Earth and Planetary Science Letters 309 (2011) 268–279
Contents lists available at ScienceDirect
Earth and Planetary Science Letters
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e p s l
Building an island-arc crustal section: Time constraints from a LA-ICP-MS
zircon study
Delphine Bosch a,⁎, Carlos J. Garrido b, d, Olivier Bruguier a, Bruno Dhuime c, Jean-Louis Bodinier a,
Jose A. Padròn-Navarta d, Béatrice Galland a
a
Université de Montpellier 2, CNRS-UMR 5243, Géosciences Montpellier, cc 49, 34095 Montpellier cedex 05, France
Instituto Andaluz de Ciencias de la Tierra (IACT), Facultad de Ciencias, 18002 Granada, Spain
Department of Earth Sciences, Room N° G43, University of Bristol, Wills Memorial Building, Queen's road, Bristol BS8 1RJ, England, United Kingdom
d
Department Mineralogia y Petrologia, Facultad de Ciencias, Universidad de Granada, 18002 Granada, Spain
b
c
a r t i c l e
i n f o
Article history:
Received 2 August 2010
Received in revised form 12 July 2011
Accepted 14 July 2011
Editor: R.W. Carlson
Keywords:
in situ U–Pb
island-arc
crustal growth
LA-ICPMS
zircon
Kohistan
a b s t r a c t
Geochronological studies of samples from continuous crustal sections of fossil intra-oceanic arcs are
paramount for determining the precise timing of major processes that took place during the building of an
oceanic island-arc crust. In this study, laser ablation ICP-MS U–Pb zircon (and conventional Rb–Sr) analyses
pinpoint the timing of major events responsible for crustal growth of the Kohistan paleo-island arc (Northern
Pakistan). Inheritance in magmatic rocks is evidence for recycling of intra-arc material and indicates that the
Kohistan arc may have formed as early as 135 ± 4 Ma. One older inherited grain indicates that the arc
developed on a young oceanic lithosphere whose remnants include a c. 175 Ma old component. An interesting
consequence is that intra-oceanic arc magmas can yield inherited zircons, which can be much older than the
arc system itself without requiring recycling back into the mantle by subduction processes. In the case of the
Kohistan arc, ante-arc Neotethys oceanic lithosphere was tapped by arc magmas during their way upward
into the arc crust. The oldest magmatic ages measured in this study fall in the range 101–102 Ma (100.9 ± 0.6
and 102.1 ± 0.4 Ma). This period corresponds to arc build-up and thickening of the arc crust. Leucogranitic
melts dated at 89.9 ± 0.4 Ma and 90.9 ± 1.0 Ma are considered as produced by dehydration/melting reaction
accompanying granulitisation of the thickened lower arc crust. A maximum age for this event is 97.7 ± 0.7 Ma,
the age of emplacement of a garnet meta-tonalite affected by granulite facies metamorphism. Magmatic
activity was still ongoing during and after the granulitisation as testified by emplacement of a diorite and a
gabbro dated at 89.1 ± 0.6 and 88.2 ± 0.9 Ma respectively. The upper part of the metaplutonic sequence
contains diorite samples dated at 84.6 ± 0.5 Ma and 84.3 ± 0.5 Ma and as young as 81.1 ± 0.7 Ma. During this
period, and based on the present day outcropping sequences, the mean crustal growth rate varied from 32 to
65 km3/km/Ma (volume per unit width along the strike of the arc) which is comparable to the range for
present day arcs of the Western Pacific region. The correspondence of crustal growth rates observed for
present-day island arcs and the good preservation of the crustal section in the Kohistan arc, make this area an
exceptional natural laboratory to study arc related processes and to check models of continental crust
formation by arc accretion.
© 2011 Elsevier B.V. All rights reserved.
1. Introduction
Recent studies on continental crustal growth processes and
oceanic island-arc formation have revealed that, during the past
~ 500 million years, accretion of island arcs to existing continents is
one of the main geological processes that can contribute to
continental crustal growth (Suyehiro et al., 1996; Takahashi et al.,
1998). In order to better understand and investigate this mechanism,
precise constraints on the nature and duration of the successive
⁎ Corresponding author. Tel.: + 33 467 143 267; fax: + 33 467 143 603.
E-mail address: [email protected] (D. Bosch).
0012-821X/$ – see front matter © 2011 Elsevier B.V. All rights reserved.
doi:10.1016/j.epsl.2011.07.016
processes occurring during building of a complete oceanic-island arc
system are essential. However, available information often derives
either from interpretation of discontinuous outcrops in modern arcs
formed at different times or from data extrapolated from theoretical
experiments (Tatsumi and Suzuki, 2009). In addition, studies of
modern oceanic arcs are often hampered by limited access to
outcrops, by a still active volcanism, and by hidden relationships
between the different parts of the arc system.
Conversely, fossil arcs obducted onto continental margins potentially offer an opportunity to sample continuous crustal sections and
thus, make it possible to determine precisely the chronology and the
nature of the various processes/events occurring during the arc
building stages. Such outcrops, however, are scarce worldwide and
D. Bosch et al. / Earth and Planetary Science Letters 309 (2011) 268–279
only two known examples have preserved a full oceanic arc crustal
section, i.e. from their mantle roots to the upper volcano-sedimentary
levels. These are the Jurassic Talkeetna arc in south-central Alaska (e.g.
Debari and Coleman, 1989; Greene et al., 2006; Hacker et al., 2008;
Rioux et al., 2010) and the Cretaceous Kohistan arc in Northern
Pakistan (e.g. Bard et al., 1980; Treloar et al., 1996). Through the whole
Kohistan arc complex, it is possible to identify and characterise the
main magmatic and metamorphic episodes responsible for the arc
growth. Recent petrological and geochemical analyses have revealed
multi-stage processes for the formation of this arc section indicating
that island arc growth is more complex than previously thought
(Jagoutz, 2010; Petterson, 2010 and references therein). This includes
participation of more than one source magma, local dehydrationmelting metamorphic reactions, magmatic underplating, variation of
degree of partial melting in the mantle wedge and changes in magma
source (e.g. Bignold et al., 2006; Ewart et al., 1998; Hawkesworth et al.,
1993; Peate et al., 1997). These events took place at specific moments
during the building of the Kohistan arc section, partly in response to
changes in the subduction regime parameters but also in the pressure/
thermal conditions inside the newly formed crust (e.g. Dhuime et al.,
2009; Garrido et al., 2006). In order to constrain the chronology of the
building of the Kohistan arc section, high precision geochronological
data are necessary. Such data however remain scarce (Anczkiewicz et
al., 2002; Anczkiewicz and Vance, 1997; Bouilhol et al., 2010; Dhuime
et al., 2007; Schaltegger et al., 2002; Yamamoto et al., 2005; Yamamoto
and Nakamura, 1996) and it is thus difficult to establish critical timing
relationships between events occurring during arc growth such as
major magmatic pulses, high-grade granulite metamorphism or
intracrustal differentiation episodes.
