Provenance of Ordovician clastic sequences of the San Rafael Block (Central Argentina), with emphasis on the Ponón Trehué Formation

Gondwana Research 19 (2011) 275–290

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Provenance of Ordovician clastic sequences of the San Rafael Block (Central Argentina), with emphasis on the Ponón Trehué Formation P. Abre a,⁎, C. Cingolani b, U. Zimmermann a,d, B. Cairncross a, F. Chemale Jr.

c

a

Department of Geology, University of Johannesburg, PO Box 524, Auckland Park 2006, Johannesburg, South Africa Centro de Investigaciones Geológicas, CONICET-Universidad Nacional de La Plata, Calle 1 no. 644, B1900TAC La Plata, Argentina Núcleo de Geociencias, Universidade Federal do Sergipe, Brazil d Present address: Department of Petroleum Engineering, University of Stavanger, 4036 Stavanger, Norway b c

a r t i c l e

i n f o

Article history: Received 23 May 2009 Received in revised form 26 April 2010 Accepted 23 May 2010 Available online 4 June 2010 Keywords: Cuyania terrane Provenance Geochemistry Isotope geochemistry Detrital zircon dating Ordovician Ponón Trehué and Pavón Formations

a b s t r a c t The Ordovician Ponón Trehué Formation is the only early Palaeozoic sedimentary sequence known to record a primary contact with the Grenvillian-age basement of the Argentinean Cuyania terrane, in its southwards extension named the San Rafael block. Petrographic and geochemical data indicate contributions from a dominantly upper continental crustal component and a subordinated depleted component. Nd isotopes indicate εNd of − 4.6, ƒSm/Nd − 0.36 and TDM 1.47 Ga in average. Pb-isotope ratios display average values for 206 Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb of 19.15, 15.69 and 38.94 respectively. U–Pb detrital zircon ages from the Ponón Trehué Formation cluster around values of 1.2 Ga, indicating a main derivation from a local basement source (Cerro La Ventana Formation). The Upper Ordovician Pavón Formation records a younger episode of clastic sedimentation within the San Rafael block, and it shows a more complex detrital zircon age population (peaks at 1.1 and 1.4 Ga as well as Palaeoproterozoic and Neoproterozoic detrital grains). Detailed comparison between the two Ordovician clastic units indicates a shift with time in provenance from localized basement to more regional sources. Middle to early Upper Ordovician age is inferred for accretion of the Cuyania terrane to the proto-Andean margin of Gondwana. © 2010 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1. Introduction The Cuyania terrane (Fig. 1) in central Argentina is characterized by a Mesoproterozoic (Grenvillian-age) basement with depleted Pb isotopic signatures and Mesoproterozoic Nd model ages resembling basement rocks of the same age from Laurentia (Ramos, 2004; Sato et al., 2004 and references therein). Several authors have proposed para-autochthonous (Aceñolaza et al., 2002; Finney et al., 2005) versus allochthonous (e. g. Ramos et al., 1986; Dalla Salda et al., 1992; Cingolani et al., 1992; Dalziel et al., 1994; Astini et al., 1995; Thomas and Astini, 1996; Keller, 1999) geotectonic models for the early Palaeozoic evolution of the Cuyania terrane. No consensus has been reached since all the stratigraphical, palaeontological, palaeomagnetic, and isotopic data available thought to support the palaeogeographic proximity of this terrane to Laurentia (e.g. Benedetto, 1993; Lehnert and Keller, 1993; Buggisch et al., 1993; Ramos et al., 1998; Thomas et al., 2001; Rapalini and Cingolani, 2004), can also be interpreted as indicating a Gondwanan signature (Aceñolaza et al., 2002; Finney, 2007). ⁎ Corresponding author. Tel.: +27 727195504; fax: +54 221 4215677. E-mail addresses: [email protected] (P. Abre), [email protected] (C. Cingolani), [email protected] (F. Chemale).

The evolution of the Cuyania terrane concept could be summarized as follows: The term ‘Precordillera’ was first used for a physiographic province within the deformed Andean foreland of western Argentina. Mainly Palaeozoic sedimentary rocks are exposed in a Cenozoic thinskinned fold-thrust belt (Fig. 1) generated by flat slab subduction of the Nazca Plate (28° to 33° SL). This geological province is characterized by the extent of its fossil-rich Cambrian–Middle Ordovician carbonate platform, unique in South America and was called the Precordillera s.st. The Precordillera terrane concept was coined later on (Ramos et al., 1986; Ramos, 1988) to name an early Palaeozoic Laurentian derived exotic block, with a Grenvillian-age basement (registered in xenoliths in Tertiary volcanic rocks), that was accreted to the pre-Andean Gondwana margin (Pampia terrane). Strong evidence that the Grenvillian-age basement and the early Palaeozoic carbonate and siliciclastic cover extended further to the East and South of the original proposed terrane, lead Ramos et al. (1998) to introduce the concept of a greater Cuyania composite terrane. This terrane incorporates the early Palaeozoic Precordillera s.st. as well as its southern extension into the San Rafael and Las Matras blocks, along with adjacent parts of the Western Pampeanas Ranges (e.g. Pie de Palo Range) that comprise a Grenvillian-age basement (Ramos, 2004). However, some authors as Finney, (2007) described the Cuyania terrane as the ‘greater Precordillera’ or ‘Precordillera terrane’. Whether minor ranges from

1342-937X/$ – see front matter © 2010 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. doi:10.1016/j.gr.2010.05.013

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Fig. 1. Satellite image based map showing Cuyania terrane boundaries as dashed lines and blocks boundaries as continuum lines (based on Ramos et al., 2000; Astini and Dávila, 2004; Porcher et al., 2004). All the entities forming the Cuyania terrane develop a Grenvillian-age basement characterized by Nd, Sr and Pb depleted isotopic signatures (Ramos, 2004; Sato et al., 2004). Righter inlet: location of neighbouring terranes.

the Western Pampeanas Ranges form as well part of this crustal block is still under debate (e.g. Umango Range; Porcher et al., 2004). The western boundary of the Cuyania terrane coincides with the western boundaries of the Precordillera s.st., and the San Rafael and the Las Matras blocks, being delimited by the Chilenia terrane

(Ramos et al., 1996; Fig. 1). Despite the disagreement regarding the geotectonic evolution of the Cuyania terrane, provenance analyses of its early Palaeozoic sedimentary record using modern techniques are scarce (e.g. Loske, 1994; Cingolani et al., 2003; Naipauer, 2007; Gleason et al., 2007; Naipauer et al., 2010).

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Two Ordovician units are found within the San Rafael block, named the Ponón Trehué (Darriwilian to Sandbian) and Pavón (Sandbian) Formations. The main purpose of this paper is to examine newly obtained data from the Ponón Trehué Formation using petrography, whole-rock geochemistry and isotope geochemistry (Sm–Nd and Pb–Pb systematics). U–Pb detrital zircon laser ablation dating from the Ponón Trehué and Pavón Formations is also presented. Comparison of the provenance indicators of the Ponón Trehué Formation, which is the only unit that directly contacts the Cerro La Ventana Formation (basement of the San Rafael block), with the Pavón Formation, reveals important information of the sources for the clastic deposition during the proposed Ordovician accretion of the Cuyania terrane. 2. Geological setting of the Ponón Trehué Formation The Darriwilian to Sandbian Ponón Trehué Formation is a fossilrich, carbonate-siliciclastic sequence unconformably overlying Mesoproterozoic basement of the Cerro La Ventana Formation (Heredia, 1996; Cingolani and Varela, 1999; Beresi and Heredia, 2003; Cingolani et al., 2005; Heredia, 2006). It outcrops near the southern end of the San Rafael block (Cuyania terrane), Mendoza Province, central-western Argentina (Fig. 2). It was subdivided into

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two members, which comprise Conodont Biozones (Heredia, 2006). The lower member (Peletay) is composed of conglomerates and conglomeratic arkoses, limestones, quartz–arenites and black shales. The upper member (Los Leones) is composed of mudstones, siltstones, arenites and conglomeratic arenites. Blocks of limestones, granitoids, gneisses and amphibolites are common in the upper part of the younger member (Bordonaro et al., 1996). The limestone blocks bearing invertebrate fossils resemble the Lower/Middle Ordovician La Silla and San Juan Formations of the Precordilleran platform (Heredia, 2006), while the crystalline clast and block compositions resemble Cerro La Ventana Formation basement, indicating substantial reworking of local lithologies. The continental sedimentary Carboniferous arkosic sandstones overlie the Ponón Trehué Formation through either an unconformity or a fault contact (Fig. 3). The Ponón Trehué Formation is exposed in three areas (Fig. 3). The outcrops located to the south of the Ponón Trehué Creek (Fig. 3, areas 2 and 3) were synchronously named as Ponón Trehué Formation (Criado Roqué and Ibáñez, 1979) and as Lindero Formation (Núñez, 1979). Bordonaro et al. (1996) defined the Ponón Trehué Formation as to the outcrops located on the southern edge of the Ponón Trehué Creek as well as the outcrop located 1.5 km to the north of the mentioned creek (Fig. 3, areas 1 and 2). The name Lindero Formation

Fig. 2. Geological sketch map of the San Rafael Block (simplified from Dessanti, 1956; González Díaz, 1972; Núñez, 1979). See the location of the main outcrops of the Pavón and Ponón Trehué Formations. On the righter side, the integrated lithostratigraphic column chart of the San Rafael Block is shown. CLV: Cerro La Ventana Formation; PT: Ponón–Trehué Formation (Darriwilian–Sandbian); P: Pavón Formation (Sandbian). PT is the only unit in contact with the basement of the Cuyania terrane.

