29), 3 points
0.17
0.02497
(0.20)
0.1698
(1.62)
0.0493
0.15
0.02498
(0.16)
0.1690
(0.38)
0.0491
0.13
0.02505
(0.19)
0.1725
(1.00)
0.0499
0.16
0.02508
(0.19)
0.1700
(1.46)
0.0492
0.15
0.02495
(0.24)
0.1699
(1.60)
0.0494
158.30 ± 0.16 Ma 95% confidence (MSWD 0.58), 3 points
0.16
0.02485
(0.22)
0.1683
(2.31)
0.0491
0.13
0.02487
(0.13)
0.1687
(0.86)
0.0492
0.11
0.02484
(0.23)
0.1658
(2.21)
0.0484
159.01 ± 0.20 Ma 95% confidence (MSWD 1.3), 4 points
0.17
0.02500
(0.46)
0.1746
(4.47)
0.0507
0.16
0.02509
(0.31)
0.1790
(2.23)
0.0517
0.16
0.02499
(0.17)
0.1792
(1.12)
0.0520
0.14
0.02496
(0.16)
0.1743
(1.39)
0.0507
0.13
0.02496
(0.41)
0.1755
(4.97)
0.0510
158.22 ± 0.29 Ma 95% confidence (MSWD 2.4), 4 points
0.13
0.02481
(0.15)
0.1642
(0.73)
0.0480
0.13
0.02488
(0.12)
0.1684
(0.45)
0.0491
0.13
0.02484
(0.12)
0.1686
(0.42)
0.0492
0.14
0.02493
(0.98)
0.1711
(2.50)
0.0498
0.13
0.02507
(0.19)
0.1713
(1.00)
0.0495
158.25 ± 0.46 Ma 95% confidence (MSWD 1.5), 3 points
0.12
0.02485
(0.18)
0.1689
(1.47)
0.0493
0.14
0.02482
(0.22)
0.1684
(1.53)
0.0492
0.09
0.02488
(0.18)
0.1677
(1.59)
0.0489
158.21 ± 0.17 Ma 95% confidence (MSWD 0.13), 3 points
0.17
0.02484
(0.15)
0.1679
(1.0)
0.0490
0.13
0.02486
(0.16)
0.1676
(0.4)
0.0489
0.12
0.02486
(0.54)
0.1700
(0.97)
0.0496
0.17
Pb
206
Pb
208
(0.48)
(0.85)
(0.87)
(0.21)
(1.51)
(0.34)
(0.92)
(1.36)
(1.49)
(2.15)
(0.79)
(2.06)
(4.15)
(2.07)
(1.03)
(1.29)
(4.62)
(0.67)
(0.41)
(0.38)
(2.15)
(0.92)
(1.37)
(1.42)
(1.48)
(0.9)
(0.3)
(0.76)
(0.5)
(% error)
TABLE 1. CHEMICAL ABRASION–THERMAL IONIZATION MASS SPECTROMETRY U-Pb ZIRCON DATA
Pb
U
age
(Ma)
159.21
159.19
159.24
159.29
159.00
159.05
159.49
159.67
158.87
158.26
158.37
158.15
159.15
159.77
159.13
158.91
158.95
158.00
158.40
158.18
158.71
159.63
158.23
158.04
158.43
158.15
158.27
158.32
155.60
238
206
± 0.13
± 0.31
± 0.23
± 0.33
± 0.31
± 0.26
± 0.30
± 0.30
± 0.37
± 0.34
± 0.21
± 0.36
± 0.74
± 0.49
± 0.27
± 0.26
± 0.66
± 0.24
± 0.19
± 0.19
± 1.55
± 0.30
± 0.29
± 0.34
± 0.29
± 0.23
± 0.26
± 0.86
± 1.19
err
Pb
U
age
(Ma)
159.74
159.90
160.66
159.47
159.25
158.52
161.55
159.44
159.30
157.94
158.31
155.78
163.44
167.24
167.39
163.14
164.18
154.36
158.02
158.16
160.39
160.51
158.45
158.00
157.45
157.55
157.36
159.39
155.56
235
207
Pb
Pb
age
(Ma)
0.57
0.49
0.55
0.72
0.64
0.49
0.51
0.60
0.53
0.77
0.54
0.69
0.72
0.57
0.56
0.66
0.87
0.46
0.46
0.47
0.53
0.50
0.62
0.57
0.61
0.56
0.52
0.63
0.83
Rho
(continued)
167.69
170.43
181.67
162.21
162.95
150.67
192.03
155.99
165.69
153.11
157.40
119.98
226.08
274.25
285.95
225.04
240.14
98.84
152.44
157.99
185.15
173.53
161.69
157.33
142.85
148.59
143.63
175.24
154.96
206
207
Assembly of a tilted batholith
1739
1740
Weight
(µg)
(ppm) Sample
Pb
(pg)
biotite two-pyroxene diorite
1.0
362.5
9.9
11.8
101.0
2.7
15.0
76.4
2.1
6.3
89.4
2.4
12.4
43.9
1.2
22.9
38.6
1.0
10
32
31
15
15
24
Pb*
Pbc
4.6
7.5
2.3
1.4
9.5
10.5
2.9
5.4
18.9
2.2
14.9
10.5
30.8
Pbc
(pg)
4.4
2.5
2.2
11.2
3.1
4.9
3.9
1.8
1.7
14.2
1.0
1.4
0.8
0.66
0.66
0.61
0.67
0.64
0.59
0.62
0.57
0.65
0.65
0.56
0.59
0.59
Th
U
333
1128
148
888
637
1855
291
469
154
102
593
647
192
Pb
204
Pb
206
Pb
U
(rad.)