In this paper we present new geochronological data for samples
located at different levels of the Kohistan arc crustal section. Combined
with existing geochemical constraints (Dhuime et al., 2009), these new
ages highlight the complexity of arc accretion processes and allow
pinpointing the duration of events active during the building of an intraoceanic island-arc crustal section.
2. Geological setting and sampling
The Kohistan Arc Complex (KAC) crops out in northern Pakistan
(Fig. 1a). It represents the exhumed section of a Cretaceous intraoceanic arc formed during the northward subduction of the Neotethys
lithosphere beneath the Karakoram (e.g. Bard, 1983; Burg et al., 1998;
Schaltegger et al., 2002). The arc was subsequently sutured to the
Karakoram between 102 Ma and 85–75 Ma (for a review, see
Petterson, 2010), possibly at c. 85 Ma (Treloar et al., 1996). The KAC
became then the Andean-type margin of Eurasia until collision with
India that occurred at around 50 Ma (Hodges, 2000). From bottom to
top of the section (Fig. 1b) and through a SW/NE transect, the KAC
section can be subdivided into six main petrological formations (see
Zeilinger, 2002 for structural details).
The 3 km-thick Jijal ultramafic–mafic complex (Bard, 1983; Burg
et al., 1998; Garrido et al., 2006; Jan, 1979; Miller and Christensen,
1994) represents the roots of the KAC. The rocks, mainly peridotites
and pyroxenites, represent the MOHO transition zone and resulted
from melt-rock reaction between the sub-arc mantle and incoming
melts. Pyroxenes from six clinopyroxenite samples yield a Sm/Nd age
of 118 ± 7 Ma (Dhuime et al., 2007), interpreted as a minimum age
for the incipient arc building stage. The top of the Jijal section
(Fig. 1c–d) is formed by coarse-grained garnet-bearing rocks, which
suffered high-pressure granulitic metamorphism (T = 700–950 °C,
N1 GPa), during the period ranging from 91.0 ± 6.3 Ma to 95.7 ±
2.7 Ma (Anczkiewicz et al., 2002; Padron-Navarta et al., 2008;
Schaltegger et al., 2002; Yamamoto, 1993; Yamamoto and Nakamura,
1996; Yoshino et al., 1998). Field and petrological studies for this
transition zone indicate that the formation of mafic garnet granulites
was associated with amphibole dehydration melting of a gabbro-
269
noritic protolith (Bard, 1983; Garrido et al., 2006; Yamamoto, 1993;
Yamamoto and Nakamura, 2000; Yamamoto and Yoshino, 1998). Just
above, the intra-oceanic arc crustal section is represented by a thick
sequence of both metaplutonic and metavolcanic rocks (Fig. 1c–d),
composed by the Patan, Kiru and Kamila sequences (e.g., Treloar et al.,
1996, Zeilinger, 2002). This part of the crustal section shows a relatively
large range of ages (76–111 Ma) depending on the petrographical
nature of the samples and the radiometric methods used (Anczkiewicz
and Vance, 2000; Schaltegger et al., 2002; Yamamoto et al., 2005). The
intrusive Chilas Complex (Fig. 1b) was dated at 86 Ma (Schaltegger et al.,
2002). The central part of the Kohistan arc is made of the large Kohistan
batholith with ages ranging from 112 to 26 Ma (Coward et al., 1986;
George et al., 1993; Heuberger et al., 2007; Jagoutz et al., 2009; Petterson
and Windley, 1985).
In this contribution, nine samples have been studied for U–Th–Pb
geochronology and trace elements of zircons, and one sample for Rb–
Sr geochronology. They were sampled at various levels of the arc
crust, from the bottom to the top of the arc section (Fig. 1c–d). The
studied samples have previously been analysed for isotopes and
geochemistry (Dhuime et al., 2009; Garrido et al., 2006) and their
detailed petrographic descriptions are available in Supplementary
data-S1.
3. Analytical techniques
U–Pb and trace element laser ablation ICP-MS analyses and Rb–Sr
conventional analyses are available in Supplementary data-S2 and
detailed analytical techniques in Supplementary data-S3. Ages
discussed in the text are reported at the ±2σ level.
4. Results
4.1. Jijal metatonalite KG-06
Zircons extracted from this metatonalite are translucent and
elongated. They display sub-euhedral shapes and rounded terminations suggesting a metamorphic corrosion. Overall, zircons are
characterised by moderate Th and U contents (often b200 ppm) and
by moderate to high Th/U ratios (0.32–1.73). Pinned to the Stacey and
Kramers (1975) present-day common Pb composition (Fig. 2a),
thirty-one analyses (#1 to #31, Supplementary data-S2-Table 1) out
of thirty-nine define an age of 97.7 ± 0.7 Ma (MSWD = 2.2). There is
no clear relationship between Th/U ratio and ages, which would have
indicated recrystallisation processes during the high-grade event.
Other analyses (#32 to #39) have 206Pb/ 238U apparent ages spreading
between 124 Ma and 107 Ma (Fig. 2a; Supplementary data-S2-Table
1). They are interpreted as inherited grains from the source regions or
as xenocrysts snatched from neighbouring rocks during ascent of the
magma within the arc crust. It is noteworthy that the oldest
concordant grain (#37) yields an age of 115.6 ± 2.2 Ma (2σ) in
agreement with the age of the Jijal pyroxenites (Dhuime et al., 2007).
Trace elements and Ti concentrations were subsequently measured
on 10 zircons previously analysed for U–Pb and belonging to the main
~98 Ma old population. The REE patterns (Fig. 3a) are overall typical of
magmatic zircons (e.g. Hoskin and Ireland, 2000) with low La
contents (LaN ≪ 10 − 3) and prominent positive Ce anomalies. REE
patterns are characterised by steep HREE slopes (YbN/GdN = 19–45)
and moderate to prominent negative Eu anomalies (Eu/Eu* = EuN/
[SmNxGdN] 0.5 = 0.10–0.48) suggesting the zircons crystallised in a Eu
depleted environment, i.e. coeval or after plagioclase crystallisation.
The analysed zircons lack the depletion in HREE as would be expected
if they had crystallised in equilibrium with a large amount of garnet.