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Fig. 3. Detailed geology of the Ponón Trehué area, with outcrops of the Ponón Trehué Formation in the central-northern part (areas named 1, 2 and 3). On the righter side, the lithostratigraphic section of the Ponón Trehué Formation at the La Tortuga section is shown. Samples used for the present study (numbered to the right) were taken at the lower part of the section where the platform deposits are represented; the upper part of the section is composed by olistostromic-type deposits. Modified from Núñez (1979), Astini (2002) and Heredia (2006).

was retained for the outcrops to the south of the Ponón Trehué Creek (also south of the Cerro Lindero; Fig. 3, area 3). Subsequent studies allowed Heredia (1996, 2006) to establish that all the abovementioned outcrops belong to an olistostromic sequence and the term Ponón Trehué Formation has since been used and is here applied. The olistostromes are related to extensional tectonics interpreted as a response to flexural subsidence of the Precordillera carbonate platform starting in late middle Ordovician (Darriwilian) and may indicate the time of accretion of Cuyania to the proto-Andean margin of Gondwana (Astini, 2002).

3. Sampling and analytical techniques Sampling for all analyses of the Ponón Trehué Formation was done at La Tortuga section (Fig. 3 and area 3; 35° 10′ 53″ S–68° 18′ 13″ W). Detrital zircons of the Pavón Formation were obtained from a subfeldspathic–arenite taken from outcrops located on the eastern side of the Cerro Bola region (Fig. 2). Samples for geochemical analyses were prepared and milled at CIG (La Plata, Argentina). Major and selected trace elements (Ni, V, Cu, Ga, Sr, Y, Zr, Zn, Nb, Rb, Ba and Pb) were measured on fusion beads (using a 50/50 lithium metaborate/

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lithium tetraborate as flux) and on pressed powder tablets (using 8:4 g ratio between sample and Herzog binder pellets), using a Phillips wavelength-dispersive X-Ray Fluorescence spectrometer at SPECTRAU (Central Analytical Facilities of the University of Johannesburg, South Africa). Detection limits using the XRF are Si2O = 250 ppm, Al 2 O 3 = 145 ppm, CaO = 40 ppm, MgO and Na 2 O = 65 ppm, K2O = 47 ppm, MnO = 15 ppm, TiO2 = 20 ppm, Fe2O3 = 150 ppm, P2O5 = 9 ppm, Ni, Rb, Sr, Cu, Y and Zr = 1 ppm, V = 3 ppm, Nb and Pb = 0.5, Ba= 13 ppm. The loss on ignition was calculated by weight difference prior and after heating 1 g of sample powder for 2 h at 1100 °C in an electric oven. Other trace elements (Sc, Cr, Co, Cs, Hf, Ta, W, Th, U, As and Sb) and rare earth elements (La, Ce, Nd, Sm, Eu, Tb, Yb and Lu) were obtained by INAA (Instrumental Neutron Activation Analysis) at ACME Laboratories (Canada). Detection limits for elements analyzed by INNA are Sc, Sm and Sb = 0.1 ppm, Cr and Nd= 5 ppm, Co, Cs, Hf and W = 1 ppm, Ta, U, As, La and Tb= 0.5 ppm, Th, Eu and Yb= 0.2 ppm, Ce = 3 ppm, and Lu= 0.05 ppm. Errors are at 1-sigma. For the isotopic determinations of Sm–Nd, whole-rock powders were spiked with mixed 149Sm–150Nd tracer and dissolved in Teflon vial using an HF–HNO3 mixture and 6 N HCl until complete material dissolution. Column procedures used cationic AG-50W-X8 (200– 400 mesh) resin in order to separate REE, followed by Sm and Nd separation using anionic politeflon HDEHP LN-B50-A (100–200 µm) resin according to Patchett and Ruiz (1987). Each sample was dried to a solid and then loaded with 0.25 N H3PO4 on appropriated filament (single Ta for Sm and triple Ta–Re–Ta for Nd). Isotopic ratios were measured in static mode with a VG Sector 54 multi-collector mass spectrometer at the Laboratório de Geologia Isotópica of Universidade Federal do Rio Grande do Sul (LGI-UFRGS, Brazil). 100–120 ratios with a 0.5 to 1-volt 144Nd beam were normally collected. Nd ratios were normalized to 146Nd/144Nd = 0.72190. All analyses were adjusted for variations instrumental bias due to periodic adjustment of collector positions as monitored by measurements of our internal standards. Measurements for the Spex 143Nd/144Nd = 0.511130 ± 0.000010. Total blanks average were b150 pg for Sm and b500 pg for Nd. Correction for blank was insignificant for Nd isotopic compositions and generally insignificant for Sm/Nd ratios. For the Pb isotopic measurements, an aliquot of 1 ml from dissolved whole-rock samples used for Sm–Nd analysis was taken. Pb was extracted with ion exchange techniques, with AG-1 X 8, 200– 400 mesh, anion resin. Each sample was dried to a solid and added a solution of HNO3 with 50 ppb Tl in order to correct the Pb fractionation during the analyses (Tanimizu and Ishikawa, 2006). Isotopic Pb compositions were obtained at LGI-UFRGS with a Neptune MC-ICP-MS in static mode, with collecting of 60 ratios of Pb isotopes. The obtained values of NBS 981 common Pb standard during the analyses were in agreement with the NIST values. Sandstones were crushed and sieved to less than 100 µm to obtain zircon crystals. Through hydraulic processes the heaviest fraction was separated, which was treated with bromoform (δ = 2.89) and methylene iodide (δ = 3.32) to obtain a fraction enriched in zircons, followed by an electromagnetic separation with a Frantz Isodynamic Separator. The final selection of individual crystals was done by handpicking under a binocular microscope. For U–Pb dating all zircons were mounted in epoxy in 2.5 cm-diameter circular grain mounts and polished down until the zircons were revealed. Cathodoluminescence images of each individual zircon grains were obtained with a SEMBSE-EDS (JEOL JSM-5600 with a tungsten filament and EDS analyses were done using a Noran X-ray detector and Noran Vantage software; the system was set at 15 keV, a working distance of 20 mm and a live time of 60 s per spot). Zircon grains were dated with a laser ablation microprobe (New Wave UP213) coupled to a MC-ICP-MS (Neptune) at LGI-UFRGS. Isotope data were acquired using static mode with spot sizes of 25 and 15 µm. Laser-induced elemental fractional and instrumental mass discrimination were corrected by the reference zircon GJ-1 (Simon et al., 2004), following the measurement of two

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GJ-1 analyses to every five sample zircon spots. The external errors were calculated after propagation error of the GJ-1 mean and the individual sample zircon (or spot). The laser operating conditions were: laser output power of 6 J/cm2 and a shot repetition rate of 10 Hz. The cup configuration of the MC-ICP-MS Neptune was: Faradays 206Pb, 208Pb, 232Th, 238U, MIC's 202Hg, 204Hg + 204Pb, 207Pb. The gas input included a coolant flow (Ar) at 15 l/min, an auxiliary flow (Ar) at 0.8 l/min and a carrier flow of 0.75 l/min (Ar) + 0.45 l/min (He); the acquisition was at 50 cycles of 1.048 s. 4. Petrography and geochemistry The samples studied from the Ponón Trehué Formation are claystones, siltstones and fine-grained sublith- and subfeldspathicarenites (Dott, 1964). The sandstones are moderately sorted and comprise a small amount of clay-rich matrix. Monocrystalline quartz is subrounded to subangular. Polycrystalline quartz, also subangular to subrounded, with no sutured boundaries is sometimes present. Kfeldspar is usually totally replaced by chlorite or clay minerals and subrounded. Sedimentary lithoclasts derived from siltstones, carbonates, mudstones and chert are present. Calcite cement is present as well as very scarce detrital muscovite lamellae. Zircon, apatite, chromian spinel, tourmaline, rutile, Fe-oxides (including hematite) and other opaque minerals comprise the heavy minerals fraction of the Ponón Trehué Formation. In the QFL ternary diagram, samples from the Ponón Trehué Formation plot in the recycled orogen field (Fig. 4). The alteration of feldspars to clay minerals (point-counted as matrix) resulted in a displacement towards the Q apex, when plotting the samples analyzed. Therefore, interpretations of tectonic setting using the QFL diagram needs to be taken with caution. Elements concentrated in mafic (Sc, Cr, and Co) and in silicic (La, Th, and REE) rocks, REE (rare earth elements) patterns and the character of the Eu-anomaly have been used for provenance and tectonic setting determinations (Taylor and McLennan, 1985; McLennan and Taylor, 1991), being particularly useful in cases were petrographical data are not conclusive. The signature of the source rock may be modified by weathering, hydraulic sorting and diagenesis (Nesbitt and Young, 1982; Cullers et al., 1987, Nesbitt et al., 1996). Therefore, the determination of the effects that these factors had on the geochemical composition of sedimentary rocks provides evidence for the correct interpretation of the provenance (e.g. Cox and Lowe,

Fig. 4. Ternary diagram (Dickinson and Suczek, 1979; Dickinson et al., 1983) showing a provenance from recycled orogen for the Ponón Trehué Formation. Qt = total quartz; F = K-feldspar + plagioclase; L = total lithoclasts (Lm + Lv + Ls), and their range of variation is: Qt = 79–83%, F = 8–11% and L = 6–12%. The shadow area represents the data spread of the Pavón Formation (data from Cingolani et al., 2003).