(% error)
Pb
Pb
(rad.)
206
207
(% error)
(4.33)
(1.32)
(10.4)
(1.6)
(2.4)
(0.9)
235
207
159.28 ± 0.17 Ma 95% confidence (MSWD 1.1), 5 points
0.22
0.02504
(1.96)
0.1774
(5.04)
0.0514
0.21
0.02502
(0.18)
0.1718
(1.42)
0.0498
0.20
0.02515
(0.93)
0.1771
(11.2)
0.0511
0.22
0.02503
(0.18)
0.1740
(1.7)
0.0504
0.20
0.02497
(0.23)
0.1702
(2.6)
0.0494
0.19
0.02537
(0.37)
0.1746
(1.0)
0.0499
(% error)
Corrected atomic ratios
(1.57)
(2.04)
(7.20)
(4.74)
(2.45)
(2.33)
(8.50)
Pb
238
U
(rad.)
206
158.32 ± 0.32 Ma 95% confidence (MSWD 1.5), 6 points
0.20
0.02489
(0.21)
0.1715
(1.68)
0.0500
0.18
0.02480
(0.27)
0.1614
(2.18)
0.0472
0.21
0.02476
(0.68)
0.1673
(7.72)
0.0490
0.21
0.02491
(0.42)
0.1710
(5.09)
0.0498
0.18
0.02505
(0.34)
0.1730
(2.65)
0.0501
0.19
0.02489
(0.26)
0.1703
(2.50)
0.0496
0.18
0.02483
(0.69)
0.1624
(9.11)
0.0474
Pb
206
Pb
208
Pb
U
age
(Ma)
159.40
159.31
160.15
159.40
159.00
161.48
158.50
157.91
157.64
158.60
159.51
158.46
158.13
238
206
± 3.12
± 0.29
± 1.50
± 0.29
± 0.37
± 0.60
± 0.33
± 0.43
± 1.08
± 0.67
± 0.54
± 0.41
± 1.09
err
Pb
U
age
(Ma)
165.84
160.98
165.56
162.87
159.58
163.38
160.75
151.93
157.09
160.29
162.04
159.69
152.79
235
207
Pb
Pb
age
(Ma)
258.79
185.67
243.62
213.64
168.21
191.00
194.11
59.71
148.82
185.37
199.14
178.12
70.79
206
207
0.53
0.61
0.87
0.71
0.77
0.48
0.58
0.58
0.78
0.84
0.62
0.70
0.89
Rho
Note: MSWD—mean square of weighted deviates. Sample: s—single grain; all multipoint dates are weighted mean 206Pb/238U dates; asterisk indicates sample excluded from weighted mean calculations. Weight
represents estimated weight after first step of CA-TIMS (chemical abrasion–thermal ionization mass spectrometry) zircon dissolution and is only approximate. U and Pb concentrations are based on this weight and
are useful for internal comparisons only. Picograms of sample and common Pb from the second dissolution step are measured directly and are accurate. Sample Pb: sample Pb [radiogenic (rad.) + initial] corrected
for laboratory blank. cPb: total common Pb. All assigned to laboratory blank unless >3 pg. Pb*/Pbc: radiogenic Pb to total common Pb (blank + initial). Corrected atomic ratios: 206Pb/204Pb corrected for mass
discrimination and tracer, all others corrected for blank, mass discrimination, tracer and initial Pb; values in parentheses are 2σ errors in percent. Rho: 206Pb/238U vs. 207Pb/235U error correlation coefficient. Zircon
dissolution and chemistry were adapted from methods developed by Krogh (1973), Parrish et al. (1987), and Mattinson (2005). All zircons were chemically abraded (CA-TIMS). Final dissolutions were spiked
with a mixed 205Pb/233U/235U tracer (ET535). Pb and UO2 from zircons were loaded onto single rhenium filaments with silica gel without any ion-exchange cleanup; isotopic compositions were measured in single
Daly-photomultiplier mode on a Micromass Sector 54 mass spectrometer at the University of Wyoming. Mass discrimination for Pb was 0.220 ± 0.10 %/amu for Daly analyses based on replicate analyses of NIST
SRM 981. U fractionation was determined internally during each run. Procedural blanks ranged from 3 to 0.7 pg Pb during the course of the study. U blanks were consistently <0.2 pg. Isotopic composition of the
Pb blank was measured as 18.463 ± 0.84, 15.686 ± 0.47, and 38.226 ± 1.2 for 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb, respectively. Concordia coordinates, intercepts, and uncertainties were calculated using
MacPBDAT and ISOPLOT programs (based on Ludwig ,1988, 1991); initial Pb isotopic compositions for zircon were estimated by the Stacey and Kramers (1975) model. The decay constants used by MacPBDAT
are those recommended by the Subcommission on Geochronology (Steiger and Jäger, 1977): 0.155125 x 10-9/yr for 238U, 0.98485 x 10-9/yr for 235U, and present-day 238U/235U = 137.88.