Considering that garnet constitutes one of the main metamorphic
minerals of this rock, this indicates that garnet grew after zircon. The
age of c. 98 Ma is thus taken as our best estimate for magmatic
crystallisation of the zircon in the tonalitic magma and is also an upper
270
Tajikistan
UM-135
UM-134
N
LA
E
W
o
KA
S
ME TA-DIORITE
UM-133
MI
35 18’N
China
Afghanistan
ME TA-GABBRO
Dasu
AMPHIBOLITE
Kamila
GRANITE
Iran
Indi a
60°E
RU
a
KI
KOHISTAN ARC
D
U
ME TA-GABBRO +
GABBRO & DIORITES
IN
N
TA
KG-31
PA
DIORITE
Karakoram plate
GRANITE
KG-18
KG-17
rn
he
KG-06
JIJ
t
Su
T)
(M K
ure
7821
Rakaposhi
No
rt
AL
KH-12a, 12b
Patan
Gilgi t
o
35 06’N
DIORITES
Kalam
Chilas
ME TA-GABBRO +
GABBRO & DIORITES
Dasu
KK
H
Patan
METAPLUTONIC COMPLEX
0
80°E
70°E
Indus Sutu re
Jijal
5000 m
(M
)
MT
M
M
T
Sapat
Jijal
Indian plate
Mingora
JIJAL COMPLEX
1:250.000
N
ULTRAMAFITES
50 Km
c
73o60’E
SW
4000
3000
2000
MMT
(Main Mantle Thrust)
72o54’E
GRANULITES
JIJAL COMPLEX
ULTRAMAFIC SECTION
MOHO
73o18’E
b
NE
METAPLUTONIC COMPLEX
MAFIC SECTION
PATAN
KG17 KG18
KH-04-12
KG06
Jijal
KIRU
KG31
KG37
KAMILA
UM-01-135
UM-01-133 UM-01-134
Kiru
Dasu
1000
Sarangar
Gabbro
0
m
d
2500 m
Fig. 1. a: Geographical position of the studied area; b: simplified geological map of the Kohistan island arc complex (modified after Burg et al., 1998); c, d: simplified geological map (c) and cross-section (d) of the Jijal-Dasu transect along the
Indus valley (KKH, Karakoram Highway) (modified after Zeilinger, 2002) showing the location of the samples studied in the present work. White stars: U–Pb analysed sample, dark star: Rb–Sr analysed sample.
D. Bosch et al. / Earth and Planetary Science Letters 309 (2011) 268–279
Kiru
Ocean
AMPHIBOLITES
KG-37
S
35o12’N
30°N
Pakistan
D. Bosch et al. / Earth and Planetary Science Letters 309 (2011) 268–279
0.12
0.10
KG06
and predated the granulitic event. Since rutile is metamorphic, we
conclude that the temperature recorded for this mineral is related to its
growth during the metamorphic event at c. 700 °C. This is in the lower
range for temperature estimate of the granulitic event (700–950 °C after
Yamamoto 1993) that affected the tonalite after its emplacement.
4.2. Grt-rich leucogranites KH04-12a and KH04-12b
These leucogranites are intrusive into the 99 Ma-old Sarangar
gabbros (Schaltegger et al., 2002), and have been sampled close to the
contact with the thick garnet-granulite unit. Both samples yield
translucent colourless grains with euhedral to sub-euhedral shapes.
0.075
Metatonalite
97.7±0.7 Ma
0.065
207Pb/206Pb
207Pb/206Pb
limit for the metamorphic event affecting this rock. Ti-in-zircon
thermometry (Ferry and Watson, 2007) yields 636–688 °C (Supplementary data-S2-Table 2) with a weighted mean temperature of 661± 15 °C
(MSWD =0.4, n =10) when individual measurements are assigned a
±25 °C uncertainty. Rutile was also analysed for its trace element
content and Zr was similarly used as a thermometer to precise the
temperature conditions of rutile crystallisation. Zr content in the six
analysed rutile grains ranges from 464 to 659 ppm, which yields tightly
grouped temperatures in the range 679–711 °C. When a ±25 °C
uncertainty is applied to each measurement, the weighted mean
temperature is 698 ± 20 °C (MSWD= 0.3, n = 6). From REE measurements it has been concluded that zircon in the metatonalite is igneous
271
KG17
Gabbro
to common Pb
(MSWD = 2.2; n = 31)
102.1±0.4 Ma
(MSWD = 1.5; n = 18)
0.08
0.055
109.6±1.4 Ma
120
0.06
0.045
a
130
0.04
46
120
50
110
54
58
62
66
70
100
90
80
d
90
100
110
74
60
70
80
238U/206Pb
238U/206Pb
0.12
KH04-12a
0.075
Leucogranite
0.065
89.9±0.5 Ma
0.10
207Pb/206Pb
207Pb/206Pb
0.14
KG18
to common Pb
Gabbro
100.9±0.6 Ma
(MSWD = 0.7; n = 22)
(MSWD = 1.2; n = 15)
0.055
0.08
98.1±1.4 Ma
(MSWD = 2.9; n = 16)
0.06
120
b
220
0.045
180
0.04
25
100
90
80
e
140
35
110
100
45
55
65
75
60
85
70
238U/206Pb
80
238U/206Pb
207Pb/206Pb
0.15
KH04-12b
Leucogranite
0.07
207Pb/206Pb
0.08
0.19
KG31
Diorite
89.1±0.6 Ma
(MSWD = 0.7; n = 11)
0.06
0.11
97.4±0.9 Ma
90.9±1.0 Ma
(MSWD = 1.6; n = 27)
0.07
(MSWD = 0.8; n = 5)
0.05
c
112
0.03
56
108
94
104
60
100
64
96
92
68
88
72
84
76
238U/206Pb
80
0.04
92
90
88
f
68.5
70.5
72.5
238U/206Pb
Fig. 2. Tera-Wasserburg concordia diagrams for LA-ICP-MS zircon analyses from rocks of the Kohistan Arc Complex. a) KG-06 metatonalite; b) KH04-12a leucogranite; c) KH04-12b
leucogranite; d) KG-17 gabbro (Patan complex); e) KG-18 gabbro (Patan complex); f) KG-31 diorite (Kiru complex); g) UMO-133 diorite (Kamila complex); h) UMO-134 diorite
(Kamila complex); i) UMO-135 diorite (Kamila complex). All ages have been anchored to a present-day common Pb composition taken from the model of Stacey and Kramers (1975)
and equivalent to a value of 0.84208 ± 5% (207Pb/206Pb). Error crosses are ±1σ.
272
D. Bosch et al. / Earth and Planetary Science Letters 309 (2011) 268–279
UM01-133
mmon Pb
to SK co
207Pb/206Pb
0.085
Diorite
84.3±0.5 Ma
0.065
(MSWD = 1.1; n = 19)
0.045
110
100
90
80
g
75
65
85
Pb
mmon
UM01-134
Diorite
to SK co
0.085
207Pb/206Pb
238U/206Pb
84.6±0.5 Ma
0.065
(MSWD = 1.1; n = 11)
0.045
110
100
90
80
h
65
75
85
comm
on Pb
UM01-135
Diorite
to SK
0.085
207Pb/206Pb
238U/206Pb
0.065
81.1±0.7 Ma
(MSWD = 0.8; n = 23)
0.045
110
100
90
80
i
65
75
85
238U/206Pb
Fig. 2 (Continued).