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1995). The chemical composition of all samples from the Ponón Trehué Formation analyzed is shown in Table 1. Weathering effects on sedimentary rocks can be quantitatively assessed using the Chemical Index of Alteration (CIA; Nesbitt and Young, 1982). The CIA values for sandstones range from 69 to 79 and those for mudstones are between 74 and 76 (Table 1). The distribution of samples within the A–CN–K space (Fig. 5a) cannot be easily explained by a normal weathering path (the field of vertical lines in Fig. 5a indicates the predicted weathering trend for the average upper crustal composition; Bock et al., 1998). Diagenetic processes or mixing of sources should therefore explain the behaviour of samples from the Ponón Trehué Formation. Thus, certain provenance discriminators based on elements subject to remobiliza-

tion during diagenesis cannot be used in this case (e.g. Roser and Korsch, 1986). The trace elements are considered useful for provenance determination, because they tend to reflect source rock compositions. In particular, the concentrations of high field strength elements such as Zr, Nb, Hf, Y, Th and U are the most useful (Taylor and McLennan, 1985). The REE represent reliable provenance indicators, as they would reflect the average REE composition of the source material (Taylor and McLennan, 1985; McLennan, 1989). During weathering, there is a tendency for an elevation in the ratio between Th and U to above upper crustal igneous values, due to the oxidation of U4+ to the more soluble U6+ (McLennan, 1989). Most of the samples from the Ponón Trehué Formation display Th/U ratios

Table 1 Chemical analyses of the Ponón Trehué Formation. Major elements are in % whereas trace and REE are in ppm. N denotes normalized to chondrite values. Aver: average; SD: standard deviation; b.d.l.: below detection limits. Eu/Eu* = EuN / (0.67SmN + 0.33TbN) and CeN/Ce* = CeN / (0.67 LaN + 0.33 NdN), where subscript N denotes values normalized to chondrite. UCC: upper continental crust; data from McLennan et al. (2006). Mudstones

Sandstones

Sample

CT3

CT7

CT8

CT6

CT1

CT2

CT4

CT5

Aver

SD

SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Σ CIA Sc Cr Ni Cs Rb Ba Hf Th U V Sr Y Zr Nb Pb Zn La Ce Nd Sm Eu Tb Yb Lu ΣREE Eu/Eu* Ce/Ce* K/Cs Th/Sc Zr/Sc Th/U La/Sc La/Th Cr/V Y/Ni Ti/Nb LaN/YbN LaN/SmN TbN/YbN Cr/Th Ti/Zr

74.36 1.42 13.00 3.62 0.01 1.04 0.27 0.62 2.93 0.17 3.26 100.70 75 14.3 240 47 3 110 1109 10 11 2.7 123 45 39.6 319 25 11.3 65 39 95 41 10.7 1.9 1.6 5.0 0.89 195.09 0.54 1.1 8107 0.77 22.31 4.07 2.73 3.55 1.95 0.84 343 5.27 2.29 1.37 21.82 26.63

59.11 1.54 20.28 5.76 0.04 1.44 0.21 0.64 4.75 0.07 4.67 98.52 76 20.1 150 49 5 178 1145 7 14 1.5 158 66 38.6 253 27 11.9 86 43 85 34 8.1 1.3 1.2 5.3 0.88 178.78 0.49 0.94 7879 0.70 12.59 9.33 2.14 3.07 0.95 0.79 337.4 5.48 3.34 0.97 10.71 36.50

60.30 1.46 18.32 8.09 0.04 1.42 0.26 1.29 3.81 0.11 5.00 100.09 74 19.3 150 75 5 153 763 7 14 4.3 148 85 38.2 272 25 19.9 117 42 83 36 8.6 1.5 1.0 5.2 0.84 178.14 0.56 0.93 6322 0.73 14.09 3.26 2.18 3.00 1.01 0.51 342.9 5.46 3.07 0.82 10.71 32.07

70.61 1.40 13.23 6.03 0.08 1.23 0.16 0.85 2.78 0.08 3.33 99.78 74 14.3 150 45 3 107 756 9 11 4.0 119 45 25.9 299 24 14.0 84 31 70 31 6.4 1.1 1.1 4.3 0.8 145.7 0.51 1.03 7698 0.77 20.91 2.75 2.17 2.82 1.26 0.58 353.1 4.87 3.05 1.09 13.64 28.11

78.99 0.42 7.77 3.09 0.02 0.52 0.55 0.48 2.40 0.40 3.52 98.15 69 10.0 80 87 3 83 4969 3 7.4 3.2 193 70 23.4 110 10 39.6 154 30 59 20 5.6 1 1 2.6 0.4 119.6 0.52 0.96 6627 0.74 11.00 2.31 3.00 4.05 0.41 0.27 263.2 7.80 3.37 1.64 10.81 22.98

83.85 0.19 3.59 1.82 0.06 0.23 2.67 1.02 1.04 0.17 3.54 98.16

75.32 1.39 10.86 4.31 0.04 0.91 0.17 0.78 2.34 0.09 2.79 98.98 74 12.9 150 31 3 89 964 9 11 3.1 117 38 25.7 295 23 10.2 69 31 71 25 6.7 1.2 1.2 4.6 0.79 141.49 0.53 1.09 6461 0.85 22.87 3.55 2.40 2.82 1.28 0.83 358.1 4.55 2.91 1.12 13.64 28.29

77.67 1.32 8.67 4.83 0.03 0.81 0.21 0.37 1.50 0.11 3.74 99.26 79 11.1 170 25 2 61 575 12 120 3.0 115 36 30.3 402 22 13.7 60 31 75 22 7.2 1.2 1.3 4.8 0.83 143.33 0.49 1.17 6234 1.08 36.22 4.00 2.79 2.58 1.48 1.21 359.9 4.36 2.71 1.16 14.17 19.67

72.53 1.14 11.96 4.69 0.04 0.95 0.56 0.76 2.69 0.15 3.73 99.20 74.4 13.26 153.75 49.56 3.43 102.1 1763.1 7.25 10.51 2.98 132 53.25 28.02 250.68 20.39 17.03 85.88 32.63 70.75 27.63 7.00 1.21 1.20 4.11 0.70 145.08 0.52 1.02 7047 0.82 19.17 3.89 2.60 3.21 1.25 0.64 315.46 5.80 3.00 1.17 16.67 26.86

8.22 0.49 5.15 1.83 0.02 0.40 0.80 0.28 1.11 0.10 0.69 0.87 2.6 4.81 40.91 20.92 1.05 43.3 1557.2 3.42 3.23 0.87 31.26 16.84 11.41 106.30 7.21 9.00 31.92 8.64 19.12 8.96 2.19 0.38 0.19 1.39 0.24 40.14 0.02 0.09 751 0.12 7.81 2.18 0.43 0.49 0.44 0.33 63.57 1.45 0.34 0.25 8.69 5.36

4.1 140 23 b.d.l. 36 3824 1 3.7 2.0 83 41 2.4 55 7 15.6 55 14 28 12 2.7 0.5 b.d.l. 1.1 0.2 58.5 0.94 0.90 13.41 1.85 3.41 3.78 1.69 0.10 165.9 8.60 3.26 37.84 20.71

UCC 66 0.5 15.2 4.5 0.06 2.2 4.2 3.9 3.4 0.16 100.1 50 13.6 83 44 4.6 112 550 5.8 10.7 2.8 107 350 22 190 12 17 71 30 64 26 4.5 0.9 0.64 2.2 0.32 128.54 0.63 1.07 6136 0.79 14 3.8 2.2 2.8 0.77 0.5 9.3 4.2 1.24 7.76 12.9

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Fig. 5. Geochemical data of the Ponón Trehué Formation (PT) were plotted on several diagrams. Data from Pavón Formation are from Cingolani et al. (2003). a) A–CN–K diagram based on molecular proportions. The CIA scale is shown on the left. This index uses molecular proportions as follows: CIA = (Al2O3 / (Al2O3 + CaO* + Na2O + K2O)) × 100, where CaO* refers to the calcium associated with silicate minerals. The average upper continental crust according to Taylor and McLennan (1985), as well as idealized mineral compositions. Field of vertical lines indicates the predicted weathering trend for the average upper crustal composition. b) Plot of Th/U versus Th based on McLennan et al. (1993). c) Th/Sc versus Zr/Sc diagram after McLennan et al. (2003). d) Chondrite normalized REE patterns for the Ponón Trehué Formation. PAAS = post-Archaean Australian shales pattern (Nance and Taylor, 1976) is draw for comparison. Continuous grey lines (QC1 and QR1) correspond to the samples with the highest and the lowest sum of REE from the Pavón Formation (data from Cingolani et al., 2003). Chondrite normalization factors are those listed by Taylor and McLennan (1985).

typical for upper crustal derived rocks (Fig. 5b). Loss of U due to weathering is evident for sample CT7 (U concentration of 1.5 ppm; Table 1). Low Th/U ratios due to U gain are also observed (Fig. 5b). The Zr/Sc ratio reflects reworking because Zr is strongly enriched in zircon whereas Sc is not, and zircon can be easily recycled (McLennan et al., 1993). The Th/Sc ratio indicates the degree of igneous differentiation processes since Th is preferentially partitioned into melts during crystallization, and as a result, it is enriched in felsic rather than mafic sources, whereas Sc is concentrated in mafic igneous rocks (Feng and Kerrich, 1990; McLennan et al., 1993). In Fig. 5c, the Ponón Trehué Formation shows a cluster of data indicating that processes of recycling were not important and that the source had dominantly a typical upper crustal composition. However, the high Sc content of certain samples indicates a subordinated mafic input. The input of a mafic source could be discriminated using the Y/Ni and Cr/V ratios (McLennan et al., 1993). Cr/V ratios higher, and Y/Ni ratios lower than the UCC, along with enrichment in elements such as Sc, Cr and V compared with the UCC values might indicate the influence of a source with a composition less evolved than the average UCC for the Ponón Trehué Formation (Table 1). An ophiolitic source can be neglected.