WCB 10510
sA
sB
sC
sD
sE
sG*
U
(ppm)
two-pyroxene quartz-bearing diorite
1.7
443.0
13.9
23
3.4
213.1
5.9
20
2.7
70.7
2.2
6
4.5
130.7
5.3
24
3.8
290.5
7.7
29
11.3
173.1
4.6
52
1.7
254.6
6.7
11
Western mafic selvage
WCB 2408
sA
sB
sC
sH
sI*
sK
sL
Southern mafic selvage
Sample
TABLE 1. CHEMICAL ABRASION–THERMAL IONIZATION MASS SPECTROMETRY U-Pb ZIRCON DATA (continued )
Coint et al.
Geosphere, December 2013
upper zone, zircon is elongated to equant and
varies in size between 100 and 200 µm. Truncated CL zoning is visible in some CL images,
suggesting that some zircon cores may be inherited. Ages from structurally lower to higher
samples are 158.22 ± 0.29 Ma, 158.25 ± 0.46
Ma, and 158.21 ± 0.17 Ma.
Samples from the central zone were collected
in Cuddihy Lakes basin (Fig. 2). Sample WCB4909 is a biotite hornblende quartz diorite collected in the zone of sheets described previously
(Fig. 3E), whereas Z5 is a tonalite that was
interpreted to be the youngest unit in the area
because it cuts older sheets. Zircon in central
zone samples is equant (100 and 300 µm) and
displays sharply oscillatory zoning. As in the
upper zone, truncation of CL zones in some zircon suggests that some cores might be inherited
or that Zr concentrations in the magma fluctuated, resulting in partial dissolution of already
grown crystals. Sample Z5 is 159.01 ± 0.20 Ma
and sample WCB-4909 is 158.30 ± 0.16 Ma;
these data belie the field observation that Z5 is
the youngest part of the central zone. Moreover,
overlap of the age of Z5 with ages from the
lower zone and overlap of the age of WCB-4909
with ages from the upper zone strongly suggest
that the central zone contains magma batches
from both zones of the batholith.
Two samples from mafic selvages were dated.
Zircon in these rocks is prismatic and displays
simple CL zones. Sample WCB-2408, from
the southern selvage, yielded an age of 158.32
± 0.32 Ma and sample WCB-10510, from
the western selvage, gave an age of 159.28
± 0.17 Ma; these ages overlap the upper and
lower zone ages, respectively.
Zircon from the late granite in the southern
part of the batholith (MMB-377) is prismatic to
equant. CL zones in the late granite crystals are
sharp and narrow and vary from 100 to 250 µm.
The late granite yielded a zircon age of 155.60 ±
1.19 Ma based on a single zircon analysis.
Bulk-Rock Major and Trace Elements
The following discussion builds on previous work (Barnes, 1983; Barnes et al., 1986a,
1990) with the addition of many new major and
trace element data. The complete data set is
available in the Supplemental Table1.
Samples from the lower zone of the WCb
(excluding felsic dikes) range from 46 to 56
wt% SiO2 (Figs. 8A−8F). In contrast, upper
1
Supplemental Table. Bulk-rock geochemistry.
If you are viewing the PDF of this paper or reading it offline, please visit http://dx.doi.org/10.1130
/GES00930.S1 or the full-text article on www.gsapubs
.org to view the Supplemental Table.
