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Zircon U-Pb ages and Hf isotopes for the Diablillos Intrusive Complex, Southern Puna, Argentina: Crustal evolution of the Lower Paleozoic Orogen, Southwestern Gondwana margin
Accepted Manuscript Zircon U-Pb ages and Hf isotopes for the Diablillos Intrusive Complex, Southern Puna, Argentina: Crustal evolution of the Lower Paleozoic Orogen, Southwestern Gondwana margin Agustín Ortiz, Natalia Hauser, Raúl Becchio, Néstor Suzaño, Alexis Nieves, Alfonso Sola, Marcio Pimentel, Wolf Reimold PII:
S0895-9811(16)30170-5
DOI:
10.1016/j.jsames.2017.09.031
Reference:
SAMES 1796
To appear in:
Journal of South American Earth Sciences
Received Date: 8 September 2016 Revised Date:
9 August 2017
Accepted Date: 24 September 2017
Please cite this article as: Ortiz, Agustí., Hauser, N., Becchio, Raú., Suzaño, Né., Nieves, A., Sola, A., Pimentel, M., Reimold, W., Zircon U-Pb ages and Hf isotopes for the Diablillos Intrusive Complex, Southern Puna, Argentina: Crustal evolution of the Lower Paleozoic Orogen, Southwestern Gondwana margin, Journal of South American Earth Sciences (2017), doi: 10.1016/j.jsames.2017.09.031. This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.
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U-Pb zircon ages for the Diablillos Intrusive Complex (CID), Southern Puna, Argentina
Sample D-13-01 Monzogranite
Hf isotopes on zircon grains from the CID, Southern Puna and contiguous geological provinces
TDM=0.85Ga
TDM=1.03Ga
a
~515-520Ma Magmatic activity summit
~490Ma Concordia Age = 521± 4Ma (2s, decay-const. errs included) MSWD (of concordance) = 3.6, Probability (of concordance) = 0.057
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All samples
~530-540Ma
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Sample D-13-02 Granodiorite
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(f)
ɛHf
Mean = 515±6 [1.1%] MSWD = 8.9
Inherited zircons ~600-630Ma
~1000-1100Ma
a
5G
1.5 M=
TD
(a)
(b)
-2.0 EHf(t)
DM CHUR Monzogranite Granodiorite
TDM=1.87Ga
TD
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Mean = 517±3 [0.68%] MSWD = 3.6
TDM=1.4Ga
.2G
1 M=
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Sample D-13-40 Diorite
~1.4 Ga ~1.2 Ga
CID
Diorite Metadacite, Northern Puna (Hauser et al. 2011) Granitoids, Northern Puna (Bahlburg et al. 2016) Porphyry dacite, Tastil batholith, Cordillera Oriental (Hauser et al. 2011) Guasayán pluton, Sierras Pampeanas (Dahlquist et al. 2016)
-3.0 EHf(t) +3.0 EHf(t)
~1.5 Ga
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Zircon U-Pb ages and Hf isotopes for the Diablillos Intrusive Complex,
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Southern Puna, Argentina: Crustal evolution of the Lower Paleozoic
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orogen, Southwestern Gondwana margin
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Agustín Ortiza,b*, Natalia Hausera, Raúl Becchiob, Néstor Suzañob, Alexis Nievesb,
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Alfonso Solab, Marcio Pimentela, Wolf Reimolda,c,d
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a
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Brasília, DF.
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b
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Geochronology Laboratory, Instituto de Geociências, Universidade de Brasília, 70910 900
GEONORTE - INENCO (Universidad Nacional de Salta – CONICET), Av. Bolivia 5150,
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A4400FVY, Salta, Argentina.
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c
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d
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*Corresponding author: [email protected]. Telephone: +54-387-4255441.
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Humboldt Universitäet zu Berlin, Unter den Linden 6, 10099 Berlin, Germany.
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Museum für Naturkunde – Leibniz-Institute for Evolution and Biodiversity Research, Invalidenstrasse 43, 10115 Berlin, Germany.
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Abstract The evolution of the rocks of the Lower Paleozoic Orogen in Puna, at the
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Southwestern Gondwana margin, has been widely debated. In particular, the scarce amount
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of geological and geochemical data available for the Diablillos Intrusive Complex, Eastern
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Magmatic Belt, Southern Puna, require a further study for new evidence towards the
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understanding of sources, magmatic processes and emplacement of magmas, in order to
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better comprehend the crustal evolution in this setting. We present new combined U–Pb
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and Hf isotope analyses on zircon by LA-MC-ICP-MS from monzogranite, granodiorite
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and diorite rocks of the Diablillos Intrusive Complex. We obtained
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206
Pb/238U concordant
weighted average ages of 517±3 Ma and 515±6 Ma for the monzogranite and diorite,
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respectively, and a concordant age of 521±4 Ma for the granodiorite. These ages permit to
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constrain the climax of magmatic activity in the Diablillos Complex around ~515-520 Ma,
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while the emplacement of the complex took place between ~540 Ma and 490 Ma
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(representing a ca. 50 Ma magmatic event). Major and trace element data, initial
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values varying from 0.70446 to 0.71278, positive and negative ɛNd(t) values between +2.5
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and -4, as well as ɛHf(t) for zircon data between + 3 and -3 indicate that the analyzed
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samples represent contaminated magmas. The ɛHf(t) and the ɛNd(t) values for this complex
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specify that these rocks are derived from interaction of a dominant Mesoproterozoic
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crystalline and/or a metasedimentary source and juvenile mantle-derived magmas, with a
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TDM model age range of ~1.2-1.5 Ga, with later reworking during lower Paleozoic times.
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The combined data obtained in this contribution together with previous data, allow us to
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suggest that the formation of the Eastern Magmatic Belt of the Puna was part of a long-
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lived magmatic event during Early Paleozoic times. Whereby the granitoids of the Eastern
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Sr/86Sr
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Magmatic Belt formed through intra-crustal recycling at an active continental margin, with
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minor contributions from juvenile material in the back-arc setting.
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Key words: Southwestern Gondwana; Puna Argentina; Lower Paleozoic Orogen; Zircon
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U-Pb geochronology; Hf isotopes.
1. Introduction
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The Neoproterozoic to Early Paleozoic rocks (present-day Eastern Cordillera and
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Puna) in the Central Andes, NW Argentina (Fig. 1), formed at the Southwestern Gondwana
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margin. Their evolution has been widely debated and traditionally two orogenic events
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have been distinguished: the Pampean (Upper Precambrian–Lower Cambrian) and the
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Famatinian (Upper Cambrian–Lower Silurian) stages involving subduction processes with
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formation of magmatic arcs, followed by accretion/collision of para-auhocthonous and
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allochthonous terranes (e. g., Ramos et al., 1986; Ramos, 1988, 2008; Loewy et al., 2004;
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Rapela et al., 2007; Cordani et al., 2009; Collo et al., 2009; Drobe et al., 2009; Hauser
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et al., 2011). Due to the lack of evidence and indicators for collisional processes - such as
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suture zones, presence of ophiolites related to oceanic crust, or high P/low T metamorphism
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indicative of subduction-accretion regimes, alternative models, for the evolution of the
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Puna and Eastern Cordillera were suggested. This include a geodynamic evolution
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dominated by intra-crustal recycling at an active continental margin, with minor
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contributions from juvenile materials in the back-arc setting (e.g., Bahlburg and Hervé,
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1997; Becchio et al., 1999; Bock et al., 2000; Lucassen et al., 2000; Coira and
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Koukharsky (2002), Kleine et al., 2004; Büttner et al., 2005; Coira et al. (2009),
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Zimmermann et al., 2014; Rapela et al., 2016; Bahlburg et al., 2016). In particular, the magmatic record in the Puna region (currently a high-plateau area,
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4000 meter above sea level) (Fig. 1), has been associated with the Famatinian orogeny, that
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was initiated in the Lower Tremadoc (Moya et al., 1993), reaching its peak in the Arenig
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(Rapela et al., 1992; Coira et al., 1999). It includes two N-S trending magmatic belts of
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~300 km extension: an Eastern Eruptive Belt (Méndez et al., 1973) represented by
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calcalkaline and peraluminous granitoids, metadacites and metarhyolites, and a Western
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Eruptive Belt (Palma et al., 1986) characterized by metaluminous I-type granitoids (Fig.
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1). Taking into account that so-called eruptive belts also include plutonic and subvolcanic
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rocks, we prefer to use the term Magmatic Belts.
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The origin of the Eastern and Western Magmatic Belts has been variably interpreted
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as two magmatic arcs related to the accretion of allochthonous, parautochthonous, or
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autochthonous terranes along the southwestern margin of Gondwana (e.g., Coira et al.,
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1982; Ramos, 1986; Conti et al., 1996). Other authors (e.g., Dalziel and Forsythe, 1985;
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Ramos 1988; Rapela et al., 1992) interpreted these magmatic belts as two magmatic arcs
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related to subduction but without accretion processes. Or these magmatic belts were
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assigned to an early to middle Ordovician back-arc setting (Coira et al., 1999, 2009), or
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they were considered as part of late stages of the Oclóyic orogeny (Bahlburg, 1990). Coira
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et al. (1999) proposed an active arc setting for the magmatism of the Western Magmatic
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Belt (Upper Cambrian-Lower Arenig). They proposed for the Eastern Belt an occurrence of
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an oblique convergence regime at that time, generating important strike faults that to the
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north resulted in partial melting of the crust and in a subduction zone to the south. These
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processes were based on models involving collision of terranes, and outcrops of oceanic
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suture zones proposed by Ramos et al., 1986; Ramos, 1988; Loewy et al., 2004; Rapela
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et al., 2007; Ramos, 2008, among others. Recently, Zimmermann et al. (2014)
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demonstrated that this alleged ophiolite suture did not exist, thus, refuting the assumption
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of plate boundaries and the collision models. Instead, these authors strengthened the models
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of geodynamic evolution of the Magmatic Belts in the Puna.
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In the past years, the advances of zircon isotope geochemistry have provided a
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formidable means to date magmatic and metamorphic events by U-Pb geochronology (e.g.,
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Hoskin and Schaltegger, 2003), and to employ zircon successfully for the tracing of
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petrogenetic processes in metamorphic and magmatic rocks (Hawkesworth and Kemp,
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2006; Castiñeiras et al., 2011). Some recent compositional investigations of zircon have
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not been directly related to geochronology, but instead based on the ability of zircon to
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record petrogenetic processes in igneous systems through analysis of Hf isotopes (e.g., Zeh
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et al. 2007; Dong et al. 2013; Li et al. 2014; and references therein). This isotope system
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provides a useful tool to discriminate the origin of rocks in general, and may potentially
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reveal the nature of the source rock (e.g., Guo et al., 2007; Willner et al., 2008; Liu et al.,
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2009; Miskjovic and Schaltegger, 2009; Claiborne et al., 2010; Nardi et al., 2012;
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Bahlburg et al., 2009, 2016). Thus, providing important parameters for modelling crustal
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evolution (e.g., Hawkesworth and Kemp, 2006).
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In the Eastern Magmatic Belt, Becchio et al. (2011) defined the Diablillos Intrusive
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Complex (CID by its Spanish name Complejo Intrusivo Diablillos; Fig. 2). This complex is
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mainly composed of granitoids, including diorites, and mafic dykes. Recently, Suzaño et
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al. (2015; 2017) proposed that the main process in the evolution of this Complex was
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magma mixing between melts of granitic and dioritic composition. Until now, there is only
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one age datum for the emplacement of the CID, a U-Pb weighted average age of 501 Ma
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with a 1σ error of ± 17 Ma (Suzaño et al., 2015). Moreover, the magma sources for this
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complex have only been inferred, based on Sr-Nd analyses from nearby area (e.g.,
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Viramonte et al., 2007), possibly involving juvenile mantle or a reworked crustal source,
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or both.
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The scarce amount of geological and geochemical data available for this sector of
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the Southern Puna makes the CID an essential place to study and provide new evidence for
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the development of the large volumes of magmatic rocks in the Eastern Magmatic Belt. In
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this contribution, we have obtained three new zircon U-Pb ages by Laser Ablation Multi-
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Collector Inductively Coupled Plasma Mass Spectrometry (LA-MC-ICP-MS) for diorite,
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granodiorite, and monzogranite rocks from the Diablillos Intrusive Complex. Additionally,
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we present the first Lu-Hf isotope data for zircon from this complex, and for the Southern
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Puna, as well as new geochemical and Sr-Nd isotope data for the igneous rocks of the CID.
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We thus, contribute to the understanding of sources, magmatic processes and emplacement
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of magmas during Cambrian-Ordovician times in this sector of the Puna, helping to unravel
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the geotectonic setting of the Lower Paleozoic Orogen including the Eastern Magmatic Belt
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at the Southwestern Gondwana margin.
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2. Geological setting
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The Southern Puna basement in the surroundings of the Centenario, Ratones and
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Diablillos salars (Fig. 1) comprises medium- to high-grade Neoproterozoic-Lower
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Cambrian metamorphic rocks of the Pachamama Formation (Hongn and Seggiaro, 2001)
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that include mica schist and
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calcsilicates and amphibolites (Fig. 1). In addition, these salars are surrounded by low- to
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medium-grade Neoproterozoic-Lower Cambrian rocks (Fig. 1) consisting of slates,
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phyllites, metaquartzites, and sillimanite schists that constitute the Rio Blanco
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Metamorphic Complex (Hongn and Seggiaro, 2001). The metamorphic peak in Puna, of
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high T and low to medium P occurred at ~510 Ma (Sm-Nd isochron of plagioclase,
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clinopyroxene, and garnet, and U-Pb in titanites from calcsilicate layers, Becchio et al.
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(1999) and Lucassen and Becchio, 2003). The metamorphic rocks are intruded by
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Cambrian-Ordovician granitoids of the Oire Eruptive Complex (CEO; Oire Formation by
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Turner, 1961, redefined as a complex by Blasco and Zappettini, 1995) from the Eastern
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Magmatic Belt. The granitoids of the CEO comprise fine- to coarse-grained, variably
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equigranular or porphyritic granites and granodiorites that occur over a ~100 km long N-S
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trending belt to the east of the Centenario, Ratones and Diablillos salars (Fig. 1). Further to
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the east, these granitoids gradually change to orthogneiss with different degrees of
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mylonitization.
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both ortho- and paragneisses that are intercalated with
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The Ordovician basement is composed of low-grade marine metasediments, with
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intercalated volcanic rocks of the Falda Ciénaga Formation (Llanvirn-Llandeilian,
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Aceñolaza et al., 1976) that occur at the western edge of the Centenario salar (Filo
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Copalayo, Fig. 1) and at Filo de Oire Grande. In tectonic contact, on the Filo Copalayo
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(Mon and Hongn, 1996), medium- to high-grade metamorphic units of mica schist and
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granitic-granodioritic-tonalitic orthogneiss crop out (Fig. 1). They have an apparent
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crystallization age of 467±10 Ma (U-Pb age on zircon, Domínguez et al., 2006). At the
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eastern edge of the Centenario salar, a volcano-sedimentary unit (Cochinoca- Escaya
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Magmatic-Sedimentary Complex, Coira et al., 2004) is composed of slates,
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metapsammites, and low- to medium-grade mica schist (Fig. 1). This sequence is
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intercalated with the products of bimodal magmatism: metarhyolite, metadacite, and
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metabasalt (Fig. 1). These metavolcanic rocks are intruded by syenogranites and
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leucogranites of the CEO, whereby hornfels and mottled schist were generated. Viramonte
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et al. (2007) defined U-Pb crystallization ages for metarhyolites of 485-472 Ma and
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concluded that the plutonic units of the CEO were emplaced between 485 and 462 Ma.
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More recently, farther north, Insel et al. (2012) dated these granitoids at 491 to 481 Ma (U-
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Pb zircon age).
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2.1. Local Geology, Field Relations and Petrography
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The Diablillos Intrusive Complex (CID) was emplaced in the Inca Viejo Range of
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the Southeastern Magmatic Belt. The Inca Viejo Range, located to the west of the Filo de
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Oire Grande, is a mountain chain that separates the Ratones and Centenario salars to the
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northwest from the Diablillos salar to the southeast (Fig. 2, 3a).
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Based on geographic location and macroscopic characteristics of the rocks, Becchio
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et al. (2011) associated the granitoids of the CID as cogenetic with the granitoids of the
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CEO (Fig. 2). Furthermore, Suzaño et al. (2015; 2017), based on petrogenetic models,
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proposed a common and cogenetic origin for both complexes.
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Recently, Suzaño et al. (2015) identified the main lithologies of the CID, namely monzogranite, granodiorite and tonalite, diorite, and mafic dykes (Fig. 2).
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On the eastern side of the Inca Viejo range, the rocks of the CID intrude low- grade
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metamorphic rocks of the Rio Blanco Metamorphic Complex (Fig. 2; Hongn and
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Seggiaro, 2001). On the western side, the monzogranite intrudes moderate grade
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metamorphic rocks, such as mica schists, and para- as well as ortho-gneisses that have been
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included with the Pachamama Formation (Fig. 2; Hongn and Seggiaro, 2001). In addition,
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subvolcanic bodies of the Ratones Andesites and Inca Viejo Formation of Miocene age
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(González, 1984) crop out to the north and south of the range. They are associated with
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hydrothermal alteration and mineral concentrations of economic interest. Lastly, the
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basement rocks and the Neogene subvolcanic intrusives are covered by Quaternary
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volcanics of the Incahuasi Formation (Aceñolaza et al., 1976), and by alluvial/fluvial
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materials including the occurrences of salars that are of particular economic interest for
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lithium and borate brines.
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The recognized units in the Diablillos Intrusive Complex are:
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Monzogranite crops out on the northern and eastern sides of the range and intrudes
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phyllites, slates, and mottled slates of the Rio Blanco Metamorphic Complex to the east
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(Fig. 3e), as well as mica schist, para- and ortho-gneiss of the Pachamama Formation to the
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west. The monzogranite unit has an irregular shape. In the central part of the range, it
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occurs in transitional contact with granodiorite and tonalite (Fig. 2). The monzogranite is medium- to coarse-grained and has a porphyritic texture. It
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varies from a light to dark grey color, according to biotite content in the matrix. The
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porphyritic monzogranites have euhedral to subhedral K-feldspar (1 to 3 cm large) and
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plagioclase (1 to 1.5 cm) phenocrysts embedded in a coarse-grained matrix composed by
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K-feldspar, plagioclase, blue quartz, and abundant biotite. Occasionally, a slight magmatic
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foliation is observed by the preferred orientation of Mafic Magmatic Enclaves (MME) and
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syn-magmatic pegmatite dykes. On the northern side of the range, MME are rare. On the
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eastern side of the Inca Viejo Range, this rocks form lobate and elongate (less than 700 m
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long) bodies and include abundant enclaves of various compositions, such as MME,
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granodiorite, tonalite, and diorite in the form of dismembered bodies of irregular shape
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(Fig. 2). The dismembered diorite bodies (up to 5 m long), granodiorite, and tonalite
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enclaves, and the MME have a heterogeneous distribution and have rounded or elliptical
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shapes, with varying sizes up to 1.5 m. The contacts between the monzogranite and mafic
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dismembered bodies and MME are sharp (Fig. 3b, c), and sharp, transitional, or lobate
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toward granodiorite and tonalite enclaves (Fig. 3b, d).
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Some subhedral or subrounded feldspar phenocrysts seem to migrate mechanically
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from the monzogranite to the granodiorite enclaves (Fig. 3d). In areas where the
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monzogranites are intruded by the mafic dismembered bodies, subrounded K-feldspar
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phenocrysts, plagioclase and quartz are incorporated into the mafic bodies, showing typical
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magma mixing-mingling textures, such as mantled plagioclase and ocelli quartz (Hibbard,
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1991).
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The modal composition shown in Fig. 4 classifies sample D-13-01 as a
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monzogranite with ~55 vol.% K-feldspar, ~24 vol.% plagioclase, and ~21 vol.% quartz. In
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thin section (Fig. 5a), this rock has a holocrystalline, porphyritic texture of medium to
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coarse grain size. It is formed mainly by K-feldspar and plagioclase phenocrysts, with K-
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feldspar, quartz, plagioclase and abundant biotite in the matrix. Accessory minerals are
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muscovite, apatite, titanite, zircon, and undifferentiated opaque minerals.
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The 0.5 to 1 cm phenocrysts are microperthite, and are subhedral with irregular rims
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and inclusions of euhedral plagioclase, quartz, biotite and muscovite; some K-feldspar is
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slightly altered to phyllosilicates. The plagioclase phenocrysts (1-4 mm) are oligoclase to
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andesine (by the Michael-Levy method) and are overgrown by K-feldspar; other
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phenocrysts are zoned. Some quartz has undulous extinction and forms triple joints in the
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matrix, owing to the occasionally observed ductile foliation. Biotite crystals are subhedral,
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frequently forming assemblages with zircon and titanite. Muscovite is subhedral and occurs
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as fine- grained aggregates (Fig. 5a).
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Granodiorite: This unit, together with the tonalite, is the most widespread unit in
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the Inca Viejo Range, occurring over an area ~7 km long and 2 km wide. The granodiorite
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has dark to light grey color, crops out on the mid-eastern side of the range, has a
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transitional contact with the monzogranite to the east, and is intercalated with the tonalite to
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the west. On the northeastern side of the range, the granodiorites intrude phyllites, slates
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and mottled slates of the Rio Blanco Metamorphic Complex.
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The granodiorite is porphyritic. It is characterized by blue quartz and subhedral
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(subrounded) K-feldspar and plagioclase phenocrysts (up to 2 cm in size) (Fig. 3c, d). They
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are immersed in a medium- to fine-grained matrix composed of the same minerals plus
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biotite. Abundant concentrations of MME (similar to those in the monzogranite) of 7 to 50
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cm size, with sharp to transitional contacts were observed in some parts of the granodiorite
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(Fig. 3c). In addition, tonalite enclaves and minor meta-wacke xenoliths (probably from the
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Rio Blanco Metamorphic Complex) were identified occasionally.
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The modal composition (Fig. 4) classifies sample D-13-02 as a granodiorite with
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~40 vol.% plagioclase, ~35 vol.% K-feldspar, and ~25 vol.% quartz. At the microscopic
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scale, the granodiorite (Fig. 5b) has a holocrystalline porphyritic texture with plagioclase,
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microcline and quartz as phenocrysts. The matrix is formed by medium- to fine-grained
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quartz, plagioclase, K-feldspar, abundant biotite, and muscovite. Accessory minerals are
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zircon, titanite, apatite, and opaques.
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Phenocrysts of K-feldspar (microcline, 0.5-1 cm) are surrounded by, and have in
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their cores and along their rims inclusions of, euhedral plagioclase, quartz and biotite. Some
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K-feldspar and plagioclase are altered to sericite in their cores (Fig. 5b). In some cases,
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these K-feldspar phenocrysts have a plagioclase rim defining a rapakivi texture, typically
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formed in a magma-mixing environment (Hibbard, 1991). The plagioclase phenocrysts
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(0.2-1 cm) are subhedral. Less frequently, zoned plagioclase phenocrysts are mantled with
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biotite and overgrown by plagioclase. The quartz phenocrysts (4 cm) are subrounded and
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some are closely aggregated with matrix. Biotite is abundant; ilmenite crystals may be
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partially replaced by titanite along the rims.
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Tonalite: was mapped together with the Granodiorite due to their close relationship
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(Suzaño et al., 2015). The tonalite crops out on the top of the Inca Viejo Range. It forms a
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~50-70 m wide continuous body, in gradational contact with the granodiorite to the east. It
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has a dark grey color and varies from equigranular and fine-grained to slightly porphyritic,
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with quartz (2 cm) and plagioclase (2-2.5 cm) phenocrysts. The matrix is composed of
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plagioclase, quartz, and abundant biotite. This rock also has MME (less than 2 m wide)
7
with sharp contacts (Fig. 3d), small granodiorite enclaves (5-7 cm), and some cross-cutting,
8
narrow (<30 cm) pegmatite dykes.
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The modal composition (Fig. 4) classifies sample D-13-15 as a tonalite with ~48
10
vol.% plagioclase, ~24 vol.% K-feldspar, and ~28 vol.% quartz. In thin section, the tonalite
11
(Fig. 5c) is porphyritic to equigranular, with some small (3 mm) plagioclase phenocrysts.
12
The matrix has medium to fine grain size and is composed of plagioclase, poikilitic quartz,
13
and abundant biotite; accessory minerals are zircon, apatite, and opaques.
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Some plagioclase phenocrysts have a biotite rim and may carry biotite inclusions
15
(Fig. 5c); some plagioclase phenocrysts are slightly altered to sericite. Biotite, in some
16
instances, forms crystal aggregates with opaque minerals and has euhedral epidote
17
inclusions. The poikilitic quartz has inclusions of plagioclase, biotite (Fig. 5c), acicular
18
apatite, fine-grained epidote, and amphibole.
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Diorite: represents the least widespread unit in the complex; it is distributed as
20
dismembered bodies and as MME in the monzogranite, granodiorite and tonalite facies. On
21
the eastern side of the range, it crops out as globular bodies up to 200 m wide and 400 m
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long (Fig. 2). At the southern edge of the range diorite occurs as dismembered bodies and
2
MME. Contacts with monzogranite are sharp, but transitional and lobate with the
4
granodiorite and tonalite. There are places where the diorites have monzogranite enclaves
5
with sharp contacts and irregular or tabular shapes (up to 1 m size) (Fig. 3b, c). This rock
6
has a dark green color, with a fine-grained, equigranular to porphyritic texture. It has
7
mantled plagioclase phenocrysts up to 1 cm in size, and blue ocelli pheno-xenocrysts of
8
quartz immersed in an aphanitic, dark green matrix.
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The modal composition for sample D-13-40 is ~96.5 vol.% plagioclase, ~1.4 vol.%
10
K-feldspar, and ~2.1 vol.% Quartz, which allows to classify it as a diorite (Fig. 4). The very
11
fine grained matrix is composed of plagioclase, amphibole, quartz, biotite, apatite, epidote
12
and opaque minerals (Fig. 5d). Several disequilibrium features are observed in the
13
plagioclase, such as individual crystals with poikilitic texture, zoned crystals with
14
polysynthetic twinning, and porphyritic agglomerates with quartz exsolution (Fig. 5d).
15
Some plagioclase phenocrysts have hornblende rims, and some of these rims are altered to
16
biotite. The quartz pheno-xenocrysts are rare, anhedral, and subrounded; others have
17
reaction rims displaying hornblende and biotite along grain margins.
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Mafic dykes: crop out on the northwestern side of the Inca Viejo Range, up to 40 m
19
thick. These concordant dykes intrude the monzogranite of the CID and the medium grade
20
metamorphic rocks, generally along a NE-SW strike trend. In addition, these dykes intrude
21
the granitic rocks of the CEO in the southern Filo de Oire Grande, to the east of the Inca
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1
Viejo Range (Fig. 2). The dykes have fine-grained, equigranular to porphyritic textures
2
with plagioclase phenocrysts in a fine-grained, dark green matrix. The monzogranites and the mafic bodies are intruded by syn-magmatic concordant
4
and discordant tabular pegmatite dykes (<2m wide) composed of K-feldspar, quartz,
5
muscovite, tourmaline and epidote. Both dyke types – the diorite and the pegmatite
6
varieties - represent later stages of the magmatic event of the CID (Suzaño et al., 2015).
7
3. Analytical Procedures
8
3.1. Geochemical analyses
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Sample preparation for major and trace element analyses of five samples from the
10
CID and three from the CEO were performed at the Geochemistry Laboratory of Geonorte
11
Institute, Salta National University, Argentina. This consisted of washing the field samples,
12
cutting off weathered and veined portions, crushing, quartering, sieving, and grinding the
13
samples with a Herzog mill with oscillatory rings of tungsten carbide. Major elements, Ni
14
and Sc (ppm) were analyzed by ICP-OES and trace element analyses were carried out by
15
ICP-MS at the ACME Laboratories in Canada (Table 1). After the analyses Tungsten was
16
disregarded due to contamination by the tungsten carbide mill. For information about
17
standards, limits of detection, etc., refer to http://acmelab.com/services/.
19
20
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3.2. U-Pb and Lu-Hf isotopes on zircon and Sr-Nd isotopes on whole rock samples
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After crushing, grinding and panning approximately 5 kg of each of the three
2
selected samples (monzogranite, granodiorite and diorite), zircon concentrates were
3
extracted with a Frantz isodynamic separator at the Geochemistry Laboratory in the
4
Geonorte Institute, Salta National University, Argentina. Final purification was achieved by
5
hand-picking of zircon with an acupuncture needle using a binocular microscope at the
6
Geochronology Laboratory of the University of Brasilia, Brazil. In order to target zircon
7
grains for isotope analysis, cathodoluminescence (CL) images of polished mounts were
8
obtained with a JEOL Quantas 450 Scanning Electron Microscope at this same laboratory.
