Andean Geology 37 (2): 375-397. July, 2010
formerly Revista Geológica de Chile
www.scielo.cl/andeol.htm

 

Permian depositional age of metaturbidites of the Duque de York Complex, southern Chile: U-Pb SHRIMP data and palynology

Edad pérmica de sedimentación de las metaturbiditas del Complejo Duque de York, sur de Chile: datos mediante U-Pb SHRIMP y palinología

 

Fernando A. Sepúlveda1, Sylvia Palma-Heldt2, Francisco Hervé3, C. Mark Fanning4

1  Servicio Nacional de Geología y Minería, Av. Santa María 0104, Santiago, Chile. fsepulveda@sernageomin.cl; fersepul@ing.uchile.cl
2  Departamento de Ciencias de la Tierra, Universidad de Concepción, Barrio Universitario S/N°, Concepción, Chile. sypalma@udec.cl
3  Departamento de Geología, Universidad de Chile, Casilla 13518, Correo 21, Santiago, Chile. fherve@ing.uchile.cl
4  Research School of Earth Sciences, The Australian National University, Milis Road, Canberra, ACT 0200, Australia. Mark.Fanning@anu.edu.au

ABSTRACT. The Duque de York Complex (DYC) is part of the low grade metamorphic accretionary complexes of the pre-Andean Patagonian 'basement'. It is a sedimentary succession exposed along the western margin of southernmost South America. New U-Pb zircon ages and palynological data restrict the maximum depositional age of the DYC to the limit between the early Permian (Kungurian) and the middle Permian (Roadian). The palynological association recorded in the DYC, characterized mainly by Gymnospermopsida pollen, indicates a humid environment of forest with an under-growth of ferns. Regional paleogeographic correlations point out that an interpretation of DYC as an autochthonous terrane cannot be discarded, contrasting with previous hypotheses which suggest an allochthonous character for this complex.

Keywords: Palynology, U-Pb dating, Duque de York Complex, Terrones, Gondwana, Chile.

RESUMEN. El Complejo Duque de York (CDY) forma parte de los complejos metamórficos acrecionarios del 'basamento' pre-Andino de la Patagonia, correspondiendo a una sucesión sedimentaria que añora a lo largo del margen occidental austral de Sudamérica. Nuevas edades U-Pb en circón, en combinación con información palinológica, permiten acotar la maxima edad de depósito posible del CDY al límite entre el Pérmico temprano (Kunguriano) y el Pérmico medio (Roadiano). La asociación palinológica registrada en el CDY está caracterizada por Gymnospermopsida, e indica un ambiente húmedo de bosque con sotobosque de heléchos. Las correlaciones paleogeográñcas apuntan a que la condición de terreno autóctono del CDY no puede ser descartada, lo que se contrapone a hipótesis anteriores, las que sugieren un carácter alóctono para este complejo.

Palabras clave: Palinología, Dotación U-Pb, Complejo Duque de York, Terrenos, Gondwana, Chile.

1.      Introduction

The Duque de York Complex (DYC) is one of the metamorphic complexes that form the pre-Andean Patagonian 'basement' rocks, which crop out ex-tensively along the western edge of South America south of 50°S. These rocks have been generally considered as part of the late Paleozoic-early Me-sozoic accretionary prism built at the paleo-Pacific (Panthalassan) margin of Gondwana (e.g., Hervé et al., 1981; Forsythe, 1982). The accretionary orogenic belt that formed on this margin is one of the largest known orogenic belts in Earth history, and now oc-cupies the eastern third of Australia, New Zealand, West Antarctica, the Transantarctic Mountains and large parts of southern South America (Vaughan et al., 2005). This orogenic belt has been termed in two different ways depending on the interval considered: the Proterozoic and Paleozoic Térra Australis orogen (Cawood, 2005), and the Paleozoic and Mesozoic Australides (Vaughan et al., 2005). It has been re-garded as a collage of accreted terranes-terranes being fault-bounded packages of rocks of regional extent characterized by a geological history that differs from that of neighboring terranes (Howell et al., 1985; Vaughan et al., 2005).

The DYC corresponds to a widespread low grade-metasedimentary succession that crops out along the Madre de Dios and Diego de Almagro archipelagos (50°00'-51°50'S) and at the Ramírez, Contreras and Desolación islands (51°50'-53°00'S), in southern Chile (Fig. 1A). Forsythe andMpodozis (1979,1983) interpreted the DYC as a continent-derived detrital succession that was deposited over two coeval late Paleozoic exotic oceanic units, as they approached the continental margin of Gondwana. The DYC together with the oceanic units, defíned as the Madre de Dios Accretionary Complex (MDAC) by Thomson and Hervé (2002), were then tectonically amalgamated to the forearc of this margin by subduction processes, resulting in an intricate tectonic interweaving. These complexes were intruded by the South Patagonian Batholith (SPB) in the Early Cretaceous (Halpern, 1973; Duhart et al., 2003; Hervé et al., 2007a).