Assembly of a tilted batholith
25
5
MgO
lower zone
lower zone pyroxenite/melagabbro
central zone
upper zone
late-stage granite: upper zone
andesitic roof-zone dikes
dacitic roof-zone dikes
Mg-rich roof-zone dikes
20
15
20
K2O
Al2O3
4
15
3
10
10
2
5
1
5
B
A
1.2
C
1400
TiO2
700
Ba (ppm)
1.0
Sr (ppm)
1200
600
1000
500
800
400
600
300
0.8
0.6
0.4
400
0.2
200
biotite
fractionation (?)
100
200
D
0
30
F
E
30
La (ppm)
6
25
25
5
20
20
4
15
15
3
10
10
2
5
1
5
allanite
fractionation
G
Yb (ppm)
zircon
accumulation
1.5
150
0.3
1.0
100
0.2
J
55
65
SiO2
75
andesitic roof-zone dikes
dacitic roof-zone dikes
Mg-rich roof-zone dikes
P2O5
200
lower zone
lower zone pyroxenite/melagabbro
central zone
upper zone
late-stage granite: upper zone
apatite ± hornblende
fractionation
0.5
Zr (ppm)
2.0
0.5
I
H
250
2.5
45
Sm (ppm)
Y (ppm)
50
0.1
zircon
fractionation
0
45
L
0.4
55
65
SiO2
K
75
0
45
55
65
75
SiO2
Figure 8. Bulk-rock compositional data. (A–D and L) Major element variation diagrams as a function of the SiO2 content.
(E–K) Trace element variations as a function of the SiO2 content. Colored trends outline the consequence of specific mineral
fractionation on the bulk-rock data.
Geosphere, December 2013
1741
Coint et al.
zone samples range from 52 to 74 wt% SiO2
and central zone samples are, with one exception, in the 51–54 wt% range (Figs. 8A−8F).
Lower zone samples are distinct in having
higher MgO contents and lower TiO2, K2O,
P2O5 Sr, Zr, Hf, and rare earth element (REE)
contents than upper zone samples with similar
silica contents. Moreover, lower zone samples
show crudely increasing TiO2 (Fig. 8D) and
P2O5 (Fig. 8L) with increasing SiO2, whereas
these oxide abundances decrease in upper zone
samples. Pyroxenite-melagabbro blocks and
intrusions in the lower zone have higher MgO
(Fig. 8A) and CaO (not shown) and lower
Al2O3 and Sr (Figs. 8C, 8F) than the majority of lower zone samples, an indication that
most pyroxenite and melagabbro samples are
pyroxene cumulates. In most major element
plots, upper zone samples show smooth, nearly
linear variation as a function of SiO2 contents
(Figs. 8A−8D). In contrast, several trace element trends show more complicated variation with SiO2. For example, the abundances
of Zr and Hf broadly increase until ~60 wt%
SiO2 content (Fig. 8K; Hf not shown), and Ba
increases to ~64 wt% SiO2 content (Fig. 8E).
These increases are followed by decreasing
concentrations at higher SiO2 contents. Lanthanum abundances increase broadly to 62–63
wt% SiO2 content, and then remain constant
or decrease (Fig. 8G); however, Sm, Yb, and
Y abundances decrease throughout the silica
range of upper zone samples (Figs. 8H–8J).
The late-stage granitic intrusion in the southern part of the upper zone is distinct in having higher Sr, Zr, Hf, Ba, and Sm than most
samples from the upper zone.
Compositions of central zone rocks generally overlap with the low SiO2 end of the upper
zone and in several cases extend the upper
zone trend to lower SiO2 contents. Central zone
rocks are overall slightly enriched in P2O5, Sm,
and Yb compared to upper zone samples at the
same SiO2 content.
Basaltic roof-zone dikes overlap in composition with the lower zone and are most similar
to high-MgO lower zone samples, with some
overlap with pyroxenite and melagabbro blocks
(Fig. 8). In contrast, the andesitic, dacitic, and
rhyodacitic roof-zone dikes overlap with samples from the central and upper zones. Barnes et
al. (1986a, 1990) interpreted the andesitic roofzone dikes to be equivalent to lower zone rocks
on the basis of similar mineral assemblages.
However, this interpretation is contradicted
by the broad overlap of andesitic roof-zone
dike compositions with central and upper zone
compositions and the lack of correlation is confirmed by trace element analysis of augite from
the dikes and lower zone (Coint, 2012).