Some grains yield a euhedral core showing embayment related to a
magmatic resorption (Supplementary data-S4A-B). U–Pb analyses
show that the two samples have a very close age distribution (Fig. 2b–
c), apart from inherited grains, which have ages of ~ 175 Ma and
~ 135 Ma (only for sample KH04-12a) or around 105–95 Ma (both
samples). Among the forty-two analyses performed on zircons from
sample KH04-12a, eleven define an age of 98.1 ± 1.4 Ma
(MSWD = 2.9) and twenty-two show a tight cluster yielding an age
of 89.9 ± 0.5 Ma (MSWD = 0.7). Sample KH04-12b also shows two
main groups defining ages of 97.4 ± 0.9 Ma (MSWD = 1.6, n = 27) and
90.9 ± 1.0 Ma (MSWD = 0.8, n = 5). For KH04-12a no significant
variation can be detected between the Th/U ratios measured for
both groups of zircons, ranging from 0.15 to 0.63 for the c. 90 Ma old
population and from 0.16 to 0.54 for the c. 98 Ma old population
(Supplementary data-S2-Table 1). Conversely, the Th/U ratios
measured for zircons from KH04-12b are slightly distinct between
the two groups of grains. Old grains have Th/U ratios ranging from
0.23 to 0.76 whereas younger grains tend to have lower Th/U values
ranging from 0.06 to 0.43. The cathodoluminescence images of zircons
from KH04-12a and KH04-12b display oscillatory zoned cores
surrounded by zoned overgrowth (Supplementary data-S4). Such
structures are consistent with the age distribution observed in both
samples and tentatively suggest that the oldest ages at c. 97–98 Ma
are related to inheritance from magmatic protoliths subjected to
partial melting which is dated at c. 90 Ma. The main group of inherited
grains has ages comparable to the dated metatonalite KG-06 (97.7 ±
0.7 Ma) or to the 99 Ma Sarangar gabbro (Schaltegger et al., 2002),
both units being intruded by the felsic veins.
In sample KH04-12a one big zircon (N200 μm), in textural contact
with garnet in a thick-section (Fig. 4), was investigated in situ for
U–Pb and trace element (including Ti) analyses. Seven spots were
performed for U–Pb analyses (Supplementary data-S2-Table 1) and
they display a complex age distribution. Four data points are close to
concordant at 91–93 Ma, with the youngest concordant point (#72-7)
providing an age of 90.9 ± 1.8 Ma (2σ), identical to the 89.9 ± 0.5 Ma
age obtained for out of context grains from this sample. It is likely that,
except spot #72-7, all analyses contain a small amount of inherited
lead whose age in this grain reaches a value of 118.4 ± 4.0 Ma (2σ).
This substantiates the hypothesis mentioned above that the leucogranitic magma crystallised at c. 90 Ma. The most precise age of 89.9 ±
0.5 Ma is thus adopted as our best estimate for crystallisation of
leucogranite KH04-12a, while older ages date inherited material
which is dominated by c. 97 Ma-old zircon grains but includes older
grains, up to 175 Ma-old.
Trace elements for the zircon grain observed in thick section
have been analysed at the location site of spot #72-7 (see Fig. 4). The
rare-earth element pattern shows a low LREE content, a positive
Ce anomaly and a negative Eu anomaly. The latter indicates crystallisation from an environment depleted in Eu due to feldspar
crystallisation. The REE pattern is also characterised by a flat HREE
pattern (YbN/GdN = 0.8), which is consistent with crystallisation in
equilibrium with the neighbouring garnet. The Ti content of the grain
is 9.2 ppm which translates into a temperature of 739 °C. Zircons
mounted in epoxy resin and belonging to the main magmatic
population were also analysed for their trace element contents
(Fig. 3b). For both samples (KH04-12a and 12b), REE patterns are
characterised by a flat to depleted HREE (YbN/GdN ranging from 0.6 to
3.9) indicating crystallisation in equilibrium with garnet, and by
variable Eu anomalies (Eu/Eu* ranging from 0.26 to 0.97). The Ti
contents of the grains range from 4.9 to 6.6 and to 4.5 to 7.0 for KH0412a and KH04-12b respectively which corresponds to Ti-in-zircon
temperatures of 697 ± 20 °C (n = 6; MSWD = 0.13) and 696 ± 22 °C
(n = 5; MSWD = 0.37). These weighted means are not significantly
different from the single value of 739 ± 25 °C obtained for the grain
analysed in the thick section and are interpreted as corresponding to
the zirconium saturation temperature of the leucogranitic magma
and crystallisation of the magmatic zircons in equilibrium with
garnet.
4.3. Patan gabbros KG-17 and KG-18
Zircons are translucent, euhedral with no visible inclusions (Supplementary data-S4C). In the concordia diagram (Fig. 2d–e), zircons from
both samples define a very restricted distribution with ages of
respectively, 102.1±0.4 Ma (n=18, MSWD=1.5) and 100.9±0.6 Ma
(n=15, MSWD=1.2). This falls in the same age range as that obtained for
the crustal section of Sapat (Bouilhol et al., 2010). No inheritance has been
detected except for one grain from sample KG-17 (#124-2, Supplementary data-S2-Table 1) that provided a slightly older 206Pb/ 238U
D. Bosch et al. / Earth and Planetary Science Letters 309 (2011) 268–279
10000
1000
a
b
Chondrite normalized
Chondrite normalized
1000
100
10
1
0.1
Metatonalite KG06
(Jijal Complex)
0.01
0.001
10000
Ce
Pr
Nd Sm Eu Gd
Tb
Dy
Ho
Er
Tm Yb
Leucogranite KH04-12a ( ) & KH04-12b ( )
(Sarangar)
La
Chondrite normalized
Chondrite normalized
0.1
Ce
Pr
Nd Sm Eu
Gd
Tb
Dy
Ho
Er
Tm Yb
Lu
d
1000
10
1
Gabbro KG17
(Patan sequence)
100
10
1
0.1
Gabbro KG18
(Patan sequence)
0.01
0.001
0.01
La
Ce
Pr
Nd Sm Eu Gd
Tb
Dy
Ho
Er
Tm Yb
Lu
La
1000
e
1000
Ce
Pr
Nd Sm Eu Gd
Tb
Dy
Ho
Er
Tm Yb
Lu
f
100
Chondrite normalized
Chondrite normalized
1
10000
c
0.1
100
10
1
0.1
Diorite KG31
(Kiru sequence)
0.01
10
1
0.1
Diorite UMO133
(Kamila sequence)
0.01
0.001
0.001
La
Ce
Pr
Nd Sm Eu Gd
Tb
Dy
Ho
Er
Tm Yb
La
Lu
Ce
Pr
Nd Sm Eu
Gd
Tb
Dy
Ho
Er
Tm Yb
Lu
10000
g
1000
1000
Chondrite normalized
Chondrite normalized
10
Lu
100
10000
100
0.01
La
1000
10000
273
100
10
1
0.1
Diorite UMO134
(Kamila sequence)
0.01
h
100
10
1
0.1
Diorite UMO135
(Kamila sequence)
0.01
0.001
0.001
La
Ce
Pr
Nd Sm Eu Gd
Tb
Dy
Ho
Er
Tm Yb
Lu
La
Ce
Pr
Nd Sm Eu Gd
Tb
Dy
Ho
Er
Tm Yb
Lu
Fig. 3. Chondrite normalised REE patterns for zircons from different samples. a) KG-06 metatonalite; b) KH04-12a and KH04-12b leucogranite; c) KG-17 gabbro (Patan complex); d) KG-18
gabbro (Patan complex); e) KG-31 diorite (Kiru complex); f) UMO-133 diorite (Kamila complex); g) UMO-134 diorite (Kamila complex); h) UMO-135 diorite (Kamila complex).