FIG. 1. A. Sketch map showing the distribution of the DYC; B. Geological map of the Madre de Dios Archipelago (after Forsythe and Mpodozis, 1983; Lacassie et al., 2006; Sepúlveda et al., 2008). 1. Quaternary deposits; 2. South Patagonian Batholith (SPB); 3. Tarlton Limestones (TL); 4. Denaro Complex (DC); 5. Duque de York Complex (DYC); 6. Unmapped basement; 7. SiU. Sampling sites are marked with black circles; all for palynology except FO04-21; FO04-21 and FO04-22 for U-Pb Shrimp dating.

For a long time, the accretionary complexes that compose the MDAC have been considered of exotic or allochthonous origin, or at least, as suspect terranes. Terranes are 'suspect' if there is doubt about their paleogeographical setting with respect to adjacent terranes or continental margin (Coney et al., 1980; Coombs, 1997), and may be described as 'exotic', 'far-travelled' or 'allochthonous' (all meaning about the same thing) if there is sufficient evidence that they originated far from their present locations, often assumed to be hundreds or thousands of kilometers away (Vaughan et al., 2005). The consideration of the MDAC as suspect and potentially exotic is based mainly on its fossil content and the inferred depositional setting for them (in the case of the TL and the coeval DC) (e.g., Ling et al., 1985; Ramos, 1988) and on the impossibility to find a contem-poraneous magmatic are as the source of Permian zircons for the DYC anywhere at a similar latitude in southern Patagonia (Hervé et al., 2003; Hervé and Mpodozis, 2005; Hervé et al., 2006). The latter has led to propose that the deposition of the DYC took place at high southern latitudes along the Antarctic sector of the Gondwana margin (Lacassie, 2003; Lacassie et al., 2006). However, derivation of the DYC from lower and warmer latitudes is a hypothesis that cannot be ruled-out.

The lack of index fossils in the DYC has prevented an aecurate determination of the depositional age of this complex. In consequence, a late early Permian maximum depositional age has been established by the use of the youngest detrital zircon population in these metasediments (Hervé et al., 2003). Nevertheless, the use of the youngest detrital zircon age or population in a sediment as a limit for the age of deposition has been questionable for both geological (there is no necessary connection between the timing of zircon-generating events in a source region and the age of final deposition of a sediment eroded from this source) and statistical reasons (Andersen, 2005). In this context, this paper presents the first palynological study in rocks of the DYC and it also corresponds to the first record of late Paleozoic palynomorphs in Chile. The aim of this work is to restrict the age range of the DYC by the combination of the palynological results with new U-Pb SHRIMP ages in key samples, and also gives a revisión of the paleoenvironmental and geochronological data regarding the place and timing of deposition of the DYC. The conclusions derived from this work al-low giving some new considerations focused on the supposed allochthonous character of the MDAC, and especially of the DYC.

2.      Geological background

The rocks of the studied area were first recognized by Cecioni (1955, 1956), who determined the presence of upper Paleozoic sediments in the Patagonian archipelago, distinguishing fusulinids as well as the low grade-metamorphic character of these rocks. The geology of the Madre de Dios Archipelago was studied in detail by Forsythe and Mpodozis (1979,1983) and Mpodozis and Forsythe (1983), who distinguished three metamorphic com-plexes that made up the MDAC (Fig. 1B):

    a. the Denaro Complex (DC), formed by tho-leiitic basalts with E- and N-MORB signatures (Hervé et al., 1999; Sepúlveda et al., 2008), banded radiolarian and metalliferous cherts, pelites and calcarenites. This complex represents fragments of ocean floor and its sedimentary cover (late Carboniferous-early Permian, according to Ling et al., 1985),

    b.  the Tarlton Limestones (TL), formed by fusulinid-bearing massive limestones, deposited in an intra-oceanic carbonate platform during Middle Pennsylvanian-early Permian times (Cecioni, 1956; Douglass andNestell, 1972, 1976),

    c. the Duque de York Complex (DYC), formed by a thick succession of greywackes, pelites, and minor conglomerates of continental provenance, and deposited on top of DC and TL. Based on field observation, Faúndez et al. (2002) described the DYC as being formed by metaturbidites. Also, it has been indicated that this succession has early Permian radiolarian cherts at Desolación Island (A. Yoshiaki, written communication, 2002; in Hervé et al., 2007b). Owing to the accretionary processes, most contacts among these units are of tectonic origin (Fig. 2), but few examples of depositional contacts nave been recognized. In fact, the preserved stratigraphic relations confirm that the DYC has both unconformable and conformable depositional contacts with the other complexes (e.g., Lacassie et al., 2006).


  FIG. 2. Outcrops at Madre de Dios Island, where tectonic contact (dashed line) between deformed metasediments of the DYC (brown) and massive limestones of the TL (white) is partly observed.