9
The grains were polished and cleaned with 3% nitric acid.
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U-Pb and Lu-Hf analyses were performed on zircon grains from the three selected
11
samples, using a Thermo-Fisher Neptune MC-ICP-MS coupled with a Nd:YAG UP213
12
New Wave laser ablation system at the Laboratory of Geochronology of the University of
13
Brasilia. For the U-Pb analyses (Table 2) the diameter of the laser beam was 30 µm. A
14
frequency of 7 Hz was used, at 80% energy and 40 counts per second. GJ-1 zircons were
15
used as standards to normalize isotopic fractionation during isotope analysis. To monitor
16
precision and accuracy over the course of this study, the TEMORA reference zircon was
17
measured as an unknown obtaining a
18
session, 206Pb/238U age of 409 ± 3 Ma for the diorite session, and a 206Pb/238U age of 409 ±
19
5 Ma for the granodiorite session. Common lead
20
(204Hg+204Pb) masses. In none of the analyzed zircon grains a common Pb correction was
21
necessary due to the very low signal of mass
22
Reported errors are propagated by quadratic addition [(2SD^2+2SE^2)1/2] of external
23
reproducibility and within-run precision. The external reproducibility is represented by the
206
Pb/238U age of 409 ± 4 Ma for the monzogranite
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204
204
Pb was monitored using the
Pb (< 30 cps) and high
206
202
Hg and
Pb/204Pb ratios.
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207
Pb/206Pb
standard deviation (SD) obtained from repeated analyses (n = 20, ~1.1 % for
2
and up to ~2 % for 206Pb/238U) of the standard zircon GJ-1, performed during the analytical
3
sessions, and the within-run precision is the standard error (SE) calculated for each
4
analysis. The U-Pb analyses on zircon grains were carried out using the standard-sample
5
bracketing method (Albarède et al., 2004). Two to four unknowns were analyzed between
6
GJ-1 analyses, and 206Pb/207Pb and 206Pb/238U ratios were time corrected. On smaller zircon
7
grains single spot laser-induced fractionation of the 206Pb/238U ratio was corrected using the
8
linear regression method (Košler et al., 2002). The raw data were processed off-line and
9
reduced using an Excel worksheet (Buhn et al., 2009). Concordia diagrams (2σ error
10
ellipses) and Concordia ages were calculated using the Isoplot/Ex software (Ludwig,
11
2003).
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Lu-Hf isotopes (Table 3) were analyzed on selected zircon grains previously
13
analyzed for U-Pb isotopes. When possible, the two spot analyses were placed as close as
14
possible in order to analyze portions of the zircon grain with the same U and Pb isotopic
15
characteristics. Lu-Hf isotopic data were collected for 40-50 seconds ablation time and with
16
a 40 µm diameter spot size, at 85% energy. The signals of the interference-free isotopes
17
171
18
interferences of
19
calculated using the isotopic abundances of Lu and Hf proposed by Chu et al. (2002). The
20
contemporaneous measurements of
21
bias of Yb using a
22
isotope ratios were normalized to the
23
detailed description of the procedures and methods is given by Matteini et al. (2010).
173
Yb and
175
176
Lu were monitored during analysis in order to correct for isobaric
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Yb,
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Yb and
173
176
Lu on the
171
176
Hf signal. The
Yb and
176
Yb and
176
Lu contribution was
173
Yb provide a method to correct for mass-
Yb/171Yb normalization factor of 1.132685 (Chu et al., 2002). Hf 179
Hf/177Hf value of 0.7325 (Patchett, 1983). A
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The εHf(t) values were calculated using the decay constant λ=1.865*10-11 proposed
1
176
Lu/177Hf and
176
Hf/177Hf CHUR values of 0.0332 and
by Scherer et al. (2006), and the
3
0.282772 proposed by Blichert-Toft and Albarède (1997). A two-stage TDM age was
4
calculated from the initial Hf isotopic composition of the zircon, using an average crustal
5
Lu/Hf ratio (Gerdes and Zeh, 2009; Nebel et al., 2007). The initial Hf composition of
6
zircon represents the
7
namely the possibly concordant U-Pb age previously obtained on the same crystal. Two-
8
stage depleted mantle Hf model ages (TDM Hf) were calculated using
9
and
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Hf/177Hf value calculated at the time of zircon crystallization,
176
Lu/177Hf=0.0384
Hf/177Hf=0.28325 for the depleted mantle (Chauvel and Blichert-Toft, 2001) and
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176
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176
11
Wedepohl, 1995). Before Hf isotope measurements on zircon, replicate analyses of a 200
12
ppb Hf JMC 475 standard solution doped with Yb (Yb/Hf=0.02) were carried out
13
(176Hf/177Hf=0.282162±13 2s, n=4). During the analytical session replicate analyses of GJ-1
14
standard zircon were done obtaining a
15
agreement with the reference value for the GJ standard zircon obtained by Morel et al.
16
(2008). The εHf values for each single grain were recalculated to the U-Pb age previously
17
obtained on the same zircon grain. For slightly discordant analyses, the
18
207
19
respectively.
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Hf/177Hf ratio of 0.282006±16 2σ (n =25), in
206
Pb/238U and
Pb/206Pb ages obtained were used for younger (<1 Ga) and older (>1 Ga) grains,
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Lu/177Hf value of 0.0113 for the average crust (Taylor and McLennan, 1985;
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Bulk rock Sm-Nd isotopic analysis of eight selected samples (Table 5) was carried
21
out at the Geochronology Laboratory of the University of Brasília, Brazil. Sample
22
dissolution was done in Teflon Savillex beakers or in Parr-type Teflon bombs. Sm and Nd
23
extraction from whole-rock powders followed the technique of Gioia and Pimentel (2000).
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149
Sm–150Nd spike was used. Sr, Sm and Nd samples were loaded onto Re
A mixed
2
filaments of a double filament assembly. Sm and Nd isotopic analyses were carried out
3
using a Finnigan Triton mass spectrometer. TDM values were calculated according to the
4
two-stage model as presented by Liew and Hofmann (1988).
5
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3.3. Microprobe analysis
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Microprobe analysis (Table 4) was performed on zircon grains previously analyzed
10
by the U-Pb laser ablation method using a JEOL JXA-8230 Superprobe at the University of
11
Brasilia, Brazil. Concentrations for HfO2, SiO2, ZrO2, Y2O3, UO2 and ThO2 were
12
determined. The conditions for analysis included 20kV energy, 10nÅ of current for Si, Zr
13
and Hf, and 100nÅ for U, Th and Y. Counting times on peaks were 10 s for Si and Zr, and
14
60 s for the other elements. Counting times on background were 5 s for Si and Zr, and 30 s
15
for the other elements.
17
18
19
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4. Results
8
4.1. Geochemistry
9
4.1.1. Major, trace and rare earth elements
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Major and trace element compositions of one monzogranite, a granodiorite, a
12
tonalite, and two diorite samples from the CID and a monzogranite, a granodiorite and a
13
diorite mafic dyke sample from the CEO are compiled in Table 1.
14
In the TAS diagram for granitoids (Middlemost, 1994) (Fig. 6a), the monzogranite unit is
15
represented by samples D-13-01 (CID) and CB-13-05 (CEO); the former plots into the
16
quartz monzonite field. However, we continue to refer to this sample as monzogranite
17
based on its modal composition and field evidence. We assume that the low silica content
18
shown by this sample is due to the mixed nature of samples. We confirmed this assumption
19
based on our field relationships and geochemistry of the CID and by previous studies by
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Suzaño et al. (2015; 2017), which defined that the main magmatic process in the formation
2
of this complex was magma mixing. The granodiorite and tonalite units are represented by samples D-13-02
4
(granodiorite, CID) and D-13-15 (tonalite, CID) that plot into the granodiorite and tonalite
5
fields, respectively; BA-13-07 (granodiorite, CEO) plots into the granodiorite field as well.
6
Finally, the diorite unit is represented by samples D-13-40 (CID), CB-13-02 (CID) and CB-
7
13-04 (CEO), all plotting into the quartz diorite field.
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In the A/CNK vs A/NK plot (Shand, 1943) (Fig. 6b) the diorite samples fall into
9
the metaluminous field representing I-type rocks; in contrast, the tonalite, granodiorite and
10
monzogranite samples lie in the peraluminous field representing S-type granitoids. There is
11
one sample (D-13-15, tonalite) that falls into the metaluminous field, near the boundary to
12
the peraluminous field. In the AFM diagram (Irvine and Baragar, 1971) (Fig. 6c), the
13
intermediate rocks of the CID and CEO show a calc-alkaline trend, whereas the mafic rocks
14
show a slight tendency towards a more primitive origin, with tholeiitic characteristics
15
(sample CB-13-04, dioritic basic dyke).
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The Harker diagrams (Harker, 1909) illustrate two contrasting sample groups (Fig.
17
7): on the one hand, the basic plutonic rocks and on the other, the intermediate plutonic
18
rocks. Al2O3 contents of all samples do not vary significantly (~14 to 16 wt.%). The SiO2
19
content in the diorite varies from ~50 to 54 wt.%; the granodiorite and tonalite show SiO2
20
values of ~60-65 wt.%; and the monzogranite has ~61 to 69 wt.% SiO2. The mafic rocks
21
have values of FeOt (~8-10 wt.%), MgO (~4-8 wt.%), CaO (~7-11 wt.%), Na2O (~2-3
22
wt.%) and K2O (~0.5-2.5 wt.%); the felsic rocks have contents of FeOt (~4-8 wt.%), MgO
23
(~2-4 wt.%), CaO (~1.5-4 wt.%), Na2O (~2.5-3.5 wt.%) and K2O (~2-5 wt.%). The two
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main groups have regular variation trends. MgO, FeOt, TiO2 and CaO are negatively
2
correlated with increasing SiO2, whereas Na2O and K2O show positive correlations (Fig. 7).
3
Trace element versus silica variation diagrams (Fig. 8) show decreasing Sr with
4
increasing SiO2, suggesting plagioclase fractionation. Rb and Ba display positive
5
correlation trends with increasing silica. In the same way, the Zr content in the felsic rocks
6
shows the same behavior with increasing SiO2 due to zircon saturation. Sample CB-13-04
7
(Diorite mafic dyke) has 94 ppm Ni and is the most primitive rock of the basic group (in the
8
AFM diagram, Fig. 6c, with slight tholeiitic character).
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In the upper crust-normalized trace element spider diagram (Taylor and
10
McLennan, 1985) (Fig. 9a), the intermediate rocks show enrichment in the Large-Ion
11
Lithophile Elements (LILE) (i.e., Ce, Rb, Th and K) compared to the mafic rocks. Besides,
12
the felsic group shows a slight depletion in High Field Strength Elements (HFSE) (i.e., Nb,
13
P, Zr and Ti), suggesting the contamination with mafic magmas. In addition, this group
14
shows negative anomalies for Sr and Ba, suggesting plagioclase fractional crystallization
15
during magma evolution (Dong et al., 2013). In Fig. 9a the mafic rocks show a minor
16
depletion of LILE and strong depletion of HFSE compared to the intermediate rocks, with
17
the exception of sample CB-13-04 (CEO, diorite mafic dyke) that has a stronger depletion
18
of the LILE and HFSE.
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The chondrite-normalized REE patterns (Fig. 9b) for the intermediate rocks show
20
enrichment of the LREE (LaN/YbN=3.91 - 11.84) and a negative slope for the HREE,
21
compared to the mafic rocks including those of Zimmermann et al. (2014) from the
22
Western Magmatic Belt, with LaN/YbN=1.62-3.67 and that have a crustal signature as
ACCEPTED MANUSCRIPT
indicated by Dong et al. (2013). In addition, the intermediate rocks display a discernible
2
negative Eu anomaly (Eu/Eu*= 0.50-0.70), indicative of plagioclase fractionation. The
3
mafic rocks show a slight depletion of the HREE compared to the felsic rocks, with the
4
exception of sample CB-13-04 (CEO, diorite mafic dyke) with even stronger depletion of
5
the LREE (Fig. 9b).
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4.2. Zircon U-Pb geochronology
2
4
Three samples representing the main units of the CID, namely monzogranite, granodiorite and diorite, were selected for U-Pb geochronology.
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All three samples show similar populations of colorless to pinkish-brown zircon
6
grains of 100-300 µm sizes. Cathodoluminescence (CL) images show that the grains
7
have variable shapes from elongate euhedral and subhedral, to bipyramidal to
8
subrounded. Some grains display internal magmatic zonation, whereas others have
9
inherited zircon cores (Fig. 10b, d, f). Some grains from the granodiorite have
10
irregularly shaped dark cores (Fig. 10d). Microprobe analyses of zircon grains from
11
monzogranite, granodiorite and diorite indicate variable Th/U ratios from 0.14 to 2.0,
12
0.14 to 1.2, 0.52 to 2.1, respectively (Table 4), which are all typical for a magmatic
13
origin (Hoskin and Schaltegger, 2003).
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14
Sample D-13-01, Monzogranite:
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15
A subset of forty-five zircon grains out of fifty-three in total, with a concordance
17
of 100 ± 2%, were considered for age calculation and display an age distribution from
18
~490 Ma to 540 Ma and a weighted average
19
10a), which is interpreted as the emplacement age of this unit. The eight zircon grains
20
left apart from the age calculation yielded inherited ages ranging from ~615 to 1037 Ma
21
(Table 2).
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16
206
Pb/238U age of 517 ± 3 Ma (inset Fig.
22
The Probability Density Plot (PDP, Fig. 10g) resolves various zircon
23
populations with similar elongate euhedral and subhedral bipyramidal crystal shapes
ACCEPTED MANUSCRIPT (Fig. 10b). The PDP displays a small, four grain population of ~490 Ma age; a major
2
population (27 grains) of ~515-520 Ma age, a significant (13 grains) population of
3
~530-540 Ma age, and another (3 grains) of ~600-630 Ma age. In addition, there is a
4
group of three zircon core analyses representing a ~1000-1100 Ma population.
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5
7
Sample D-13-02, Granodiorite:
Analyses from eight zircon grains, with concordance of 100 ± 3%, out of fifteen 206
SC
6
Pb/238U age of 521±4 Ma (Fig. 10c), assumed to be
grains were used to calculate a
9
the emplacement age of this unit. The seven zircon grains left out from the age
10
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calculation gave inherited ages from ~577 to 1364 Ma (Table 2). The PDP (Fig. 10h) characterizes several zircon populations with similar
12
mineral characteristics (elongate euhedral and subhedral bipiramidals, Fig. 10d) and
13
displays two data points of ~490 Ma age, a significant population of ~515-520 Ma ages
14
(5 grains), a grain of ~540 Ma age, an older population of ~600 Ma (5 grains), and a
15
single older zircon grain of ~1300 Ma age.
17
18
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Sample D-13-40, Diorite: Twenty seven zircon grains out of a total of twenty-nine grains (Table 2), gave
19
analyses with a concordance of 100±2% and showed an age distribution from ~490 Ma
20
to 540 Ma (Fig. 10e). The
21
10e), which is interpreted as the emplacement age of this unit. The two zircon grains left
22
out from the age calculation gave inherited ages from ~650 to 1000 Ma (Table 2).
206
Pb/238U weighted average age is 515±6 Ma (inset Fig.
ACCEPTED MANUSCRIPT In the PDP (Fig. 10i) several zircon populations with varied ages and similar
2
habit characteristics (elongate euhedral and subhedral bipiramidal to subrounded, Fig.
3
10f) are distinguished. First, a five grain population of ~490 Ma occurs; second, a large
4
population of ~515-520 Ma ages (15 grains); a significant population of ~530-540 Ma
5
ages (7 grains); one grain of ~630 Ma, and one at ~1100 Ma.
6
4.3. Lu-Hf isotope data for zircon
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7
8
10
Lu-Hf analyses (Table 3) were performed on zircon grains previously analyzed by the U-Pb method.
11
Sample D-13-01, Monzogranite:
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Eight zircon crystals with varied U-Pb ages were analyzed in this sample,
14
showing 176Lu/177Hf and 176Hf/177Hf ratios between 0.001257 - 0.003581 and 0.282285 -
15
0.282465 respectively (Table 3).
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Zircon grains with ages between ~515 and 530 Ma show ɛHf(t) from -2.6 to
17
+0.4 (Fig. 11), with Hf TDM model ages between 1.5 and 1.3 Ga. Three Neoproterozoic
18
inherited zircon crystals with ages between 616 and 703 Ma show positive to negative
19
ɛHf(t) between +1.2 (age=616 Ma) and -1.9 (age=703 Ma) with Hf TDM model ages
20
between 1.4 and 1.6 Ga. Two zircon cores with Mesoproterozoic ages (1034 and 1056
21
Ma) gave positive ɛHf(t) values of 7.0 and 5.9 (Fig. 11), with Hf TDM model ages
22
between 1.3 and 1.5 Ga.
23 24
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Sample D-13-02, Granodiorite:
2
Eight zircon crystals with different U-Pb ages were analyzed for Lu-Hf isotopes
3
in this sample, disregarding one analysis (Grain N° Z8c) due to an anomalous ɛHf(t).
4
Otherwise, this sample yielded
5
176
Lu/177Hf ratios of 0.000979 - 0.004253 and
Hf/177Hf ratios of 0.282341 - 0.282536.
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Zircon grains from this sample with ages between ~485 to 540 Ma show ɛHf(t)
7
values from +2.1 and -2.2 (Fig. 12), with Hf TDM model ages between 1.2 and 1.6 Ga.
8
Also, this sample presents a negative maximum ɛHf(t) of -3.8 (Age=538 Ma), and a
9
positive ɛHf(t) maximum of +4.2 (Age=654 Ma) (Fig. 12), and Hf TDM model ages of 1.6 Ga and 1.2 Ga, respectively.
11 12
Sample D-13-40, Diorite:
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Eight zircon crystals with varied U-Pb ages were analyzed in this sample,
14
yielding 176Lu/177Hf ratios between 0.001075 - 0.005571 and 176Hf/177Hf ratios between
15
0.282192 - 0.282706 (Table 3).
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Zircon grains with ages between ~518 and 536 Ma show ɛHf(t) values from -9.5
17
to +9.1 (Fig. 11), with Hf TDM model ages between 1.9 and 0.9 Ga. A younger zircon
18
(Z48) of 495 Ma age yielded a positive ɛHf(t) of +2.1 and Hf TDM model age of 1.2 Ga,
19
and an inherited zircon crystal with a Mesoproterozoic age (1115 Ma) a positive ɛHf(t)
20
of +4.0 and Hf TDM model age of 1.6 Ga.
21
4.4. Sr-Nd isotope data
22
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Whole-rock Sr and Nd isotope analyses on eight samples (monzogranite,
2
granodiorite, tonalite, and diorite) from the CID and CEO were recalculated to 518 Ma,
3
the proposed average age of emplacement (Fig. 12; Table 4). The two monzogranites (D-13-01, CID; CB-13-05, CEO) have
4
147
Sm/144Nd
ratios of 0.12 – 0.13 and 143Nd/144Nd ratios of 0.511745 and 0.511833, respectively. The
6
εNd(t) values are -4.3 (CB-13-05) and -2.6 (D-13-01), with TDM model ages of 1.6 and
7
1.3 Ga, respectively (Fig. 12a). The
8
12b).
Sr/86Sr values are 0.71065 and 0.71278 (Fig.
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87
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5
147
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The two granodiorites (D-13-02, CID; BA-13-07, CEO) and one tonalite (D-13-
9
Sm/144Nd ratios between 0.11 and 0.15 and
143
Nd/144Nd
15, CID) sample yielded
11
ratios of 0.511722 to 0.511860 (Table 4). The εNd (t) values range from -4.8 (BA-13-
12
07), over -3.4 (D-13-02), to -2.1 (D-13-15), with TDM model ages of 1.38 Ga, 1.4 Ga and
13
1.55 Ga, respectively (Fig. 12a). These samples gave
14
0.71526 (Table 4) (Fig. 12b).
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10
87
Sr/86Sr values of 0.71112 -
The three diorites (D-13-40 and CB-13-02, CID; CB-13-04 mafic dyke, CEO)
15
147
Sm/144Nd ratios between 0.15 and 0.18 and
143
Nd/144Nd ratios between
yielded
17
0.511942 and 0.512088 (Table 4). These samples have ɛNd(t) values of +2.3 (CB-13-
18
04), +1.8 (CB-13-02), and -0.5 (D-13-40), with TDM model ages of 1.5 Ga, 1.3, and 1.4
19
Ga, respectively (Fig. 12a). The 87Sr/86Sr values range from 0.70446 to 0.70569 (Table
20
4) (Fig. 12b).
22 23 24 25
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2
5. Discussion
3
5.1. Emplacement age of the Diablillos Intrusive Complex
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A weighted average 206Pb/238U age of 517±3 Ma (Fig. 10a) was obtained for the
6
monzogranite, 515±6 Ma (Fig. 10e) for the diorite, and a concordant U-Pb age of 521±4
7
Ma for the granodiorite (Fig. 10c). The ages of the three samples overlap within error,
8
and are considered a good estimate of the emplacement age of this complex.
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5
The age variations in the Concordia diagrams for the three samples (Fig. 10a, e)
10
could reflect analytical scatter; but the age range is similar in the three analyzed
11
samples, regardless of which analytical session is scrutinized. No evidence for Pb loss
12
was observed either. Likewise, the possibility of analytical scatter in the reference
13
material analyzed with the samples was also disregarded, due to the similar data
14
obtained on the TEMORA reference zircon for each analyzed sample: a
15
of 409 ± 4 Ma for the monzogranite, a 206Pb/238U age of 409 ± 3 Ma for the diorite, and
16
a 206Pb/238U age of 409 ± 5 Ma for the granodiorite.
206
Pb/238U age
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Furthermore, a combination of Pb loss and inheritance of older Pb could provide
18
a possible explanation for the age scatter (Castiñeiras et al., 2010). Excluding the
19
oldest grains with comparatively higher Th/U ratios, the remaining slight variation in
20
the Th/U ratios of the samples suggests that the inherited component involved in the
21
analyzed zircon crystals is insignificant.
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22
Another possibility for the observed age variation is a true difference in age, for
23
example, when a sequence of different geological processes, such as metamorphism or
24
magmatism, occurred over a limited time span (Castiñeiras et al., 2010). Hoskin and
25
Schaltegger (2003) and Willner et al. (2008), among others, concluded that zircon
ACCEPTED MANUSCRIPT crystals of Th/U ratios ≥0.1-0.5 have a magmatic origin, whereas zircon crystals with a
2
metamorphic origin should have Th/U ratios <0.1. Hence, the slight variation observed
3
in the zircon fractionation indices in the analyzed samples, namely Th/U = 0.14 to 2
4
(Table 4), may suggest that zircon crystals from the CID developed during a magmatic
5
event.
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6
On the other hand, a long-lived magma chamber, which could have formed the
7
CID, could also explain most likely the continuous age distribution for these three rock
8
types
9
recrystallization/overgrowth of older cores -not entirely crystallized, at lower levels of
10
the magma chamber- during the intrusion of new pulses of hot magma. Resulting in
11
minor mixing of the new zircon crystals from newly intruded magma with older zircon
12
cores (Dong et al., 2013; Folkes et al., 2011; Castiñeiras et al., 2010).
11
a,
e).
This
variation
could
be
explained
by
the
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(Fig.
The observed internal zonation, irregular dark cores or irregular shapes with
14
irregular dissolution zonation displayed by some zircon grains (Fig. 10b, d, f), support
15
these mixing ideas. Such zonations indicate periods of zircon undersaturation followed
16
by zircon-saturation in a melt (e.g., Vavra 1994; de la Rosa et al., 2002; Zeck and
17
Williams, 2002). In addition, Hoskin and Schaltegger (2003) concluded that such
18
crystals preserved compositional information for the melt prior to, and after, dissolution.
19
Consequently, they recorded a change in composition related to an event such as
20
magma-mixing.
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One way of comparing the individual age distributions of the three analyzed
22
samples is by comparing Figures 10g, h and i, the Probability Density Plots for specific
23
zircon age populations. In Fig. 10j the combined data base for the three rock types is
24
considered.
ACCEPTED MANUSCRIPT 1
Clearly, the zircon population at ~490 Ma (Fig. 10j) is present in the data for all
2
three samples but the meaning of this age is unclear as the number of data for this
3
combined population is limited (10 zircon grains only). A possible interpretation could
4
be that this age group may indicate the last magmatic activity in the complex. The largest population at ~515-520 Ma (39 zircon grains) (Fig. 10j) could
6
represent the peak of magmatic activity related to the CID, whereas the population at
7
~530-540 Ma (32 zircon grains) (Fig. 10j) may represent zircon antecrysts
8
(incorporated from previous crystallization events in the magmatic system - Schmitt,
9
2011). This would mean that they represent early magmatic activity in the CID.
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As a result, after analyzing the age populations in the rock units of the CID, we
11
interpret that this intrusive complex was emplaced between 540 and 490 Ma, possibly
12
over a ca. 50 Ma magmatic event.
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The oldest zircon populations of Late Neoproterozoic ages (~580/600/630 Ma)
14
and Mesoproterozoic ages (1000-1300 Ma) (Fig. 10j) signify inherited xenocrysts from
15
the source rock. These data - mostly from the cores of analyzed zircon grains - along
16
with similar age groups presented by Bahlburg et al. (2016) and Bahlburg and
17
Berndt (2016) from the northern Puna (~550-700 Ma; 900-1200 Ma), illustrate the
18
presence of Neoproterozoic sources in the region, i.e., possibly the metasedimentary
19
Puncoviscana Formation. Finally, the population of Mesoproterozoic age represents the
20
Grenvillian/Sunsas orogeny. This indicates that the unknown basement of the Puna is
21
formed either by the Puncoviscana Formation as an erosional product of these orogenic
22
belts or that the Puna including its Puncoviscana basement is underlain by crust evolved
23
during these orogenies (Bahlburg et al., 2016).
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5.2. Magma sources for the Diablillos Intrusive Complex
4
5.2.1. Geochemistry and Sr-Nd isotopes
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Based on field observation, petrography, and geochemistry - including Sr-Nd
7
isotope analysis, two groups of rocks can be distinguished in the CID and CEO: the
8
intermediate group (monzogranite, granodiorite and tonalite rocks) with silica contents
9
between 69.99 and 59.65 wt.% and the mafic group (diorites) with silica contents between 54.50 and 50.27 wt.%.
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The monzogranite and granodiorite-tonalite rocks are enriched in LILE, have
12
negative anomalies for Ba, Sr, and Eu compared to the dioritic rocks (Fig. 9a), and are
13
enriched in the LREE compared to the mafic rocks (Fig. 9b). The intermediate-felsic
14
samples show slightly negative εNd(t) values, between -4.8 (Granodiorite from CEO,
15
BA-13-07) and -2.1 (tonalite from CID, D-13-15) and initial
16
0.71526 (granodiorite from CEO, BA-13-07) and 0.71064 (monzogranite from CEO,
17
CB-13-05) (Fig. 12b). They display TDM model ages between 1.6 and 1.3 Ga (Fig. 12a),
18
suggesting that Mesoproterozoic crust was involved in the generation of these rocks.
Sr/86Sr values between
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The diorites show an enrichment in incompatible elements (i.e., Ce, Rb, Th and
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20
K) (Fig. 9a) and a slight depletion of the HREE (Fig. 9b), compared to the intermediate
21
rocks, these agrees with a crust-contaminated mafic magma source for the mafic rocks
22
(Kleine et al., 2004). However, there is the exception of sample CB-13-04 (CEO,
23
diorite mafic dyke), with a more primitive origin (Fig. 9b). It was recognized by
24
Suzaño et al. (2015; 2017) that this dyke is related to a later stage of the evolution of
25
the CID. These basic dykes have the same REE pattern as the gabbros from the Ojo de
ACCEPTED MANUSCRIPT 1
Colorados Complex in the Western Magmatic Belt, Puna (Zimmermann et al., 2014),
2
which suggests the involvement of analogous mantle-derived sources. Contamination of mafic magma with crust is also evident in the Sr-Nd isotope
3
87
Sr/86Sr values of 0.70446 (diorite from CID,
data where the diorites show low initial
5
CB-13-02) to 0.70523 (diorite from CEO, CB-13-04) (Fig. 12b). They also show
6
positive to slightly negative ɛNd(t) values between +2.3 (diorite from CEO, CB-13-04)
7
and -0.5 (diorite from CID, D-13-04), and gave Mesoproterozoic TDM model ages (Fig.
8
12a).
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9
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4
Likewise, the negative correlation trend towards the felsic rocks (Fig. 12b), 87
Sr/86Sr and low
143
Nd/144Nd ratios of the
evidenced by the moderate to high
11
intermediate rocks, and the moderate to low 87Sr/86Sr and low
12
mafic rocks from the CID, support contamination of primary magma with crustal
13
material and suggest a magma-mixing origin of the CID. These conclusions are in
14
agreement with field relationships and petrographic findings (Fig. 3 and 5), such as
15
subrounded phenocrysts that migrated mechanically from the monzogranites to the
16
granodiorite enclaves, subrounded K-feldspar phenocrysts, mantled and zoned
17
plagioclase, inter alia.