The chondrite normalized REE patterns for mudstones and sandstones of the Ponón Trehué Formation (Fig. 5d) show a moderately enrichment in LREE, a negative Eu-anomaly and a suspected rather flat HREE. The samples are enriched in HREE compared with the PAAS, which is considered in turn to reflect the average composition of the post-Archaean upper crust. Sample CT2 shows the effects of dilution in the REE concentration due to high silica (Table 1). 5. Isotope geochemistry 5.1. Sm–Nd The Sm/Nd ratio is primarily modified during processes of mantle– crust differentiation allowing the estimation of the model age or the time at which the initial magma was separated from the upper mantle (Nelson and DePaolo, 1988). In the case of sediments, it can only approximate the average crustal residence age of the protoliths. Results from the five selected samples from the Ponón Trehué Formation analyzed for the present provenance study are shown in Table 2. The εNd (t) values where t = 462 Ma (depositional age) are of −4.47 ± 0.39 in average; ƒSm/Nd has an average value of −0.37 ± 0.02 and the average TDM age is 1.44 ± 0.078 Ga.

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Table 2 Sm–Nd and Pb–Pb isotopes for the Ponón Trehué Formation. TDM calculated according to DePaolo (1981). n.d.: no data. The εNd indicates the deviation of the 143Nd/144Nd value of the sample from that of CHUR (DePaolo and Wasserburg 1976), whereas the ƒSm/Nd is the fractional deviation of the sample 147Sm/144Nd from a chondritic reference. εNd (0) = {[(143Nd / 144 Nd)sample (t= 0)/ 0.512638]− 1}*10,000. εNd (t) = {[(143Nd/ 144Nd)sample (t) / (143Nd / 144Nd)CHUR (t)]− 1}* 10,000. ƒSm/Nd = (147Sm / 144Nd)sample / (147Sm/ 144Nd)CHUR− 1. εNd (t) = {[(143Nd / 144Nd)sample (t) / (143Nd/144Nd)CHUR (t)]− 1} * 10,000. 143Nd/144NdCHUR= 0.512638. 147Sm/144NdCHUR = 0.1967. Pb data are corrected for mass fractionation. Sample

CT1

CT3

CT4

CT5

CT6

CT8

Age (Ma) Sm (ppm) Nd (ppm) 147 Sm/144Nd 143 Nd/144Nd Error (ppm) εNd (0) εNd (t) TDM (Ma) ƒSm/Nd 206 Pb/204Pb Error (ppm) 207 Pb/204Pb Error (ppm) 208 Pb/204Pb Error (ppm) 208 Pb/206Pb Error (ppm) 207 Pb/206Pb error (ppm)

462 4.29 21.88 0.11858 0.512230 18 − 7.96 − 3.95 1298 − 0.40 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.

462 6.22 28.87 0.13016 0.512203 13 − 8.48 − 4.56 1522 − 0.34 19.303141 14 15.706133 10 38.989437 16 2.0198499 17 0.81364848 14

462 4.50 21.56 0.12613 0.512173 36 − 9.07 − 4.91 1504 − 0.36 19.173025 22 15.703028 23 38.970501 23 2.0325439 15 0.81902006 16

462 4.54 21.68 0.12676 0.512217 14 − 8.20 − 4.08 1438 − 0.36 19.160300 31 15.703960 38 38.975361 40 2.0341752 11 0.81960700 3

462 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 19.100324 16 15.689129 19 38.833531 17 2.0331038 8 0.82139061 9

462 5.47 27.32 0.12102 0.512161 14 − 9.31 − 4.84 1442 − 0.38 19.027669 5 15.666211 5 38.948940 14 2.0469290 23 0.82333189 34

The TDM ages are comparable to TDM ages for Mesoproterozoic and supracrustal younger rocks of the Cuyania terrane (Kay et al., 1996; Cingolani et al., 2003; Sato et al., 2004; Cingolani et al., 2005; Gleason et al., 2007). They are also within the range of variation of the TDM ages of Neoproterozoic to Palaeozoic sequences in northwestern Argentina, including the igneous rocks of the Pampia terrane (Rapela et al., 1998; Bock et al., 2000; Lucassen et al., 2000). Similar data are known from the basement of Chilenia (Bahlburg et al., 2001), from Mesoproterozoic rocks from Antarctica, Falklands/Malvinas plateau and Natal–Namaqua Metamorphic belt (Wareham et al., 1998) and the Western Pampeanas Ranges (Vujovich et al., 2005). Fig. 6a shows the relationship between εNd (t) and Th/Sc ratio of samples from the Ponón Trehué Formation where it is seen that the εNd (t) values obtained are neither typical of UCC nor of a juvenile input and the Th/Sc ratio is indicative of a source less fractionated than the UCC but not clearly mafic in composition. The same can be deduced using the plot of ƒSm/Nd versus εNd (see Fig. 6b), where the ƒSm/Nd values could be assigned to an old upper crust or an arc component but the εNd (t) values are between the two fields. εNd (t) values for the Ordovician units of the San Rafael block are similar to those from other Ordovician sedimentary rocks from the Cuyania terrane (Abre, 2007; Gleason et al., 2007), as well as from rocks of the Famatinian arc (Pankhurst et al., 1998; Fig. 7). εNd values are within the range of variation of the Laurentian Grenville crust (Patchett and Ruiz, 1989) and the Central and Southern domains of the Arequipa–Antofalla Basement (sensu Loewy et al., 2004). Nd data presented by Cingolani et al. (2005) for the Cerro La Ventana Formation show εNd values in the range of variation of data calculated to the time of deposition of the Ponón Trehué Formation (Fig. 7).

5.2. Pb–Pb The use of Pb isotopes is an established tool for evaluating the provenance of clastic sedimentary rocks (Hemming and McLennan, 2001). Pb isotopes from five selected samples from the Ponón Trehué Formation were analyzed for the present provenance study. The 206 Pb/204Pb ratio ranges from 19.028 to 19.303, the 208Pb/204Pb ratio ranges from 38.83 to 38.99, whereas the radiogenic 207Pb/204Pb ratio is in between 15.66 and 15.71 (Table 2). On an uranogenic-Pb diagram (Fig. 8a), samples from the Ponón Trehué Formation plot slightly above the Stacey and Kramers (1975)

second stage Pb evolution curve for average crust. They overlap the field of Proterozoic rocks of the Southern and Central domains of the Arequipa–Antofalla Basement (Aitcheson et al., 1995; Tosdal, 1996; Loewy et al., 2004). A similar behaviour is observed for the samples on a thorogenic (208Pb/204Pb) versus 206Pb/204Pb present-day composition (Fig. 8b). In both diagrams, it is also evident that the Pb system of the Ponón Trehué Formation differs consistently from the Grenvillian xenoliths of the inferred basement of the Cuyania terrane (Kay et al., 1996), Proterozoic rocks from Eastern North America, rocks from the Northern domain of the Arequipa–Antofalla Basement (Tosdal, 1996; Loewy et al., 2004) and from Mesoproterozoic rocks of the Natal– Namaqua Metamorphic belt, Falkland/Malvinas Microplate and Antarctica (Wareham et al., 1998). Compared with supracrustal rocks of the Cuyania terrane (Gleason et al., 2007) the data here presented have higher 208Pb/204Pb and lower 206Pb/204Pb ratios (Fig. 8).

6. Comparison with the Pavón Formation The early Upper Ordovician (Sandbian) Pavón Formation crops out in the central area of the San Rafael block (Fig. 2). The unit appears in isolated outcrops dismembered by Tertiary tectonism; it is neither in contact with the Grenvillian-age basement nor with the Ponón Trehué Formation and it is covered by Upper Palaeozoic volcaniclastic rocks. It is a sandy marine turbidite sequence composed of arenites, wackes and pelites and bearing Climacograptus bicornis Biozone (Cuerda and Cingolani, 1998; Manassero et al., 1999). Geological details regarding the Pavón Formation can be found in Cingolani et al. (2003). The Pavón Formation comprises feldspathic- and quartz–wackes and subordinated subfeldspathic- and sublith-arenites. The sandstone composition of the Pavón Formation indicates in QFL diagrams of Dickinson et al. (1983), a provenance from recycled orogen and continental block (Fig. 4; Cingolani et al., 2003). Noteworthy is the presence of detrital chromian spinels derived from host rocks emplaced within mid-ocean ridge and intraplate environments (Abre, 2007; Abre et al., 2009). Geochemical data for the Pavón Formation can be found in Cingolani et al. (2003). The CIA values for this unit indicate intermediate to advanced weathering conditions, but samples follow a general weathering trend parallel to the A–CN boundary (Fig. 5a). Th and U concentrations and Th/U and Zr/Sc ratios are similarly variable compared with the Ponón Trehué Formation

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chromian spinels confirm inputs from a depleted source of an uncertain age. Detrital zircon ages also suggest that the source of the Pavón Formation is not as restricted as that of the Ponón Trehué Formation (see below). Sources that could have contributed to this less negative εNd value are found within the Central and Southern domains of the Arequipa–Antofalla Basement (Fig. 7; Loewy et al., 2004). Sm–Nd data from the Western Pampeanas Ranges are scarce and ages of the analyzed rocks are not available (Porcher et al., 2004; Vujovich et al., 2005), being therefore not possible to calculate their εNd at the time of deposition of the Pavón Formation. However, the positive εNd values found within the Umango Range (εNd(t = 1.1Ga) between +1.4 and + 4.2; Porcher et al., 2004) are worth noting as probable contributors of detritus (Fig. 7). 7. Detrital zircon dating of the Ponón Trehué and Pavón Formations

Fig. 6. a) εNd (t) versus Th/Sc ratios and b) ƒSm/Nd versus εNd (t). In both diagrams it is shown that the Ponón Trehué Formation shows a similar distribution compared with the Pavón Formation (data from Cingolani et al., 2003), although one sample from the latter unit has an εNd (t) of − 0.4 and a low Th/Sc ratio (0.43) which is indicating the input from a juvenile source. A ƒSm/Nd of − 0.34 tends to indicate that the Sm–Nd system from this sample is not fractionated.