1742
DISCUSSION
Temporal Relationships Between Zones of
the Batholith
Geochronology has played an important role
in our understanding of the timing and pace of
pluton assembly (Coleman et al., 2004; Glazner
et al., 2004; Miller et al., 2007; Schaltegger et
al., 2009; Memeti et al., 2010; Paterson et al.,
2011; Schoene et al., 2012). Such data have
been used to assess whether multiple magma
batches were emplaced into the middle to upper
crust and solidified to form large, cohesive plutonic masses (e.g., Coleman et al., 2004; Matzel
et al., 2006). Several assumptions are made in
regard to these data. The first is that the youngest age provides the minimum crystallization
age for a given sample, which is then assumed
to represent a specific magma batch (e.g., Coleman et al., 2004). Older ages are interpreted as
resulting from inheritance, either from the same
magmatic system (antecrysts) or from the host
rocks (xenocrysts) (Miller et al., 2007). The
second assumption is related to whether intrusions assembled by multiple batch emplacement
can remain partially molten for long durations.
Intrusions of similar spatial scale and emplacement depth as the WCb are inferred to cool
through the solidus in <0.5 m.y. to >1 m.y. (e.g.,
Paterson et al., 2011). Geochronology studies
of large intermediate to silicic ignimbrites also
indicate a span of ages of ~100–300 k.y. for the
respective magma chambers (e.g., Brown and
Fletcher, 1999; Vazquez and Reid, 2002; Bachmann et al., 2007a, 2007b). Therefore, if ages
of nonxenocrystic zircons span more than 1–2
m.y., they are commonly interpreted as antecrysts inherited from older magma batches that
were reworked during pluton assembly.
As discussed here and elsewhere (Barnes,
1983; Barnes et al., 1986a), field observations
indicate that the three zones of the WCb have
gradational contacts with one another, thereby
suggesting that the three zones were comagmatic. The contact between rocks of the upper
zone and the mafic selvage along the western
margin of the pluton is also gradational. Therefore, it is expected that the ages of all of these
units should be within 100–300 k.y. In all zones,
antecrysts were identified and excluded from the
calculations (black rectangles in Fig. 9B). The
available data (Table 1; Fig. 9) show that when
individual sample dates are considered, overlap
exists between ages of samples from the lower,
one central zone sample (Z5), plus the western
mafic selvage. These overlapping ages are distinct from ages of upper zone, the other central
zone sample (WCB-4909), and the southern
mafic selvage. The lower and upper zone ages
Geosphere, December 2013
do not overlap, the latter being at least 110 k.y.
younger at the 95% confidence level, probably
~800 k.y. younger. Clearly, the lower and upper
zone magmas were emplaced relatively quickly
compared to many plutons for which high-precision age data are available (e.g., North Cascades plutons: Matzel et al., 2006; Adamello:
Schoene et al., 2012; Tuolumne: Memeti et al.,
2010). For example, assembly of the upper zone
occurred over as few as 100 k.y. (to 500 k.y.
maximum if errors are taken into account). It is
therefore not possible to use the age data alone
to determine the extent of interactions between
the different magma batches and whether two or
more batches joined to form a single connected
magma body.
Ages of the two samples from the central
zone are distinct (159.01 ± 0.20 Ma and 158.30
± 0.16 Ma); the first is coeval with the lower
zone samples, whereas the second is coeval
with the upper zone. The difference in ages of
the two central zone samples is consistent with
the sheeted nature of the zone (Fig. 3F). Evidently, despite the intense mingling and mixing
that occurred in this part of the pluton, geochronologic evidence for batch-wise emplacement is
preserved in the field and in the U-Pb ages. The
lobate contacts between sheets (e.g., Fig. 3F),
abundant mutually intrusive contacts between
sheets suggests, and the intensity of magmatic
deformation in much of the central zone indicate
that some or all of the magma batches in this
zone remained at magmatic conditions for time
intervals to ~300 k.y.
In contrast to the gradational contact relations
between the main zones of the pluton, the latestage granite has sharp, clear intrusive contacts
with upper zone rocks and its U-Pb (zircon) age
is distinctly younger, 155.60 ± 1.19 Ma.
In summary, the age data indicate that (1) the
upper and lower zones of the WCb have distinct ages, but that these distinctions are blurred
in the central zone, the latter evidently formed
from magmas from both upper and lower zones;
(2) gradational contacts between all zones, and
the presence of temporally distinct sheets in the
central zone, attest either to long-lived interstitial melts throughout the batholith or to reheating (defrosting) of lower zone and mafic selvage
rocks as upper zone magmas were emplaced.