19 20 21
Nd/144Nd ratios of the
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5.2.2. Crustal evolution evidences from Hf isotopes
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Zircon with a Hf isotopic composition similar to depleted mantle could not be
22
identified; closest to this comes a zircon grain from the diorite sample with an ɛHf(536)
23
value of +9.1, with a TDM age of 0.9 Ga (Fig. 11). The two Mesoproterozoic inherited
24
cores of zircon in the monzogranite sample have juvenile magma component signatures,
25
with ɛHf(t) values of +7.0 and +5.9 and TDM ages of 1.3 and 1.4 Ga, respectively (Fig.
ACCEPTED MANUSCRIPT 11). A zircon from the granodiorite gave a ɛHf(654) value of +4.2 and TDM age of 1.2
2
Ga. On the other hand, zircon grains indicating magmas with an extensively reworked
3
crustal source are exemplified in the diorite sample with ɛHf(519) of -9.5, with a TDM
4
age of 1.9 Ga, and ɛHf(521) of -4.1, with a TDM age of 1.6 Ga. And in the granodiorite
5
sample with a grain characterized by ɛHf(538) of -3.8, giving a TDM age of 1.5 Ga (Fig.
6
11).
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The presence of slightly positive and negative ɛHf(t) values between +3 and -3
8
(Fig. 11) for the three analyzed samples emplaced between ca. 490-540 Ma (Fig. 10j)
9
indicates that the rocks from the Diablillos Intrusive Complex derived from an
10
interaction of a dominant Mesoproterozoic crustal source (crystalline and/or
11
metasedimentary, most frequent TDM age at ~1.4 Ga), and juvenile mantle-derived
12
magma, that were reworked during lower Paleozoic times. In addition, the Sm-Nd
13
whole rock isotope data (Fig. 12a) agree with this interpretation, whereby similar ɛNd(t)
14
(+2.5 to -4) and Nd TDM model ages between 1.3 to 1.6 Ga, with a most frequently
15
observed TDM age of ~1.4 Ga, were obtained.
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The zircon grain from the diorite with an ɛHf(536) positive value of 9.1 indicates
17
that new juvenile magma was formed at TDM = 0.9 Ga (Fig. 11). In the Northeastern
18
Magmatic Belt of the Puna, Hauser et al. (2011) analyzed a Mesoproterozoic zircon
19
grain in a metadacite, with an ɛHf(1070) of +13.0 value and a TDM of 1.0 Ga (Fig. 11).
20
Consequently, a magmatic event took place in the magmatic arc at the SW Gondwana
21
margin between ca. 0.9 and 1.0 Ga (around the Meso/Neoproterozoic boundary). This
22
lends further support to the models of depleted mantle evolution by Condie et al.
23
(2005; 2009) and Andersen et al. (2007), for which only little evidence for juvenile
24
magmatic rocks has actually been recognized worldwide for the ca. 1-1.20 Ga period.
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ACCEPTED MANUSCRIPT The samples from the CID present similar ɛHf(t) and TDM ages to those
2
previously obtained for the ca. 480 Ma subvolcanic rocks from the Northeastern
3
Magmatic Belt of the Puna (Hauser et al., 2011), as well as for the granitoids of ca. 470
4
Ma age analysed by Bahlburg et al. (2016) in the same region (Fig. 11). Additionally,
5
in the Centenario salar and nearby areas (Fig. 1) of the Southern Puna, Viramonte et al.
6
(2007) obtained whole rock ɛNd(t) data ranging from +2.5 to -7 and Nd (TDM) ages
7
between 1.5 to 1.7 Ga for granitoids (460-475 Ma) and metavolcanics (485 Ma) (Fig.
8
12a), similar to the ɛNd(t) (+2.5 to -4) and model Nd (TDM) ages of this study (1.3 to 1.6
9
Ga) (Fig. 12a). Furthermore, Kleine et al., (2004), presented similar Nd results in the
10
Complejo Igneo Pocitos of the Western Eruptive Belt. Similarly to the CID rocks, the
11
samples of Viramonte et al. (2007) and Kleine et al., (2004) have low
12
isotope ratios, indicative of recycling of older crustal material without development of
13
great volumes of new crust (Damm et al., 1990; Becchio et al., 1999; Bock et al.,
14
2000; Lucassen et al., 2000). These similarities between the CID and nearby areas
15
support the interaction of a dominant Mesoproterozoic crustal protolith and juvenile
16
mantle-derived magmas for the evolution of the CID and magmatic complexes in the
17
Centenario salar area in the Southeastern and Northern Magmatic Belt, as well as in the
18
Western Magmatic Belt.
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Similarly, the apparent magma source similarities shown for the Southern and
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20
Northern Puna (Fig. 11) indicate that the evolution of the Eastern Magmatic Belt was
21
analogous in both sectors, as already discussed by Viramonte et al. (2007) and Coira
22
et al. (2009).
23
The whole rock ɛNd(t) (Fig. 12a) and zircon ɛHf(t) values (Fig. 11) presented in
24
this study and by Hauser et al. (2011), Bahlburg et al. (2016) and Dahlquist et al.
25
(2016) show that the interaction of a Mesoproterozoic crustal (crystalline and/or a
ACCEPTED MANUSCRIPT metasedimentary) source of ~1.4 Ga and juvenile mantle-derived magmas, that were
2
reworked during lower Paleozoic times in the Eastern Magmatic Belt, extended not only
3
to the Cordillera Oriental geological province, east of the magmatic belt, but also to the
4
Northern Sierras Pampeanas (Fig. 13) south of the Eastern Magmatic Belt. Thus, this
5
Mesoproterozoic event represents a regional episode that could be correlated with the
6
Grenvillian/Sunsas Orogenic Belt (Condie et al., 2005, 2009; Bahlburg et al., 2016;
7
Bahlburg and Berndt, 2016; Augustsson et al., 2016).
10
5.3. Eastern Magmatic Belt, Lower Paleozoic Arc, SW Gondwana margin: A longlived magmatic event?
11
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Gehrels et al. (2009) determined that the Coast Mountain batholith at the
13
northwestern coast of North America was active nearly continuously from ca. 170 Ma to
14
ca. 50 Ma, but that there were dramatic variations in magmatic flux (low and high level
15
of magmatic activity) during this time. More recently, Cope, (2017) in northeast China,
16
determined cyclic magmatism with magmatic flare-ups and lulls (De Silva et al., 2015;
17
Ducea et al., 2015), along the northern and eastern margins of the North China block
18
during late Paleozoic time (~60 Ma long) and Mesozoic time (~50 Ma long),
19
respectively.
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As stated in the previous section, the magma source similarities of both the
21
North and Southeastern Magmatic Belt of the Puna, and the zircon U-Pb age ranges
22
presented in this contribution and in other recent studies, allow us to suggest that, the
23
magmatism in the Eastern Magmatic Belt in the Puna region represent a long-lived
24
magmatic event. We reach to this statement analyzing the zircon U-Pb age ranges
25
presented in this contribution (CID: 540-490 Ma), and in other recent studies, such as,
ACCEPTED MANUSCRIPT 1
those of the granitoids of 491 to 481 (Insel et al., 2012); granitoids and metavolcanics
2
of 485 to 460 Ma in the Southeastern Magmatic Belt in Puna (Viramonte et al., 2007).
3
Furthermore, the ca. 480 Ma (metavolcanics, Hauser et al., 2011), and 469 to 445 Ma
4
granitoids (Bahlburg et al., 2016) in the Northeastern Magmatic Belt in Puna, Hence, the long-lived magmatic event, started at ca. 540 Ma (Diablillos Intrusive
6
Complex) and lasted until 460 Ma in the Southern Puna (Salar Centenario and nearby
7
areas), with proven younger activity in the Northern Puna, until ca. 440 Ma (Fig. 13,
8
inset a).
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5
The mentioned episodes of low and high level magmatic activity, underline the
10
long-lived magmatic event in the Eastern Magmatic Belt, where the early stages of
11
~540-530 Ma magmatism in the CID and the peak magmatic activity (~520-515 Ma) in
12
the CID would represent a lull episode, whereas the most voluminous magmatic activity
13
registered in the Eastern Magmatic Belt between 485-470 Ma (Viramonte et al., 2007;
14
Hauser et al., 2011; Bahlburg et al., 2016) would represent a flare-up episode (Fig.
15
13, inset a).
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Zimmermann et al. (2014) presented for the Western Magmatic Belt, Southern
17
Puna, U-Pb zircon ages for gabbroic bodies of 543 Ma (Complejo Ojo de Colorados),
18
highlighting the early stages of the magmatic event that took place in both, the Eastern
19
and Western Magmatic belts in Southern Puna at ~540 Ma. In addition, Dahlquist et al.
20
(2016) defined a ~533 Ma U-Pb age for the Guasayán Pluton in the Northern Sierras
21
Pampeanas (near the boundary of the Puna and Cordillera Oriental, Fig. 13, inset b).
22
Aparicio González et al. (2011) presented a ~533 Ma U-Pb age for granitic dykes in
23
the Sierra de Mojotoro, Cordillera Oriental (Fig. 13, inset b), and there is the ~530 to
24
~515 Ma U-Pb age range by Hauser et al. (2011) for the Tastil batholith, Cordillera
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ACCEPTED MANUSCRIPT 1
Oriental (Fig. 13, inset b). Furthermore, these granitoids present similar REE patterns to
2
those of the Eastern Magmatic Belt (Puna), as illustrated in Fig. 9b. These evidences for a number of contiguous regions suggests that the likely
4
long-lived magmatic event in the Eastern Magmatic Belt of the Puna, proposed in this
5
contribution, was a regional event. It could have evolved in a way analogous to the
6
model of Lucassen et al. (2000), who proposed for the Central Andean basement a
7
prolonged evolution, where the Pampean and Famatinian cycles are not distinct events
8
but belong to a single, non-differentiable event from ~540 to 440 Ma. Whereby, the
9
granitoids of the Eastern Magmatic Belt formed in a geodynamic evolution, dominated
10
by intra-crustal recycling at an active continental margin, with minor contributions from
11
juvenile materials in the back-arc setting as stated by Becchio et al., (1999), Bock et al.
12
(2000), Lucassen et al. (2000), Coira and Koukharsky (2002), Kleine et al. (2004),
13
Coira et al. (2009), Bahlburg et al., (2016).
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6. Conclusions
M AN U
9
SC
8
The new data and observations reported here provide a contribution to the
11
understanding of the magmatic activity in the Diablillos Intrusive Complex of the
12
Eastern Magmatic Belt in the Southern Puna. U-Pb data on zircon indicate that the
13
emplacement of the complex took place between ~540 Ma and 490 Ma (representing a
14
ca. 50 Ma magmatic event). The main magmatic activity related to this complex was
15
seemingly at ~515-520 Ma.
The combined major and trace, including rare earth, element data, as well as 87
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10
Sr/86Sr and ɛNd(t) and ɛHf(t) on zircon results indicate that all analyzed
whole rock
18
samples represent crustally contaminated magmas.
19
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17
The ɛHf(t) values of + 3 and -3 and the ɛNd(t) values of +2.5 and -4 indicate that
20
the rocks from the Diablillos Intrusive Complex derived from an interaction of a
21
dominant Mesoproterozoic crustal (crystalline and/or a metasedimentary) source and
22
juvenile mantle-derived magmas - with a TDM model age range of ~1.2-1.5 Ga- that
23
were subsequently reworked during lower Paleozoic times.
ACCEPTED MANUSCRIPT These data, along with the whole rock ɛNd(t) and zircon ɛHf(t) values, of this
2
study and from other authors (e.g. Viramonte et al., 2007; Hauser et al., 2011, and
3
Bahlburg et al., 2016; among others) demonstrate that the interaction of a crustal
4
source of ~1.4 Ga and juvenile mantle-derived magmas in the Eastern Magmatic Belt is
5
related to a similar evolution in the Western Magmatic Belt, Cordillera Oriental and
6
Northern Sierras Pampeanas, emphasizing that the reworking process during lower
7
Paleozoic times was a regional episode.
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The magma source similarities and the zircon U-Pb age range presented in this
9
and previous studies (Viramonte et al., 2007; Hauser et al., 2011; Bahlburg et al.,
10
2016) of both the Northern and Southern Puna allows to propose that the Eastern
11
Magmatic Belt in the Puna region was possibly created in a long-lived magmatic event,
12
initiated at ca. 540 Ma with the intrusion of the Diablillos Intrusive Complex, and
13
lasting to 460 Ma in the Southestern Magmatic Belt of the Puna, with even younger
14
activity in the Northern Magmatic Belt of the Puna, until ca. 440 Ma. Whereby, the
15
granitoids of the Eastern Magmatic Belt, formed in a geodynamic evolution dominated
16
by intra-crustal recycling at an active continental margin, with minor contributions from
17
juvenile materials in the back-arc setting.
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8
Further detailed study, involving additional geochronology not only on the
19
plutonic rocks of the CID but also on the basement rocks from the study area and
20
environs may provide a test for the existence of such a long-lived magmatic event in the
21
Eastern Magmatic Belt of the Puna.
22 23 24 25
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7. Acknowledgements
SC
8
Financial support from the CNPq program and from CONICET is
10
acknowledged. We thank Alejandro Nieva (GEONORTE-Universidad Nacional de
11
Salta) for help with sample preparation for chemical analysis. And we would like to
12
thank Luciana Pereira and Bárbara Lima (Laboratório de geocronologia - Universidade
13
de Brasília) for technical assistance with the CL imaging and isotope analyses. We
14
acknowledge the Rodinia Lithium-Silver Standard Company of Salta for field and
15
lodging assistance. WUR’s research is supported by the German Science Foundation
16
(DFG) and Museum für Naturkunde Berlin. Finally yet importantly, we want to extend
17
our gratitude to reviewers Dr. Beatriz Coira, Dr. Juan Otamendi and Dr. Victor Ramos
18
for taking the time and effort necessary while revising the manuscript, their comments
19
and suggestions helped improving the final version of this work.
21
22
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8. References
7
Aceñolaza, F., Toselli, A., González, O. 1976. Geología de la región comprendida entre
8
el Salar de Hombre Muerto y Antofagasta de la Sierra, Provincia de Catamarca.
9
Revista de la Asociación Geológica Argentina 31, 127- 136.
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6
Albarède, F., Telouk, S., Blichert-Toft, J., Boyet, M., Agranier, A., Nelson, B., 2004.
11
Precise and accurate isotopic measurements using multiple-collector ICPMS.
12
Geochimica and Cosmochimica Acta 68 (12), 2725–2744.
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the Neoproterozoic to Early Paleozoic metasedimentary rocks of the Sierras
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Pampeanas: Paleogeographic and tectonic implications. Gondwana Research 32,
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M AN U
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Schmitt, A., 2011. Uranium series accessory crystal dating of magmatic processes. Annual Review of Earth and Planetary Sciences 39, 321–349. Shand, S. J. 1943. Eruptive Rocks: Their Genesis, Composition, and Classification, with
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Scherer, E., Münker, C., Mezger, K. 2006. Calibration of the lutetium–hafnium clock.
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9
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Suzaño, N. Becchio, R., Nieves, A., Sola, A., Ortiz, A. 2015. Mezcla de magmas en el
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arco magmático Famatiniano del noroeste de Argentina: ejemplo en el Complejo
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Intrusivo Diablillos. Puna Austral. Revista Mexicana de Ciencias Geológicas 32, (3), 433-454.
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22
The role of magma mixing in the evolution of the Early Paleozoic calc-alkaline
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American Earth Sciences, 76, 25-46.
6
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Evolution. Blackwell, Oxford. 312 pp. Turner, J. C., 1961. Estratigrafía del Nevado de Cachi y adyacencias. Acta Geológica Lilloana, 3:191-226. Tucumán.
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Taylor, S.R. and McLennan, S.M. 1985. The Continental Crust: its Composition and
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Viramonte, J. M., Becchio, R. A., Viramonte, J. G., Pimentel, M. M., Martino, R. D.
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Wedepohl, K.H. 1995. The compositions of the continental crust. Geochimica et
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Wetherill, G. W. 1956. An interpretation of the Rhodesia and Witwatersrand age
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Cosmochima Acta 59, 1217–1232.
patterns. Geochimica et Cosmochimica Acta 9(5), 290-292.
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14
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Whitney, D. L. and Evans, B.W. 2010. Abbreviations for names of rock-forming minerals.American Mineralogist 95(1), 185.
19
Willner, A.P., Gerdes, A., Masonne, H.J., 2008. History of crustal growth and recycling
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revealed by a U–Pb and Lu–Hf isotope study of detrital zircon from late
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Paleozoic accretionary systems. Chemical Geology 253, 114–129.
ACCEPTED MANUSCRIPT 1
Zappettini, E., Coira, B., Santos, J. O. 2008. Edad U/Pb de la Formación Chañi: granito
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Zeh, A., Gerdes, A., Klemd, R., Barton Jackson, M., J. R. 2007. Archaean to
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Proterozoic Crustal Evolution in the Central Zone of the Limpopo Belt (South
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Africa-Botswana): Constraints from Combined U-Pb and Lu-Hf Isotope of
Zircon.
M AN U
SC
7
10
Analyses
Journal
11
doi:10.1093/petrology/egm032.
of
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(8),
1605-1639.
Zimmermann, U., Bahlburg, H., Mezger, K., Berndt, J., Kay, S. M. 2014. Origin and
13
age of ultramafic rocks and gabbros in the southern Puna of Argentina: an
14
alleged
15
Sciences 103(4), 1023-1036.
18
19
20
21
22
suture
EP
17
Ordovician
AC C
16
TE D
12
revisited. International
Journal
of
Earth
ACCEPTED MANUSCRIPT 1
2
3
9. Appendix
5
Table 1. Whole rock geochemical analyses for major and trace element data for the CID
6
and CEO, Eastern Magmatic Belt, Southern Puna. Major element data in wt.%, trace
7
element data in ppm.
Petrology/ Element SiO2
CB-13-05
D-13-02
BA-13-07
D-13-15
CB-13-04
D-13-40
CB-13-02
(CID)
(CEO)
(CID)
(CEO)
(CID)
(CEO)
(CID)
(CID)
Tonalite
Diorite mafic dyke
Diorite
Diorite
59.65
50.27
54.5
50.93
Monzogranite Monzogranite Granodiorite Granodiorite 60.93
69.99
61.69
1.06
0.55
1.08
16.53
13.95
15.93
Fe2O3
7.26
4
7.39
MgO
2.73
1.6
2.98
MnO
0.1
0.07
CaO
2.36
1.56
Na2O
3.07
2.56
K 2O
4.6
3.84
P 2O 5
0.18
0.15
LOI
0.9
Sum
99.73
Cr
100
62.05 1.03
1.69
0.83
1.31
1.82
16.69
15.01
16.13
15.8
14.54
6.74
9.12
9.1
8.87
11.13
2.74
3.26
8.08
6.09
6.31
TE D
TiO2 Al2O3
0.12
0.08
0.14
0.15
0.14
0.17
3.16
3.5
4.84
11.26
7.56
9.05
3.04
3.37
2.65
2.36
2.35
2.59
3.15
2.68
2.15
0.28
1.54
1.21
0.19
0.24
0.08
0.18
0.24
1
0.8
1
1.1
1.4
1.7
99.78
99.74
99.74
99.76
99.76
99.75
99.74
100
100
100
100
600
300
300
EP
0.07
1.5
125
67
143
144
235
245
205
303
61.5
67.9
58
41
52.6
63.3
56.1
53
23
<20.0
25
21
26
94
67
35
23.1
16.2
20.2
20.6
19.8
13.8
16.7
17.7
173.8
168.5
145.7
126.8
86.1
9
74.5
47.7
110.9
89.5
116.3
201.5
126.1
104.4
141.1
155.6
Y
40.9
31.8
39.5
31
41.5
21.7
31.8
34.1
Zr
286.7
197.1
281.3
231.9
254.3
57.4
170.3
190.4
Nb
15.4
12.4
16.3
13.4
11.2
2.1
7
6.1
Cs
6.7
13.3
5.5
10.2
5.7
1.1
6.3
4.4
Ba
474
342
324
355
190
39
140
92
Hf
7.8
5.4
7
6
6.6
1.6
4.3
4.4
Ta
1
1.2
1.1
1.1
0.9
0.1
0.5
0.3
Co Ni Ga Rb Sr
AC C
V
SC
D-13-01
M AN U
Sample
RI PT
4
ACCEPTED MANUSCRIPT Th
21.7
13.5
18.6
20.7
7.2
0.9
5.9
2.6
U
1.7
2.3
1.4
1.6
1.6
0.5
1.5
0.6
Be
3
2
2
<1.0
2
<1.0
<1.0
3
As
6.7
4.2
7.8
8
6.9
2.4
19.5
26.5
Au
<0.5
1.2
0.9
1.3
<0.5
0.9
0.9
0.8
0.2
1
<0.1
0.1
<0.1
<0.1
0.1
<0.1
14.6
3
31.1
16.4
37.2
95.2
83.6
73.5
Pb
3.8
1.2
3.1
3.3
2.2
0.8
3
1.3
Sb
<0.1
<0.1
<0.1
<0.1
<0.1
Sc
18
10
22
15
28
Sn
4
4
3.0.
3
3
Ti
0.9
0.7
0.8
0.6
0.4
Zn
107
62
104
98
74 26.2
32.7
44.2
52.9
68.9
94.4
102.6
Pr
13.2
7.8
10.8
12.1
Nd
48.8
30.7
44.9
47.1
Sm
10.3
6.1
8.6
Eu
1.7
0.9
1.8
Gd
9.4
5.9
8.4
Tb
1.4
0.9
1.4
Dy
8.2
5.8
7.8
Ho
1.6
1.2
1.6
Er
4.2
3.3
4.6
Tm
0.6
0.5
0.7
Yb
3.8
3.2
Lu
0.6
0.5
LaN/YbN
9.4
6.8
Eu/Eu*
0.5
0.5
3 4 5
<0.1
0.1
37
30
42
<1.0
2
2
<0.1
0.3
0.1
6
43
21
5.9
18.5
12.3
57.8
12
39.6
30.9
7.3
1.7
5.08
4.2
32.4
8.1
23.6
21.1
7.3
2.2
5.1
5.2
1.7
1.8
0.8
1.5
1.7
7.5
8.1
3.3
6
6.7
1.1
1.4
0.6
1
1.1
5.7
7.6
4.2
5.9
6.8
1.2
1.7
0.9
1.3
1.5
3.1
4.5
2.5
3.6
4
0.5
0.7
0.4
0.5
0.6
TE D
8.8
4.4
3
4.5
2.5
3.4
3.6
0.7
0.5
0.7
0.4
0.5
0.6
6.7
11.9
3.9
1.6
3.7
2.3
0.6
0.7
0.7
0.9
0.8
0.9
EP
2
AC C
1
<0.1
SC
53.1 111.6
M AN U
La Ce
RI PT
Bi Cu
ACCEPTED MANUSCRIPT
Table 2. Results of in situ U-Pb isotope analysis of zircon from rocks of the Diablillos Intrusive Complex, Eastern Magmatic Belt, Southern
Monzogranite
Isotopic Ratios
Apparent Ages
D-13-01
0.206 0.058
1.4
0.625
1.9
0.079
1.3
517.5
0.057
1.5
0.626
2.3
0.079
1.7
502.6
0.136 0.058
1.2
0.633
1.8
0.08
1.3
0.057
1.7
0.632
2.1
0.08
1.3
0.268 0.059
1.4
0.654
1.8
0.081
1.2
Z36
0.058
2.4
0.65
3
0.081
1.9
Z66
0.059
1.5
0.658
2.2
0.081
1.6
Z60C
0.058
2.4
0.649
3.3
0.081
2.2
Z19
0.057
1.1
0.644
1.5
0.082
1
Z68
0.059
1.3
0.661
1.8
0.082
Z60R
0.058
1.5
0.659
2
Z27R
0.059
2.1
0.667
2.5
Z40
0.356 0.058
1.4
0.658
1.9
Z64
0.058
1.9
0.661
2.6
Z65
0.058
1.9
0.662
2.3
Z6
0.057
1.4
0.652
1.8
Z15
0.058
1.2
0.657
Z3
0.058
8.5
Z30B
0.058
Z18
0.057
Z51
Z14
Pb/235U 1σ(%)
206
Pb/238U 1σ(%)
207
Pb/206Pb 1σ(Ma)
207
Pb/235U 1σ(Ma)
206
Pb/238U 1σ(Ma) Rho Conc(%)
30.6
493.2
7.5
488
6.3
0.68 98.9
32.9
493.6
8.9
491.7
8.1
0.75 99.6
25.6
497.8
6.9
494.2
6.2
0.74 99.3
505.7
36.9
497.3
8.2
495.4
6
0.69 99.6
559.7
30.7
511
7.4
500.1
5.7
0.63 97.9
536
51.9
508.3
12
502.2
9.2
0.62 98.8
556.2
33.7
513.3
9
503.7
7.9
0.72 98.1
520.8
52.2
507.8
13
504.9
11
0.68 99.4
500.2
24.8
505
6
506.1
4.9
0.65 100
1.1
555.4
29.1
515.2
7.1
506.2
5.5
0.63 98.3
0.082
1.4
547.3
32.2
514.3
8.1
506.9
6.6
0.67 98.6
0.082
1.4
569.9
44.8
518.6
10
507.1
6.8
0.67 97.8
0.082
1.3
540.3
31.4
513.6
7.7
507.6
6.1
0.65 98.8
0.082
1.8
547.6
41
515.4
11
508.1
8.9
0.69 98.6
0.082
1.3
546.1
41
516
9.2
509.2
6.2
0.66 98.7
0.083
1.1
503.2
30.8
509.9
7.2
511.4
5.5
0.65 100
1.6
0.083
1.1
517.8
25.4
512.8
6.5
511.6
5.5
0.68 99.8
0.659
8.6
0.083
1.1
517.9
187
514.1
35
513.2
5.3
0.63 99.8
1.2
0.662
1.6
0.083
1
527.2
25.4
516
6.3
513.4
5.1
0.65 99.5
1.3
0.65
1.7
0.083
1
486.1
29.5
508.8
6.7
513.8
4.9
0.58 101
0.058
1.5
0.659
1.8
0.083
1
513.4
32.5
514
7.3
514.1
5.1
0.55 100
Z37
0.059
1.7
0.672
2.3
0.083
1.6
553.7
36.9
522.2
9.4
515.1
7.8
0.68 98.6
Z45
0.059
2.3
0.678
2.9
0.083
1.7
564.9
51
525.5
12
516.5
8.4
0.69 98.3
Z24B
0.059
1.5
0.678
2
0.084
1.3
557.4
31.7
525.3
8
517.9
6.5
0.66 98.6
Z67R
TE D
514.3
Z17
AC C
Z9
Pb/206Pb
EP
Z5C
207
SC
207
Th/U
M AN U
1σ(%)
Grain N°
RI PT
Puna. R: rim analysis. C: core analysis.