(Fig. 5b and c), whereas the Th/Sc ratios include also lower values. Cr/ V and Y/Ni ratios suggest the influence of a mafic source. For both units the depleted component could be represented by the detrital spinels (Abre, 2007; Abre et al., 2009). REE patterns of the Pavón Formation are also similar to those from the Ponón Trehué Formation (Fig. 5d). The Pavón Formation shows similar εNd (t), ƒSm/Nd and TDM values (Cingolani et al., 2003) compared with the Ponón Trehué Formation (Fig. 6), besides a juvenile input represented by a εNd value of −0.4 (Fig. 6a). Pb data from the Pavón Formation are not available. The Pavón Formation foreland basin resulted from the accretion of the Cuyania terrane to Gondwana, and received a main sedimentary input from the local Grenvillian basement known as the Cerro La Ventana Formation (Cingolani et al., 2003). The range of variation of εNd values of the Cerro La Ventana Formation are in the range of variation of data calculated to the time of deposition of the Pavón Formation (Fig. 7). The exemption is the − 0.4 value (Fig. 7), implying contributions from other sources besides the Cerro La Ventana Formation. Geochemistry (Cr/V and Y/Ni) and the presence of detrital

U–Pb dating of detrital zircons is a powerful tool that allows the identification of mainly felsic to intermediate provenance components in clastic sedimentary rocks (e.g., Fedo et al., 2003; Veevers and Saeed, 2009; Kuznetsov et al., 2010). Zircons were obtained for the Ponón Trehué (sample CT1; Fig. 3) and the Pavón Formations (sample QMOTO1; map on Cingolani et al., 2003). Detrital zircons with Th/U ratios indicative of a magmatic origin (Th/U ratio more than 0.2; Vavra et al., 1999; Hoskin and Schaltegger, 2003) were observed for both units, with only a very few exceptions of metamorphic derived zircon grains (Tables 3 and 4). Cathodoluminescence images from both units show that most of the zircon grains are subhedral and display oscillatory magmatic zoning, whereas only a few have patchy metamorphic zoning (e.g. the zircon C-III-85 from Fig. 9a). Detrital zircon ages of the Ponón Trehué Formation (n = 38) display a main probability peak at 1213.4 Ma, and clusters at 1164 and 1066 Ma; only one discordant grain has a younger age of 834 Ma (Fig. 9a). The very narrow range of detrital zircon ages implies a restricted provenance. The low Zr/Sc ratio of that specific sample, which is the less recycled sandstone from this unit, implies a local source. The detrital zircon dating of the Pavón Formation (n = 51) indicates a main probability peak at 1106.3 Ma and clusters at 1058.8 and 1369 Ma, whereas two grains are Neoproterozoic (634 and 615 Ma) and one grain has a Palaeoproterozoic age of 1652 Ma (Fig. 9b). These detrital zircon ages indicate sources of wider age range compared to the Ponón Trehué Formation source. Upper Ordovician sandstones of the Cuyania terrane showed a broader U–Pb detrital zircon spectrum with dominant peaks between 1.0 and 1.5 Ga, although the 1.2 Ga peak present in the Ponón Trehué Formation is not evident (Gleason et al., 2007). Noteworthy is the 1.4 Ga peak found by these authors in the Estancia San Isidro Formation (Middle Ordovician; Precordillera s.st.) which matches a similar cluster observed in the Pavón Formation. The 1.4 Ga peak is also found in the Cambrian La Laja and Cerro Totora Formations (Cuyania terrane), but these units completely lacks of 1.0 to 1.2 Ga zircons (Thomas et al., 2004; Finney et al., 2005). Minor population of 600–700 Ma detrital zircons is also referred by Finney (2007) and Gleason et al. (2007) in Upper Ordovician siliciclastics of the Cuyania terrane. The detrital zircon population could comprise zircon grains provided by recycling of older sedimentary sequences of the Cuyania terrane, particularly when detritus had been deposited by turbidite currents. Mesoproterozoic rocks that could have contributed to the more important cluster of detrital zircons are present in several neighbouring areas. Mesoproterozoic ages within the basement of the Cuyania terrane are found at: the Cerro La Ventana Formation, San Rafael block (1.1–1.2 Ga; Cingolani and Varela, 1999; Cingolani et al., 2005), the Pie de Palo Range (1.0–1.2 Ga; McDonough et al., 1993) and the Umango, Maz and Espinal Ranges (1.0–1.2 Ga; Varela and Dalla Salda, 1992; Varela et al., 1996; Casquet et al., 2006; Rapela et al., 2010). The

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Fig. 7. εNd versus age of the Ponón Trehué and Pavón (data from Cingolani et al., 2003) Formations and of areas evaluated as probable sources. CHUR: Chondritic Uniform Reservoir; U: Umango Range. A: granulitic and amphibolitic, and G: acidic-gneissic (εNd (t) = + 7), xenoliths from the inferred basement of the Cuyania terrane (Kay et al., 1996).

Arequipa–Antofalla Basement shows ages between 0.97 and 1.6 Ga (Bahlburg and Hervé, 1997 and references therein). The Amazon craton (1.0–1.55 Ga; Loewy et al., 2004; Schwartz and Gromet 2004 and references therein), the Grenville Province of Laurentia (0.9– 1.3 Ga; summary from Carrigan et al., 2003) and the basement rocks of the Chilenia terrane (Ramos and Basei, 1997) also display Mesoproterozoic ages. Palaeoproterozoic rocks are also found in the Arequipa–Antofalla Basement (1.9–2.0 Ga; Bahlburg and Hervé, 1997 and references therein), the Amazon craton (1.8–1.95 Ga; see Loewy et al., 2004; Schwartz and Gromet 2004 and references therein), within Laurentia (1.6–1.8 Ga; summary from Carrigan et al., 2003) but also at certain areas of the Western Pampeanas Ranges (1.8–1.9 Ga; Maz Range; Casquet et al., 2006). Some of these areas can be ruled out as sources based on Sm–Nd, Pb–Pb and palaeocurrents data as explained in the discussion. The Neoproterozoic zircons could be linked mainly to the Pampean/Brazilian Orogen. However, zircons in the 600–700 Ma age range are recorded in Cuyania Palaeozoic sedimentary units (Finney, 2007; Gleason et al., 2007). Noteworthy is the absence of detrital zircons of the Famatinian cycle (Upper Cambrian–Lower Devonian). The Mesoproterozoic detrital zircon ages found are consistent with recycled portions of the Pie de Palo basement (Finney, 2007; Gleason et al., 2007 and references therein), thought to represent exposed basement of the Cuyania terrane north of the San Rafael block (Fig. 1). 8. Discussion 8.1. Provenance indicators and source rocks

Fig. 8. a) 207Pb/204Pb versus 206Pb/204Pb present-day ratios and b) 208Pb/204Pb versus 206 Pb/204Pb present-day ratios. Solid circles represent samples from the Ponón Trehué Formation; SK: Stacey and Kramers reference line.

Petrographical, geochemical and isotopic analyses show that the source rocks for the Ordovician sedimentary sequences of the San Rafael block have a dominant unrecycled UCC composition, but an input from a less fractionated component is also evident. The main geochemical differences between the Ponón Trehué and the Pavón Formations are the general wider variability of geochemical proxies for the later (Fig. 5), indicating a broader range of provenance composition. Zircon age spectra for both units (Fig. 9) constrain the age of the main sources to the Mesoproterozoic. However, the Pavón

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Table 3 Detrital zircon dating of the Ponón Trehué Formation. Isotopic ratios corrected for fractionation by comparison with standard GJ-1. Concordant ages calculated using Isoplot/Ex (Ludwig, 2001); in italics, zircon spots for which the 207Pb/206Pb was used. Errors are in 1-sigma for ratios (in %) and ages (absolute). Isotopic ratios Sample

232

Th/238U

A-I-01 A-I-01b A-I-03 A-I-02a A-I-03 A-I-05 A-06 A-I-10 A-I-12 A-I-15 A-I-19 B-II-25 B-II-29 B-II-33 B-II-34 B-II-40 B-II-43 B-II-48 B-II-54 B-II-55 B-II-56 C-III-63 B-II-61 B-II-65 C-III-66 C-III-69 B-II-75 C-III-73 B-II-85 D-IV-109a D-IV-109b D-IV-112 D-IV-116 D-IV-118 D-IV-120 D-IV-126 D-IV-131 D-IV-115

0.49 0.45 0.48 0.33 0.16 0.50 0.41 0.31 0.16 0.16 0.48 0.26 0.49 0.40 0.47 0.42 0.49 0.23 0.47 0.27 0.34 0.15 0.17 0.37 0.49 0.25 0.34 0.37 0.36 0.45 0.47 0.32 0.53 0.34 0.37 0.58 0.76 0.43