Internal Fabric Formation and Batholith
Assembly
Host Rock–Pluton Relationship and Pluton
Emplacement
Host-rock displacement and emplacement
of magmas into the growing WCb occurred
by multiple mechanisms that varied in magnitude and significance over the duration of
Assembly of a tilted batholith
5
km
Mafic selvage
gd-gr
Upper zone
3
Kbar
ton-gd
10
qd-ton
Central zone
A
Lower zone
di-qd
7
2-mica
granite
Granite
Tonalite
Diorite
Granodiorite
Qtz diorite
Basalt-gabbro
206
Pb/
238
U date (Ma)
WCB-2408
WCB-1408
WCB-5109
WCB-4909
Z5
WCB-7809
WCB-7909
WCB-2308
MMB-377
161
WCB-10510
Late granite
B 163
162
20
160
159
158
157
156
155
154
153
Late Upper
granites zone
Central
zone
Lower
zone
Mafic selvages
Figure 9. Schematic cross section of the Wooley Creek batholith
(WCb) system associated with chemical abrasion–isotope dilution–
thermal ionization mass spectrometry (CA-ID-TIMS) data available for the different zones. The data are presented in Table 1. (A)
Schematic cross section of the WCb displaying rock type distribution throughout the intrusion. Abbreviations are as in color-coded
key. (B) CA-ID-TIMS 206Pb/238U dates (Ma) plotted as a function of
their location in the pluton. Gray rectangles represent the calculated
weighted average age for each sample (2σ error). Red bars are the
data included into the previous calculations, whereas the black bars
correspond to the grains interpreted as antecrysts.
batholith assembly. The presence or absence
of emplacement-related deformation within the
contact aureole may be due to the lithologically
diverse and structurally complex host rocks (see
discussion of Geologic Setting) as well as the
nature of heat transfer in the aureole. Because
of the discontinuous and/or disrupted nature
of the host-rock terranes, it is difficult to quantify deformation within the aureole. However,
detailed mapping within the aureole indicates
that emplacement-related deformation along
the northeastern contact includes a component
of distributed ductile flow in a zone as much as
1 km wide (Donato et al., 1982; Barnes, 1983).
Geosphere, December 2013
In contrast, along the southern and southwestern margins, dynamothermal metamorphism is
commonly absent, although a thermal aureole
is well developed and reaches at least 200 m in
width (e.g., Fig. 6); in these regions, host-rock
structures are discordant to the intrusion contact.
In some areas, particularly along the northeastern and eastern margins, lit-par-lit dike–
host rock contact zones several meters to
hundreds of meters wide are consistent with
magma emplacement by diking. These zones
are broadly concordant with the elongate eastern and western margins of the batholith (Fig.
5) and with the overall elongation of the zones
within the batholith. Thus, it is possible that the
overall assembly of the WCb was facilitated by
emplacement of north-trending (current geographic reference frame) elongate batches of
magma. Exposures of the southern WCb contacts along the Salmon River (Fig. 6) display a
variety of features that are consistent with this
interpretation, including elongate outcrop- to
map-scale north-trending intrusive sheets and
host-rock screens. Contacts between diorite and
granodiorite intrusive units in this region, also
north trending, are locally lobate and indicative
of synmagmatic recharge of new magmas into
preexisting crystal-rich mushes (e.g., Fig. 7).
The presence of xenoliths adjacent to hostrock contacts indicates that stoping occurred during assembly of the batholith. However, xenoliths
are also present in the interior of the batholith,
some in dense swarms (Fig. 5D; Barnes, 1983).
Work in progress demonstrates that these xenoliths are can be related to the three host-rock terranes on the basis of their geochemical signatures
(Barnes et al., 2011). Because these xenolith
swarms occur in the interior of the batholith as
well as near the margins, it is evident that host
rocks were incorporated into the growing batholith throughout much of its assembly.
Relationship Between Magma Batches and
Magmatic Foliation
Foliation within the intrusion is of magmatic
origin except along the northeastern contact,
where near solidus and subsolidus deformation overprinted a magmatic fabric within a few
hundred meters of the contact (Fig. 5). These
protomylonitic fabrics are attributed to local
mechanical coupling between the crystallizing
lower zone and the host rocks (i.e., ballooning). Along the western, eastern, and southern
margins of the intrusion the trends of magmatic
foliation are variably discordant to the contact.
Magmatic foliations in the central zone and
most of the lower zone strike broadly northsouth, parallel to the elongation of the batholith.
In contrast, magmatic foliations in the upper
zone display a variable orientation with a weakly
1743
Coint et al.
defined concentric pattern (Fig. 5). Moreover, in
the northwestern part of the pluton, magmatic
foliations in the upper and adjacent lower zone
rocks strike approximately east-west, locally
transect gradational compositional boundaries,
and are at high angles to the batholith–host rock
contact (Fig. 5). In general, foliation in all parts
of the batholith is formed by alignment of plagioclase and either pyroxenes (lower zone) or
hornblende (central and upper zones). In contrast, biotite is rarely oriented parallel to magmatic foliation in the upper and central zones,
and in the lower zone only thin, poikilitic biotite
crystals are in the foliation. These observations
lead to the interpretation of magmatic fabric formation in the lower zone by flow and displacement of preexisting magmas during emplacement of new magmas into the growing batholith.