D. Bosch et al. / Earth and Planetary Science Letters 309 (2011) 268–279
100 µm
0.019
Grt
206Pb/238U
274
KH04-12a
120
118.4±4.0 Ma
110
0.017
100
0.015
90
207Pb/235 U
#72-2 & 72-7
0.095
0.105
0.115
0.125
Zrn
Sample/Chondrite
100
10
1
Leucogranite KH04-12a
Zircon
0.1
Garnet
La
Ce
Pr
Nd Sm Eu Gd Tb
Dy Ho
Er
Tm Yb
Lu
Fig. 4. Composite figure: a) a natural light photograph of a thick section from sample KH04-12a showing a c. 200 μm zircon grain in contact with garnet; b) a concordia diagram for
LA-ICP-MS analyses of this zircon grain; c) Chondrite normalised REE patterns of the zircon grain and neighbouring garnet.
apparent age of 109.6±1.4 Ma (2σ). Both ages are interpreted in terms of
magmatic crystallisation of the gabbroic magmas and, inheritance set
apart, are the oldest obtained in this study. The slightly older age of
analysis #124-2 is close to the zircon age of a dioritic dyke dated at
Magmatic ages
107.7 ± 1.8 Ma in the higher, Kiru sequence (Yamamoto et al., 2005).
Other records of ages older than 102 Ma have been detected in mafic
samples, such as a granulite from the Jijal unit (Yamamoto and
Nakamura, 2000), the clinopyroxenites from the base of Jijal
this study
Inheritance
Chilas
Kamila
sequence
Magmatic ages
Metamorphic ages
litterature
Inheritance
Kiru
sequence
Patan/Sarrangar
sequence
Jijal Complex
Collision with
Neotethys
180
Age (Ma)
160
Granulitic
event
Beginning of arc build-up, crustal
growth and thickening
140
120
100
Early calc-alkaline
magmatism (Kohistan batholith)
EURASIA
INDIA
(minimum age)
80
60
40
Major thermal
modification
Fig. 5. Distribution of ages along the crustal section of the arc including the Jijal Complex, the Patan, Kiru and Kamila sequences and the Chilas Complex. Data from the literature are
reported after: Anczkiewicz and Vance (1997); Anczkiewicz et al. (2002); Dhuime et al. (2007); Schaltegger et al. (2002); Yamamoto and Nakamura (1996), (2000); and Yamamoto
et al., 2005.
D. Bosch et al. / Earth and Planetary Science Letters 309 (2011) 268–279
(Dhuime et al., 2007) and a foliated amphibolite xenolith in a massive
granitic body intrusive into the Kamila sequence (Yamamoto et al.,
2005).
Zircons from both samples provide REE patterns (Fig. 3c, d) with
pronounced positive Ce anomalies and negative Eu anomalies (0.33–
0.41 for KG-17 and 0.32–0.64 for KG-18), the latter being consistent
with zircon crystallisation after plagioclase. For sample KG-17, the Ti
content of the analysed grains is variable, ranging from 2.6 to 7.2 ppm
which indicates crystallisation in the temperature interval of 632–
716 °C (n = 5). The analysis (#Zr4) with the highest Ti content
(7.2 ppm) also yields a high Ca content suggesting the laser beam
struck an inclusion. This analysis was removed from the Ti dataset
which restricts the crystallisation interval to 632–690 °C (n = 4).
The mean temperature for these four analyses is 654 ± 24 °C
(MSWD = 1.2) which we interpret as corresponding to the crystallisation of the zircons in the gabbro KG-17. Analysed zircons from KG18 have almost the same range of Ti content (1.8–8.6 ppm) and a
crystallisation temperature interval of 610–731 °C. A weighted mean
of all data yields a mean temperature of 668 ± 28 °C (MSWD = 2.4).
The high MSWD value suggests a scattering of the data. However,
examination of the dataset does not reveal any problem in the
measurements (such as the occurrence of Ti-rich inclusions) and we
conclude that the c. 120 °C temperature range is real and is likely to
reflect the crystallisation interval of magmatic zircon in this rock,
which in turn suggests a rather slow cooling rate of the magma.
4.4. Kiru diorite KG-31 and gabbro KG-37
Zircons from diorite KG-31 are euhedral and translucent. The LAICP-MS U–Pb analyses performed on eleven zircon grains (Fig. 2f)
define an age of 89.1 ± 0.6 Ma (MSWD = 0.7, n = 11). Zircons show
REE patterns (Fig. 3e) characterised by pronounced negative Eu
anomalies (EuN/Eu*N = 0.35–0.77) and steep HREE slopes (YbN/
GdN = 31–58). The Ti content of these grains is very homogeneous,
between 6.8 and 9.3 ppm (n = 5) which provides a fairly restricted
range of temperature of 710–740 °C. The mean Ti-in-zircon temperature is 725 ± 22 °C (MSWD = 0.3), which is interpreted as corresponding to the crystallisation of the zircon in the dioritic magma
during a rather short temperature interval.
No zircons were found from gabbro KG-37, but Rb-Sr analyses
attempted on the whole rock and on pure clinopyroxene and
plagioclase fractions define an internal isochron (Supplementary
data-S6 and S2-Table 3) yielding an age of 88.2 ± 0.9 Ma.
4.5. Kamila diorites UM01-133, UM01-134, UM01-135
These three dioritic samples have been collected in the upper level of
the Kamila section. The diorite UM01-135 has been sampled close to the
contact with the Chilas Complex. Zircons from these three samples
exhibit euhedral shapes consistent with a magmatic crystallisation
(Supplementary data-S4D). They have simple age distributions with
concordia intercepts at 84.3 ± 0.5 Ma, 84.6 ± 0.5 Ma and 81.1 ± 0.7 Ma,
for UM01-133, -134 and -135 respectively (Fig. 2g, h, i; Supplementary
data-S2-Table 1). These ages are the youngest obtained during this study
and are broadly coeval with or slightly younger than the neighbouring
Chilas intrusion dated at 85.7 ± 0.15 Ma (Schaltegger et al., 2002). Trace
element analyses provide REE patterns typical of magmatic zircons
(Fig. 3f, g, h) with positive Ce anomalies and large Eu anomalies (EuN/
Eu*N ranging from 0.31 to 0.63, from 0.21 to 0.24 and from 0.28 to 0.40
for UM01-133, -134 and -135 respectively). UM01-134 and UM01-135
have the flattest HREE patterns with YbN/GdN = 10–24 and 12–27
respectively, while sample UM01-133 has YbN/GdN ranging from 28 to
43. Ti-in zircon temperatures yield identical weighted mean values of
812 °C for UM01-134 and UM01-135 although zircons from UM01-135
yield more scattered values, ranging from 750 to 855 °C instead of 778 to
832 °C for UM01-134. One grain from UM01-135 (analysis #Zr6; Table
275
2) yields a core to rim variation with a Ti-in-zircon temperature of 855 °C
for the internal part and of 815 °C for the edge of the grain. A third
intermediate spot yields a temperature of 842 °C. Interestingly, this is
correlated with a decrease of the Th/U ratio from 1.13 to 0.65 and an
increase of the Eu negative anomaly (EuN/Eu*N increasing from 0.28 in
the centre to 0.35 in the edge). The decrease of the Th/U ratio is at odd
with a simple trend of magmatic differentiation, and rather suggests the
influence of a mineral phase fractionating Th against U (e.g. apatite). The
increase of the Eu negative anomaly is consistent with zircon crystallisation in an environment where the occurrence of feldspar is becoming
more and more important or with a variation of the magmatic oxidation
state (Ballard et al., 2002), where magma conditions are becoming more
reducing. The large spread of Ti-in-temperatures for this sample (about
100 °C) indicates that Zr saturation was reached or maintained during a
long temperature interval, which supports a slow cooling rate. Zircons
from sample UM01-133 yield significantly lower crystallisation temperatures from 660 to 737 °C with a weighted mean value of 705 ± 27 °C
(n= 7; MSWD = 1.4)
5. Discussion
5.1. Time constraints on the building of the Kohistan arc crustal section
5.1.1. Evidence for the inception of the building of the Kohistan arc
section and the significance of inheritance in the studied zircons
An important issue concerning the building of the KAC section is
the time at which the subduction of the Neotethys oceanic crust
responsible for the build-up of the Kohistan arc began. This has
implications on the lifetime of the arc and thus on the rate of the
processes that were active in the KAC. Although we have no direct
constraints, this can be addressed following two complementary
ways.