207

Pb/235U

1.99744 2.19197 2.13908 1.34998 2.05176 2.26971 2.17768 2.33728 2.27266 2.29091 2.33586 2.06161 2.33317 2.49116 2.28317 2.24650 2.32042 2.29776 2.20956 2.25312 2.19581 2.11401 2.17676 1.97031 2.10575 1.92456 2.21532 2.21646 2.16012 1.82395 2.05928 2.07703 1.87917 2.11591 1.52118 2.14267 2.17900 2.12890

Age (Ma) ±

206

Pb/238U

2.52 1.55 1.97 4.87 1.64 1.04 2.00 2.38 2.60 2.33 1.90 2.12 2.16 1.20 2.09 1.84 1.84 1.70 1.06 1.45 1.51 3.82 2.36 1.90 1.44 2.61 4.84 4.21 1.69 5.42 1.98 3.32 4.81 2.07 3.84 1.66 1.83 2.51

0.17479 0.19149 0.19915 0.13708 0.18954 0.20442 0.19888 0.21279 0.20182 0.20389 0.21036 0.19312 0.20909 0.22019 0.20436 0.20312 0.20716 0.20492 0.19828 0.20223 0.19473 0.18815 0.19250 0.18117 0.19080 0.16934 0.20399 0.20352 0.19643 0.16150 0.18466 0.18123 0.17799 0.20177 0.14503 0.20534 0.20721 0.19237

±

% disc.

ƒ206

206

2.20 1.01 1.39 2.64 1.44 0.43 1.60 2.30 1.92 1.86 1.63 1.85 0.60 0.73 1.79 1.63 1.31 1.06 0.82 0.97 0.46 2.35 2.28 1.25 1.15 2.27 4.59 4.02 1.03 4.22 1.66 2.32 1.22 1.43 3.47 1.03 1.10 2.28

18 11 −2 15 4 1 1 −5 4 3 −2 −1 0 −3 2 1 1 2 4 2 8 10 9 8 6 20 −3 −2 3 22 10 16 5 −8 20 − 11 − 10 6

0.001 0.001 0.001 0.102 0.006 0.001 0.001 0.000 0.003 0.003 0.001 0.000 0.005 0.012 0.000 0.000 0.014 0.009 0.001 0.001 0.000 0.009 0.002 0.008 0.001 0.011 0.001 0.001 0.003 0.024 0.007 0.024 0.058 0.000 0.007 0.000 0.003 0.000

1038 1129 1171 828 1119 1199 1169 1244 1185 1196 1231 1138 1224 1283 1199 1192 1214 1202 1166 1187 1147 1111 1135 1073 1126 1008 1197 1194 1156 965 1092 1074 1056 1185 873 1204 1214 1134

Formation shows two peaks, one at 1.1 Ga and another at 1.4 Ga (with minor contributions from Palaeoproterozoic and Neoproterozoic sources), whereas the Ponón Trehué Formation shows a main peak at 1.2 Ga. Further constraints are provided by sedimentologic characteristics which indicates for the Ponón Trehué Formation a dominant provenance from the underlying Mesoproterozoic Cerro La Ventana Formation (Heredia, 2006), while palaeocurrents point to an eastern provenance for the detritus of the Pavón Formation invalidating western sources such as the Chilenia terrane (Manassero et al., 1999; Cingolani et al., 2003). Several areas should be evaluated as sources for the Ponón Trehué and Pavón Formations in relation to the models proposed to explain the tectonic evolution of the Cuyania terrane: the Grenville Province of Laurentia, the Western Pampeanas Ranges, the basement of the Cuyania terrane, the Famatinian arc, the Arequipa–Antofalla Basement (Central Andes), Antarctica, Falklands/Malvinas Microplate and the Natal–Namaqua Metamorphic belt (e.g. Ramos et al., 1986; Dalla Salda et al., 1992; Dalziel et al., 1994; Astini et al., 1995; Aceñolaza et al. 2002; Finney et al., 2005). The subduction towards east of the Iapetus crust beneath Gondwana resulted in developing the Famatinian magmatic arc (530–460 Ma; Pankhurst et al., 1998) and the docking of the Cuyania terrane. As documented in foreland basins elsewhere, a provenance from the arc should be expected (Miall, 2000). The lack of detrital zircons within the age range of the Famatinian magmatic arc (Fig. 9) indicates that this area was not a source for the Ponón Trehué and

Pb/238U

Concordant age ±

207

Pb/235U

23 11 16 22 16 5 19 29 23 22 20 21 7 9 21 19 16 13 10 12 5 26 26 13 13 23 55 48 12 41 18 25 13 17 30 12 13 26

1115 1178 1162 868 1133 1203 1174 1224 1204 1209 1223 1136 1222 1269 1207 1196 1219 1212 1184 1198 1180 1153 1174 1105 1151 1090 1186 1186 1168 1054 1135 1141 1074 1154 939 1163 1174 1158

±

207

Pb/206Pb

28 18 23 42 19 12 23 29 31 28 23 24 26 15 25 22 22 21 13 17 18 44 28 21 17 28 57 50 20 57 23 38 52 24 36 19 21 29

1266 1270 1144 970 1160 1210 1183 1189 1238 1233 1210 1132 1220 1247 1222 1202 1227 1229 1217 1217 1240 1233 1246 1169 1198 1256 1166 1172 1191 1243 1219 1272 1110 1097 1097 1087 1102 1203

±

Ma

15 15 16 40 9 11 14 7 22 17 12 12 25 12 13 10 16 16 8 13 18 37 8 17 10 16 18 15 16 42 13 30 52 16 18 14 16 13

1266 ± 48 1270 ± 45 1165 ± 13 834 ± 20 1138 ± 11 1199 ± 4.6 1174 ± 14 1199 ± 11 1199 ± 18 1209 ± 16 1221 ± 13 1135 ± 14 1224 ± 7 1277 ± 8 1209 ± 15 1197 ± 12 1217 ± 13 1206 ± 11 1217 ± 27 1193 ± 9.5 1149 ± 100 1126 ± 23 1246 ± 25 1169 ± 59 1200 ± 17 1256 ± 59 1175 ± 26 1177 ± 22 1161 ± 10 1243 ± 100 1241 ± 44 1272 ± 94 1065 ± 12 1097 ± 60 1097 ± 69 1170 ± 100 1184 ± 100 1203 ± 46

Pavón Formations. The scarcity of Famatinian aged detrital zircons within the Ordovician record of the Cuyania terrane could indicate that the Famatinian magmatic arc was not related to the docking of the Cuyania terrane (Finney et al., 2005). However, the presence of a positive area acting as a barrier would have been enough to prevent those detritus to reach the Ordovician basin (Abre, 2007). As it is shown on Figs. 7 and 8 rocks from Antarctica, Falklands/ Malvinas Microplate and the Natal–Namaqua Metamorphic belt accomplish for Mesoproterozoic ages and have comparable TDM ages, but have different Pb-isotope compositions (Wareham et al., 1998). These differences allow discarding such areas (named SAFRAN in the para-autochthonous model; Aceñolaza et al. 2002; Finney et al., 2005) as sources. Several models assigned a Laurentian (Southern Appalachian) origin for the Cuyania terrane (e.g. Ramos et al., 1986; Dalla Salda et al., 1992; Dalziel et al., 1994; Astini et al., 1995; Thomas and Astini, 1996; Keller, 1999). The time of detaching from Laurentia varies according to different authors, but a link to Laurentia is considered even until the end of the Ordovician (e.g. Thomas and Astini, 2003). If the Cuyania terrane was attached to Laurentia during the Ordovician, then a Laurentian signature could be recorded within (and/or reworked into) the sedimentary sequences, as suggested for the Cambrian clastic record (Naipauer, 2007; Naipauer et al., 2010). The Appalachian belt comprises Grenvillian-age igneous-metamorphic rocks (and minor Palaeoproterozoic and Neoproterozoic rocks), and since it was elevated during most of the Palaeozoic, it acted as a major source (Boghossian et al., 1996). Even though these rocks could

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Table 4 Detrital zircon dating from the Pavón Formation. Isotopic ratios corrected for fractionation by comparison with standard GJ-1. Concordant ages calculated using Isoplot/Ex (Ludwig, 2001); in italics, zircon spots for which the 207Pb/206Pb were used. Errors are in 1-sigma for ratios (in %) and ages (absolute). Isotopic ratios Sample

232

Th/238U

G-VII-458 G-VII-456 G-VII-454 G-VII-449 G-VII-446 G-VII-439 G-VII-437 G-VII-436a G-VII-436b G-VII-435 G-VII-434 G-VII-432 G-VII-431 G-VII-429 G-VII-428 G-VII-426 G-VII-424 G-VII-423 G-VII-417 G-VII-415 G-VII-414 G-VII-413 G-VII-412 G-VII-411 G-VII-409 G-VII-404 G-VII-408 F-VI-402 F-VI-401 F-VI-400 F-VI-397 F-VI-396 F-VI-395 F-VI-393 F-VI-391 F-VI-390 F-VI-389 F-VI-387 F-VI-385 F-VI-384 F-VI-382 F-VI-379 F-VI-376 F-VI-375 F-VI-372 F-VI-368 F-VI-363 F-VI-359 F-VI-358 F-VI-356 F-VI-355