ACCEPTED MANUSCRIPT
0.058
1.2
0.675
2.7
0.084
2.5
544.9
25.4
523.5
11
518.6
12
0.73 99.1
Z50
0.827 0.058
1.2
0.67
1.5
0.084
1
527.7
26.2
520.9
6.2
519.4
4.8
0.61 99.7
Z16
0.058
1.7
0.669
2.2
0.084
1.4
517.4
36.7
519.8
8.8
520.4
6.8
0.62 100
Z57
0.058
1.2
0.678
1.5
0.084
0.9
547.8
25.8
525.8
6.1
520.7
4.5
0.58 99
Z69
0.058
1.2
0.677
1.5
0.084
0.9
535.7
26.4
524.8
6.1
522.3
4.5
0.57 99.5
2.091 0.058
1.2
0.674
1.7
0.085
1.2
522.4
26
523.4
7
523.6
6.2
0.71 100
Z28R
0.059
1.1
0.689
1.4
0.085
0.8
566.9
24.3
531.9
5.8
523.8
4.2
0.57 98.5
Z22
0.058
1.1
0.681
1.4
0.085
0.8
541
24.1
527.6
5.7
524.5
4.1
0.57 99.4
Z39C
0.059
2.5
0.687
3.4
0.085
2.4
558.3
52.9
530.9
14
524.5
12
0.69 98.8
Z48
0.058
1.3
0.684
1.8
0.085
1.3
544.6
27.5
529.2
7.4
525.6
6.4
0.7
Z54
0.058
1.8
0.685
2.8
0.085
2.1
542.7
Z49
0.058
1.2
0.686
1.6
0.085
1
Z55
0.059
2
0.693
3
0.085
2.3
Z23
0.174 0.058
1.3
0.685
1.7
0.085
1.1
Z52R
0.058
1.3
0.689
1.7
0.085
1.1
Z47
0.058
1.4
0.685
1.8
0.086
1.1
Z44
0.059
2.3
0.696
3.3
0.086
2.3
Z58
0.058
1.2
0.689
1.5
0.086
0.9
Z21R
0.059
2.7
0.696
2.9
0.086
Z21C
0.058
1.9
0.683
2.4
0.086
Z61
0.059
2.4
0.696
2.7
0.086
Z24
0.261 0.059
1.3
0.699
1.8
SC
99.3
529.5
11
526.5
10
0.65 99.4
26.1
530.2
6.6
527.5
5.3
0.64 99.5
563.9
43.9
534.4
13
527.6
12
0.75 98.7
538.1
28.5
530
7
528.1
5.5
0.62 99.6
547.6
27.7
532.3
6.9
528.7
5.5
0.63 99.3
534.9
30.3
530.1
7.4
529
5.8
0.62 99.8
564.5
51
536.5
14
529.9
12
0.7
541.7
25.8
532.3
6.2
530.1
4.7
0.58 99.6
0.9
561.4
59.2
536.2
12
530.3
4.7
0.6
1.6
516.2
41.1
528.5
10
531.4
8
0.64 101
1.3
549.7
51.4
536.2
11
533
6.7
0.63 99.4
0.087
1.3
551
28.3
538.4
7.6
535.5
6.6
0.66 99.4
TE D
M AN U
40.3
542
EP
Z7
RI PT
Z10
98.8 98.9
0.06
1.3
0.834
1.9
0.1
1.4
612.4
27.6
615.6
8.7
616.5
8.1
0.72 100
Z12C
0.064
1.3
0.94
2.1
0.107
1.7
734.7
28
672.7
10
654.3
10
0.78 97.3
Z67C
0.065
1.3
0.994
1.9
0.11
1.4
786.5
26.9
700.7
9.4
674.3
8.7
0.72 96.2
Z38
0.064
1.6
1.012
3.9
0.115
3.6
733.3
34
710
20
702.6
24
0.66 99
Z1R
0.07
1.1
1.417
1.6
0.147
1.1
926.1
22.8
896
9.4
883.9
9.2
0.69 98.6
Z1C
0.074
1.1
1.739
1.4
0.171
0.8
1035
22
1023
8.8
1018
7.7
0.57 98.4
Z5R
0.074
1.1
1.739
1.4
0.171
0.8
1035
22
1023
8.8
1018
7.7
0.57 98.4
Z27C
0.075
1.6
1.793
2.1
0.174
1.3
1056
32.8
1043
14
1037
13
0.62 98.2
AC C
Z31
ACCEPTED MANUSCRIPT
Isotopic Ratios 207
0.354 0.057
7.7
0.606
8.2
0.077
3
485.3
170
0.529 0.057
1.6
0.627
2.3
0.079
1.7
506.5
34.4 20.2
Th/U
Z19 Z21
Pb/235U 2σ(%)
206
Pb/238U 2σ(%)
207
Pb/206Pb 2σ(Ma)
0.058
0.9
0.659
2.2
0.083
2
524.6
Z39
0.179 0.058
1.2
0.668
2
0.084
1.7
528.3
Z66
0.058
2.5
0.674
4.7
0.084
4.1
538
Z11
0.058
1.1
0.675
4.1
0.084
3.9
530.9
207
Pb/235U 2σ(Ma)
206
Pb/238U 2σ(Ma) Rho Conc(%)
481.2
32
480.4
14
0.36 99
493.9
9.1
491.2
8.1
0.74 97
514.1
8.9
511.8
9.9
0.91 97.6
26.1
519.4
8.3
517.4
8.3
0.81 97.9
53.7
523.2
19
519.8
20
0.86 96.6
24.8
523.6
17
521.9
20
0.96 98.3
541.6
21.4
527.9
9.5
524.7
11
0.91 96.9
524.1
27.9
527.5
16
528.3
19
0.95 101
518.1
24.5
565.6
19
577.5
24
0.97 111
608.6
28.5
586.8
14
581.2
15
0.9
M AN U
Z14C
Pb/206Pb
SC
2σ(%)
Grain N°
207
Apparent Ages
RI PT
Granodiorite D-13-02
1.198 0.058
1
0.682
2.3
0.085
2.1
Z4
0.058
1.3
0.681
4
0.085
3.8
Z24
0.058
1.1
0.745
4.5
0.094
4.3
Z10
0.06
1.3
0.782
3
0.094
2.7
0.059
2.1
0.773
5.2
0.096
4.7
551.4
46.3
581.5
23
589.3
26
0.91 107
1.8
0.796
5.2
0.096
4.8
609.4
39.7
594.7
23
590.8
27
0.93 97
Z2R Z13
0.137 0.06
TE D
Z9
95.5
0.058
1.4
0.778
4.9
0.097
4.7
532.9
30
584.3
22
597.7
27
0.96 112
Z7
0.06
1.9
0.884
7.9
0.107
7.7
605.7
40.8
643.2
38
653.9
48
0.97 108
Z8C
0.087
4.8
0.941
5.2
0.078
2.1
1364
89.8
673.5
25
486.1
9.7
0.4
AC C
EP
Z48R
35.6
ACCEPTED MANUSCRIPT
Z22
0.058
Apparent Ages 207
Pb/235U 1σ(%)
206
Pb/238U 1σ(%)
207
Pb/206Pb 1σ(Ma)
207
RI PT
Diorite D-13Isotopic Ratios 40 Grain Th/U 207Pb/206Pb 1σ(%) N°
Pb/235U 1σ(Ma)
206
Pb/238U 1σ(Ma) Rho Conc(%)
2
0.632
2.4
0.079
1.4
540.9
43.3
497.2
9.6
487.8
6.6
0.57 98.1
0.778 0.058
1.5
0.63
1.7
0.079
0.8
523.8
32.8
496.1
6.7
490.1
3.9
0.46 98.8
Z35
0.079 0.058
1.6
0.635
2
0.079
1.3
528.5
34.8
499.3
8
493
6
0.61 98.7
Z48
0.058
2.5
0.64
2.6
0.08
0.9
536.4
54
502.4
10
494.9
4.5
0.49 98.5
32
505.8
1.5
0.646
1.7
0.08
0.9
541.6
6.9
498
4.5
0.52 98.4
1.433 0.058
1.6
0.647
2.2
0.081
1.5
535.8
35.4
506.6
8.8
500.1
7.2
0.67 98.7
Z40
0.058
1.4
0.648
1.6
0.081
0.9
516
29.8
507
6.6
505.1
4.5
0.46 99.6
Z56
0.058
1.4
0.65
1.5
0.082
0.7
518.8
29.8
508.7
6.2
506.5
3.6
0.45 99.6
Z38
1.104 0.058
1.5
0.663
1.8
0.083
1.1
540.6
31.9
516.6
7.4
511.2
5.4
0.58 99
Z12
1.648 0.058
1.4
0.669
1.8
0.083
1.1
547.8
31.5
520.1
7.3
513.9
5.2
0.57 98.8
Z44
0.058
1.6
0.662
1.8
0.083
0.9
523.3
34.1
516.1
7.2
514.5
4.3
0.5
Z46
0.058
1.6
0.668
1.9
0.083
Z34
2.074 0.059
1.8
0.673
2
0.083
0.058
7.8
0.668
7.8
0.083
Z49
1.627 0.059
1.7
0.673
2
0.083
Z26
0.059
1.4
0.679
1.7
0.084
Z3
0.058
1.4
0.669
1.7
Z1
0.058
1.4
0.674
1.7
Z45
0.059
2.2
0.683
Z11
0.058
1.6
0.673
Z30
0.517 0.059
1.5
0.69
Z32
0.058
1.8
0.68
Z52
0.059
1.6
Z9C
0.964 0.059
Z15 Z29 Z57
TE D
99.7
1.1
539.9
34.3
519.7
7.8
515.1
5.4
0.56 99.1
0.8
555.9
39.5
522.7
8.1
515.2
4
0.47 98.6
0.8
535.6
171
519.3
32
515.6
4
0.48 99.3
0.9
551
37.6
522.8
8
516.4
4.6
0.45 98.8
0.9
565.7
31.3
526.2
6.9
517.2
4.2
0.48 98.3
0.084
1
530
31.5
520.3
7.1
518.1
4.9
0.55 99.6
0.084
0.8
542.7
31.6
523.3
6.8
518.9
4.2
0.47 99.2
3
0.084
2
567.1
48.6
528.6
12
519.7
9.8
0.66 98.3
1.9
0.084
1.1
531.6
34.3
522.8
7.7
520.8
5.3
0.54 99.6
1.8
0.085
1.1
573.6
31.7
532.8
7.7
523.4
5.7
0.57 98.2
2.6
0.086
1.9
518
39.5
527
11
529
9.7
0.72 100
0.703
2.4
0.086
1.8
581.6
34.9
540.3
10
530.6
9.1
0.74 98.2
1.3
0.695
1.6
0.086
1
550
29.1
536
6.8
532.8
4.9
0.56 99.4
0.058
1.5
0.689
1.9
0.086
1.1
528.3
32.5
532.3
7.8
533.2
5.9
0.6
0.058
1.5
0.694
1.7
0.087
0.8
531.6
32
535.1
6.9
535.9
4.1
0.45 100
0.058
1.3
0.698
1.6
0.087
0.9
538.9
28.6
537.8
6.6
537.6
4.7
0.54 100
EP
Z21C
M AN U
0.058
Z51
AC C
Z36R
SC
Z37C
100
ACCEPTED MANUSCRIPT
0.877 0.058
1.4
0.692
1.6
0.087
0.8
518.5
29.7
534
6.6
537.7
4.3
0.49 101
Z9R
0.057
4.3
0.828
5.8
0.106
3.8
477.3
94.7
612.2
26
649.3
24
0.67 106
Z14
0.077
4.7
1.598
4.9
0.151
1.3
1115
93.9
969.3
30
906.2
11
0.61 93.5
AC C
EP
TE D
M AN U
SC
RI PT
Z20
ACCEPTED MANUSCRIPT
Table 3. Results of in situ Lu-Hf isotope analysis of zircon from rocks of the Diablillos Intrusive Complex, Eastern Magmatic Belt, Southern Puna. 2 sigma values refer to last digits of the isotope ratios. R: rim analysis. C: core analysis. Meas: measured. 176
Lu/177Hf
±2σ (176Hf/177Hf)meas ±2σ (176Hf/177Hf)t ±2σ
ɛHf (T)
±2σ
TDM (Ga) 1.482
513
0.002691
21
0.282417
22
0.28239
18
-2.61
0.1
Z49
527
0.003581
95
0.282455
35
0.282418
16
-1.3
0.06 1.422
Z47
529
0.002323
37
0.282489
31
0.282465
16
0.41
0.06 1.329
Z31
616
0.002455
50
0.282461
23
0.282432
16
1.17
0.02 1.359
Z12c
654
0.002836
77
0.282444
20
0.282408
13
1.2
0.05 1.389
Z38
703
0.001257
18
0.282308
83
0.282291
16
-1.88
0.01 1.599
Z1c
1035
0.000979
18
0.282351
21
0.282331
16
7.02
0.16 1.325
Z27c
1056
0.003054
81
0.282348
53
0.282285
16
5.86
0.17 1.463
Z48
495
0.001864
11
Z3
518
0.002691
21
Z1
519
0.001951
15
Z11
521
0.003496
22
Z15
533
0.005571
68
Z29
536
0.003236
30
Z8B
570
0.001505
Z14
1115
59
0.282534
16
2.08
0.15 1.208
0.282417
22
0.28239
16
-2.51
0.07 1.48
0.282211
42
0.282192
16
-9.51
0.07 1.865
0.282379
60
0.282343
16
-4.09
0.06 1.57
0.282476
95
0.282418
16
-1.16
0.06 1.419
0.282739
12
0.282706
16
9.07
0.12 0.854
23
0.282488
38
0.282471
16
1.53
0.04 1.301
0.001075
88
0.282217
54
0.282194
16
3.98
0.19 1.616
Granodiorite D-13-02 Grain N°
TE D
0.282552
EP
D-13-40 Grain N°
SC
Diorite
RI PT
Z3
M AN U
Monzogranite T D-13-01 (Ma) Grain N°
484
0.001997
15
0.28244
22
0.282421
16
-2.17
0.08 1.433
Z25
490
0.002629
29
0.282561
23
0.282536
16
2.05
0.08 1.206
Z21
491
0.003171
29
0.282477
24
0.282447
16
-1.1
0.08 1.381
AC C
Z5
Z39
517
0.002394
11
0.282433
28
0.282409
16
-1.84
0.07 1.443
Z4
528
0.003207
10
0.282477
33
0.282444
16
-0.37
0.06 1.371
Z1c
538
0.000979
18
0.282351
21
0.282341
16
-3.8
0.06 1.568
Z7
654
0.004253
44
0.282548
49
0.282494
16
4.2
0.01 1.222
1364
0.001668
12
0.282431
26
0.282386
16
16.4
0.3
Z8c
1.125
ACCEPTED MANUSCRIPT
Table 4. Microprobe analyses for zircon crystals from rocks of the Diablillos Intrusive Complex, Eastern Magmatic Belt, Southern Puna. Data in wt.%. Nd: not determined. Grain N°/ Monzogranite Z5c
Z7
Z9
Z15
Z23
Z24
Z40
Z50
Z67r
Z67c
D-13-01 1.556 1.423 1.592 1.596 1.72
SiO2
31.99 31.98 31.01 32.28 31.31 31.25 30.73 31.92 31.11 32.25
ZrO2
65.26 66.41 64.37 65.74 65.05 65.54 64.89 64.94 65.02 64.83
Total
98.81 99.81 96.97 99.61 98.07 98.33 97.28 98.39 97.76 98.56
Y 2O 3
0.374 0.334 0.371 0.243 0.171 0.356 0.439 0.4
UO2
0.034 0.011 0.066 0.029 0.023 0.046 0.045 0.098 0.041 0.019
ThO2
0.007 0.023 0.009 nd
Total
0.485 0.408 0.446 0.335 0.2
Th/U
0.206 2.091 0.136 nd
Granodiorite D-13-02
Z1r
HfO2
1.611 1.565 1.559 1.352 1.411 1.889 1.729 1.922 1.389 1.464
SiO2
32.64 32.32 32.94 33.17 34.12 32.2
ZrO2
64.33 64.86 65.84 62.16 63.55 64.51 64.61 64.58 64.28 63.92
Total
98.58 98.75 100.3 96.68 99.09 98.6
Y 2O 3
0.208 0.193 0.103 1.181 0.103 0.426 0.278 0.136 0.041 0.254
UO2
0.042 0.018 0.001 0.187 0.051 0.048 0.017 0.039 0.039 0.045
ThO2
nd
nd
Total
0.25
0.224 0.194 1.592 0.161 0.605 0.365 0.175 0.163 0.319
Th/U
nd
nd
Diorite D-13-40
Z9C
HfO2
0.958 1.113 1.196 1.204 1.154 1.632 1.734 1.168 1.326 1.026
SiO2
32.89 32.72 31.87 32.83 32.39 34.2
ZrO2
65.8
Total
99.65 98.3
Y 2O 3
0.235 0.817 0.237 nd
UO2
0.056 0.125 0.057 0.029 0.094 0.089 0.009 0.077 0.051 0.06
ThO2
0.054 0.206 0.05
Total
0.379 1.148 0.366 0.044 1.031 0.255 0.045 0.663 0.428 0.498
Th/U
0.964 1.648 0.877 0.517 2.074 0.079 0.778 1.104 1.627 1.433
Z9c
nd
SC
0.455 0.576 0.579 0.408 0.343
0.174 0.261 0.356 0.827 0.268 nd
Z13
Z19
Z21
Z22r
Z30
Z34
Z35
Z45c
98.83 99.96 99.95 97
1.198 0.137 0.354 0.529 nd
Z20
Z39
32.49 33.46 34.28 31.62
0.224 0.007 0.017 0.009 nd
EP
Z12
nd
0.275 0.289
M AN U
Z8c
1.525 1.632 1.48
0.004 0.012 0.016 0.081 0.011 nd
TE D
Z6
1.541 1.66
RI PT
HfO2
Z37c
Z38
0.007 0.01 0.179 0.222 Z49
Z51
32.89 31.65 32.28 34.28
AC C
64.46 65.57 66.72 64.93 61.06 67.19 65.41 65.59 65.79 98.63 100.8 98.48 96.89 101.8 98.22 99.2
101.1
0.671 0.159 0.029 0.373 0.294 0.244
0.015 0.195 0.007 0.007 0.085 0.083 0.086
ACCEPTED MANUSCRIPT
Petrology
Age (Ma)
Rb Sr (ppm) (ppm)
87
Rb/86Sr (87Sr/86Sr)meas (87Sr/86Sr)t
Monzogranite
518
173.8 110.9 4.551
0.74638±2
0.71278
Monzogranite
518
168.5 89.5
0.75103±2
0.71065
Granodiorite
518
145.7 116.3 3.635
0.73796±1
0.71112
Granodiorite
518
126.8 201.5 1.824
0.72873±3
Tonalite
518
86.1
126.1 1.839
-
D-13-40 (CID) CB-13-04 (CEO) CB-13-02 (CID)
Diorite Diorite mafic dyke
518
74.5
141.1 1.529
518
9
Diorite
518
47.7
147
Sm/144Nd (143Nd/144Nd)meas (143Nd/144Nd)t eNd(T)
TDM (Ga)
10.9
53.99 0.122
0.512250+/-17
0.511833
-2.6
1.32
3.3
15.08 0.132
0.512196+/-25
0.511745
-4.4
1.56
8.65
40.95 0.128
0.512229+/-10
0.511794
-3.4
1.43
0.71526
11.5
63.04 0.11
0.512096+/-10
0.511722
-4.8
1.38
-
8.58
34.98 0.148
0.512365+/-12
0.51186
-2.1
1.55
0.71698±3
0.70569
5.86
23.88 0.148
0.512447+/-20
0.511942
-0.5
1.38
104.4 0.249
0.70708±2
0.70524
3.01
10.11 0.18
0.512700+/-24
0.512088
2.34
1.42
155.6 0.887
0.71101±2
0.70446
6.34
23.14 0.166
0.512625+/-14
0.512061
1.82
1.32
TE D
5.47
M AN U
D-13-01 (CID) CB-13-05 (CEO) D-13-02 (CID) BA-13-07 (CEO) D-13-15 (CID)
Sm Nd (ppm) (ppm)
SC
Sample
RI PT
Table 5. Whole rock Sr-Nd isotope analyses on rocks of the CID and CEO, Eastern Magmatic Belt, Southern Puna.
AC C
EP
Elemental abundances of Rb and Sr are taken from whole rock geochemical analyses. Meas: measured.
ACCEPTED MANUSCRIPT
Figures captions:
2
Fig. 1 (2-column fitting image): Regional map, showing the main geological units in the
3
geological provinces of NW Argentina. CN: Salar Centenario. DB: Salar Diablillos. RT:
4
Salar Ratones. CB: Cerro Blanco. TA: Tajamar. AC: Acazoque. Modified after Viramonte
5
et al. (2007) and Zimmermann et al. (2014).
SC
6
RI PT
1
Fig. 2 (2-column fitting image): Detailed geological map showing the main units in the
8
Diablillos Intrusive Complex. Modified after Nieves (2014). The sampling sites for this
9
study are also shown.
M AN U
7
10
Fig. 3(2-column fitting image): (a) Panoramic view of the Diablillos salar and the Inca
12
Viejo Range. (b) Contact relationship between a porphyritic monzogranite and parts of a
13
mafic dismembered body. Hammer is 40 cm long. (c) Contact relationships between a
14
porphyritic monzogranite, a granodiorite enclave, and a mafic dismembered body. White
15
dotted lines delimit a granodiorite enclave with K-feldspar phenocrysts. Black lens cover is
16
4 cm in diameter.
17
monzogranite and the granodiorite facies, and a contact with a tonalite enclave. Black
18
arrows show K-feldspar phenocrysts that migrated from the monzogranite to the
19
granodiorite. Note the MME included into the tonalite enclave. (e) Intrusive contact,
20
delimited with a white dotted line, between the monzogranite and slates from the Rio
21
Blanco Metamorphic Complex. Hammer is 40 cm long. Dt: Diorite. Gr: Monzogranite. Gd:
22
Granodiorite. Tn: Tonalite.
EP
TE D
11
AC C
(d) Field image showing the transitional contact between the
68
ACCEPTED MANUSCRIPT
1
Fig. 4(Single-column fitting image): Q-A-P diagram showing the modal classification of
3
the analyzed samples from the CID, in comparison with the samples from Suzaño et al.
4
(2015). The modal classification was made by counting approximately 800 points in a thin
5
section.
RI PT
2
SC
6
Fig. 5 (2-column fitting image): Microphotographs of the rock types of the Diablillos
8
Intrusive Complex. (a) Monzogranite (sample D-13-01) with holocrystalline porphyritic
9
texture. It is composed mainly of Kfs and Pl phenocrysts, and the matrix is composed of
10
Kfs, Qz, and Pl. (b) Granodiorite (sample D-13-02). Holocrystalline porphyritic texture
11
with Pl, Kfs and Qz phenocrysts. The matrix is formed by fine- to medium-grained Qz, Pl,
12
Kfs. Bt is abundant. The microcline phenocrysts have Bt inclusions (Inc) along the rims. (c)
13
Tonalite (sample D-13-15) showing porphyritic to equigranular texture. Some Pl
14
phenocrysts have a rim of Bt as well as Bt inclusions (Inc). Poikilitic Qz with inclusions of
15
Bt (Inc) is shown. (d) Porphyritic diorite with Pl phenocrysts (sample D-13-40). The fine-
16
grained matrix is formed by Pl, Amp, Bt and Qz. All images taken in cross-polarized light.
17
Mineral abbreviations after Whitney and Evans (2010). Kfs: K-feldspar. Pl: Plagioclase.
18
Qz: Quartz. Bt: Biotite. Ms: Muscovite. Amp: Amphibole.
TE D
EP
AC C
19
M AN U
7
20
Fig. 6(2-column fitting image): (a) Granitoid TAS diagram (Middlemost, 1994) for the
21
various units of the CID. Rocks are classified as granites, granodiorites, tonalites, and
22
quartz diorites. (b) In the A/CNK vs A/NK plot (Shand, 1943) diorites fall into the 69
ACCEPTED MANUSCRIPT
metaluminous field, whereas the tonalites, granodiorites and monzogranites fall into the
2
peraluminous field. (c) In the AFM diagram (Irvine and Baragar 1971), the CID rocks
3
show a calc-alkaline trend. The diorites show a slight tendency towards tholeiitic
4
characteristics.
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Fig. 7(2-column fitting image): Harker diagrams for the mafic and intermediate plutonic
7
rocks. With increasing SiO2, negative correlation trends are found for MgO, FeOt, TiO2 and
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CaO, whereas the Na2O and K2O values show positive correlation with increasing SiO2.
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Data in wt.%. References same as Fig. 6.
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Fig. 8(2-column fitting image): Harker-style trace element variation diagrams, showing
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decreasing Sr with increasing SiO2. Rb and Ba show positive correlation trends with
13
increasing SiO2. The intermediate rocks show increasing values of Zr with increasing SiO2.
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Trace element data in ppm, SiO2 contents in wt.%. References same as Fig. 6.
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Fig. 9 (Single-column fitting image): (a) Upper crust-normalized trace element spider
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diagram (after Taylor and McLennan, 1985). Intermediate and felsic rocks show
18
enrichment in the LILE compared to the mafic rocks. Also, note in the more evolved group
19
depletion in the HFSE. The mafic rocks show a slight depletion in the LILE and illustrate a
20
stronger depletion in the HFSE than the intermediate rocks. (b) REE abundances
21
normalized to chondrite (Boynton, 1984). The intermediate rocks show enrichment of the
22
LREE and a decreasing HREE pattern, compared to the mafic rocks. Note the similarities
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between the mafic CEO rocks with those from the Ojo de Colorados Complex, Western
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Magmatic Belt (Zimmermann et al., 2014), Guasayán Pluton, Northern Sierras
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Pampeanas (Dahlquist et al., 2016), a metadacite from Niño Muerto, Northeastern
4
Magmatic Belt (Hauser et al., 2011), and the Tastil batholith, Cordillera Oriental (Hauser
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et al., 2011).
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Fig. 10(2-column fitting image): U-Pb isotope data. (a) Concordia diagram (Wetherill,
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1956) displaying the scatter of concordant U-Pb ages for zircon from monzogranite sample
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D-13-01. Errors are shown as ellipses at the 2σ level. Upper left inset shows the Concordia
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diagram for U-Pb ratios of all zircon grains analyzed. Lower right inset shows the weighted
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average age diagram (data in Ma) for the monzogranite. (b) CL images of representative
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zircon crystals from the monzogranite sample. Age results are shown for the spots
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analyzed. (c) Concordia diagram displaying the concordant U-Pb ages for zircon from
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granodiorite sample D-13-02. Errors are shown as ellipses at the 2σ level. Upper left inset
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shows the Concordia diagram for U-Pb ratios of all zircon analyses. (d) CL images of
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representative zircon crystals from the granodiorite sample. (e) Concordia diagram
17
displaying the scatter of concordant U-Pb ages for zircon from diorite sample D-13-40.
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Errors are shown as ellipses at the 2σ level. Upper left inset shows the Concordia diagram
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for U-Pb ratios of all zircon grains analyzed. Lower right inset shows the weighted average
20
age diagram (data in Ma) for the diorite unit. (f) CL images of representative zircon crystals
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from the diorite sample. (g) PDP of the Monzogranite. (h) PDP for the Granodiorite data.
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(i) PDP for the Diorite. (j) PDP for the combined data set from all three samples from the
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Diablillos Intrusive Complex. Note: In (b), (d), and (f), orange circles mark the spots of U-
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Pb analyses. r: rim; c: core. Light blue circles mark the locations of Hf analyses. In (g), (h),
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(i) and (j), ages for the various peak populations are given in black letters.
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Fig. 11(2-column fitting image): Plot of εHf(t) vs. U-Pb ages for zircon from the CID rocks.
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Black dotted ellipse shows that most of the analyzed samples with ages between ~490 and
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540 Ma are plotting near the CHUR line. The majority of the TDM ages are constrained
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between the red dotted lines, denoting TDM ages from 1.2 to 1.55 Ga. The red solid line
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indicates the most frequent TDM observed at 1.4 Ga. For comparison, the Ordovician
9
metadacite εHf(t) data from the Northeastern Magmatic Belt, the Cambrian Tastil batolith
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εHf(t) data from the Cordillera Oriental (Hauser et al., 2011), granitoids from the
11
Northeastern Magmatic Belt (Bahlburg et al., 2016) and the data by Dahlquist et al.
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(2016) for the Cambrian Guasayán Pluton in the Sierras Pampeanas are also shown. Inset
13
(a) ɛHf(t) histogram: The ɛHf(t) values for the selected samples range from +3.0 to -3.0,
14
whereby the value -2.0 is observed more frequently. Inset (b) shows a TDM age histogram
15
for the analyzed samples. The most frequent TDM age for the ~490-540 Ma zircon U/Pb
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ages is ~1.4 Ga.
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Fig. 12(2-column fitting image): (a) Nd isotope evolution diagram for the CID and CEO
19
rocks. TDM ages vary from ~1.3 to ~1.6 Ga (black dotted lines); the most frequent TDM age
20
is ~1.4 Ga. Note that the samples show ɛNd(t) values and TDM ages similar to the results
21
from Viramonte et al. (2007) for the Centenario salar area, Southeastern Magmatic Belt,
22
Puna, and to the data by Hauser et al. (2011) for the Northeastern Magmatic Belt, Puna,
23
and the Tastil batholith, Cordillera Oriental. (b) Nd isotopic ratios plot versus Sr isotopic 72
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ratios for the CID and CEO rocks. There is a negative correlation trend from the basic rocks
2
towards the felsic rocks (black dotted line). Felsic rocks display similar Nd and Sr ratios to
3
the rocks from the Cordillera Oriental and Eastern Magmatic Belt, Northern Puna (Hauser
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et al., 2011).