– Pyroxenites from the basal Jijal ultramafic zone have boninitic
trace element affinities (Garrido et al., 2007), a geochemical
feature ascribed to the percolation of mobile element rich magmas
into remnants of the ante-arc lithosphere located in the fore-arc
zone during the onset of the subduction zone activity (e.g. Stern,
2004). A 118 ± 7 Ma Sm/Nd internal isochron age obtained on
these pyroxenites was interpreted to date the first episodes of slab
dehydration fluids metasomatism of the lithospheric mantle
(Dhuime et al., 2007) and is therefore a minimum age for incipient
subduction of the Neotethys at the onset of the building of the
Kohistan intra-oceanic arc. The mean time interval necessary to
transfer fluids from the slab to the mantle wedge, where they
induce partial melting and the subsequent magma production is
estimated, on the basis of Be and U/Th isotopes, between 2 and
3 Ma to highly shorter times (i.e. 30 Ka) assuming that slab
dehydrates more efficiently in zones of amphibole destabilisation,
i.e. between 50 and 150 km depth (Elliott et al., 1997; Hawkesworth
et al., 1997). This time interval (b3 Ma) is well within error margin of
the Sm–Nd age (Dhuime et al., 2007) and is considered without
influence on this time constraints. On this basis, a minimum age
for initial build-up of the Kohistan arc is thus in the time interval
125–111 Ma.
– The age of inherited or xenocrystic zircons in the studied rocks
constitutes another source of information that can usefully
complement the above constraint (Fig. 5). In the crystals where
inheritance has been detected, two groups can be distinguished: a
first with ages younger than or within error to 118 ± 7 Ma, i.e. in
the age range for initiation of the subduction; and, a second group,
with ages significantly older than 118 ± 7 Ma. Noteworthy, the
first group of inherited grains are almost exclusively located in the
transition zone (petrologic Moho) between the basal mantle rocks
(Jijal complex) and the overlying crustal metaplutonic section.
These grains have been found in tonalite KG-06 (eight grains
276
D. Bosch et al. / Earth and Planetary Science Letters 309 (2011) 268–279
ranging from 124.3 ± 4.8 Ma to 107.1 ± 1.4 Ma), and leucogranite
KH04-12a (116.7 ± 4.6 Ma) and KH04-12b (118.4 ± 4.0 Ma). Ages,
in the range 110–120 Ma, have been previously obtained for
various lithologies, in particular for a gabbronorite located on top
of the Grt-granulite zone (Yamamoto and Nakamura, 2000), and
from a deformed amphibolite xenolith in a Kiru granite (Yamamoto
et al., 2005). As inherited ages from the first batch are not
significantly older than the inferred age of initial arc build-up, i.e.
118 Ma (Dhuime et al., 2007), they are thought to reflect intra-arc
assimilation in the plumbing system of the arc by magmas
participating to the construction of the arc section.
The second batch of analyses (significantly older than 118 ± 7 Ma)
is made of grains from leucogranite KH04-12a and includes one grain
dated at 174.9 ± 9.4 Ma and two grains with ages of about 130 Ma
(analyses #38 and #39 at 131.8 ± 13.0 and 134.8 ± 3.8 Ma). Analysis
#38 is within error of the first batch of analyses and cannot be
unambiguously regarded as belonging to either group. Analysis #39 is
slightly but significantly older than 118 ± 7 Ma and may indicate that
subduction and related magmatism started as early as 135 Ma. In the
Izu–Bonin–Mariana arc system, McPherson and Hall (2001) indicated
that boninitic magmatism is not necessarily a characteristic of infant
subduction zone, but reflects interaction between a subduction zone
and a thermal anomaly, thus supporting the idea that, in the case of
the Kohistan arc, boninitic material may have been produced after the
initiation of subduction.
The last analysis (174.9 ± 9.4 Ma) is significantly older and we
propose four alternative possibilities:
1) One possibility is that this grain reflects inheritance from material
from the subducting slab (oceanic lithosphere and its sedimentary
cover) that has survived both partial melting and a subsequent
travel through the mantle wedge. Recent modelling studies
(Hermann and Rubatto, 2009) indicate that accessory minerals
such as zircon and monazite are expected in the residue up to
temperatures of 850–900 °C and up to 50% partial melting
affecting plunging slabs and their sedimentary cover, thus
sequestrating their trace element contents deep into the mantle.
However the likelihood that these accessory minerals survive in
the low silica activity magmas in the mantle wedge has not yet
been demonstrated and is considered unlikely.
2) Another possibility could be that it reflects crustal inheritance
from either the Indian or Karakoram margins. Inheritance from the
Indian margin is unlikely given that collision between India and
Kohistan is thought to occur at around 50 Ma (Hodges, 2000), i.e.
well after formation of the garnet-bearing leucogranite that
contains this xenocryst. Although the age of collision between
KAC and Karakoram is not well constrained (i.e. between 102 ±
12 Ma and 75 Ma, Petterson and Windley, 1985) it is considered to
happen most likely around 85 Ma (Petterson, 2010; Treloar et al.,
1996 and references herein), which indicates that Kohistan was
still separated from Eurasia at the time the garnet-bearing
leucogranite formed (90 Ma). Moreover, the timing of calcalkaline arc volcanism in the Karakoram extends back to
~120 Ma (Heuberger et al., 2007) but not to 175 Ma.