Within the upper zone, the weakly concentric
foliation pattern is attributed to formation in a
mushy state prior to crystallization of biotite.
Moreover, the single trend defined by bulk-rock
data (Fig. 8) and the textural and mineralogical
homogeneity of the upper zone indicates chemical connection over the entire exposed area of
the upper zone. Therefore, we hypothesize that
the anastomosing and weakly concentric magmatic fabric pattern is the result of convective
overturn within the upper zone crystal mush
prior to biotite crystallization and final solidification of the WCb.
Bulk-Rock Data
Lower Zone: Crystal Accumulation and
Incremental Assembly
Compositional variation among lower zone
samples is most readily described as an array
of data points in which clear-cut trends are difficult to identify (Fig. 8). Overall, the expected
decrease in MgO and increase in K2O with
increasing SiO2 content is observed; however,
with the exception of the cluster of tonalitic
compositions, concentrations of elements such
as TiO2, P2O5, Sr, Zr, and the REEs show broad
scatter rather than distinct trends (Fig. 8). This
lack of a well-defined compositional trend in the
lower zone could be explained as the result of
variable proportions of cumulate minerals from
one sample to the next. It could also be related
to emplacement of multiple batches of intermediate magmas (basaltic andesite to andesite)
in which the batches were unrelated by in situ
magmatic processes. Field evidence clearly
shows that the lower zone formed from several
magma batches, thus part of the scatter observed
in the geochemical data must be related to the
presence of compositionally distinct magma
batches. Nevertheless, textural data (Fig. 4A)
and the high CaO contents of many lower zone
1744
samples are consistent with accumulation of
pyroxene ± plagioclase. In any case, both of
these conclusions differ from those in Barnes
(1983), wherein it was suggested that the lower
zone differentiated from a large magma batch
that was episodically recharged and mixed.
Evolution of the Upper and Central Zones by
Fractional Crystallization
Upper zone samples show somewhat less
scatter than is seen in the lower zone. In addition, for most elements the dacitic-rhyodacitic
roof dikes have similar compositions and plot
on the same trends as evolved samples of the
upper zone (i.e., from the structurally highest
levels of the pluton). As is the case for the lower
zone, the range of upper zone compositions and
the scatter observed in the bulk-rock data could
result from emplacement of multiple magma
batches of different composition or could result
from differentiation of a more or less homogeneous upper zone magma. In the latter case,
compositional scatter at any given SiO2 content
would result from variable amounts of cumulate
phases (e.g., Deering and Bachmann, 2010).
The consistency of hornblende compositions
and zoning patterns demonstrates the presence
of a large, homogeneous upper zone magma
body in which upward tonalite to granite zoning
developed as this magma cooled and crystallized, mainly by upward separation of evolved
melt-rich magma (Coint, 2012). This interpretation would explain the scatter observed within
upper zone samples as the effect of variable
proportions of cumulate minerals. It would also
explain the central zone as a transition zone
where lower zone magmas provided a mushy
structural base for the upper zone magma. The
structural setting would be an ideal trap for
injections of mafic magmas after which they
were variably deformed (synplutonic dikes) or
disrupted (enclave swarms). Mafic magmas that
rose through the central zone formed clusters of
mafic pillows or enclave swarms in the lowest
parts of the upper zone. The dacitic-rhyodacitic
roof dikes are explained as leaks from the upper
part of the magma body during differentiation.
If the upper zone was once a large homogeneous magma batch, then the variations of elements such as Zr, Hf, Sr, Ba, and the REEs have
implications for the parental magma composition and the nature of the differentiation process.
The trends for Zr, Hf, Ba, and La show changes
in slope when plotted against SiO2, and the
SiO2 value of these slope changes varies from
one element to the next (Fig. 8). Such changes
in slope are characteristic of fractional crystallization (Bowen, 1928). Moreover, the changes
may be used to approximate the composition of
the parental magma. For example, in the plot of
Geosphere, December 2013
Zr versus SiO2, the change in slope occurs at a
SiO2 value of ~60 wt% (Fig. 8). This silica value
restricts parental compositions to ≤60 wt%
SiO2. If the simplest case, i.e., that the upper
zone parent had 60 wt% SiO2 and ~125 ppm
Zr, is assumed, then the upper and central zone
samples with lower SiO2 and Zr would represent
cumulates (e.g., hornblende + plagioclase ± biotite, without cumulate zircon) from the parental
magma. The roof dikes and the two central zone
samples with SiO2 > 60% and Zr > ~125 ppm
would be differentiates related to those cumulates (Fig. 8K). In contrast, upper and central
zone samples with SiO2 < 60 wt% and Zr > 125
ppm would represent zircon + hornblende +
plagioclase ± biotite cumulates and the samples
that trend toward higher SiO2 and lower Zr contents would be differentiates.