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Fig. 13 (1.5-column fitting image): Schematic map showing the main geological provinces,
7
sampling sites, and related ages of some Lower Paleozoic granitoids in NW Argentina: the
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Complejo Ojo de Colorados (Zimmermann et al., 2014) in the Western Magmatic Belt
9
(Puna), the Tastil batholith (Hauser et al., 2011), the granitic dykes in Mojotoro (Aparicio
10
González et al., 2011), and the Nevado de Chañi and Cañañi granitoids (Zapettini et
11
al., 2008; Escayola et al., 2011, respectively) in the Cordillera Oriental. Also shown is the
12
Guasayán pluton in the Northern Sierras Pampeanas (Dahlquist et al., 2016). Inset (a)
13
shows the Probability Density Plot illustrating zircon U-Pb ages for the Eastern Magmatic
14
Belt (Puna), from the CID, and the North and Southeastern Magmatic Belt; data from
15
Viramonte et al. (2007), Hauser et al. (2011), and Bahlburg et al. (2016). Note that the
16
Lull peak at ~515-520 Ma is higher than the Flare-up peak at ~470-480 Ma because we
17
used all the ages obtained in this work, concentrated at the former peak. Inset (b) shows the
18
Probability Density Plot for zircon U-Pb ages for the Cordillera Oriental and Northern
19
Sierras Pampeanas from Hauser et al. (2011), Aparicio González et al. (2011), and
20
Dahlquist et al. (2016). CID: Diablillos Intrusive Complex. COC: Complejo Ojo de
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Colorados. CO: Cordillera Oriental. SP: Sierras Pampeanas. SS: Sierras Subandinas.
22
Modified from Aparicio González et al. (2011).
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Figures:
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Fig. 1 (2-column fitting image):
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Fig. 2 (2-column fitting image):
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Fig. 4(Single-column fitting image): 77
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Fig. 5 (2-column fitting image):
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Fig. 6(2-column fitting image):
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Highlights •
U-Pb ages on zircon indicate that the emplacement of the Diablillos Intrusive
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Complex, Southern Puna, Argentina, took place between ~540 Ma until 490 Ma. •
The magmatic activity climax of the complex was at ~515-520 Ma.
•
The ɛHf(t) and ɛNd(t) values indicates that the rocks from the Diablillos Intrusive
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Complex, derived from an interaction of a dominant crustal source and juvenile mantle-derived magmas, with a TDM model age range of ~1.2 - 1.5 Ga, reworked
•
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during lower Paleozoic times.
The interaction of a Mesoproterozoic crustal source and juvenile mantle-derived magmas in Eastern Magmatic Belt, Puna, experienced a similar evolution in the Western Magmatic Belt, Cordillera Oriental, and Northern Sierras Pampeanas,
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emphasizing that this event was a regional episode.
S0895-9811(16)30170-5
DOI:
10.1016/j.jsames.2017.09.031
Reference:
SAMES 1796
To appear in:
Journal of South American Earth Sciences
Received Date: 8 September 2016 Revised Date:
9 August 2017
Accepted Date: 24 September 2017
Please cite this article as: Ortiz, Agustí., Hauser, N., Becchio, Raú., Suzaño, Né., Nieves, A., Sola, A., Pimentel, M., Reimold, W., Zircon U-Pb ages and Hf isotopes for the Diablillos Intrusive Complex, Southern Puna, Argentina: Crustal evolution of the Lower Paleozoic Orogen, Southwestern Gondwana margin, Journal of South American Earth Sciences (2017), doi: 10.1016/j.jsames.2017.09.031. This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.
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U-Pb zircon ages for the Diablillos Intrusive Complex (CID), Southern Puna, Argentina
Sample D-13-01 Monzogranite
Hf isotopes on zircon grains from the CID, Southern Puna and contiguous geological provinces
TDM=0.85Ga
TDM=1.03Ga
a
~515-520Ma Magmatic activity summit
~490Ma Concordia Age = 521± 4Ma (2s, decay-const. errs included) MSWD (of concordance) = 3.6, Probability (of concordance) = 0.057
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All samples
~530-540Ma
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Sample D-13-02 Granodiorite
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(f)
ɛHf
Mean = 515±6 [1.1%] MSWD = 8.9
Inherited zircons ~600-630Ma
~1000-1100Ma
a
5G
1.5 M=
TD
(a)
(b)
-2.0 EHf(t)
DM CHUR Monzogranite Granodiorite
TDM=1.87Ga
TD
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Mean = 517±3 [0.68%] MSWD = 3.6
TDM=1.4Ga
.2G
1 M=
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Sample D-13-40 Diorite
~1.4 Ga ~1.2 Ga
CID
Diorite Metadacite, Northern Puna (Hauser et al. 2011) Granitoids, Northern Puna (Bahlburg et al. 2016) Porphyry dacite, Tastil batholith, Cordillera Oriental (Hauser et al. 2011) Guasayán pluton, Sierras Pampeanas (Dahlquist et al. 2016)
-3.0 EHf(t) +3.0 EHf(t)
~1.5 Ga
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Zircon U-Pb ages and Hf isotopes for the Diablillos Intrusive Complex,
2
Southern Puna, Argentina: Crustal evolution of the Lower Paleozoic
3
orogen, Southwestern Gondwana margin
4
Agustín Ortiza,b*, Natalia Hausera, Raúl Becchiob, Néstor Suzañob, Alexis Nievesb,
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Alfonso Solab, Marcio Pimentela, Wolf Reimolda,c,d
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a
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Brasília, DF.
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b
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Geochronology Laboratory, Instituto de Geociências, Universidade de Brasília, 70910 900
GEONORTE - INENCO (Universidad Nacional de Salta – CONICET), Av. Bolivia 5150,
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A4400FVY, Salta, Argentina.
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c
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d
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*Corresponding author: [email protected]. Telephone: +54-387-4255441.
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Humboldt Universitäet zu Berlin, Unter den Linden 6, 10099 Berlin, Germany.
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Museum für Naturkunde – Leibniz-Institute for Evolution and Biodiversity Research, Invalidenstrasse 43, 10115 Berlin, Germany.
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1
Abstract The evolution of the rocks of the Lower Paleozoic Orogen in Puna, at the
3
Southwestern Gondwana margin, has been widely debated. In particular, the scarce amount
4
of geological and geochemical data available for the Diablillos Intrusive Complex, Eastern
5
Magmatic Belt, Southern Puna, require a further study for new evidence towards the
6
understanding of sources, magmatic processes and emplacement of magmas, in order to
7
better comprehend the crustal evolution in this setting. We present new combined U–Pb
8
and Hf isotope analyses on zircon by LA-MC-ICP-MS from monzogranite, granodiorite
9
and diorite rocks of the Diablillos Intrusive Complex. We obtained
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206
Pb/238U concordant
weighted average ages of 517±3 Ma and 515±6 Ma for the monzogranite and diorite,
11
respectively, and a concordant age of 521±4 Ma for the granodiorite. These ages permit to
12
constrain the climax of magmatic activity in the Diablillos Complex around ~515-520 Ma,
13
while the emplacement of the complex took place between ~540 Ma and 490 Ma
14
(representing a ca. 50 Ma magmatic event). Major and trace element data, initial
15
values varying from 0.70446 to 0.71278, positive and negative ɛNd(t) values between +2.5
16
and -4, as well as ɛHf(t) for zircon data between + 3 and -3 indicate that the analyzed
17
samples represent contaminated magmas. The ɛHf(t) and the ɛNd(t) values for this complex
18
specify that these rocks are derived from interaction of a dominant Mesoproterozoic
19
crystalline and/or a metasedimentary source and juvenile mantle-derived magmas, with a
20
TDM model age range of ~1.2-1.5 Ga, with later reworking during lower Paleozoic times.
21
The combined data obtained in this contribution together with previous data, allow us to
22
suggest that the formation of the Eastern Magmatic Belt of the Puna was part of a long-
23
lived magmatic event during Early Paleozoic times. Whereby the granitoids of the Eastern
87
Sr/86Sr
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Magmatic Belt formed through intra-crustal recycling at an active continental margin, with
2
minor contributions from juvenile material in the back-arc setting.
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Key words: Southwestern Gondwana; Puna Argentina; Lower Paleozoic Orogen; Zircon
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U-Pb geochronology; Hf isotopes.
1. Introduction
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The Neoproterozoic to Early Paleozoic rocks (present-day Eastern Cordillera and
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Puna) in the Central Andes, NW Argentina (Fig. 1), formed at the Southwestern Gondwana
10
margin. Their evolution has been widely debated and traditionally two orogenic events
11
have been distinguished: the Pampean (Upper Precambrian–Lower Cambrian) and the
12
Famatinian (Upper Cambrian–Lower Silurian) stages involving subduction processes with
13
formation of magmatic arcs, followed by accretion/collision of para-auhocthonous and
14
allochthonous terranes (e. g., Ramos et al., 1986; Ramos, 1988, 2008; Loewy et al., 2004;
15
Rapela et al., 2007; Cordani et al., 2009; Collo et al., 2009; Drobe et al., 2009; Hauser
16
et al., 2011). Due to the lack of evidence and indicators for collisional processes - such as
17
suture zones, presence of ophiolites related to oceanic crust, or high P/low T metamorphism
18
indicative of subduction-accretion regimes, alternative models, for the evolution of the
19
Puna and Eastern Cordillera were suggested. This include a geodynamic evolution
20
dominated by intra-crustal recycling at an active continental margin, with minor
21
contributions from juvenile materials in the back-arc setting (e.g., Bahlburg and Hervé,
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1997; Becchio et al., 1999; Bock et al., 2000; Lucassen et al., 2000; Coira and
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Koukharsky (2002), Kleine et al., 2004; Büttner et al., 2005; Coira et al. (2009),
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Zimmermann et al., 2014; Rapela et al., 2016; Bahlburg et al., 2016). In particular, the magmatic record in the Puna region (currently a high-plateau area,
4
4000 meter above sea level) (Fig. 1), has been associated with the Famatinian orogeny, that
5
was initiated in the Lower Tremadoc (Moya et al., 1993), reaching its peak in the Arenig
6
(Rapela et al., 1992; Coira et al., 1999). It includes two N-S trending magmatic belts of
7
~300 km extension: an Eastern Eruptive Belt (Méndez et al., 1973) represented by
8
calcalkaline and peraluminous granitoids, metadacites and metarhyolites, and a Western
9
Eruptive Belt (Palma et al., 1986) characterized by metaluminous I-type granitoids (Fig.
10
1). Taking into account that so-called eruptive belts also include plutonic and subvolcanic
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rocks, we prefer to use the term Magmatic Belts.
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The origin of the Eastern and Western Magmatic Belts has been variably interpreted
13
as two magmatic arcs related to the accretion of allochthonous, parautochthonous, or
14
autochthonous terranes along the southwestern margin of Gondwana (e.g., Coira et al.,
15
1982; Ramos, 1986; Conti et al., 1996). Other authors (e.g., Dalziel and Forsythe, 1985;
16
Ramos 1988; Rapela et al., 1992) interpreted these magmatic belts as two magmatic arcs
17
related to subduction but without accretion processes. Or these magmatic belts were
18
assigned to an early to middle Ordovician back-arc setting (Coira et al., 1999, 2009), or
19
they were considered as part of late stages of the Oclóyic orogeny (Bahlburg, 1990). Coira
20
et al. (1999) proposed an active arc setting for the magmatism of the Western Magmatic
21
Belt (Upper Cambrian-Lower Arenig). They proposed for the Eastern Belt an occurrence of
22
an oblique convergence regime at that time, generating important strike faults that to the
23
north resulted in partial melting of the crust and in a subduction zone to the south. These
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processes were based on models involving collision of terranes, and outcrops of oceanic
2
suture zones proposed by Ramos et al., 1986; Ramos, 1988; Loewy et al., 2004; Rapela
3
et al., 2007; Ramos, 2008, among others. Recently, Zimmermann et al. (2014)
4
demonstrated that this alleged ophiolite suture did not exist, thus, refuting the assumption
5
of plate boundaries and the collision models. Instead, these authors strengthened the models
6
of geodynamic evolution of the Magmatic Belts in the Puna.
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In the past years, the advances of zircon isotope geochemistry have provided a
8
formidable means to date magmatic and metamorphic events by U-Pb geochronology (e.g.,
9
Hoskin and Schaltegger, 2003), and to employ zircon successfully for the tracing of
10
petrogenetic processes in metamorphic and magmatic rocks (Hawkesworth and Kemp,
11
2006; Castiñeiras et al., 2011). Some recent compositional investigations of zircon have
12
not been directly related to geochronology, but instead based on the ability of zircon to
13
record petrogenetic processes in igneous systems through analysis of Hf isotopes (e.g., Zeh
14
et al. 2007; Dong et al. 2013; Li et al. 2014; and references therein). This isotope system
15
provides a useful tool to discriminate the origin of rocks in general, and may potentially
16
reveal the nature of the source rock (e.g., Guo et al., 2007; Willner et al., 2008; Liu et al.,
17
2009; Miskjovic and Schaltegger, 2009; Claiborne et al., 2010; Nardi et al., 2012;
18
Bahlburg et al., 2009, 2016). Thus, providing important parameters for modelling crustal
19
evolution (e.g., Hawkesworth and Kemp, 2006).
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In the Eastern Magmatic Belt, Becchio et al. (2011) defined the Diablillos Intrusive
21
Complex (CID by its Spanish name Complejo Intrusivo Diablillos; Fig. 2). This complex is
22
mainly composed of granitoids, including diorites, and mafic dykes. Recently, Suzaño et
23
al. (2015; 2017) proposed that the main process in the evolution of this Complex was
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magma mixing between melts of granitic and dioritic composition. Until now, there is only
2
one age datum for the emplacement of the CID, a U-Pb weighted average age of 501 Ma
3
with a 1σ error of ± 17 Ma (Suzaño et al., 2015). Moreover, the magma sources for this
4
complex have only been inferred, based on Sr-Nd analyses from nearby area (e.g.,
5
Viramonte et al., 2007), possibly involving juvenile mantle or a reworked crustal source,
6
or both.
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The scarce amount of geological and geochemical data available for this sector of
8
the Southern Puna makes the CID an essential place to study and provide new evidence for
9
the development of the large volumes of magmatic rocks in the Eastern Magmatic Belt. In
10
this contribution, we have obtained three new zircon U-Pb ages by Laser Ablation Multi-
11
Collector Inductively Coupled Plasma Mass Spectrometry (LA-MC-ICP-MS) for diorite,
12
granodiorite, and monzogranite rocks from the Diablillos Intrusive Complex. Additionally,
13
we present the first Lu-Hf isotope data for zircon from this complex, and for the Southern
14
Puna, as well as new geochemical and Sr-Nd isotope data for the igneous rocks of the CID.
15
We thus, contribute to the understanding of sources, magmatic processes and emplacement
16
of magmas during Cambrian-Ordovician times in this sector of the Puna, helping to unravel
17
the geotectonic setting of the Lower Paleozoic Orogen including the Eastern Magmatic Belt
18
at the Southwestern Gondwana margin.
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2. Geological setting
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The Southern Puna basement in the surroundings of the Centenario, Ratones and
4
Diablillos salars (Fig. 1) comprises medium- to high-grade Neoproterozoic-Lower
5
Cambrian metamorphic rocks of the Pachamama Formation (Hongn and Seggiaro, 2001)
6
that include mica schist and
7
calcsilicates and amphibolites (Fig. 1). In addition, these salars are surrounded by low- to
8
medium-grade Neoproterozoic-Lower Cambrian rocks (Fig. 1) consisting of slates,
9
phyllites, metaquartzites, and sillimanite schists that constitute the Rio Blanco
10
Metamorphic Complex (Hongn and Seggiaro, 2001). The metamorphic peak in Puna, of
11
high T and low to medium P occurred at ~510 Ma (Sm-Nd isochron of plagioclase,
12
clinopyroxene, and garnet, and U-Pb in titanites from calcsilicate layers, Becchio et al.
13
(1999) and Lucassen and Becchio, 2003). The metamorphic rocks are intruded by
14
Cambrian-Ordovician granitoids of the Oire Eruptive Complex (CEO; Oire Formation by
15
Turner, 1961, redefined as a complex by Blasco and Zappettini, 1995) from the Eastern
16
Magmatic Belt. The granitoids of the CEO comprise fine- to coarse-grained, variably
17
equigranular or porphyritic granites and granodiorites that occur over a ~100 km long N-S
18
trending belt to the east of the Centenario, Ratones and Diablillos salars (Fig. 1). Further to
19
the east, these granitoids gradually change to orthogneiss with different degrees of
20
mylonitization.
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both ortho- and paragneisses that are intercalated with
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The Ordovician basement is composed of low-grade marine metasediments, with
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intercalated volcanic rocks of the Falda Ciénaga Formation (Llanvirn-Llandeilian,
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Aceñolaza et al., 1976) that occur at the western edge of the Centenario salar (Filo
2
Copalayo, Fig. 1) and at Filo de Oire Grande. In tectonic contact, on the Filo Copalayo
3
(Mon and Hongn, 1996), medium- to high-grade metamorphic units of mica schist and
4
granitic-granodioritic-tonalitic orthogneiss crop out (Fig. 1). They have an apparent
5
crystallization age of 467±10 Ma (U-Pb age on zircon, Domínguez et al., 2006). At the
6
eastern edge of the Centenario salar, a volcano-sedimentary unit (Cochinoca- Escaya
7
Magmatic-Sedimentary Complex, Coira et al., 2004) is composed of slates,
8
metapsammites, and low- to medium-grade mica schist (Fig. 1). This sequence is
9
intercalated with the products of bimodal magmatism: metarhyolite, metadacite, and
10
metabasalt (Fig. 1). These metavolcanic rocks are intruded by syenogranites and
11
leucogranites of the CEO, whereby hornfels and mottled schist were generated. Viramonte
12
et al. (2007) defined U-Pb crystallization ages for metarhyolites of 485-472 Ma and
13
concluded that the plutonic units of the CEO were emplaced between 485 and 462 Ma.
14
More recently, farther north, Insel et al. (2012) dated these granitoids at 491 to 481 Ma (U-
15
Pb zircon age).
18
19
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2.1. Local Geology, Field Relations and Petrography
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The Diablillos Intrusive Complex (CID) was emplaced in the Inca Viejo Range of
20
the Southeastern Magmatic Belt. The Inca Viejo Range, located to the west of the Filo de
21
Oire Grande, is a mountain chain that separates the Ratones and Centenario salars to the
22
northwest from the Diablillos salar to the southeast (Fig. 2, 3a).
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Based on geographic location and macroscopic characteristics of the rocks, Becchio
2
et al. (2011) associated the granitoids of the CID as cogenetic with the granitoids of the
3
CEO (Fig. 2). Furthermore, Suzaño et al. (2015; 2017), based on petrogenetic models,
4
proposed a common and cogenetic origin for both complexes.
6
Recently, Suzaño et al. (2015) identified the main lithologies of the CID, namely monzogranite, granodiorite and tonalite, diorite, and mafic dykes (Fig. 2).
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On the eastern side of the Inca Viejo range, the rocks of the CID intrude low- grade
8
metamorphic rocks of the Rio Blanco Metamorphic Complex (Fig. 2; Hongn and
9
Seggiaro, 2001). On the western side, the monzogranite intrudes moderate grade
10
metamorphic rocks, such as mica schists, and para- as well as ortho-gneisses that have been
11
included with the Pachamama Formation (Fig. 2; Hongn and Seggiaro, 2001). In addition,
12
subvolcanic bodies of the Ratones Andesites and Inca Viejo Formation of Miocene age
13
(González, 1984) crop out to the north and south of the range. They are associated with
14
hydrothermal alteration and mineral concentrations of economic interest. Lastly, the
15
basement rocks and the Neogene subvolcanic intrusives are covered by Quaternary
16
volcanics of the Incahuasi Formation (Aceñolaza et al., 1976), and by alluvial/fluvial
17
materials including the occurrences of salars that are of particular economic interest for
18
lithium and borate brines.
19
The recognized units in the Diablillos Intrusive Complex are:
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Monzogranite crops out on the northern and eastern sides of the range and intrudes
21
phyllites, slates, and mottled slates of the Rio Blanco Metamorphic Complex to the east
22
(Fig. 3e), as well as mica schist, para- and ortho-gneiss of the Pachamama Formation to the
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west. The monzogranite unit has an irregular shape. In the central part of the range, it
2
occurs in transitional contact with granodiorite and tonalite (Fig. 2). The monzogranite is medium- to coarse-grained and has a porphyritic texture. It
4
varies from a light to dark grey color, according to biotite content in the matrix. The
5
porphyritic monzogranites have euhedral to subhedral K-feldspar (1 to 3 cm large) and
6
plagioclase (1 to 1.5 cm) phenocrysts embedded in a coarse-grained matrix composed by
7
K-feldspar, plagioclase, blue quartz, and abundant biotite. Occasionally, a slight magmatic
8
foliation is observed by the preferred orientation of Mafic Magmatic Enclaves (MME) and
9
syn-magmatic pegmatite dykes. On the northern side of the range, MME are rare. On the
10
eastern side of the Inca Viejo Range, this rocks form lobate and elongate (less than 700 m
11
long) bodies and include abundant enclaves of various compositions, such as MME,
12
granodiorite, tonalite, and diorite in the form of dismembered bodies of irregular shape
13
(Fig. 2). The dismembered diorite bodies (up to 5 m long), granodiorite, and tonalite
14
enclaves, and the MME have a heterogeneous distribution and have rounded or elliptical
15
shapes, with varying sizes up to 1.5 m. The contacts between the monzogranite and mafic
16
dismembered bodies and MME are sharp (Fig. 3b, c), and sharp, transitional, or lobate
17
toward granodiorite and tonalite enclaves (Fig. 3b, d).
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Some subhedral or subrounded feldspar phenocrysts seem to migrate mechanically
19
from the monzogranite to the granodiorite enclaves (Fig. 3d). In areas where the
20
monzogranites are intruded by the mafic dismembered bodies, subrounded K-feldspar
21
phenocrysts, plagioclase and quartz are incorporated into the mafic bodies, showing typical
22
magma mixing-mingling textures, such as mantled plagioclase and ocelli quartz (Hibbard,
23
1991).
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The modal composition shown in Fig. 4 classifies sample D-13-01 as a
2
monzogranite with ~55 vol.% K-feldspar, ~24 vol.% plagioclase, and ~21 vol.% quartz. In
3
thin section (Fig. 5a), this rock has a holocrystalline, porphyritic texture of medium to
4
coarse grain size. It is formed mainly by K-feldspar and plagioclase phenocrysts, with K-
5
feldspar, quartz, plagioclase and abundant biotite in the matrix. Accessory minerals are
6
muscovite, apatite, titanite, zircon, and undifferentiated opaque minerals.
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The 0.5 to 1 cm phenocrysts are microperthite, and are subhedral with irregular rims
8
and inclusions of euhedral plagioclase, quartz, biotite and muscovite; some K-feldspar is
9
slightly altered to phyllosilicates. The plagioclase phenocrysts (1-4 mm) are oligoclase to
10
andesine (by the Michael-Levy method) and are overgrown by K-feldspar; other
11
phenocrysts are zoned. Some quartz has undulous extinction and forms triple joints in the
12
matrix, owing to the occasionally observed ductile foliation. Biotite crystals are subhedral,
13
frequently forming assemblages with zircon and titanite. Muscovite is subhedral and occurs
14
as fine- grained aggregates (Fig. 5a).
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Granodiorite: This unit, together with the tonalite, is the most widespread unit in
17
the Inca Viejo Range, occurring over an area ~7 km long and 2 km wide. The granodiorite
18
has dark to light grey color, crops out on the mid-eastern side of the range, has a
19
transitional contact with the monzogranite to the east, and is intercalated with the tonalite to
20
the west. On the northeastern side of the range, the granodiorites intrude phyllites, slates
21
and mottled slates of the Rio Blanco Metamorphic Complex.
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The granodiorite is porphyritic. It is characterized by blue quartz and subhedral
2
(subrounded) K-feldspar and plagioclase phenocrysts (up to 2 cm in size) (Fig. 3c, d). They
3
are immersed in a medium- to fine-grained matrix composed of the same minerals plus
4
biotite. Abundant concentrations of MME (similar to those in the monzogranite) of 7 to 50
5
cm size, with sharp to transitional contacts were observed in some parts of the granodiorite
6
(Fig. 3c). In addition, tonalite enclaves and minor meta-wacke xenoliths (probably from the
7
Rio Blanco Metamorphic Complex) were identified occasionally.
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The modal composition (Fig. 4) classifies sample D-13-02 as a granodiorite with
9
~40 vol.% plagioclase, ~35 vol.% K-feldspar, and ~25 vol.% quartz. At the microscopic
10
scale, the granodiorite (Fig. 5b) has a holocrystalline porphyritic texture with plagioclase,
11
microcline and quartz as phenocrysts. The matrix is formed by medium- to fine-grained
12
quartz, plagioclase, K-feldspar, abundant biotite, and muscovite. Accessory minerals are
13
zircon, titanite, apatite, and opaques.
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Phenocrysts of K-feldspar (microcline, 0.5-1 cm) are surrounded by, and have in
15
their cores and along their rims inclusions of, euhedral plagioclase, quartz and biotite. Some
16
K-feldspar and plagioclase are altered to sericite in their cores (Fig. 5b). In some cases,
17
these K-feldspar phenocrysts have a plagioclase rim defining a rapakivi texture, typically
18
formed in a magma-mixing environment (Hibbard, 1991). The plagioclase phenocrysts
19
(0.2-1 cm) are subhedral. Less frequently, zoned plagioclase phenocrysts are mantled with
20
biotite and overgrown by plagioclase. The quartz phenocrysts (4 cm) are subrounded and
21
some are closely aggregated with matrix. Biotite is abundant; ilmenite crystals may be
22
partially replaced by titanite along the rims.
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Tonalite: was mapped together with the Granodiorite due to their close relationship
2
(Suzaño et al., 2015). The tonalite crops out on the top of the Inca Viejo Range. It forms a
3
~50-70 m wide continuous body, in gradational contact with the granodiorite to the east. It
4
has a dark grey color and varies from equigranular and fine-grained to slightly porphyritic,
5
with quartz (2 cm) and plagioclase (2-2.5 cm) phenocrysts. The matrix is composed of
6
plagioclase, quartz, and abundant biotite. This rock also has MME (less than 2 m wide)
7
with sharp contacts (Fig. 3d), small granodiorite enclaves (5-7 cm), and some cross-cutting,
8
narrow (<30 cm) pegmatite dykes.
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The modal composition (Fig. 4) classifies sample D-13-15 as a tonalite with ~48
10
vol.% plagioclase, ~24 vol.% K-feldspar, and ~28 vol.% quartz. In thin section, the tonalite
11
(Fig. 5c) is porphyritic to equigranular, with some small (3 mm) plagioclase phenocrysts.
12
The matrix has medium to fine grain size and is composed of plagioclase, poikilitic quartz,
13
and abundant biotite; accessory minerals are zircon, apatite, and opaques.
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Some plagioclase phenocrysts have a biotite rim and may carry biotite inclusions
15
(Fig. 5c); some plagioclase phenocrysts are slightly altered to sericite. Biotite, in some
16
instances, forms crystal aggregates with opaque minerals and has euhedral epidote
17
inclusions. The poikilitic quartz has inclusions of plagioclase, biotite (Fig. 5c), acicular
18
apatite, fine-grained epidote, and amphibole.
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Diorite: represents the least widespread unit in the complex; it is distributed as
20
dismembered bodies and as MME in the monzogranite, granodiorite and tonalite facies. On
21
the eastern side of the range, it crops out as globular bodies up to 200 m wide and 400 m
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long (Fig. 2). At the southern edge of the range diorite occurs as dismembered bodies and
2
MME. Contacts with monzogranite are sharp, but transitional and lobate with the
4
granodiorite and tonalite. There are places where the diorites have monzogranite enclaves
5
with sharp contacts and irregular or tabular shapes (up to 1 m size) (Fig. 3b, c). This rock
6
has a dark green color, with a fine-grained, equigranular to porphyritic texture. It has
7
mantled plagioclase phenocrysts up to 1 cm in size, and blue ocelli pheno-xenocrysts of
8
quartz immersed in an aphanitic, dark green matrix.
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The modal composition for sample D-13-40 is ~96.5 vol.% plagioclase, ~1.4 vol.%
10
K-feldspar, and ~2.1 vol.% Quartz, which allows to classify it as a diorite (Fig. 4). The very
11
fine grained matrix is composed of plagioclase, amphibole, quartz, biotite, apatite, epidote
12
and opaque minerals (Fig. 5d). Several disequilibrium features are observed in the
13
plagioclase, such as individual crystals with poikilitic texture, zoned crystals with
14
polysynthetic twinning, and porphyritic agglomerates with quartz exsolution (Fig. 5d).
15
Some plagioclase phenocrysts have hornblende rims, and some of these rims are altered to
16
biotite. The quartz pheno-xenocrysts are rare, anhedral, and subrounded; others have
17
reaction rims displaying hornblende and biotite along grain margins.
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Mafic dykes: crop out on the northwestern side of the Inca Viejo Range, up to 40 m
19
thick. These concordant dykes intrude the monzogranite of the CID and the medium grade
20
metamorphic rocks, generally along a NE-SW strike trend. In addition, these dykes intrude
21
the granitic rocks of the CEO in the southern Filo de Oire Grande, to the east of the Inca
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1
Viejo Range (Fig. 2). The dykes have fine-grained, equigranular to porphyritic textures
2
with plagioclase phenocrysts in a fine-grained, dark green matrix. The monzogranites and the mafic bodies are intruded by syn-magmatic concordant
4
and discordant tabular pegmatite dykes (<2m wide) composed of K-feldspar, quartz,
5
muscovite, tourmaline and epidote. Both dyke types – the diorite and the pegmatite
6
varieties - represent later stages of the magmatic event of the CID (Suzaño et al., 2015).