0.58 0.66 0.44 0.78 0.44 0.67 0.86 0.26 0.50 0.90 0.68 0.47 0.61 0.51 0.29 0.89 0.45 0.72 0.40 0.44 0.43 0.40 0.71 0.69 0.41 0.45 0.96 0.59 1.46 0.87 0.51 0.49 0.46 0.55 0.32 0.56 0.42 0.66 0.55 0.49 0.51 0.44 0.46 0.67 0.97 0.36 0.50 0.02 0.52 0.39 0.29

207

Pb/235UU

1.81152 1.83079 1.98432 2.86412 2.90877 1.96370 2.09490 2.02791 2.03482 2.57362 1.78238 3.04117 1.87367 2.79883 1.82389 1.98214 2.90639 2.96160 1.87287 2.73001 1.95331 2.45500 2.35359 1.90578 2.16132 1.83630 2.07574 4.08749 1.87702 2.00281 1.90289 1.98944 1.92244 1.94213 1.95343 2.37360 1.82514 1.83003 2.00671 1.99576 0.87311 2.02631 1.93860 2.94807 1.87010 1.94540 2.73140 0.82952 1.95880 1.93119 1.87360

Age (Ma) ±

206

Pb/238

2.31 3.89 1.66 2.09 2.44 1.52 2.97 2.74 1.95 2.77 4.06 1.70 2.57 2.18 1.65 3.10 2.81 2.52 2.61 2.63 1.99 2.61 1.64 1.96 2.70 2.58 2.36 1.89 2.48 2.60 2.04 2.14 2.48 2.60 2.64 3.01 1.93 1.75 2.54 2.23 2.50 1.53 2.50 1.62 3.22 1.82 2.34 2.59 2.52 3.42 2.94

0.17440 0.17622 0.18713 0.23598 0.24167 0.18602 0.19529 0.19107 0.19228 0.21294 0.17078 0.25658 0.18004 0.23342 0.17585 0.16918 0.24215 0.24593 0.17763 0.22594 0.18756 0.21043 0.20218 0.17675 0.20232 0.17594 0.19212 0.29270 0.17151 0.18453 0.18074 0.18645 0.18328 0.18404 0.18551 0.21921 0.17191 0.17121 0.19225 0.18650 0.10328 0.19214 0.18484 0.23155 0.18052 0.18206 0.23087 0.10018 0.18562 0.18232 0.18019

206

±

% disc.

ƒ206

1.88 1.74 1.17 1.38 1.95 1.12 1.92 1.84 0.89 1.71 2.72 1.37 1.73 1.31 1.16 0.84 1.47 1.39 1.46 1.76 1.11 1.51 0.76 1.19 1.56 1.40 1.55 1.35 1.99 2.17 1.67 1.27 1.24 0.93 1.19 2.51 1.17 1.25 1.73 1.30 1.45 1.11 1.41 1.38 1.25 1.17 1.37 1.73 1.77 1.89 2.05

4 3 1 1 −2 1 −1 −1 −2 9 6 − 10 1 1 3 23 −3 −4 5 4 −2 6 9 9 −5 4 2 0 14 6 3 3 1 2 1 − 10 9 10 −4 3 3 −2 0 9 0 5 0 −2 1 3 1

0.0003 0.0046 0.0013 0.0005 0.0003 0.0000 0.0008 0.0013 0.0009 0.0009 0.0015 0.0012 0.0007 0.0004 0.0005 0.0123 0.0008 0.0009 0.0008 0.0005 0.0004 0.0009 0.0001 0.0026 0.0012 0.0009 0.0005 0.0016 0.0014 0.0014 0.0001 0.0016 0.0007 0.0015 0.0003 0.0009 0.0005 0.0029 0.0012 0.0015 0.0003 0.0003 0.0005 0.0004 0.0014 0.0011 0.0006 0.0006 0.0008 0.0010 0.0007

account for the detrital zircon ages and the Sm–Nd signature (Patchett and Ruiz, 1989) of the clastic rocks here studied, the Pb isotopes do not show a Laurentian signature (Fig. 8; Kay et al., 1996). Since the basement known as the Cerro La Ventana Formation was already uplifted by the Darriwilian (Heredia, 2006) such a provenance is very likely, particularly when analyzing the Ponón Trehué Formation. The Cerro La Ventana Formation consists of mafic to intermediate gneisses grading to amphibolites and migmatites, as well as acidic to intermediate granitoids (Cingolani et al., 2005). Sm– Nd, Rb–Sr and U–Pb on zircons indicate Mesoproterozoic ages (1.1 to 1.2 Ga) and geochemical and isotopic characteristics of mantlederived rocks (Cingolani and Varela, 1999; Thomas et al., 2001; Cingolani et al., 2005). Although only some of all the basement rock types had been studied until today (diorites and tonalites), they provide the best fit regarding sedimentology (Heredia, 2006), Nd signature and zircon ages. Nd data show εNd values in the range of variation of data calculated to the time of deposition of the Ponón

Concordant age

Pb/238U

±

207

Pb/235U

1036 1046 1106 1366 1395 1100 1150 1127 1134 1244 1016 1472 1067 1352 1044 1008 1398 1417 1054 1313 1108 1231 1187 1049 1188 1045 1133 1655 1020 1092 1071 1102 1085 1089 1097 1278 1023 1019 1134 1102 634 1133 1093 1343 1070 1078 1339 615 1098 1080 1068

19 18 13 19 27 12 22 21 10 21 28 20 18 18 12 8 21 20 15 23 12 19 9 12 19 15 18 22 20 24 18 14 13 10 13 32 12 13 20 14 9 13 15 19 13 13 18 11 19 20 22

1050 1057 1110 1373 1384 1103 1147 1125 1127 1293 1039 1418 1072 1355 1054 1109 1384 1398 1072 1337 1100 1259 1229 1083 1169 1059 1141 1652 1073 1116 1082 1112 1089 1096 1100 1235 1055 1056 1118 1114 637 1124 1095 1394 1071 1097 1337 613 1101 1092 1072

±

207

Pb/206Pb

24 41 18 29 34 17 34 31 22 36 42 24 28 30 17 34 39 35 28 35 22 33 20 21 32 27 27 31 27 29 22 24 27 28 29 37 20 18 28 25 16 17 27 23 34 20 31 16 28 37 32

1077 1078 1119 1383 1367 1110 1142 1121 1115 1375 1087 1337 1081 1360 1075 1315 1362 1368 1107 1374 1083 1307 1302 1152 1134 1087 1156 1648 1181 1165 1105 1131 1097 1109 1105 1160 1121 1135 1087 1137 650 1108 1097 1474 1072 1134 1334 605 1109 1117 1080

±

Ma

15 37 13 22 20 11 26 23 19 30 33 13 20 24 13 39 33 29 24 27 18 28 19 18 25 24 21 22 17 17 13 20 24 27 26 19 17 14 20 21 13 12 23 13 32 16 25 12 20 32 23

1049 ± 15 1048 ± 16 1108 ± 11 1375 ± 12 1375 ± 12 1090 ± 7.7 1149 ± 18 1126 ± 17 1132 ± 8.9 1356 ± 34 1087 ± 40 1356 ± 34 1070 ± 16 1354 ± 14 1050 ± 10 1361 ± 66 1393 ± 17 1409 ± 16 1059 ± 14 1327 ± 19 1105 ± 11 1304 ± 45 1304 ± 45 1167 ± 35 1180 ± 16 1048 ± 13 1137 ± 14 1652 ± 15 1167 ± 35 1167 ± 35 1082 ± 14 1106 ± 12 1086 ± 12 1090 ± 9.2 1097 ± 12 1126 ± 32 1126 ± 32 1126 ± 32 1124 ± 16 1106 ± 12 634 ± 8.6 1127 ± 10 1094 ± 13 1474 ± 32 1070 ± 12 1086 ± 11 1338 ± 15 615 ± 9.9 1100 ± 16 1083 ± 18 1070 ± 18

Trehué Formation (Fig. 7; Cingolani et al., 2005). Unfortunately, the Pb composition of the Cerro La Ventana Formation remains unknown but samples are in progress. The relative wider range of detrital zircon ages and εNd values of the Pavón Formation might indicate that although the Cerro La Ventana Formation could have been an important source, another source was also providing detritus to the basin. Source areas towards the east are likely (Manassero et al., 1999; Cingolani et al., 2003), but no outcrops are known. Instead, similar Mesoproterozoic basement rocks are recorded in the northern outcrops of the Western Pampeanas Ranges (e.g. Pie de Palo, Varela and Dalla Salda, 1992; McDonough et al., 1993; Varela et al., 1996; Casquet et al., 2006). Zircon age dating from the clastic record agrees with the ages found within the Western Pampeanas Ranges. Sm–Nd data available from the Western Pampeanas Ranges are scarce, but particularly noteworthy are the positive εNd values found within the Umango Range (Fig. 7; Porcher et al., 2004, Vujovich et al., 2005). The TDM ages of the Pavón

P. Abre et al. / Gondwana Research 19 (2011) 275–290

287

Fig. 9. U–Pb distribution of analyzed detrital zircons with probability curves and concordia plot diagrams for the a) Ponón Trehué (n = 38) and b) Pavón Formations (n = 51). Histogram bars represent time intervals of 40 Ma. Isoplot/Ex (Ludwig, 2001) was used for age calculations. Representative CL microphotographs of selected zircon grains used for detrital dating show the predominance of magmatic internal textures. Bar length is 100 µm.