3) It is also possible that the 175 Ma old xenocryst reflects initiation
of the arc during the Aalenian. Although this hypothesis cannot be
ruled out, it is noteworthy that the oldest sedimentary units
preserved in the KAC are mid-Cretaceous (Petterson and Treloar,
2004) with the Yasin Group being Albian–Aptian (Pudsey et al.,
1986). This fits better with the 130–135 Ma ages provided by other
xenocrystic zircons. Assuming a 175 Ma age for initiation of the arc
requires that sedimentation from Aalenian to Aptian has not been
preserved, which is at odd with the excellent preservation state of
the whole crustal section in the KAC, or did not occur. The latter
implies that no topographic highs were created during the time
interval spanning from the Jurassic to the mid-Cretaceous.
4) Finally, such old zircons may come from fragments of the ante-arc
Neotethys oceanic lithosphere preserved in the arc section and
partly assimilated during ascent of the magmas through the arc
crust. Remnant of such an old lithosphere has been recovered as a
lherzolite lens (200 × 25 m) located at the top of the Jijal Sequence
(sample KH04-112 of Dhuime et al., 2007) and, although we have
no age constraints on this rock, we consider likely that such relics
may have provided the 175 Ma old grain. From a U–Pb systematic
point of view, an interesting consequence is that magmas formed
in an intra-oceanic arc setting can yield inherited zircons, that,
depending of the age of the oceanic lithosphere on top of which
the arc grew, can be much older (~ 40 to 60 Ma) than the arc
system itself. This material is trapped into the arc and preserved
until it is tapped by arc magmas on their way upward into the arc
crust. This does not necessarily require that such old material had
been recycled back into the mantle by subduction processes and
survived a transfer through the mantle wedge.
5.1.2. Magmatic episodes, thickening of the KAC crustal section and age
of the high-grade metamorphism
Based on an extensive geochemical study, Dhuime et al. (2009)
proposed that magmatic rocks of the Jijal sequence and metaplutonic
Complex (Patan, Kiru and Kamila sequences) belong to two main
geochemical suites (A and B), both with island arc tholeiite geochemical
features. Suite A corresponds to the volcanic arc growth and reflects
melting of the ambient asthenosphere with various degrees of slab melt/
slab fluid components. Suite B is considered as magmatic underplating
events and is characterised by an increase of a sedimentary component
compared to suite A. On the basis of isotope signatures it has been
divided into three stages (B1 to B3) reflecting, through time, an increase
of the subduction erosion of fore-arc metasediments.
The oldest evidence for magmatic activity obtained from this study
comes from the two gabbros sampled in the lower part of the crustal
section of the Patan zone. They yield ages of 101–102 Ma, which are very
close to the c. 99 Ma age of the underlying Sarangar gabbro (Schaltegger
et al., 2002). Well-constrained absolute ages, ranging from 112 to
100 Ma, have been previously published (Heuberger et al., 2007;
Yamamoto et al., 2005; Yamamoto and Nakamura, 2000) and
interpreted as corresponding to the main building stage of the arc
section and the climax of magmatism in the KAC. This suggests the
occurrence of successive stages of ponding of magmas at the base of the
crust. The two studied samples (KG-17 and KG-18) belong respectively
to type B1 and B2 of the geochemical suite B (Dhuime et al., 2009).
Intrusion of these magmas was thought to occur between 105 and
99 Ma (Dhuime et al., 2009), which is in agreement with the ages of the
two dated samples.
Zircons extracted from tonalite KG-06 yield an age of 97.7 ±
0.7 Ma, interpreted as a maximum age for high-grade granulite facies
metamorphism. This is in the upper range of the 91–96 Ma interval for
granulitisation of the crustal section of the Jijal Complex (e.g.
Anczkiewicz et al., 2002; Yamamoto, 1993). Garnet-bearing composite samples from the Jijal garnet granulite unit record the arrested
transformation of hornblende-gabbronorite to garnet granulite with
intensive melting–dehydration processes involving coeval breakdown of orthopyroxene and amphibole and the formation of garnet
and quartz (Garrido et al., 2006; Yoshino and Okudaira, 2004).
Padron-Navarta et al. (2008) proposed that the meta-gabbronorite
were transformed into high-pressure garnet granulite only with
substantial compression allowing a pressure increase from 0.5 to
1.1 GPa. Looking carefully at the distribution of ages obtained during
this study from the base of the crustal section (Jijal unit) to the top
(Patan-Dasu complex and Kamila amphibolites), a general younging
of ages can roughly be observed (Fig. 5). This supports an overall
upward vertical accretion of the arc crust with younger magmas
intruding the upper part of the crustal column. We propose that
underplating of magmas and their subsequent emplacement in the arc
D. Bosch et al. / Earth and Planetary Science Letters 309 (2011) 268–279
crust was responsible for the gradual thickening of the crustal section.
This allowed an increase of the pressure and temperature conditions
of the base of the section and triggered the melting/dehydration of the
precursor gabbronorite, thus producing garnet-bearing granulites at
depth. The fact that granulitisation lasted from 91 to 96 Ma is
regarded as reflecting the downward migration of the crust and the
progressive “per descensum” granulitisation of the crustal units now
outcropping above the Jijal ultramafics. The age of leucogranites
KH04-12a and KH04-12b (c. 90–91 Ma) indicates that highly
differentiated melts were produced at the end of (and possibly
during) the granulite-facies event. This is consistent with melting/
dehydration processes producing garnet bearing granulites at depth
and granitic material as a side-product of the reaction (Garrido et al.,
2006). Although the volume of such material is not known, these
intra-arc differentiation processes, associated with delamination of
ultramafic keels (Behn and Kelemen, 2006; Jull and Kelemen, 2001)
may drive the global composition of the arcs towards more andesitic
values, akin to bulk continental crust thus potentially helping solve
the “arc paradox” (Garrido et al., 2006; Kuno, 1968).
5.1.3. Final steps of the intra-oceanic KAC evolution: major thermal
regime modification and Chilas intrusion
Magma production at depth and underplating were still ongoing
after the granulitisation of the deep arc crust as testified by the age of
diorite KG-31 (89.1 ± 0.6 Ma) and gabbro KG-37 (88.2 ± 0.9 Ma) and
by available data such as the 91.8 ± 1.4 Ma age of a foliated diorite
near Kiru (Schaltegger et al., 2002). Samples KG-37 and KG-31 belong
respectively to type 1 and type 3 of the geochemical suite A defined by
Dhuime et al. (2009). The geochemical difference between these two
types of magma corresponds to a more pronounced influence of a
slab-fluid component in type 1 magmas whereas type 3 magmas yield
influence of a slab-melt component. Since both magmas are coeval
within errors, it indicates that melting of hydrated peridotites of the
mantle wedge triggered by the release of a slab fluid (type 1) or, at
deeper levels, of a slab melt (type 3) component, occurred
simultaneously at different levels of the subduction zone. The main
difference with the model of Dhuime et al. (2009) is that A-type
magmas are not restricted to the first stages of arc build-up and that
production of such magmas from partially molten hydrated peridotites of the mantle wedge is a persistent phenomenon that lasted at
least until 88 Ma.
One event is at odd with a continuous magmatic production and
indicates pulses in the magma generation. Indeed, at c. 86 Ma, the KAC
underwent a major event with emplacement of the huge Chilas
Complex (c. 300 × 40 km). Schaltegger et al. (2002) proposed that the
Chilas magmas tapped a mantle reservoir, which was either
metasomatically enriched or contaminated by a sedimentary component. The latter is similar to the process envisioned by Dhuime et al.