7
3. Analytical Procedures
8
3.1. Geochemical analyses
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Sample preparation for major and trace element analyses of five samples from the
10
CID and three from the CEO were performed at the Geochemistry Laboratory of Geonorte
11
Institute, Salta National University, Argentina. This consisted of washing the field samples,
12
cutting off weathered and veined portions, crushing, quartering, sieving, and grinding the
13
samples with a Herzog mill with oscillatory rings of tungsten carbide. Major elements, Ni
14
and Sc (ppm) were analyzed by ICP-OES and trace element analyses were carried out by
15
ICP-MS at the ACME Laboratories in Canada (Table 1). After the analyses Tungsten was
16
disregarded due to contamination by the tungsten carbide mill. For information about
17
standards, limits of detection, etc., refer to http://acmelab.com/services/.
19
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3.2. U-Pb and Lu-Hf isotopes on zircon and Sr-Nd isotopes on whole rock samples
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After crushing, grinding and panning approximately 5 kg of each of the three
2
selected samples (monzogranite, granodiorite and diorite), zircon concentrates were
3
extracted with a Frantz isodynamic separator at the Geochemistry Laboratory in the
4
Geonorte Institute, Salta National University, Argentina. Final purification was achieved by
5
hand-picking of zircon with an acupuncture needle using a binocular microscope at the
6
Geochronology Laboratory of the University of Brasilia, Brazil. In order to target zircon
7
grains for isotope analysis, cathodoluminescence (CL) images of polished mounts were
8
obtained with a JEOL Quantas 450 Scanning Electron Microscope at this same laboratory.
9
The grains were polished and cleaned with 3% nitric acid.
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U-Pb and Lu-Hf analyses were performed on zircon grains from the three selected
11
samples, using a Thermo-Fisher Neptune MC-ICP-MS coupled with a Nd:YAG UP213
12
New Wave laser ablation system at the Laboratory of Geochronology of the University of
13
Brasilia. For the U-Pb analyses (Table 2) the diameter of the laser beam was 30 µm. A
14
frequency of 7 Hz was used, at 80% energy and 40 counts per second. GJ-1 zircons were
15
used as standards to normalize isotopic fractionation during isotope analysis. To monitor
16
precision and accuracy over the course of this study, the TEMORA reference zircon was
17
measured as an unknown obtaining a
18
session, 206Pb/238U age of 409 ± 3 Ma for the diorite session, and a 206Pb/238U age of 409 ±
19
5 Ma for the granodiorite session. Common lead
20
(204Hg+204Pb) masses. In none of the analyzed zircon grains a common Pb correction was
21
necessary due to the very low signal of mass
22
Reported errors are propagated by quadratic addition [(2SD^2+2SE^2)1/2] of external
23
reproducibility and within-run precision. The external reproducibility is represented by the
206
Pb/238U age of 409 ± 4 Ma for the monzogranite
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204
204
Pb was monitored using the
Pb (< 30 cps) and high
206
202
Hg and
Pb/204Pb ratios.
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207
Pb/206Pb
standard deviation (SD) obtained from repeated analyses (n = 20, ~1.1 % for
2
and up to ~2 % for 206Pb/238U) of the standard zircon GJ-1, performed during the analytical
3
sessions, and the within-run precision is the standard error (SE) calculated for each
4
analysis. The U-Pb analyses on zircon grains were carried out using the standard-sample
5
bracketing method (Albarède et al., 2004). Two to four unknowns were analyzed between
6
GJ-1 analyses, and 206Pb/207Pb and 206Pb/238U ratios were time corrected. On smaller zircon
7
grains single spot laser-induced fractionation of the 206Pb/238U ratio was corrected using the
8
linear regression method (Košler et al., 2002). The raw data were processed off-line and
9
reduced using an Excel worksheet (Buhn et al., 2009). Concordia diagrams (2σ error
10
ellipses) and Concordia ages were calculated using the Isoplot/Ex software (Ludwig,
11
2003).
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Lu-Hf isotopes (Table 3) were analyzed on selected zircon grains previously
13
analyzed for U-Pb isotopes. When possible, the two spot analyses were placed as close as
14
possible in order to analyze portions of the zircon grain with the same U and Pb isotopic
15
characteristics. Lu-Hf isotopic data were collected for 40-50 seconds ablation time and with
16
a 40 µm diameter spot size, at 85% energy. The signals of the interference-free isotopes
17
171
18
interferences of
19
calculated using the isotopic abundances of Lu and Hf proposed by Chu et al. (2002). The
20
contemporaneous measurements of
21
bias of Yb using a
22
isotope ratios were normalized to the
23
detailed description of the procedures and methods is given by Matteini et al. (2010).
173
Yb and
175
176
Lu were monitored during analysis in order to correct for isobaric
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Yb and
173
176
Lu on the
171
176
Hf signal. The
Yb and
176
Yb and
176
Lu contribution was
173
Yb provide a method to correct for mass-
Yb/171Yb normalization factor of 1.132685 (Chu et al., 2002). Hf 179
Hf/177Hf value of 0.7325 (Patchett, 1983). A
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The εHf(t) values were calculated using the decay constant λ=1.865*10-11 proposed
1
176
Lu/177Hf and
176
Hf/177Hf CHUR values of 0.0332 and
by Scherer et al. (2006), and the
3
0.282772 proposed by Blichert-Toft and Albarède (1997). A two-stage TDM age was
4
calculated from the initial Hf isotopic composition of the zircon, using an average crustal
5
Lu/Hf ratio (Gerdes and Zeh, 2009; Nebel et al., 2007). The initial Hf composition of
6
zircon represents the
7
namely the possibly concordant U-Pb age previously obtained on the same crystal. Two-
8
stage depleted mantle Hf model ages (TDM Hf) were calculated using
9
and
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Hf/177Hf value calculated at the time of zircon crystallization,
176
Lu/177Hf=0.0384
Hf/177Hf=0.28325 for the depleted mantle (Chauvel and Blichert-Toft, 2001) and
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176
11
Wedepohl, 1995). Before Hf isotope measurements on zircon, replicate analyses of a 200
12
ppb Hf JMC 475 standard solution doped with Yb (Yb/Hf=0.02) were carried out
13
(176Hf/177Hf=0.282162±13 2s, n=4). During the analytical session replicate analyses of GJ-1
14
standard zircon were done obtaining a
15
agreement with the reference value for the GJ standard zircon obtained by Morel et al.
16
(2008). The εHf values for each single grain were recalculated to the U-Pb age previously
17
obtained on the same zircon grain. For slightly discordant analyses, the
18
207
19
respectively.
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Hf/177Hf ratio of 0.282006±16 2σ (n =25), in
206
Pb/238U and
Pb/206Pb ages obtained were used for younger (<1 Ga) and older (>1 Ga) grains,
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Lu/177Hf value of 0.0113 for the average crust (Taylor and McLennan, 1985;
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Bulk rock Sm-Nd isotopic analysis of eight selected samples (Table 5) was carried
21
out at the Geochronology Laboratory of the University of Brasília, Brazil. Sample
22
dissolution was done in Teflon Savillex beakers or in Parr-type Teflon bombs. Sm and Nd
23
extraction from whole-rock powders followed the technique of Gioia and Pimentel (2000).
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149
Sm–150Nd spike was used. Sr, Sm and Nd samples were loaded onto Re
A mixed
2
filaments of a double filament assembly. Sm and Nd isotopic analyses were carried out
3
using a Finnigan Triton mass spectrometer. TDM values were calculated according to the
4
two-stage model as presented by Liew and Hofmann (1988).
5
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3.3. Microprobe analysis
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Microprobe analysis (Table 4) was performed on zircon grains previously analyzed
10
by the U-Pb laser ablation method using a JEOL JXA-8230 Superprobe at the University of
11
Brasilia, Brazil. Concentrations for HfO2, SiO2, ZrO2, Y2O3, UO2 and ThO2 were
12
determined. The conditions for analysis included 20kV energy, 10nÅ of current for Si, Zr
13
and Hf, and 100nÅ for U, Th and Y. Counting times on peaks were 10 s for Si and Zr, and
14
60 s for the other elements. Counting times on background were 5 s for Si and Zr, and 30 s
15
for the other elements.
17
18
19
20
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4. Results
8
4.1. Geochemistry
9
4.1.1. Major, trace and rare earth elements
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Major and trace element compositions of one monzogranite, a granodiorite, a
12
tonalite, and two diorite samples from the CID and a monzogranite, a granodiorite and a
13
diorite mafic dyke sample from the CEO are compiled in Table 1.
14
In the TAS diagram for granitoids (Middlemost, 1994) (Fig. 6a), the monzogranite unit is
15
represented by samples D-13-01 (CID) and CB-13-05 (CEO); the former plots into the
16
quartz monzonite field. However, we continue to refer to this sample as monzogranite
17
based on its modal composition and field evidence. We assume that the low silica content
18
shown by this sample is due to the mixed nature of samples. We confirmed this assumption
19
based on our field relationships and geochemistry of the CID and by previous studies by
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Suzaño et al. (2015; 2017), which defined that the main magmatic process in the formation
2
of this complex was magma mixing. The granodiorite and tonalite units are represented by samples D-13-02
4
(granodiorite, CID) and D-13-15 (tonalite, CID) that plot into the granodiorite and tonalite
5
fields, respectively; BA-13-07 (granodiorite, CEO) plots into the granodiorite field as well.
6
Finally, the diorite unit is represented by samples D-13-40 (CID), CB-13-02 (CID) and CB-
7
13-04 (CEO), all plotting into the quartz diorite field.
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In the A/CNK vs A/NK plot (Shand, 1943) (Fig. 6b) the diorite samples fall into
9
the metaluminous field representing I-type rocks; in contrast, the tonalite, granodiorite and
10
monzogranite samples lie in the peraluminous field representing S-type granitoids. There is
11
one sample (D-13-15, tonalite) that falls into the metaluminous field, near the boundary to
12
the peraluminous field. In the AFM diagram (Irvine and Baragar, 1971) (Fig. 6c), the
13
intermediate rocks of the CID and CEO show a calc-alkaline trend, whereas the mafic rocks
14
show a slight tendency towards a more primitive origin, with tholeiitic characteristics
15
(sample CB-13-04, dioritic basic dyke).
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The Harker diagrams (Harker, 1909) illustrate two contrasting sample groups (Fig.
17
7): on the one hand, the basic plutonic rocks and on the other, the intermediate plutonic
18
rocks. Al2O3 contents of all samples do not vary significantly (~14 to 16 wt.%). The SiO2
19
content in the diorite varies from ~50 to 54 wt.%; the granodiorite and tonalite show SiO2
20
values of ~60-65 wt.%; and the monzogranite has ~61 to 69 wt.% SiO2. The mafic rocks
21
have values of FeOt (~8-10 wt.%), MgO (~4-8 wt.%), CaO (~7-11 wt.%), Na2O (~2-3
22
wt.%) and K2O (~0.5-2.5 wt.%); the felsic rocks have contents of FeOt (~4-8 wt.%), MgO
23
(~2-4 wt.%), CaO (~1.5-4 wt.%), Na2O (~2.5-3.5 wt.%) and K2O (~2-5 wt.%). The two
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main groups have regular variation trends. MgO, FeOt, TiO2 and CaO are negatively
2
correlated with increasing SiO2, whereas Na2O and K2O show positive correlations (Fig. 7).
3
Trace element versus silica variation diagrams (Fig. 8) show decreasing Sr with
4
increasing SiO2, suggesting plagioclase fractionation. Rb and Ba display positive
5
correlation trends with increasing silica. In the same way, the Zr content in the felsic rocks
6
shows the same behavior with increasing SiO2 due to zircon saturation. Sample CB-13-04
7
(Diorite mafic dyke) has 94 ppm Ni and is the most primitive rock of the basic group (in the
8
AFM diagram, Fig. 6c, with slight tholeiitic character).
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In the upper crust-normalized trace element spider diagram (Taylor and
10
McLennan, 1985) (Fig. 9a), the intermediate rocks show enrichment in the Large-Ion
11
Lithophile Elements (LILE) (i.e., Ce, Rb, Th and K) compared to the mafic rocks. Besides,
12
the felsic group shows a slight depletion in High Field Strength Elements (HFSE) (i.e., Nb,
13
P, Zr and Ti), suggesting the contamination with mafic magmas. In addition, this group
14
shows negative anomalies for Sr and Ba, suggesting plagioclase fractional crystallization
15
during magma evolution (Dong et al., 2013). In Fig. 9a the mafic rocks show a minor
16
depletion of LILE and strong depletion of HFSE compared to the intermediate rocks, with
17
the exception of sample CB-13-04 (CEO, diorite mafic dyke) that has a stronger depletion
18
of the LILE and HFSE.
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The chondrite-normalized REE patterns (Fig. 9b) for the intermediate rocks show
20
enrichment of the LREE (LaN/YbN=3.91 - 11.84) and a negative slope for the HREE,
21
compared to the mafic rocks including those of Zimmermann et al. (2014) from the
22
Western Magmatic Belt, with LaN/YbN=1.62-3.67 and that have a crustal signature as
ACCEPTED MANUSCRIPT
indicated by Dong et al. (2013). In addition, the intermediate rocks display a discernible
2
negative Eu anomaly (Eu/Eu*= 0.50-0.70), indicative of plagioclase fractionation. The
3
mafic rocks show a slight depletion of the HREE compared to the felsic rocks, with the
4
exception of sample CB-13-04 (CEO, diorite mafic dyke) with even stronger depletion of
5
the LREE (Fig. 9b).
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4.2. Zircon U-Pb geochronology
2
4
Three samples representing the main units of the CID, namely monzogranite, granodiorite and diorite, were selected for U-Pb geochronology.
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All three samples show similar populations of colorless to pinkish-brown zircon
6
grains of 100-300 µm sizes. Cathodoluminescence (CL) images show that the grains
7
have variable shapes from elongate euhedral and subhedral, to bipyramidal to
8
subrounded. Some grains display internal magmatic zonation, whereas others have
9
inherited zircon cores (Fig. 10b, d, f). Some grains from the granodiorite have
10
irregularly shaped dark cores (Fig. 10d). Microprobe analyses of zircon grains from
11
monzogranite, granodiorite and diorite indicate variable Th/U ratios from 0.14 to 2.0,
12
0.14 to 1.2, 0.52 to 2.1, respectively (Table 4), which are all typical for a magmatic
13
origin (Hoskin and Schaltegger, 2003).
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14
Sample D-13-01, Monzogranite:
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15
A subset of forty-five zircon grains out of fifty-three in total, with a concordance
17
of 100 ± 2%, were considered for age calculation and display an age distribution from
18
~490 Ma to 540 Ma and a weighted average
19
10a), which is interpreted as the emplacement age of this unit. The eight zircon grains
20
left apart from the age calculation yielded inherited ages ranging from ~615 to 1037 Ma
21
(Table 2).
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16
206
Pb/238U age of 517 ± 3 Ma (inset Fig.
22
The Probability Density Plot (PDP, Fig. 10g) resolves various zircon
23
populations with similar elongate euhedral and subhedral bipyramidal crystal shapes
ACCEPTED MANUSCRIPT (Fig. 10b). The PDP displays a small, four grain population of ~490 Ma age; a major
2
population (27 grains) of ~515-520 Ma age, a significant (13 grains) population of
3
~530-540 Ma age, and another (3 grains) of ~600-630 Ma age. In addition, there is a
4
group of three zircon core analyses representing a ~1000-1100 Ma population.
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5
7
Sample D-13-02, Granodiorite:
Analyses from eight zircon grains, with concordance of 100 ± 3%, out of fifteen 206
SC
6
Pb/238U age of 521±4 Ma (Fig. 10c), assumed to be
grains were used to calculate a
9
the emplacement age of this unit. The seven zircon grains left out from the age
10
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8
calculation gave inherited ages from ~577 to 1364 Ma (Table 2). The PDP (Fig. 10h) characterizes several zircon populations with similar
12
mineral characteristics (elongate euhedral and subhedral bipiramidals, Fig. 10d) and
13
displays two data points of ~490 Ma age, a significant population of ~515-520 Ma ages
14
(5 grains), a grain of ~540 Ma age, an older population of ~600 Ma (5 grains), and a
15
single older zircon grain of ~1300 Ma age.
17
18
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Sample D-13-40, Diorite: Twenty seven zircon grains out of a total of twenty-nine grains (Table 2), gave
19
analyses with a concordance of 100±2% and showed an age distribution from ~490 Ma
20
to 540 Ma (Fig. 10e). The
21
10e), which is interpreted as the emplacement age of this unit. The two zircon grains left
22
out from the age calculation gave inherited ages from ~650 to 1000 Ma (Table 2).
206
Pb/238U weighted average age is 515±6 Ma (inset Fig.
ACCEPTED MANUSCRIPT In the PDP (Fig. 10i) several zircon populations with varied ages and similar
2
habit characteristics (elongate euhedral and subhedral bipiramidal to subrounded, Fig.
3
10f) are distinguished. First, a five grain population of ~490 Ma occurs; second, a large
4
population of ~515-520 Ma ages (15 grains); a significant population of ~530-540 Ma
5
ages (7 grains); one grain of ~630 Ma, and one at ~1100 Ma.
6
4.3. Lu-Hf isotope data for zircon
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7
8
10
Lu-Hf analyses (Table 3) were performed on zircon grains previously analyzed by the U-Pb method.
11
Sample D-13-01, Monzogranite:
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Eight zircon crystals with varied U-Pb ages were analyzed in this sample,
14
showing 176Lu/177Hf and 176Hf/177Hf ratios between 0.001257 - 0.003581 and 0.282285 -
15
0.282465 respectively (Table 3).
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Zircon grains with ages between ~515 and 530 Ma show ɛHf(t) from -2.6 to
17
+0.4 (Fig. 11), with Hf TDM model ages between 1.5 and 1.3 Ga. Three Neoproterozoic
18
inherited zircon crystals with ages between 616 and 703 Ma show positive to negative
19
ɛHf(t) between +1.2 (age=616 Ma) and -1.9 (age=703 Ma) with Hf TDM model ages
20
between 1.4 and 1.6 Ga. Two zircon cores with Mesoproterozoic ages (1034 and 1056
21
Ma) gave positive ɛHf(t) values of 7.0 and 5.9 (Fig. 11), with Hf TDM model ages
22
between 1.3 and 1.5 Ga.
23 24
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Sample D-13-02, Granodiorite:
2
Eight zircon crystals with different U-Pb ages were analyzed for Lu-Hf isotopes
3
in this sample, disregarding one analysis (Grain N° Z8c) due to an anomalous ɛHf(t).
4
Otherwise, this sample yielded
5
176
Lu/177Hf ratios of 0.000979 - 0.004253 and
Hf/177Hf ratios of 0.282341 - 0.282536.
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176
Zircon grains from this sample with ages between ~485 to 540 Ma show ɛHf(t)
7
values from +2.1 and -2.2 (Fig. 12), with Hf TDM model ages between 1.2 and 1.6 Ga.
8
Also, this sample presents a negative maximum ɛHf(t) of -3.8 (Age=538 Ma), and a
9
positive ɛHf(t) maximum of +4.2 (Age=654 Ma) (Fig. 12), and Hf TDM model ages of 1.6 Ga and 1.2 Ga, respectively.
11 12
Sample D-13-40, Diorite:
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Eight zircon crystals with varied U-Pb ages were analyzed in this sample,
14
yielding 176Lu/177Hf ratios between 0.001075 - 0.005571 and 176Hf/177Hf ratios between
15
0.282192 - 0.282706 (Table 3).
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Zircon grains with ages between ~518 and 536 Ma show ɛHf(t) values from -9.5
17
to +9.1 (Fig. 11), with Hf TDM model ages between 1.9 and 0.9 Ga. A younger zircon
18
(Z48) of 495 Ma age yielded a positive ɛHf(t) of +2.1 and Hf TDM model age of 1.2 Ga,
19
and an inherited zircon crystal with a Mesoproterozoic age (1115 Ma) a positive ɛHf(t)
20
of +4.0 and Hf TDM model age of 1.6 Ga.
21
4.4. Sr-Nd isotope data
22
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Whole-rock Sr and Nd isotope analyses on eight samples (monzogranite,
2
granodiorite, tonalite, and diorite) from the CID and CEO were recalculated to 518 Ma,
3
the proposed average age of emplacement (Fig. 12; Table 4). The two monzogranites (D-13-01, CID; CB-13-05, CEO) have
4
147
Sm/144Nd
ratios of 0.12 – 0.13 and 143Nd/144Nd ratios of 0.511745 and 0.511833, respectively. The
6
εNd(t) values are -4.3 (CB-13-05) and -2.6 (D-13-01), with TDM model ages of 1.6 and
7
1.3 Ga, respectively (Fig. 12a). The
8
12b).
Sr/86Sr values are 0.71065 and 0.71278 (Fig.
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147
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The two granodiorites (D-13-02, CID; BA-13-07, CEO) and one tonalite (D-13-
9
Sm/144Nd ratios between 0.11 and 0.15 and
143
Nd/144Nd
15, CID) sample yielded
11
ratios of 0.511722 to 0.511860 (Table 4). The εNd (t) values range from -4.8 (BA-13-
12
07), over -3.4 (D-13-02), to -2.1 (D-13-15), with TDM model ages of 1.38 Ga, 1.4 Ga and
13
1.55 Ga, respectively (Fig. 12a). These samples gave
14
0.71526 (Table 4) (Fig. 12b).
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87
Sr/86Sr values of 0.71112 -
The three diorites (D-13-40 and CB-13-02, CID; CB-13-04 mafic dyke, CEO)
15
147
Sm/144Nd ratios between 0.15 and 0.18 and
143
Nd/144Nd ratios between
yielded
17
0.511942 and 0.512088 (Table 4). These samples have ɛNd(t) values of +2.3 (CB-13-
18
04), +1.8 (CB-13-02), and -0.5 (D-13-40), with TDM model ages of 1.5 Ga, 1.3, and 1.4
19
Ga, respectively (Fig. 12a). The 87Sr/86Sr values range from 0.70446 to 0.70569 (Table
20
4) (Fig. 12b).
22 23 24 25
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2
5. Discussion
3
5.1. Emplacement age of the Diablillos Intrusive Complex
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A weighted average 206Pb/238U age of 517±3 Ma (Fig. 10a) was obtained for the
6
monzogranite, 515±6 Ma (Fig. 10e) for the diorite, and a concordant U-Pb age of 521±4
7
Ma for the granodiorite (Fig. 10c). The ages of the three samples overlap within error,
8
and are considered a good estimate of the emplacement age of this complex.
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5
The age variations in the Concordia diagrams for the three samples (Fig. 10a, e)
10
could reflect analytical scatter; but the age range is similar in the three analyzed
11
samples, regardless of which analytical session is scrutinized. No evidence for Pb loss
12
was observed either. Likewise, the possibility of analytical scatter in the reference
13
material analyzed with the samples was also disregarded, due to the similar data
14
obtained on the TEMORA reference zircon for each analyzed sample: a
15
of 409 ± 4 Ma for the monzogranite, a 206Pb/238U age of 409 ± 3 Ma for the diorite, and
16
a 206Pb/238U age of 409 ± 5 Ma for the granodiorite.
206
Pb/238U age
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Furthermore, a combination of Pb loss and inheritance of older Pb could provide
18
a possible explanation for the age scatter (Castiñeiras et al., 2010). Excluding the
19
oldest grains with comparatively higher Th/U ratios, the remaining slight variation in
20
the Th/U ratios of the samples suggests that the inherited component involved in the
21
analyzed zircon crystals is insignificant.
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22
Another possibility for the observed age variation is a true difference in age, for
23
example, when a sequence of different geological processes, such as metamorphism or
24
magmatism, occurred over a limited time span (Castiñeiras et al., 2010). Hoskin and
25
Schaltegger (2003) and Willner et al. (2008), among others, concluded that zircon
ACCEPTED MANUSCRIPT crystals of Th/U ratios ≥0.1-0.5 have a magmatic origin, whereas zircon crystals with a
2
metamorphic origin should have Th/U ratios <0.1. Hence, the slight variation observed
3
in the zircon fractionation indices in the analyzed samples, namely Th/U = 0.14 to 2
4
(Table 4), may suggest that zircon crystals from the CID developed during a magmatic
5
event.
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6
On the other hand, a long-lived magma chamber, which could have formed the
7
CID, could also explain most likely the continuous age distribution for these three rock
8
types
9
recrystallization/overgrowth of older cores -not entirely crystallized, at lower levels of
10
the magma chamber- during the intrusion of new pulses of hot magma. Resulting in
11
minor mixing of the new zircon crystals from newly intruded magma with older zircon
12
cores (Dong et al., 2013; Folkes et al., 2011; Castiñeiras et al., 2010).
11
a,
e).
This
variation
could
be
explained
by
the
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(Fig.
The observed internal zonation, irregular dark cores or irregular shapes with
14
irregular dissolution zonation displayed by some zircon grains (Fig. 10b, d, f), support
15
these mixing ideas. Such zonations indicate periods of zircon undersaturation followed
16
by zircon-saturation in a melt (e.g., Vavra 1994; de la Rosa et al., 2002; Zeck and
17
Williams, 2002). In addition, Hoskin and Schaltegger (2003) concluded that such
18
crystals preserved compositional information for the melt prior to, and after, dissolution.
19
Consequently, they recorded a change in composition related to an event such as
20
magma-mixing.
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One way of comparing the individual age distributions of the three analyzed
22
samples is by comparing Figures 10g, h and i, the Probability Density Plots for specific
23
zircon age populations. In Fig. 10j the combined data base for the three rock types is
24
considered.
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Clearly, the zircon population at ~490 Ma (Fig. 10j) is present in the data for all
2
three samples but the meaning of this age is unclear as the number of data for this
3
combined population is limited (10 zircon grains only). A possible interpretation could
4
be that this age group may indicate the last magmatic activity in the complex. The largest population at ~515-520 Ma (39 zircon grains) (Fig. 10j) could
6
represent the peak of magmatic activity related to the CID, whereas the population at
7
~530-540 Ma (32 zircon grains) (Fig. 10j) may represent zircon antecrysts
8
(incorporated from previous crystallization events in the magmatic system - Schmitt,
9
2011). This would mean that they represent early magmatic activity in the CID.
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As a result, after analyzing the age populations in the rock units of the CID, we
11
interpret that this intrusive complex was emplaced between 540 and 490 Ma, possibly
12
over a ca. 50 Ma magmatic event.
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The oldest zircon populations of Late Neoproterozoic ages (~580/600/630 Ma)
14
and Mesoproterozoic ages (1000-1300 Ma) (Fig. 10j) signify inherited xenocrysts from
15
the source rock. These data - mostly from the cores of analyzed zircon grains - along
16
with similar age groups presented by Bahlburg et al. (2016) and Bahlburg and
17
Berndt (2016) from the northern Puna (~550-700 Ma; 900-1200 Ma), illustrate the
18
presence of Neoproterozoic sources in the region, i.e., possibly the metasedimentary
19
Puncoviscana Formation. Finally, the population of Mesoproterozoic age represents the
20
Grenvillian/Sunsas orogeny. This indicates that the unknown basement of the Puna is
21
formed either by the Puncoviscana Formation as an erosional product of these orogenic
22
belts or that the Puna including its Puncoviscana basement is underlain by crust evolved
23
during these orogenies (Bahlburg et al., 2016).
24 25
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5.2. Magma sources for the Diablillos Intrusive Complex
4
5.2.1. Geochemistry and Sr-Nd isotopes
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Based on field observation, petrography, and geochemistry - including Sr-Nd
7
isotope analysis, two groups of rocks can be distinguished in the CID and CEO: the
8
intermediate group (monzogranite, granodiorite and tonalite rocks) with silica contents
9
between 69.99 and 59.65 wt.% and the mafic group (diorites) with silica contents between 54.50 and 50.27 wt.%.
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The monzogranite and granodiorite-tonalite rocks are enriched in LILE, have
12
negative anomalies for Ba, Sr, and Eu compared to the dioritic rocks (Fig. 9a), and are
13
enriched in the LREE compared to the mafic rocks (Fig. 9b). The intermediate-felsic
14
samples show slightly negative εNd(t) values, between -4.8 (Granodiorite from CEO,
15
BA-13-07) and -2.1 (tonalite from CID, D-13-15) and initial
16
0.71526 (granodiorite from CEO, BA-13-07) and 0.71064 (monzogranite from CEO,
17
CB-13-05) (Fig. 12b). They display TDM model ages between 1.6 and 1.3 Ga (Fig. 12a),
18
suggesting that Mesoproterozoic crust was involved in the generation of these rocks.