Following the same logic, the upper zone
magma was saturated in apatite and a Ti-bearing
phase at the time of emplacement (decreasing
P2O5 and TiO2; Figs. 8D and 8L, respectively)
and became saturated in a Ba-rich phase (biotite
or K-feldspar) at ~64 wt% SiO2. We suggest that
biotite is the probable Ba-rich phase because
there is no commensurate decrease in K2O at
~64 wt% SiO2 and because textural evidence
shows K-feldspar to be the last phase to crystallize in all upper zone and roof-dike samples.
Variation in La (Fig. 8G) suggests that a light
REE–rich phase became stable at ~63 wt% SiO2
(allanite), but that the middle and heavy REEs
were compatible throughout crystallization of
the upper zone magma. These trends would
suggest that hornblende + apatite fractionation
controlled middle and heavy REE variations and
are consistent with the steady decrease in P2O5
with increasing SiO2 among the central and
upper zone samples (Fig. 8L). If zircon had controlled the bulk-rock heavy REE budget, Yb and
Y would show the same behavior as Zr, which is
not the case (Fig. 8).
CONCLUSIONS
Contrary to earlier interpretations (Barnes,
1983; Barnes et al., 1986a, 1990), the WCb
did not develop from a single big tank magma
body. Rather, the batholith consists of two petrologically and temporally distinct zones: a
lower zone that formed from multiple batches
of gabbroic through tonalitic magmas at 159.22
± 0.10 Ma and an upper zone of tonalite to granite emplaced at 158.99 ± 0.17 Ma. Compared to
the Tuolumne batholith (Coleman et al., 2004;
Paterson et al., 2011), the assembly of the
WCb was rapid, the bulk of the batholith being
emplaced in no more than 1.5 m.y. Moreover,
despite the distinct ages of upper versus lower
zones, their gradational mutual contacts and evi-
Assembly of a tilted batholith
dence for reworking of older into younger magmas (Coint, 2012) suggest that melt was present
in the oldest parts of the system as upper zone
magmas were emplaced. In the upper zone, the
upward compositional zoning from tonalite to
granite and the lack of internal intrusive contacts indicate that differentiation of this zone
occurred within a single large magma batch.
The central zone yielded ages suggestive of
interleaving and/or mixing of lower and upper
zone rocks and magmas. Injection of basaltic
magmas, now synplutonic dikes and enclave
swarms, into the central and upper zone magmas provided heat to sustain a long-lived
magma body. Moreover, the similarities in mineral assemblage and composition between the
upper zone and dacitic roof dikes indicate that
the upper zone was eruptible. The high crystallinity of these dikes (23%–54% phenocrysts)
indicates that eruption was possible when the
upper zone magma was a crystal mush.
The bulk compositions of many central zone
rocks indicate a petrologic affinity to the upper
zone, yet the ages of central zone rocks indicate
affinities to both upper and lower zone magmatism. The age data indicate that the central zone
is truly a transition zone that contains vestiges
of structurally high levels of the lower zone into
which numerous sheets of upper zone–like magmas were emplaced. The central zone may therefore be interpreted as part of the feeder system
to the upper zone; if so, perhaps all of the upper
zone magmas were emplaced in small increments, but these small incrementally emplaced
magmas were homogenized by convective mixing prior to development of upward zoning. The
thermal energy to drive convection is thought
to come from repeated intrusions of mafic and
intermediate magma that are locally preserved
as synplutonic dikes and enclave swarms.
The WCb is an excellent example of a magmatic system that was emplaced incrementally.
In the lower and central zones individual batches
are preserved, whereas in the upper zone they
homogenized to form a large volume of intermediate composition eruptible mush. Therefore,
timing of emplacement is not the only controlling factor on whether a large interconnected
magma chamber may form. In the case of the
WCb, homogenization of upper zone magmas and possibly triggering of roof-zone dike
emplacement probably required the added heat
provided by emplacement of mafic magmas into
the center of the batholith.
ACKNOWLEDGMENTS
We thank Monika Leopold, Samantha Buck, and
Brendan Hargrove for assistance in the field, and Dave
Atwood and Glenna Atwood for their hospitality. We
appreciate the helpful comments and advice of Olivier
Bachmann, Scott Paterson, an anonymous reviewer,
and A.E. Mike Williams. This work was supported
by National Science Foundation grant EAR-0838342
to Yoshinobu and C. Barnes, grant EAR-0838546 to
Chamberlain, a 2009 Geological Society of America
Penrose grant to Coint, and the Texas Tech University
Department of Geosciences.
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