(2009) for the B suite. Age constraints on B1 or B2 series are given by
samples KG-17 and KG-18 dated respectively at 102 and 101 Ma. B3
(UMO-133; -134 and -135) is bracketed between 85 and 81 Ma. The
gap in age (101 to 85 Ma) is consistent with distinct magmatic pulses.
Chilas magmatic rocks have isotopic signatures (εNdt = + 4.7–+6.1;
Jagoutz et al., 2006) more mantellic than the contemporaneous B3
suite (εNdt = +1.9–+3.5 and 87Sr/ 86Srt = 0.70378–0.70415; Dhuime
et al., 2009) implying a greater contribution of asthenospheric
material. Since the Chilas complex is related to intra-arc extension
(Burg et al., 1998) associated to an increase of the magmatic
productivity we propose, by comparison with numerical models,
that production of the Chilas magmas was related to lower crustal
delamination (Jull and Kelemen, 2001) and/or thermo-mechanical
erosion (Arcay et al., 2006) responsible for a major mantle upwelling.
The B3 suite (85–81 Ma) originated from a reservoir similar to B1 and
B2 suites (c. 100 Ma), only modified by a higher contribution of
sediments. This suggests that the conditions for B magma production
were still ongoing during and after the delamination event or thermo-
277
mechanical erosion. In agreement with the model of Behn et al.
(2007), this indicates that foundering of arc lower crust and the
correlated upwelling of asthenospheric material can produce complicated spatial patterns where foundering influenced areas alternate
with coherent regions.
5.2. Estimation of the rate of crustal formation during the intra-oceanic
evolution of the KAC
Available geochronological constraints indicate that subduction
related magmatism and crustal growth started at least at 118 ± 7 Ma
(Dhuime et al., 2007) but possibly as early as 135 Ma, based on single
zircon spot ages of 132 Ma and 135 Ma (this study). Taking the age of
tonalite KG06 (97.7 ± 0.7 Ma) into account, this constitutes a 20 to
37 Ma period for crustal growth before granulitisation, which is dated at
91–96 Ma (e.g. Anczkiewicz et al., 2002; Yamamoto, 1993). Granulitisation requires that the roots of the arc were buried to depth exceeding
25–30 km and most likely 30–35 km (Garrido et al., 2007). Considering
that the length of the arc is approximately 330 km along the strike
direction, and its width about 40 km, the rate of crustal generation
varies from 32 to 60 km3/km/Ma (volume per unit width along the
strike direction of the arc) for a subduction initiation at 135 or 118 Ma
respectively, and for a mean crustal thickness of 30 km. The large
volume of the Chilas Complex is evidence that magmatic productivity
greatly increased at the end of the intra-oceanic evolution of this arc
section. If the Chilas complex is taken into account (and assuming a
crustal thickness of only 30 km), then the average rate of crust
production during the period 135/118 Ma to 81 Ma ranges from 44 to
65 km3/km/Ma (volume per unit width along the strike direction of the
arc), which compares very well with the present day average crustal
production for the Aleutians (55–82 km 3/km/Ma after Hoolbrook et al.,
1999; and 59–61 km 3/km/Ma after Dimalanta et al., 2002), the Izu arc
(66 km 3/km/Ma after Suyehiro et al., 1996; and 56–60 km3/km/Ma
after Dimalanta et al., 2002), the Marianas (44–50 km3/km/Ma after
Dimalanta et al., 2002), the Tonga (56 km3/km/Ma after Dimalanta et al.,
2002) and to a lesser degree to the New Hebrides (87–95 km3/km/Ma
after Dimalanta et al., 2002). The Kohistan batholith has not been taken
into account in this calculation since only two ages are older than 85 Ma
(Heuberger et al., 2007; Petterson and Windley, 1985) thus suggesting
that the magmatic production associated with the batholith during the
intra-oceanic evolution of the arc can be neglected with regard to the
large amount of magmas produced by the Jijal, Patan, Kiru, Kamila and
Chilas units. The good agreement of crustal growth rates between fossil
and present-day island arcs suggests that processes operating in
nowadays subduction zones were already active since the Cretaceous
and have not drastically changed in the Cenozoic.
6. Conclusion
Geochronological results from this study indicate that the Kohistan
volcanic arc may have started forming as early as 135 Ma. This is
consistent with the oldest identified calc-alkaline magmatic activity in
the Karakoram active margin (c. 120 Ma, after Heuberger et al., 2007)
indicating that contractional stress already existed at that time in the
Neotethyan realm, offshore of Asia. This implies the occurrence of two
coexisting subduction zones, one under the Kohistan arc (ocean–
ocean) and a second under the southern margin of Eurasia (ocean–
continent), in agreement with the model envisioned by Burg et al.
(1998). It is tentatively proposed that the arc developed onto an
oceanic lithospheric basement as old as 175 Ma. Along the arc lifetime, production of magmas and their underplating led to crustal
thickening which culminated in the granulitisation of the lower arc
crust after 98 Ma, and based on literature data, between 91 and 96 Ma.
This c. 5 Ma time interval is thought to reflect the progressive
granulitisation of the lowermost arc-crust during bottom-to-top
growth of the arc and burial of older arc intrusions toward the arc
278
D. Bosch et al. / Earth and Planetary Science Letters 309 (2011) 268–279
root. Leucogranitic veins dated at 90–91 Ma were produced during
this event and are related to dehydration/melting reactions affecting
hornblende-gabbronorite protoliths at depth. We propose that the
granulitisation of the lower arc-crust produced a dense keel of garnet
granulite and pyroxenite (Ducea and Saleeby, 1998), which delaminated in the upper mantle and was responsible for upwelling of
asthenospheric melts generating the Chilas magmatic complex. Lastly,
magmas with geochemical signatures akin to those dated at 101–
102 Ma, were also produced between 85 and 81 Ma (i.e. after
emplacement of the Chilas complex), suggesting that delamination
processes left mantle areas unaffected by foundering of the dense arc
crust, and still influenced by subduction- and mantle wedge-related
dynamics. During the period extending from 135 or 118 Ma to 81 Ma
the average crustal growth rate of 44–65 km 3/Ma/km is close to those
calculated from present-day arcs of the Western Pacific region. The
excellent state of preservation of the island arc crustal section in the
Kohistan arc complex and the similarity of the processes operating in
present-day arcs should stimulate further studies on these materials
as well as on other well preserved fossil occurrences such as the
Jurassic Talkeetna arc (Rioux et al., 2007, 2010).
Acknowledgements
This work has benefited from a financial support from the CNRSINSU (DYETI Program Thème IV) in 2008 to D.B. We thank J.P. Burg
and G. Zeilinger for providing the “KG-xx” samples. We would like to
thank the anonymous reviewers whose critical comments significantly helped to improve the manuscript.
Appendix A. Supplementary data
Supplementary data to this article can be found online at doi:10.
1016/j.epsl.2011.07.016.
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