Sr/86Sr values between
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The diorites show an enrichment in incompatible elements (i.e., Ce, Rb, Th and
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19
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11
20
K) (Fig. 9a) and a slight depletion of the HREE (Fig. 9b), compared to the intermediate
21
rocks, these agrees with a crust-contaminated mafic magma source for the mafic rocks
22
(Kleine et al., 2004). However, there is the exception of sample CB-13-04 (CEO,
23
diorite mafic dyke), with a more primitive origin (Fig. 9b). It was recognized by
24
Suzaño et al. (2015; 2017) that this dyke is related to a later stage of the evolution of
25
the CID. These basic dykes have the same REE pattern as the gabbros from the Ojo de
ACCEPTED MANUSCRIPT 1
Colorados Complex in the Western Magmatic Belt, Puna (Zimmermann et al., 2014),
2
which suggests the involvement of analogous mantle-derived sources. Contamination of mafic magma with crust is also evident in the Sr-Nd isotope
3
87
Sr/86Sr values of 0.70446 (diorite from CID,
data where the diorites show low initial
5
CB-13-02) to 0.70523 (diorite from CEO, CB-13-04) (Fig. 12b). They also show
6
positive to slightly negative ɛNd(t) values between +2.3 (diorite from CEO, CB-13-04)
7
and -0.5 (diorite from CID, D-13-04), and gave Mesoproterozoic TDM model ages (Fig.
8
12a).
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9
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4
Likewise, the negative correlation trend towards the felsic rocks (Fig. 12b), 87
Sr/86Sr and low
143
Nd/144Nd ratios of the
evidenced by the moderate to high
11
intermediate rocks, and the moderate to low 87Sr/86Sr and low
12
mafic rocks from the CID, support contamination of primary magma with crustal
13
material and suggest a magma-mixing origin of the CID. These conclusions are in
14
agreement with field relationships and petrographic findings (Fig. 3 and 5), such as
15
subrounded phenocrysts that migrated mechanically from the monzogranites to the
16
granodiorite enclaves, subrounded K-feldspar phenocrysts, mantled and zoned
17
plagioclase, inter alia.
19 20 21
Nd/144Nd ratios of the
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5.2.2. Crustal evolution evidences from Hf isotopes
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Zircon with a Hf isotopic composition similar to depleted mantle could not be
22
identified; closest to this comes a zircon grain from the diorite sample with an ɛHf(536)
23
value of +9.1, with a TDM age of 0.9 Ga (Fig. 11). The two Mesoproterozoic inherited
24
cores of zircon in the monzogranite sample have juvenile magma component signatures,
25
with ɛHf(t) values of +7.0 and +5.9 and TDM ages of 1.3 and 1.4 Ga, respectively (Fig.
ACCEPTED MANUSCRIPT 11). A zircon from the granodiorite gave a ɛHf(654) value of +4.2 and TDM age of 1.2
2
Ga. On the other hand, zircon grains indicating magmas with an extensively reworked
3
crustal source are exemplified in the diorite sample with ɛHf(519) of -9.5, with a TDM
4
age of 1.9 Ga, and ɛHf(521) of -4.1, with a TDM age of 1.6 Ga. And in the granodiorite
5
sample with a grain characterized by ɛHf(538) of -3.8, giving a TDM age of 1.5 Ga (Fig.
6
11).
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The presence of slightly positive and negative ɛHf(t) values between +3 and -3
8
(Fig. 11) for the three analyzed samples emplaced between ca. 490-540 Ma (Fig. 10j)
9
indicates that the rocks from the Diablillos Intrusive Complex derived from an
10
interaction of a dominant Mesoproterozoic crustal source (crystalline and/or
11
metasedimentary, most frequent TDM age at ~1.4 Ga), and juvenile mantle-derived
12
magma, that were reworked during lower Paleozoic times. In addition, the Sm-Nd
13
whole rock isotope data (Fig. 12a) agree with this interpretation, whereby similar ɛNd(t)
14
(+2.5 to -4) and Nd TDM model ages between 1.3 to 1.6 Ga, with a most frequently
15
observed TDM age of ~1.4 Ga, were obtained.
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7
The zircon grain from the diorite with an ɛHf(536) positive value of 9.1 indicates
17
that new juvenile magma was formed at TDM = 0.9 Ga (Fig. 11). In the Northeastern
18
Magmatic Belt of the Puna, Hauser et al. (2011) analyzed a Mesoproterozoic zircon
19
grain in a metadacite, with an ɛHf(1070) of +13.0 value and a TDM of 1.0 Ga (Fig. 11).
20
Consequently, a magmatic event took place in the magmatic arc at the SW Gondwana
21
margin between ca. 0.9 and 1.0 Ga (around the Meso/Neoproterozoic boundary). This
22
lends further support to the models of depleted mantle evolution by Condie et al.
23
(2005; 2009) and Andersen et al. (2007), for which only little evidence for juvenile
24
magmatic rocks has actually been recognized worldwide for the ca. 1-1.20 Ga period.
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ACCEPTED MANUSCRIPT The samples from the CID present similar ɛHf(t) and TDM ages to those
2
previously obtained for the ca. 480 Ma subvolcanic rocks from the Northeastern
3
Magmatic Belt of the Puna (Hauser et al., 2011), as well as for the granitoids of ca. 470
4
Ma age analysed by Bahlburg et al. (2016) in the same region (Fig. 11). Additionally,
5
in the Centenario salar and nearby areas (Fig. 1) of the Southern Puna, Viramonte et al.
6
(2007) obtained whole rock ɛNd(t) data ranging from +2.5 to -7 and Nd (TDM) ages
7
between 1.5 to 1.7 Ga for granitoids (460-475 Ma) and metavolcanics (485 Ma) (Fig.
8
12a), similar to the ɛNd(t) (+2.5 to -4) and model Nd (TDM) ages of this study (1.3 to 1.6
9
Ga) (Fig. 12a). Furthermore, Kleine et al., (2004), presented similar Nd results in the
10
Complejo Igneo Pocitos of the Western Eruptive Belt. Similarly to the CID rocks, the
11
samples of Viramonte et al. (2007) and Kleine et al., (2004) have low
12
isotope ratios, indicative of recycling of older crustal material without development of
13
great volumes of new crust (Damm et al., 1990; Becchio et al., 1999; Bock et al.,
14
2000; Lucassen et al., 2000). These similarities between the CID and nearby areas
15
support the interaction of a dominant Mesoproterozoic crustal protolith and juvenile
16
mantle-derived magmas for the evolution of the CID and magmatic complexes in the
17
Centenario salar area in the Southeastern and Northern Magmatic Belt, as well as in the
18
Western Magmatic Belt.
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Similarly, the apparent magma source similarities shown for the Southern and
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1
20
Northern Puna (Fig. 11) indicate that the evolution of the Eastern Magmatic Belt was
21
analogous in both sectors, as already discussed by Viramonte et al. (2007) and Coira
22
et al. (2009).
23
The whole rock ɛNd(t) (Fig. 12a) and zircon ɛHf(t) values (Fig. 11) presented in
24
this study and by Hauser et al. (2011), Bahlburg et al. (2016) and Dahlquist et al.
25
(2016) show that the interaction of a Mesoproterozoic crustal (crystalline and/or a
ACCEPTED MANUSCRIPT metasedimentary) source of ~1.4 Ga and juvenile mantle-derived magmas, that were
2
reworked during lower Paleozoic times in the Eastern Magmatic Belt, extended not only
3
to the Cordillera Oriental geological province, east of the magmatic belt, but also to the
4
Northern Sierras Pampeanas (Fig. 13) south of the Eastern Magmatic Belt. Thus, this
5
Mesoproterozoic event represents a regional episode that could be correlated with the
6
Grenvillian/Sunsas Orogenic Belt (Condie et al., 2005, 2009; Bahlburg et al., 2016;
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Bahlburg and Berndt, 2016; Augustsson et al., 2016).
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5.3. Eastern Magmatic Belt, Lower Paleozoic Arc, SW Gondwana margin: A longlived magmatic event?
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Gehrels et al. (2009) determined that the Coast Mountain batholith at the
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northwestern coast of North America was active nearly continuously from ca. 170 Ma to
14
ca. 50 Ma, but that there were dramatic variations in magmatic flux (low and high level
15
of magmatic activity) during this time. More recently, Cope, (2017) in northeast China,
16
determined cyclic magmatism with magmatic flare-ups and lulls (De Silva et al., 2015;
17
Ducea et al., 2015), along the northern and eastern margins of the North China block
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during late Paleozoic time (~60 Ma long) and Mesozoic time (~50 Ma long),
19
respectively.
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As stated in the previous section, the magma source similarities of both the
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North and Southeastern Magmatic Belt of the Puna, and the zircon U-Pb age ranges
22
presented in this contribution and in other recent studies, allow us to suggest that, the
23
magmatism in the Eastern Magmatic Belt in the Puna region represent a long-lived
24
magmatic event. We reach to this statement analyzing the zircon U-Pb age ranges
25
presented in this contribution (CID: 540-490 Ma), and in other recent studies, such as,
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those of the granitoids of 491 to 481 (Insel et al., 2012); granitoids and metavolcanics
2
of 485 to 460 Ma in the Southeastern Magmatic Belt in Puna (Viramonte et al., 2007).
3
Furthermore, the ca. 480 Ma (metavolcanics, Hauser et al., 2011), and 469 to 445 Ma
4
granitoids (Bahlburg et al., 2016) in the Northeastern Magmatic Belt in Puna, Hence, the long-lived magmatic event, started at ca. 540 Ma (Diablillos Intrusive
6
Complex) and lasted until 460 Ma in the Southern Puna (Salar Centenario and nearby
7
areas), with proven younger activity in the Northern Puna, until ca. 440 Ma (Fig. 13,
8
inset a).
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The mentioned episodes of low and high level magmatic activity, underline the
10
long-lived magmatic event in the Eastern Magmatic Belt, where the early stages of
11
~540-530 Ma magmatism in the CID and the peak magmatic activity (~520-515 Ma) in
12
the CID would represent a lull episode, whereas the most voluminous magmatic activity
13
registered in the Eastern Magmatic Belt between 485-470 Ma (Viramonte et al., 2007;
14
Hauser et al., 2011; Bahlburg et al., 2016) would represent a flare-up episode (Fig.
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13, inset a).
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Zimmermann et al. (2014) presented for the Western Magmatic Belt, Southern
17
Puna, U-Pb zircon ages for gabbroic bodies of 543 Ma (Complejo Ojo de Colorados),
18
highlighting the early stages of the magmatic event that took place in both, the Eastern
19
and Western Magmatic belts in Southern Puna at ~540 Ma. In addition, Dahlquist et al.
20
(2016) defined a ~533 Ma U-Pb age for the Guasayán Pluton in the Northern Sierras
21
Pampeanas (near the boundary of the Puna and Cordillera Oriental, Fig. 13, inset b).
22
Aparicio González et al. (2011) presented a ~533 Ma U-Pb age for granitic dykes in
23
the Sierra de Mojotoro, Cordillera Oriental (Fig. 13, inset b), and there is the ~530 to
24
~515 Ma U-Pb age range by Hauser et al. (2011) for the Tastil batholith, Cordillera
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Oriental (Fig. 13, inset b). Furthermore, these granitoids present similar REE patterns to
2
those of the Eastern Magmatic Belt (Puna), as illustrated in Fig. 9b. These evidences for a number of contiguous regions suggests that the likely
4
long-lived magmatic event in the Eastern Magmatic Belt of the Puna, proposed in this
5
contribution, was a regional event. It could have evolved in a way analogous to the
6
model of Lucassen et al. (2000), who proposed for the Central Andean basement a
7
prolonged evolution, where the Pampean and Famatinian cycles are not distinct events
8
but belong to a single, non-differentiable event from ~540 to 440 Ma. Whereby, the
9
granitoids of the Eastern Magmatic Belt formed in a geodynamic evolution, dominated
10
by intra-crustal recycling at an active continental margin, with minor contributions from
11
juvenile materials in the back-arc setting as stated by Becchio et al., (1999), Bock et al.
12
(2000), Lucassen et al. (2000), Coira and Koukharsky (2002), Kleine et al. (2004),
13
Coira et al. (2009), Bahlburg et al., (2016).
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6. Conclusions
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The new data and observations reported here provide a contribution to the
11
understanding of the magmatic activity in the Diablillos Intrusive Complex of the
12
Eastern Magmatic Belt in the Southern Puna. U-Pb data on zircon indicate that the
13
emplacement of the complex took place between ~540 Ma and 490 Ma (representing a
14
ca. 50 Ma magmatic event). The main magmatic activity related to this complex was
15
seemingly at ~515-520 Ma.
The combined major and trace, including rare earth, element data, as well as 87
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Sr/86Sr and ɛNd(t) and ɛHf(t) on zircon results indicate that all analyzed
whole rock
18
samples represent crustally contaminated magmas.
19
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The ɛHf(t) values of + 3 and -3 and the ɛNd(t) values of +2.5 and -4 indicate that
20
the rocks from the Diablillos Intrusive Complex derived from an interaction of a
21
dominant Mesoproterozoic crustal (crystalline and/or a metasedimentary) source and
22
juvenile mantle-derived magmas - with a TDM model age range of ~1.2-1.5 Ga- that
23
were subsequently reworked during lower Paleozoic times.
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2
study and from other authors (e.g. Viramonte et al., 2007; Hauser et al., 2011, and
3
Bahlburg et al., 2016; among others) demonstrate that the interaction of a crustal
4
source of ~1.4 Ga and juvenile mantle-derived magmas in the Eastern Magmatic Belt is
5
related to a similar evolution in the Western Magmatic Belt, Cordillera Oriental and
6
Northern Sierras Pampeanas, emphasizing that the reworking process during lower
7
Paleozoic times was a regional episode.
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The magma source similarities and the zircon U-Pb age range presented in this
9
and previous studies (Viramonte et al., 2007; Hauser et al., 2011; Bahlburg et al.,
10
2016) of both the Northern and Southern Puna allows to propose that the Eastern
11
Magmatic Belt in the Puna region was possibly created in a long-lived magmatic event,
12
initiated at ca. 540 Ma with the intrusion of the Diablillos Intrusive Complex, and
13
lasting to 460 Ma in the Southestern Magmatic Belt of the Puna, with even younger
14
activity in the Northern Magmatic Belt of the Puna, until ca. 440 Ma. Whereby, the
15
granitoids of the Eastern Magmatic Belt, formed in a geodynamic evolution dominated
16
by intra-crustal recycling at an active continental margin, with minor contributions from
17
juvenile materials in the back-arc setting.
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Further detailed study, involving additional geochronology not only on the
19
plutonic rocks of the CID but also on the basement rocks from the study area and
20
environs may provide a test for the existence of such a long-lived magmatic event in the
21
Eastern Magmatic Belt of the Puna.
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7. Acknowledgements
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Financial support from the CNPq program and from CONICET is
10
acknowledged. We thank Alejandro Nieva (GEONORTE-Universidad Nacional de
11
Salta) for help with sample preparation for chemical analysis. And we would like to
12
thank Luciana Pereira and Bárbara Lima (Laboratório de geocronologia - Universidade
13
de Brasília) for technical assistance with the CL imaging and isotope analyses. We
14
acknowledge the Rodinia Lithium-Silver Standard Company of Salta for field and
15
lodging assistance. WUR’s research is supported by the German Science Foundation
16
(DFG) and Museum für Naturkunde Berlin. Finally yet importantly, we want to extend
17
our gratitude to reviewers Dr. Beatriz Coira, Dr. Juan Otamendi and Dr. Victor Ramos
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for taking the time and effort necessary while revising the manuscript, their comments
19
and suggestions helped improving the final version of this work.
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Analyses
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doi:10.1093/petrology/egm032.
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18
19
20
21
22
suture
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17
Ordovician
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revisited. International
Journal
of
Earth
ACCEPTED MANUSCRIPT 1
2
3
9. Appendix
5
Table 1. Whole rock geochemical analyses for major and trace element data for the CID
6
and CEO, Eastern Magmatic Belt, Southern Puna. Major element data in wt.%, trace
7
element data in ppm.
Petrology/ Element SiO2
CB-13-05
D-13-02
BA-13-07
D-13-15
CB-13-04
D-13-40
CB-13-02
(CID)
(CEO)
(CID)
(CEO)
(CID)
(CEO)
(CID)
(CID)
Tonalite
Diorite mafic dyke
Diorite
Diorite
59.65
50.27
54.5
50.93
Monzogranite Monzogranite Granodiorite Granodiorite 60.93
69.99
61.69
1.06
0.55
1.08
16.53
13.95
15.93
Fe2O3
7.26
4
7.39
MgO
2.73
1.6
2.98
MnO
0.1
0.07
CaO
2.36
1.56
Na2O
3.07
2.56
K 2O
4.6
3.84
P 2O 5
0.18
0.15
LOI
0.9
Sum
99.73
Cr
100
62.05 1.03
1.69
0.83
1.31
1.82
16.69
15.01
16.13
15.8
14.54
6.74
9.12
9.1
8.87
11.13
2.74
3.26
8.08
6.09
6.31
TE D
TiO2 Al2O3
0.12
0.08
0.14
0.15
0.14
0.17
3.16
3.5
4.84
11.26
7.56
9.05
3.04
3.37
2.65
2.36
2.35
2.59
3.15
2.68
2.15
0.28
1.54
1.21
0.19
0.24
0.08
0.18
0.24
1
0.8
1
1.1
1.4
1.7
99.78
99.74
99.74
99.76
99.76
99.75
99.74
100
100
100
100
600
300
300
EP
0.07
1.5
125
67
143
144
235
245
205
303
61.5
67.9
58
41
52.6
63.3
56.1
53
23
<20.0
25
21
26
94
67
35
23.1
16.2
20.2
20.6
19.8
13.8
16.7
17.7
173.8
168.5
145.7
126.8
86.1
9
74.5
47.7
110.9
89.5
116.3
201.5
126.1
104.4
141.1
155.6
Y
40.9
31.8
39.5
31
41.5
21.7
31.8
34.1
Zr
286.7
197.1
281.3
231.9
254.3
57.4
170.3
190.4
Nb
15.4
12.4
16.3
13.4
11.2
2.1
7
6.1
Cs
6.7
13.3
5.5
10.2
5.7
1.1
6.3
4.4
Ba
474
342
324
355
190
39
140
92
Hf
7.8
5.4
7
6
6.6
1.6
4.3
4.4
Ta
1
1.2
1.1
1.1
0.9
0.1
0.5
0.3
Co Ni Ga Rb Sr
AC C
V
SC
D-13-01
M AN U
Sample
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4
ACCEPTED MANUSCRIPT Th
21.7
13.5
18.6
20.7
7.2
0.9
5.9
2.6
U
1.7
2.3
1.4
1.6
1.6
0.5
1.5
0.6
Be
3
2
2
<1.0
2
<1.0
<1.0
3
As
6.7
4.2
7.8
8
6.9
2.4
19.5
26.5
Au
<0.5
1.2
0.9
1.3
<0.5
0.9
0.9
0.8
0.2
1
<0.1
0.1
<0.1
<0.1
0.1
<0.1
14.6
3
31.1
16.4
37.2
95.2
83.6
73.5
Pb
3.8
1.2
3.1
3.3
2.2
0.8
3
1.3
Sb
<0.1
<0.1
<0.1
<0.1
<0.1
Sc
18
10
22
15
28
Sn
4
4
3.0.
3
3
Ti
0.9
0.7
0.8
0.6
0.4
Zn
107
62
104
98
74 26.2
32.7
44.2
52.9
68.9
94.4
102.6
Pr
13.2
7.8
10.8
12.1
Nd
48.8
30.7
44.9
47.1
Sm
10.3
6.1
8.6
Eu
1.7
0.9
1.8
Gd
9.4
5.9
8.4
Tb
1.4
0.9
1.4
Dy
8.2
5.8
7.8
Ho
1.6
1.2
1.6
Er
4.2
3.3
4.6
Tm
0.6
0.5
0.7
Yb
3.8
3.2
Lu
0.6
0.5
LaN/YbN
9.4
6.8
Eu/Eu*
0.5
0.5
3 4 5
<0.1
0.1
37
30
42
<1.0
2
2
<0.1
0.3
0.1
6
43
21
5.9
18.5
12.3
57.8
12
39.6
30.9
7.3
1.7
5.08
4.2
32.4
8.1
23.6
21.1
7.3
2.2
5.1
5.2
1.7
1.8
0.8
1.5
1.7
7.5
8.1
3.3
6
6.7
1.1
1.4
0.6
1
1.1
5.7
7.6
4.2
5.9
6.8
1.2
1.7
0.9
1.3
1.5
3.1
4.5
2.5
3.6
4
0.5
0.7
0.4
0.5
0.6
TE D
8.8
4.4
3
4.5
2.5
3.4
3.6
0.7
0.5
0.7
0.4
0.5
0.6
6.7
11.9
3.9
1.6
3.7
2.3
0.6
0.7
0.7
0.9
0.8
0.9
EP
2
AC C
1
<0.1
SC
53.1 111.6
M AN U
La Ce
RI PT
Bi Cu
ACCEPTED MANUSCRIPT
Table 2. Results of in situ U-Pb isotope analysis of zircon from rocks of the Diablillos Intrusive Complex, Eastern Magmatic Belt, Southern
Monzogranite
Isotopic Ratios
Apparent Ages
D-13-01
0.206 0.058
1.4
0.625
1.9
0.079
1.3
517.5
0.057
1.5
0.626
2.3
0.079
1.7
502.6
0.136 0.058
1.2
0.633
1.8
0.08
1.3
0.057
1.7
0.632
2.1
0.08
1.3
0.268 0.059
1.4
0.654
1.8
0.081
1.2
Z36
0.058
2.4
0.65
3
0.081
1.9
Z66
0.059
1.5
0.658
2.2
0.081
1.6
Z60C
0.058
2.4
0.649
3.3
0.081
2.2
Z19
0.057
1.1
0.644
1.5
0.082
1
Z68
0.059
1.3
0.661
1.8
0.082
Z60R
0.058
1.5
0.659
2
Z27R
0.059
2.1
0.667
2.5
Z40
0.356 0.058
1.4
0.658
1.9
Z64
0.058
1.9
0.661
2.6
Z65
0.058
1.9
0.662
2.3
Z6
0.057
1.4
0.652
1.8
Z15
0.058
1.2
0.657
Z3
0.058
8.5
Z30B
0.058
Z18
0.057
Z51
Z14
Pb/235U 1σ(%)
206
Pb/238U 1σ(%)
207
Pb/206Pb 1σ(Ma)
207
Pb/235U 1σ(Ma)
206
Pb/238U 1σ(Ma) Rho Conc(%)
30.6
493.2
7.5
488
6.3
0.68 98.9
32.9
493.6
8.9
491.7
8.1
0.75 99.6
25.6
497.8
6.9
494.2
6.2
0.74 99.3
505.7
36.9
497.3
8.2
495.4
6
0.69 99.6
559.7
30.7
511
7.4
500.1
5.7
0.63 97.9
536
51.9
508.3
12
502.2
9.2
0.62 98.8
556.2
33.7
513.3
9
503.7
7.9
0.72 98.1
520.8
52.2
507.8
13
504.9
11
0.68 99.4
500.2
24.8
505
6
506.1
4.9
0.65 100
1.1
555.4
29.1
515.2
7.1
506.2
5.5
0.63 98.3
0.082
1.4
547.3
32.2
514.3
8.1
506.9
6.6
0.67 98.6
0.082
1.4
569.9
44.8
518.6
10
507.1
6.8
0.67 97.8
0.082
1.3
540.3
31.4
513.6
7.7
507.6
6.1
0.65 98.8
0.082
1.8
547.6
41
515.4
11
508.1
8.9
0.69 98.6
0.082
1.3
546.1
41
516
9.2
509.2
6.2
0.66 98.7
0.083
1.1
503.2
30.8
509.9
7.2
511.4
5.5
0.65 100
1.6
0.083
1.1
517.8
25.4
512.8
6.5
511.6
5.5
0.68 99.8
0.659
8.6
0.083
1.1
517.9
187
514.1
35
513.2
5.3
0.63 99.8
1.2
0.662
1.6
0.083
1
527.2
25.4
516
6.3
513.4
5.1
0.65 99.5
1.3
0.65
1.7
0.083
1
486.1
29.5
508.8
6.7
513.8
4.9
0.58 101
0.058
1.5
0.659
1.8
0.083
1
513.4
32.5
514
7.3
514.1
5.1
0.55 100
Z37
0.059
1.7
0.672
2.3
0.083
1.6
553.7
36.9
522.2
9.4
515.1
7.8
0.68 98.6
Z45
0.059
2.3
0.678
2.9
0.083
1.7
564.9
51
525.5
12
516.5
8.4
0.69 98.3
Z24B
0.059
1.5
0.678
2
0.084
1.3
557.4
31.7
525.3
8
517.9
6.5
0.66 98.6
Z67R
TE D
514.3
Z17
AC C
Z9
Pb/206Pb
EP
Z5C
207
SC
207
Th/U
M AN U
1σ(%)
Grain N°
RI PT
Puna. R: rim analysis. C: core analysis.