Formation are comparable to TDM ages for Mesoproterozoic xenoliths interpreted as the basement of the Cuyania terrane (Kay et al., 1996). The Western Pampeanas Ranges and the Arequipa–Antofalla Basement might constitute a single crustal block para-autochthonous with respect to the pre-Famatinian margin of Gondwana (Casquet et al., 2006), then a minor contribution from the Central and Southern Domains can also be invoked to explain the Sm–Nd signature of the Pavón Formation (Fig. 7). However, a northern provenance is not supported by palaeocurrents, and immaturity of the sedimentary rocks including the subhedral character of the zircons tends to dismiss long transports. 8.2. Tectonic implications The evidence here presented undoubtedly link the Mesoproterozoic Cerro La Ventana Formation as a provenance component to the Ordovician sedimentary deposition within the San Rafael block. However, extensive petrographical, geochemical and isotopic studies of all the rock types comprised in the Cerro La Ventana Formation are needed to further support this. Considering the current information available, another eastern source such as the Western Pampeanas Ranges has to be invoked to understand the provenance of the Pavón Formation. These differences might indicate that even though the Cerro La Ventana Formation was a source since the Darriwilian providing detritus to a restricted extensional basin, towards younger ages (Sandbian) other sources such as the Western Pampeanas Ranges were available to provide detritus to the Pavón Formation foreland basin (Fig. 10). Although certain parts of the Western Pampeanas Ranges (Umango Range) have a Laurentian signature and are therefore

part of the Cuyania basement, other areas (Maz and Espinal ranges) are interpreted as representing the active Gondwana margin during the Lower Palaeozoic (Porcher et al., 2004). Moreover, the Pie de Palo Range is considered autochthonous to Gondwana prior to the Lower Palaeozoic (Casquet et al., 2006). In this scenario, the depositional basins resulted from the extensional regime (Astini, 2002; Cingolani et al., 2003) that followed the accretion during the Middle Ordovician of the Cuyania terrane to Gondwana as depicted on Fig. 10. A para-autochthonous or allochthonous origin for the Cuyania terrane cannot be resolved, until more data from the probable sources are added. The isotopic signatures from source rocks proposed in the para-autochthonous to Gondwana models are not imprinted within the sedimentary successions here studied. Because a Laurentian derivation for the Neoproterozoic to Cambrian successions of the Cuyania terrane was deduced (Astini et al., 1995; Naipauer, 2007; Naipauer et al., 2010), it could be expected that the same rocks could have provided detritus to the Ordovician record. Typical Laurentian signatures however are not shown by the provenance indicators applied to the Ordovician successions of the San Rafael block, particularly concerning Pb–Pb composition (Figs. 7 and 8). Considering that the Cerro La Ventana Formation and the Western Pampeanas Ranges are the most probable source areas (Fig. 10), they would have been uplifted since at least the Darriwilian in order to provide detritus to the Ponón Trehué and Pavón Formations. If the Western Pampeanas Ranges were para-autochthonous to Gondwana (Casquet et al., 2006), then the Cuyania terrane might have collided at least immediately before the beginning of the Ordovician clastic deposition (Darriwilian; Fig. 10). In such geotectonic scenario, the Western Pampeanas Ranges not only provided detritus to the basin

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Fig. 10. Interpretative schematic cross sections showing the geotectonic evolution of the Cuyania terrane. a) Sedimentary record of Cambrian–Early Ordovician age is not recorded within the San Rafael block; the Cuyania terrane was approaching Gondwana as a result of the closure of the Iapetus Ocean. The Famatinian arc was active and volcaniclastic sediments were deposited (Pankhurst et al., 1998). b) The Ponón Trehué Formation received an input restricted to the Cerro La Ventana Formation; it is the only sedimentary unit from the Cuyania terrane in contact with the basement. c) Progressive subsidence of the basin; the uplift of the Western Pampeanas Ranges provided other eastern source for the Pavón Formation (besides the Cerro La Ventana Fm.) and prevented input of detrital material from the Famatinian arc. The basins were generated as a response of the extension that followed the accretion of the Cuyania terrane to Gondwana (Astini, 2002; Cingolani et al., 2003; Heredia, 2006). The Ponón Trehué and the Pavón Formations are not in contact.

(Pavón Formation), but acting as a positive area prevented that detrital material from the Famatinian arc reached the basin. 9. Conclusions 1) Petrography, geochemistry and isotope geochemistry indicate that the Ponón Trehué Formation had a dominantly local, basementderived provenance that is consistent with all that is known of the

Mesoproterozoic Cerro La Ventana Formation, with which it is in contact. 2) The Pavón Formation shows similar but wider ranges of variation of all the provenance indicators, implying that it received more regional sediment dispersal. Particularly, Sm–Nd isotopes and detrital zircon dating indicate external to the San Rafael block sources (besides the Cerro La Ventana Formation), most probably from the nearby Western Pampeanas Ranges.

P. Abre et al. / Gondwana Research 19 (2011) 275–290

3) The lack of Famatinian detrital zircons observed for both units is explained by the presence of a positive area that not only prevented the arrival of arc-derived detritus but has also acted as source (Western Pampeanas Ranges). 4) The Cuyania terrane might have accreted to Gondwana before the deposition of the units here studied (Darriwilian), in order to receive the input from the Western Pampeanas Ranges.

Acknowledgements P. Abre thanks the Faculty of Sciences (University of Johannesburg) for the financial support and G. Blanco for extensive discussions. Fieldwork was financed by CONICET and ANPCYT (PICT 07-10829, Argentina). Sm–Nd and Pb–Pb data were possible thanks to the Postgraduate students grant system of the International Association of Sedimentologists and the staff of the LGI-UFRGS (Brazil), particularly Prof. Kawashita, K. and Prof. Dussin, I. Zircon dating was financed by the National Research Foundation (NRF), South Africa. Any opinion, findings and conclusions or recommendations expressed in this material are those of the authors and therefore the NRF does not accept any liability in regard thereto. J. Maggi, N. Uriz, P. Frigerio and C. Pereyra (La Plata, Argentina) are thanked for their helpfulness with XRD data, fieldworks and laboratory assistance. Special thanks to S. Heredia (San Juan, Argentina) for discussions with one of us (C.A.C.) on Ponón Trehué geology. Extensive revision by J.D. Gleason, P. Mueller and Editor E. Tohver helped to improve strongly this manuscript and is deeply acknowledged. References Abre, P., 2007. Provenance of Ordovician to Silurian clastic rocks of the Argentinean Precordillera and its geotectonic implications. PhD Thesis. University of Johannesburg, South Africa. Unpublished. Abre, P., Cingolani, C., Zimmermann, U., Cairncross, B., 2009. Detrital chromian spinels form Upper Ordovician deposits in the Precordillera terrane, Argentine: a mafic crust input. Journal of South America Earth Sciences Special Issue on Mafic and Ultramafic complexes in South America and the Caribbean10.1016/j. jsames.2009.04.005. Aceñolaza, F.G., Miller, H., Toselli, A.J., 2002. Proterozoic–Early Paleozoic evolution in western South America — a discussion. Tectonophysics 354, 121–137. Aitcheson, S.J., Harmon, R.S., Moorbath, S., Schneider, A., Soler, P., Soria-Escalante, E., Steele, G., Swainbank, I., Wörner, G., 1995. Pb isotopes define basement domains of the Altiplano, central Andes. Geology 23, 555–558. Astini, R.A., 2002. Los conglomerados basales del Ordovícico de Ponón Trehué (Mendoza) y su significado en la historia sedimentaria del terreno exótico de Precordillera. Revista de la Asociación Geológica Argentina 57, 19–34. Astini, R.A., Dávila, F.M., 2004. Ordovician back arc foreland and Ocloyic thrust belt development on the Western Gondwana margin as a response to Precordillera terrane accretion. Tectonics 23, TC400810.1029/2003TC001620. Astini, R.A., Benedetto, J.L., Vaccari, N.E., 1995. The Early Paleozoic evolution of the Argentine Precordillera as a Laurentian rifted, drifted and collided terrane: a geodynamic model. Geological Society of America Bulletin 107, 253–273. Bahlburg, H., Hervé, F., 1997. Geodynamic evolution and tectonostratigraphic terranes of northwestern Argentina and northern Chile. Geological Society of America Bulletin 109, 869–884. Bahlburg, H., Hervé, F., Bock, B., 2001. New Nd isotope data on the ?Devonian Chilenia terrane in northern Chile: implications for the Paleozoic accretionary history of the Andean Gondwana margin. XI Congreso Latinoamericano de Geología y III Congreso Uruguayo de Geología, p. 225. Benedetto, J.L., 1993. La hipótesis de la aloctonía de la Precordillera argentina: un test estratigráfico y biogeográfico. XII Congreso Geológico Argentino y II Congreso de Exploración de Hidrocarburos, Actas III, pp. 375–384. Beresi, M.S., Heredia, S., 2003. First occurrence of Epiphyton, cyanobacteria from the Middle Ordovician of the Ponón Trehué, Mendoza province, Argentina. In: Albanesi, G.L., Beresi, M.S., Peralta, S.H. (Eds.), Ordovician from the Andes: INSUGEO, Serie Correlación Geológica, 17, pp. 257–261. Bock, B., McLennan, S.M., Hanson, G.N., 1998. Geochemistry and provenance of the Middle Ordovician Austin Glen Member (Normanskill Formation) and the Taconian Orogeny in New England. Sedimentology 45, 635–655. Bock, B., Bahlburg, H., Wörner, G., Zimmermann, U., 2000. Tracing crustal evolution in the Southern Central Andes from Late Precambrian to Permian with geochemical and Nd and Pb isotope data. Journal of Geology 108, 515–535. Boghossian, N.D., Patchett, P.J., Ross, G.M., Gehrels, G.E., 1996. Nd isotopes and the source of sediments in the Miogeocline of the Canadian Cordillera. Journal of Geology 104, 259–277.

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