ACCEPTED MANUSCRIPT
0.058
1.2
0.675
2.7
0.084
2.5
544.9
25.4
523.5
11
518.6
12
0.73 99.1
Z50
0.827 0.058
1.2
0.67
1.5
0.084
1
527.7
26.2
520.9
6.2
519.4
4.8
0.61 99.7
Z16
0.058
1.7
0.669
2.2
0.084
1.4
517.4
36.7
519.8
8.8
520.4
6.8
0.62 100
Z57
0.058
1.2
0.678
1.5
0.084
0.9
547.8
25.8
525.8
6.1
520.7
4.5
0.58 99
Z69
0.058
1.2
0.677
1.5
0.084
0.9
535.7
26.4
524.8
6.1
522.3
4.5
0.57 99.5
2.091 0.058
1.2
0.674
1.7
0.085
1.2
522.4
26
523.4
7
523.6
6.2
0.71 100
Z28R
0.059
1.1
0.689
1.4
0.085
0.8
566.9
24.3
531.9
5.8
523.8
4.2
0.57 98.5
Z22
0.058
1.1
0.681
1.4
0.085
0.8
541
24.1
527.6
5.7
524.5
4.1
0.57 99.4
Z39C
0.059
2.5
0.687
3.4
0.085
2.4
558.3
52.9
530.9
14
524.5
12
0.69 98.8
Z48
0.058
1.3
0.684
1.8
0.085
1.3
544.6
27.5
529.2
7.4
525.6
6.4
0.7
Z54
0.058
1.8
0.685
2.8
0.085
2.1
542.7
Z49
0.058
1.2
0.686
1.6
0.085
1
Z55
0.059
2
0.693
3
0.085
2.3
Z23
0.174 0.058
1.3
0.685
1.7
0.085
1.1
Z52R
0.058
1.3
0.689
1.7
0.085
1.1
Z47
0.058
1.4
0.685
1.8
0.086
1.1
Z44
0.059
2.3
0.696
3.3
0.086
2.3
Z58
0.058
1.2
0.689
1.5
0.086
0.9
Z21R
0.059
2.7
0.696
2.9
0.086
Z21C
0.058
1.9
0.683
2.4
0.086
Z61
0.059
2.4
0.696
2.7
0.086
Z24
0.261 0.059
1.3
0.699
1.8
SC
99.3
529.5
11
526.5
10
0.65 99.4
26.1
530.2
6.6
527.5
5.3
0.64 99.5
563.9
43.9
534.4
13
527.6
12
0.75 98.7
538.1
28.5
530
7
528.1
5.5
0.62 99.6
547.6
27.7
532.3
6.9
528.7
5.5
0.63 99.3
534.9
30.3
530.1
7.4
529
5.8
0.62 99.8
564.5
51
536.5
14
529.9
12
0.7
541.7
25.8
532.3
6.2
530.1
4.7
0.58 99.6
0.9
561.4
59.2
536.2
12
530.3
4.7
0.6
1.6
516.2
41.1
528.5
10
531.4
8
0.64 101
1.3
549.7
51.4
536.2
11
533
6.7
0.63 99.4
0.087
1.3
551
28.3
538.4
7.6
535.5
6.6
0.66 99.4
TE D
M AN U
40.3
542
EP
Z7
RI PT
Z10
98.8 98.9
0.06
1.3
0.834
1.9
0.1
1.4
612.4
27.6
615.6
8.7
616.5
8.1
0.72 100
Z12C
0.064
1.3
0.94
2.1
0.107
1.7
734.7
28
672.7
10
654.3
10
0.78 97.3
Z67C
0.065
1.3
0.994
1.9
0.11
1.4
786.5
26.9
700.7
9.4
674.3
8.7
0.72 96.2
Z38
0.064
1.6
1.012
3.9
0.115
3.6
733.3
34
710
20
702.6
24
0.66 99
Z1R
0.07
1.1
1.417
1.6
0.147
1.1
926.1
22.8
896
9.4
883.9
9.2
0.69 98.6
Z1C
0.074
1.1
1.739
1.4
0.171
0.8
1035
22
1023
8.8
1018
7.7
0.57 98.4
Z5R
0.074
1.1
1.739
1.4
0.171
0.8
1035
22
1023
8.8
1018
7.7
0.57 98.4
Z27C
0.075
1.6
1.793
2.1
0.174
1.3
1056
32.8
1043
14
1037
13
0.62 98.2
AC C
Z31
ACCEPTED MANUSCRIPT
Isotopic Ratios 207
0.354 0.057
7.7
0.606
8.2
0.077
3
485.3
170
0.529 0.057
1.6
0.627
2.3
0.079
1.7
506.5
34.4 20.2
Th/U
Z19 Z21
Pb/235U 2σ(%)
206
Pb/238U 2σ(%)
207
Pb/206Pb 2σ(Ma)
0.058
0.9
0.659
2.2
0.083
2
524.6
Z39
0.179 0.058
1.2
0.668
2
0.084
1.7
528.3
Z66
0.058
2.5
0.674
4.7
0.084
4.1
538
Z11
0.058
1.1
0.675
4.1
0.084
3.9
530.9
207
Pb/235U 2σ(Ma)
206
Pb/238U 2σ(Ma) Rho Conc(%)
481.2
32
480.4
14
0.36 99
493.9
9.1
491.2
8.1
0.74 97
514.1
8.9
511.8
9.9
0.91 97.6
26.1
519.4
8.3
517.4
8.3
0.81 97.9
53.7
523.2
19
519.8
20
0.86 96.6
24.8
523.6
17
521.9
20
0.96 98.3
541.6
21.4
527.9
9.5
524.7
11
0.91 96.9
524.1
27.9
527.5
16
528.3
19
0.95 101
518.1
24.5
565.6
19
577.5
24
0.97 111
608.6
28.5
586.8
14
581.2
15
0.9
M AN U
Z14C
Pb/206Pb
SC
2σ(%)
Grain N°
207
Apparent Ages
RI PT
Granodiorite D-13-02
1.198 0.058
1
0.682
2.3
0.085
2.1
Z4
0.058
1.3
0.681
4
0.085
3.8
Z24
0.058
1.1
0.745
4.5
0.094
4.3
Z10
0.06
1.3
0.782
3
0.094
2.7
0.059
2.1
0.773
5.2
0.096
4.7
551.4
46.3
581.5
23
589.3
26
0.91 107
1.8
0.796
5.2
0.096
4.8
609.4
39.7
594.7
23
590.8
27
0.93 97
Z2R Z13
0.137 0.06
TE D
Z9
95.5
0.058
1.4
0.778
4.9
0.097
4.7
532.9
30
584.3
22
597.7
27
0.96 112
Z7
0.06
1.9
0.884
7.9
0.107
7.7
605.7
40.8
643.2
38
653.9
48
0.97 108
Z8C
0.087
4.8
0.941
5.2
0.078
2.1
1364
89.8
673.5
25
486.1
9.7
0.4
AC C
EP
Z48R
35.6
ACCEPTED MANUSCRIPT
Z22
0.058
Apparent Ages 207
Pb/235U 1σ(%)
206
Pb/238U 1σ(%)
207
Pb/206Pb 1σ(Ma)
207
RI PT
Diorite D-13Isotopic Ratios 40 Grain Th/U 207Pb/206Pb 1σ(%) N°
Pb/235U 1σ(Ma)
206
Pb/238U 1σ(Ma) Rho Conc(%)
2
0.632
2.4
0.079
1.4
540.9
43.3
497.2
9.6
487.8
6.6
0.57 98.1
0.778 0.058
1.5
0.63
1.7
0.079
0.8
523.8
32.8
496.1
6.7
490.1
3.9
0.46 98.8
Z35
0.079 0.058
1.6
0.635
2
0.079
1.3
528.5
34.8
499.3
8
493
6
0.61 98.7
Z48
0.058
2.5
0.64
2.6
0.08
0.9
536.4
54
502.4
10
494.9
4.5
0.49 98.5
32
505.8
1.5
0.646
1.7
0.08
0.9
541.6
6.9
498
4.5
0.52 98.4
1.433 0.058
1.6
0.647
2.2
0.081
1.5
535.8
35.4
506.6
8.8
500.1
7.2
0.67 98.7
Z40
0.058
1.4
0.648
1.6
0.081
0.9
516
29.8
507
6.6
505.1
4.5
0.46 99.6
Z56
0.058
1.4
0.65
1.5
0.082
0.7
518.8
29.8
508.7
6.2
506.5
3.6
0.45 99.6
Z38
1.104 0.058
1.5
0.663
1.8
0.083
1.1
540.6
31.9
516.6
7.4
511.2
5.4
0.58 99
Z12
1.648 0.058
1.4
0.669
1.8
0.083
1.1
547.8
31.5
520.1
7.3
513.9
5.2
0.57 98.8
Z44
0.058
1.6
0.662
1.8
0.083
0.9
523.3
34.1
516.1
7.2
514.5
4.3
0.5
Z46
0.058
1.6
0.668
1.9
0.083
Z34
2.074 0.059
1.8
0.673
2
0.083
0.058
7.8
0.668
7.8
0.083
Z49
1.627 0.059
1.7
0.673
2
0.083
Z26
0.059
1.4
0.679
1.7
0.084
Z3
0.058
1.4
0.669
1.7
Z1
0.058
1.4
0.674
1.7
Z45
0.059
2.2
0.683
Z11
0.058
1.6
0.673
Z30
0.517 0.059
1.5
0.69
Z32
0.058
1.8
0.68
Z52
0.059
1.6
Z9C
0.964 0.059
Z15 Z29 Z57
TE D
99.7
1.1
539.9
34.3
519.7
7.8
515.1
5.4
0.56 99.1
0.8
555.9
39.5
522.7
8.1
515.2
4
0.47 98.6
0.8
535.6
171
519.3
32
515.6
4
0.48 99.3
0.9
551
37.6
522.8
8
516.4
4.6
0.45 98.8
0.9
565.7
31.3
526.2
6.9
517.2
4.2
0.48 98.3
0.084
1
530
31.5
520.3
7.1
518.1
4.9
0.55 99.6
0.084
0.8
542.7
31.6
523.3
6.8
518.9
4.2
0.47 99.2
3
0.084
2
567.1
48.6
528.6
12
519.7
9.8
0.66 98.3
1.9
0.084
1.1
531.6
34.3
522.8
7.7
520.8
5.3
0.54 99.6
1.8
0.085
1.1
573.6
31.7
532.8
7.7
523.4
5.7
0.57 98.2
2.6
0.086
1.9
518
39.5
527
11
529
9.7
0.72 100
0.703
2.4
0.086
1.8
581.6
34.9
540.3
10
530.6
9.1
0.74 98.2
1.3
0.695
1.6
0.086
1
550
29.1
536
6.8
532.8
4.9
0.56 99.4
0.058
1.5
0.689
1.9
0.086
1.1
528.3
32.5
532.3
7.8
533.2
5.9
0.6
0.058
1.5
0.694
1.7
0.087
0.8
531.6
32
535.1
6.9
535.9
4.1
0.45 100
0.058
1.3
0.698
1.6
0.087
0.9
538.9
28.6
537.8
6.6
537.6
4.7
0.54 100
EP
Z21C
M AN U
0.058
Z51
AC C
Z36R
SC
Z37C
100
ACCEPTED MANUSCRIPT
0.877 0.058
1.4
0.692
1.6
0.087
0.8
518.5
29.7
534
6.6
537.7
4.3
0.49 101
Z9R
0.057
4.3
0.828
5.8
0.106
3.8
477.3
94.7
612.2
26
649.3
24
0.67 106
Z14
0.077
4.7
1.598
4.9
0.151
1.3
1115
93.9
969.3
30
906.2
11
0.61 93.5
AC C
EP
TE D
M AN U
SC
RI PT
Z20
ACCEPTED MANUSCRIPT
Table 3. Results of in situ Lu-Hf isotope analysis of zircon from rocks of the Diablillos Intrusive Complex, Eastern Magmatic Belt, Southern Puna. 2 sigma values refer to last digits of the isotope ratios. R: rim analysis. C: core analysis. Meas: measured. 176
Lu/177Hf
±2σ (176Hf/177Hf)meas ±2σ (176Hf/177Hf)t ±2σ
ɛHf (T)
±2σ
TDM (Ga) 1.482
513
0.002691
21
0.282417
22
0.28239
18
-2.61
0.1
Z49
527
0.003581
95
0.282455
35
0.282418
16
-1.3
0.06 1.422
Z47
529
0.002323
37
0.282489
31
0.282465
16
0.41
0.06 1.329
Z31
616
0.002455
50
0.282461
23
0.282432
16
1.17
0.02 1.359
Z12c
654
0.002836
77
0.282444
20
0.282408
13
1.2
0.05 1.389
Z38
703
0.001257
18
0.282308
83
0.282291
16
-1.88
0.01 1.599
Z1c
1035
0.000979
18
0.282351
21
0.282331
16
7.02
0.16 1.325
Z27c
1056
0.003054
81
0.282348
53
0.282285
16
5.86
0.17 1.463
Z48
495
0.001864
11
Z3
518
0.002691
21
Z1
519
0.001951
15
Z11
521
0.003496
22
Z15
533
0.005571
68
Z29
536
0.003236
30
Z8B
570
0.001505
Z14
1115
59
0.282534
16
2.08
0.15 1.208
0.282417
22
0.28239
16
-2.51
0.07 1.48
0.282211
42
0.282192
16
-9.51
0.07 1.865
0.282379
60
0.282343
16
-4.09
0.06 1.57
0.282476
95
0.282418
16
-1.16
0.06 1.419
0.282739
12
0.282706
16
9.07
0.12 0.854
23
0.282488
38
0.282471
16
1.53
0.04 1.301
0.001075
88
0.282217
54
0.282194
16
3.98
0.19 1.616
Granodiorite D-13-02 Grain N°
TE D
0.282552
EP
D-13-40 Grain N°
SC
Diorite
RI PT
Z3
M AN U
Monzogranite T D-13-01 (Ma) Grain N°
484
0.001997
15
0.28244
22
0.282421
16
-2.17
0.08 1.433
Z25
490
0.002629
29
0.282561
23
0.282536
16
2.05
0.08 1.206
Z21
491
0.003171
29
0.282477
24
0.282447
16
-1.1
0.08 1.381
AC C
Z5
Z39
517
0.002394
11
0.282433
28
0.282409
16
-1.84
0.07 1.443
Z4
528
0.003207
10
0.282477
33
0.282444
16
-0.37
0.06 1.371
Z1c
538
0.000979
18
0.282351
21
0.282341
16
-3.8
0.06 1.568
Z7
654
0.004253
44
0.282548
49
0.282494
16
4.2
0.01 1.222
1364
0.001668
12
0.282431
26
0.282386
16
16.4
0.3
Z8c
1.125
ACCEPTED MANUSCRIPT
Table 4. Microprobe analyses for zircon crystals from rocks of the Diablillos Intrusive Complex, Eastern Magmatic Belt, Southern Puna. Data in wt.%. Nd: not determined. Grain N°/ Monzogranite Z5c
Z7
Z9
Z15
Z23
Z24
Z40
Z50
Z67r
Z67c
D-13-01 1.556 1.423 1.592 1.596 1.72
SiO2
31.99 31.98 31.01 32.28 31.31 31.25 30.73 31.92 31.11 32.25
ZrO2
65.26 66.41 64.37 65.74 65.05 65.54 64.89 64.94 65.02 64.83
Total
98.81 99.81 96.97 99.61 98.07 98.33 97.28 98.39 97.76 98.56
Y 2O 3
0.374 0.334 0.371 0.243 0.171 0.356 0.439 0.4
UO2
0.034 0.011 0.066 0.029 0.023 0.046 0.045 0.098 0.041 0.019
ThO2
0.007 0.023 0.009 nd
Total
0.485 0.408 0.446 0.335 0.2
Th/U
0.206 2.091 0.136 nd
Granodiorite D-13-02
Z1r
HfO2
1.611 1.565 1.559 1.352 1.411 1.889 1.729 1.922 1.389 1.464
SiO2
32.64 32.32 32.94 33.17 34.12 32.2
ZrO2
64.33 64.86 65.84 62.16 63.55 64.51 64.61 64.58 64.28 63.92
Total
98.58 98.75 100.3 96.68 99.09 98.6
Y 2O 3
0.208 0.193 0.103 1.181 0.103 0.426 0.278 0.136 0.041 0.254
UO2
0.042 0.018 0.001 0.187 0.051 0.048 0.017 0.039 0.039 0.045
ThO2
nd
nd
Total
0.25
0.224 0.194 1.592 0.161 0.605 0.365 0.175 0.163 0.319
Th/U
nd
nd
Diorite D-13-40
Z9C
HfO2
0.958 1.113 1.196 1.204 1.154 1.632 1.734 1.168 1.326 1.026
SiO2
32.89 32.72 31.87 32.83 32.39 34.2
ZrO2
65.8
Total
99.65 98.3
Y 2O 3
0.235 0.817 0.237 nd
UO2
0.056 0.125 0.057 0.029 0.094 0.089 0.009 0.077 0.051 0.06
ThO2
0.054 0.206 0.05
Total
0.379 1.148 0.366 0.044 1.031 0.255 0.045 0.663 0.428 0.498
Th/U
0.964 1.648 0.877 0.517 2.074 0.079 0.778 1.104 1.627 1.433
Z9c
nd
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0.455 0.576 0.579 0.408 0.343
0.174 0.261 0.356 0.827 0.268 nd
Z13
Z19
Z21
Z22r
Z30
Z34
Z35
Z45c
98.83 99.96 99.95 97
1.198 0.137 0.354 0.529 nd
Z20
Z39
32.49 33.46 34.28 31.62
0.224 0.007 0.017 0.009 nd
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Z12
nd
0.275 0.289
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Z8c
1.525 1.632 1.48
0.004 0.012 0.016 0.081 0.011 nd
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Z6
1.541 1.66
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HfO2
Z37c
Z38
0.007 0.01 0.179 0.222 Z49
Z51
32.89 31.65 32.28 34.28
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64.46 65.57 66.72 64.93 61.06 67.19 65.41 65.59 65.79 98.63 100.8 98.48 96.89 101.8 98.22 99.2
101.1
0.671 0.159 0.029 0.373 0.294 0.244
0.015 0.195 0.007 0.007 0.085 0.083 0.086
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Petrology
Age (Ma)
Rb Sr (ppm) (ppm)
87
Rb/86Sr (87Sr/86Sr)meas (87Sr/86Sr)t
Monzogranite
518
173.8 110.9 4.551
0.74638±2
0.71278
Monzogranite
518
168.5 89.5
0.75103±2
0.71065
Granodiorite
518
145.7 116.3 3.635
0.73796±1
0.71112
Granodiorite
518
126.8 201.5 1.824
0.72873±3
Tonalite
518
86.1
126.1 1.839
-
D-13-40 (CID) CB-13-04 (CEO) CB-13-02 (CID)
Diorite Diorite mafic dyke
518
74.5
141.1 1.529
518
9
Diorite
518
47.7
147
Sm/144Nd (143Nd/144Nd)meas (143Nd/144Nd)t eNd(T)
TDM (Ga)
10.9
53.99 0.122
0.512250+/-17
0.511833
-2.6
1.32
3.3
15.08 0.132
0.512196+/-25
0.511745
-4.4
1.56
8.65
40.95 0.128
0.512229+/-10
0.511794
-3.4
1.43
0.71526
11.5
63.04 0.11
0.512096+/-10
0.511722
-4.8
1.38
-
8.58
34.98 0.148
0.512365+/-12
0.51186
-2.1
1.55
0.71698±3
0.70569
5.86
23.88 0.148
0.512447+/-20
0.511942
-0.5
1.38
104.4 0.249
0.70708±2
0.70524
3.01
10.11 0.18
0.512700+/-24
0.512088
2.34
1.42
155.6 0.887
0.71101±2
0.70446
6.34
23.14 0.166
0.512625+/-14
0.512061
1.82
1.32
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5.47
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D-13-01 (CID) CB-13-05 (CEO) D-13-02 (CID) BA-13-07 (CEO) D-13-15 (CID)
Sm Nd (ppm) (ppm)
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Sample
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Table 5. Whole rock Sr-Nd isotope analyses on rocks of the CID and CEO, Eastern Magmatic Belt, Southern Puna.
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Elemental abundances of Rb and Sr are taken from whole rock geochemical analyses. Meas: measured.
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Figures captions:
2
Fig. 1 (2-column fitting image): Regional map, showing the main geological units in the
3
geological provinces of NW Argentina. CN: Salar Centenario. DB: Salar Diablillos. RT:
4
Salar Ratones. CB: Cerro Blanco. TA: Tajamar. AC: Acazoque. Modified after Viramonte
5
et al. (2007) and Zimmermann et al. (2014).
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Fig. 2 (2-column fitting image): Detailed geological map showing the main units in the
8
Diablillos Intrusive Complex. Modified after Nieves (2014). The sampling sites for this
9
study are also shown.
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Fig. 3(2-column fitting image): (a) Panoramic view of the Diablillos salar and the Inca
12
Viejo Range. (b) Contact relationship between a porphyritic monzogranite and parts of a
13
mafic dismembered body. Hammer is 40 cm long. (c) Contact relationships between a
14
porphyritic monzogranite, a granodiorite enclave, and a mafic dismembered body. White
15
dotted lines delimit a granodiorite enclave with K-feldspar phenocrysts. Black lens cover is
16
4 cm in diameter.
17
monzogranite and the granodiorite facies, and a contact with a tonalite enclave. Black
18
arrows show K-feldspar phenocrysts that migrated from the monzogranite to the
19
granodiorite. Note the MME included into the tonalite enclave. (e) Intrusive contact,
20
delimited with a white dotted line, between the monzogranite and slates from the Rio
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Blanco Metamorphic Complex. Hammer is 40 cm long. Dt: Diorite. Gr: Monzogranite. Gd:
22
Granodiorite. Tn: Tonalite.
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(d) Field image showing the transitional contact between the
68
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1
Fig. 4(Single-column fitting image): Q-A-P diagram showing the modal classification of
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the analyzed samples from the CID, in comparison with the samples from Suzaño et al.
4
(2015). The modal classification was made by counting approximately 800 points in a thin
5
section.
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2
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Fig. 5 (2-column fitting image): Microphotographs of the rock types of the Diablillos
8
Intrusive Complex. (a) Monzogranite (sample D-13-01) with holocrystalline porphyritic
9
texture. It is composed mainly of Kfs and Pl phenocrysts, and the matrix is composed of
10
Kfs, Qz, and Pl. (b) Granodiorite (sample D-13-02). Holocrystalline porphyritic texture
11
with Pl, Kfs and Qz phenocrysts. The matrix is formed by fine- to medium-grained Qz, Pl,
12
Kfs. Bt is abundant. The microcline phenocrysts have Bt inclusions (Inc) along the rims. (c)
13
Tonalite (sample D-13-15) showing porphyritic to equigranular texture. Some Pl
14
phenocrysts have a rim of Bt as well as Bt inclusions (Inc). Poikilitic Qz with inclusions of
15
Bt (Inc) is shown. (d) Porphyritic diorite with Pl phenocrysts (sample D-13-40). The fine-
16
grained matrix is formed by Pl, Amp, Bt and Qz. All images taken in cross-polarized light.
17
Mineral abbreviations after Whitney and Evans (2010). Kfs: K-feldspar. Pl: Plagioclase.
18
Qz: Quartz. Bt: Biotite. Ms: Muscovite. Amp: Amphibole.
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Fig. 6(2-column fitting image): (a) Granitoid TAS diagram (Middlemost, 1994) for the
21
various units of the CID. Rocks are classified as granites, granodiorites, tonalites, and
22
quartz diorites. (b) In the A/CNK vs A/NK plot (Shand, 1943) diorites fall into the 69
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metaluminous field, whereas the tonalites, granodiorites and monzogranites fall into the
2
peraluminous field. (c) In the AFM diagram (Irvine and Baragar 1971), the CID rocks
3
show a calc-alkaline trend. The diorites show a slight tendency towards tholeiitic
4
characteristics.
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1
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Fig. 7(2-column fitting image): Harker diagrams for the mafic and intermediate plutonic
7
rocks. With increasing SiO2, negative correlation trends are found for MgO, FeOt, TiO2 and
8
CaO, whereas the Na2O and K2O values show positive correlation with increasing SiO2.
9
Data in wt.%. References same as Fig. 6.
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6
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Fig. 8(2-column fitting image): Harker-style trace element variation diagrams, showing
12
decreasing Sr with increasing SiO2. Rb and Ba show positive correlation trends with
13
increasing SiO2. The intermediate rocks show increasing values of Zr with increasing SiO2.
14
Trace element data in ppm, SiO2 contents in wt.%. References same as Fig. 6.
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Fig. 9 (Single-column fitting image): (a) Upper crust-normalized trace element spider
17
diagram (after Taylor and McLennan, 1985). Intermediate and felsic rocks show
18
enrichment in the LILE compared to the mafic rocks. Also, note in the more evolved group
19
depletion in the HFSE. The mafic rocks show a slight depletion in the LILE and illustrate a
20
stronger depletion in the HFSE than the intermediate rocks. (b) REE abundances
21
normalized to chondrite (Boynton, 1984). The intermediate rocks show enrichment of the
22
LREE and a decreasing HREE pattern, compared to the mafic rocks. Note the similarities
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between the mafic CEO rocks with those from the Ojo de Colorados Complex, Western
2
Magmatic Belt (Zimmermann et al., 2014), Guasayán Pluton, Northern Sierras
3
Pampeanas (Dahlquist et al., 2016), a metadacite from Niño Muerto, Northeastern
4
Magmatic Belt (Hauser et al., 2011), and the Tastil batholith, Cordillera Oriental (Hauser
5
et al., 2011).
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1
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Fig. 10(2-column fitting image): U-Pb isotope data. (a) Concordia diagram (Wetherill,
8
1956) displaying the scatter of concordant U-Pb ages for zircon from monzogranite sample
9
D-13-01. Errors are shown as ellipses at the 2σ level. Upper left inset shows the Concordia
10
diagram for U-Pb ratios of all zircon grains analyzed. Lower right inset shows the weighted
11
average age diagram (data in Ma) for the monzogranite. (b) CL images of representative
12
zircon crystals from the monzogranite sample. Age results are shown for the spots
13
analyzed. (c) Concordia diagram displaying the concordant U-Pb ages for zircon from
14
granodiorite sample D-13-02. Errors are shown as ellipses at the 2σ level. Upper left inset
15
shows the Concordia diagram for U-Pb ratios of all zircon analyses. (d) CL images of
16
representative zircon crystals from the granodiorite sample. (e) Concordia diagram
17
displaying the scatter of concordant U-Pb ages for zircon from diorite sample D-13-40.
18
Errors are shown as ellipses at the 2σ level. Upper left inset shows the Concordia diagram
19
for U-Pb ratios of all zircon grains analyzed. Lower right inset shows the weighted average
20
age diagram (data in Ma) for the diorite unit. (f) CL images of representative zircon crystals
21
from the diorite sample. (g) PDP of the Monzogranite. (h) PDP for the Granodiorite data.
22
(i) PDP for the Diorite. (j) PDP for the combined data set from all three samples from the
23
Diablillos Intrusive Complex. Note: In (b), (d), and (f), orange circles mark the spots of U-
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1
Pb analyses. r: rim; c: core. Light blue circles mark the locations of Hf analyses. In (g), (h),
2
(i) and (j), ages for the various peak populations are given in black letters.
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Fig. 11(2-column fitting image): Plot of εHf(t) vs. U-Pb ages for zircon from the CID rocks.
5
Black dotted ellipse shows that most of the analyzed samples with ages between ~490 and
6
540 Ma are plotting near the CHUR line. The majority of the TDM ages are constrained
7
between the red dotted lines, denoting TDM ages from 1.2 to 1.55 Ga. The red solid line
8
indicates the most frequent TDM observed at 1.4 Ga. For comparison, the Ordovician
9
metadacite εHf(t) data from the Northeastern Magmatic Belt, the Cambrian Tastil batolith
10
εHf(t) data from the Cordillera Oriental (Hauser et al., 2011), granitoids from the
11
Northeastern Magmatic Belt (Bahlburg et al., 2016) and the data by Dahlquist et al.
12
(2016) for the Cambrian Guasayán Pluton in the Sierras Pampeanas are also shown. Inset
13
(a) ɛHf(t) histogram: The ɛHf(t) values for the selected samples range from +3.0 to -3.0,
14
whereby the value -2.0 is observed more frequently. Inset (b) shows a TDM age histogram
15
for the analyzed samples. The most frequent TDM age for the ~490-540 Ma zircon U/Pb
16
ages is ~1.4 Ga.
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17
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4
18
Fig. 12(2-column fitting image): (a) Nd isotope evolution diagram for the CID and CEO
19
rocks. TDM ages vary from ~1.3 to ~1.6 Ga (black dotted lines); the most frequent TDM age
20
is ~1.4 Ga. Note that the samples show ɛNd(t) values and TDM ages similar to the results
21
from Viramonte et al. (2007) for the Centenario salar area, Southeastern Magmatic Belt,
22
Puna, and to the data by Hauser et al. (2011) for the Northeastern Magmatic Belt, Puna,
23
and the Tastil batholith, Cordillera Oriental. (b) Nd isotopic ratios plot versus Sr isotopic 72
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ratios for the CID and CEO rocks. There is a negative correlation trend from the basic rocks
2
towards the felsic rocks (black dotted line). Felsic rocks display similar Nd and Sr ratios to
3
the rocks from the Cordillera Oriental and Eastern Magmatic Belt, Northern Puna (Hauser
4
et al., 2011).
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1
5
Fig. 13 (1.5-column fitting image): Schematic map showing the main geological provinces,
7
sampling sites, and related ages of some Lower Paleozoic granitoids in NW Argentina: the
8
Complejo Ojo de Colorados (Zimmermann et al., 2014) in the Western Magmatic Belt
9
(Puna), the Tastil batholith (Hauser et al., 2011), the granitic dykes in Mojotoro (Aparicio
10
González et al., 2011), and the Nevado de Chañi and Cañañi granitoids (Zapettini et
11
al., 2008; Escayola et al., 2011, respectively) in the Cordillera Oriental. Also shown is the
12
Guasayán pluton in the Northern Sierras Pampeanas (Dahlquist et al., 2016). Inset (a)
13
shows the Probability Density Plot illustrating zircon U-Pb ages for the Eastern Magmatic
14
Belt (Puna), from the CID, and the North and Southeastern Magmatic Belt; data from
15
Viramonte et al. (2007), Hauser et al. (2011), and Bahlburg et al. (2016). Note that the
16
Lull peak at ~515-520 Ma is higher than the Flare-up peak at ~470-480 Ma because we
17
used all the ages obtained in this work, concentrated at the former peak. Inset (b) shows the
18
Probability Density Plot for zircon U-Pb ages for the Cordillera Oriental and Northern
19
Sierras Pampeanas from Hauser et al. (2011), Aparicio González et al. (2011), and
20
Dahlquist et al. (2016). CID: Diablillos Intrusive Complex. COC: Complejo Ojo de
21
Colorados. CO: Cordillera Oriental. SP: Sierras Pampeanas. SS: Sierras Subandinas.
22
Modified from Aparicio González et al. (2011).
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1 2 3
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Figures:
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Fig. 1 (2-column fitting image):
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74
2
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Fig. 2 (2-column fitting image):
75
1
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Fig. 3(2-column fitting image): 76
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Fig. 4(Single-column fitting image): 77
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Fig. 5 (2-column fitting image):
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Fig. 6(2-column fitting image):
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Fig. 7(2-column fitting image):
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Fig. 8(2-column fitting image): 80
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Fig. 9 (Single-column fitting image): 81
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Fig. 10(2-column fitting image): 82
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Fig. 11(2-column fitting image):
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Fig. 12(2-column fitting image):
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Fig. 13 (1.5-column fitting image): 84
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Highlights •
U-Pb ages on zircon indicate that the emplacement of the Diablillos Intrusive
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Complex, Southern Puna, Argentina, took place between ~540 Ma until 490 Ma. •
The magmatic activity climax of the complex was at ~515-520 Ma.
•
The ɛHf(t) and ɛNd(t) values indicates that the rocks from the Diablillos Intrusive
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Complex, derived from an interaction of a dominant crustal source and juvenile mantle-derived magmas, with a TDM model age range of ~1.2 - 1.5 Ga, reworked
•
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during lower Paleozoic times.
The interaction of a Mesoproterozoic crustal source and juvenile mantle-derived magmas in Eastern Magmatic Belt, Puna, experienced a similar evolution in the Western Magmatic Belt, Cordillera Oriental, and Northern Sierras Pampeanas,
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emphasizing that this event was a regional episode.