All these units where metamorphosed under the conditions of pumpellyite-actinolite facies in a frontal accretionary wedge (Sepúlveda, 2004; Sepúlveda et al., 2008) during Middle Triassic to earliest Jurassic times, as indicated by in situ Ar/Ar UV-LAMP ages on phengites (Willner et al., 2009). Thomson and Hervé (2002) used zircon fission track data to point out that the metamorphism that affects the DYC, and also the underlying TL and DC, took place during or before the earliest Jurassic (ca. 195 Ma). This information constrains the minimum probable age of deposition for the DYC, and demonstrates that, in this area, the metamorphism occurred prior to the emplacement of the SPB in the Early Cretaceous. The isotopic ages of the SPB, in the outcrops adjacent to the contact with the MDAC, are 133-112 Ma (Rb-Sr whole rock and biotite isochron, Halpern, 1973), 130-143 Ma (K-Ar biotite, Duhart et al., 2003) and ca. 133 Ma (U-Pb SHPJMP zircon, Hervé et al., 2007a).

U-Pb SHRIMP detrital zircon ages from sand-stones of the DYC reveal that the youngest main opulation, and henee the maximum possible depositional age, is late early Permian (ca. 270 Ma) (Hervé et al., 2003). The geochemical study of Lacassie et al. (2006), complementing the data and refming the conclusions of Faúndez et al. (2002), indicates that the DYC sandstones and mudstones had their source in a volcanic are of granodioritic average composition located relatively proximal to the depositional basin, and whose plutonic roots had been exposed by erosión. Also, they propose that the DYC was deposited in a tectonic setting corresponding to an active continental margin, possibly located along the Antarctic segment of the Panthalassan Gondwana margin.

2.1.      Paleogeographic setting

The fusulinid fauna in the TL shows that these carbonate sedimentary rocks must have been deposited in marine warm water (Douglass and Nestell, 1976). Similarities between the fossil content of the TL with those of the backarc marine carbonate deposits of the late Paleozoic Copacabana Formation in Perú and Bolivia (Cabrera La Rosa and Petersen, 1936; Chamot, 1965; Mamet, 1996), indicates that deposition of TL occurred in low latitude zones (ca. 20°S) during the late Carboniferous-early Permian (Lacassie, 2003). However, recent paleogeographic recon-structions for those periods (Torsvik and Cocks, 2004; Veevers, 2004; Cocks and Torsvik, 2006; Cawood and Buchan, 2007) locate the portion of the Gondwana margin where the Madre de Dios archipelago is presently situated at a high southern latitude, well outside the tropical zone where the TL is likely to have been deposited. Also, it is indicated that the late Paleozoic Ice Age in Gondwana was active between the Carboniferous and the early Permian (Isbell et al., 2003; Isbell et al., 2005; López-Gamundí, 2005; Buatois et al., 2006; Fielding et al., 2008; Rocha Campos et al., 2008). These faets, together with the contemporaneity of the TL with the ocean floor deposit of the DC (Ling et al., 1985; Ling and Forsythe, 1987) lead to the conclusión that the MDAC represents an allochthonous or exotic terrane derived from lower latitudes and accreted via subduction processes to Gondwana (Ramos, 1988). The timing of the accretion of these units would be bracketed between the maximum age of deposition of the DYC (ca. 270 Ma; Hervé et al., 2003) and the minimum age of metamorphism (195 Ma; Thomson and Hervé, 2002).

The apparent lack of a Permian magmatic are in southernmost Patagonia allowed Lacassie (2003) and Lacassie et al. (2006), following Hervé et al. (2000) and Cawood et al. (2002), to propose that the accretion of the TL and the DC would have oceurred against the Antarctic-Australian segment of the Gondwana margin, from where both would have been displaced by dextral translation, together with the DYC, as a coherent block to their current position. In addition, Lacassie et al. (2006) show that the DYC metasediments share important petrographic, geochemical and geochronological characteristics with metaturbidites present in the Rakaia Terrane in New Zealand and with the east-ern (Triassic) Le May Group in Alexander Island. These similarities point towards similar igneous sources for the three successions, suggesting that they were coevally deposited along the same active continental margin (Lacassie et al., 2006). This margin was probably located along the Antarctic sector of the Panthalassan Gondwana margin, as favored by the studies of Willan (2003) for the source area of the Le May Group, and of Wandres et al. (2004) and Wandres and Bradshaw (2005) for the source area of the Rakaia terrane. The last two studies indicate that the origin of the Permian detritus in the Rakaia terrane would be in the igneous rocks of the Amundsen and Ross Provinces, East Antarctica, which during the Permian were close to 60°S (Veevers, 2004; Cawood and Buchan, 2007). If the source of Permian detritus was the same for these three successions (DYC, Rakaia Terrane and Le May Group), this would imply dextral strike-slip displacement of the MDAC along the SW Gondwana margin from these high latitude to its present position.

On the other hand, paleomagnetic information on the TL and the DC demónstrate that, after Early Cretaceous remagnetization produced by the thermal influence of the SPB, both units underwent a counter-clockwise rotation of ca. 117° with an inappreciable latitudinal change (Rapalini et al., 2001). This evidence, coupled to the structural data of Forsythe and Mpodozis (1979, 1983), allowed Rapalini et al. (2001) to propose that the former units have been accreted to the Gondwana margin from the NW rather than from the SW, as had been previously considered (Forsythe and Mpodozis, 1983; Ling and Forsythe, 1987). That agrees with the early hypothesis of Ozawa and Kanmera (1984), which suggested the north-western Pacific area for the origin of the exotic oceanic units of the MDAC, and is also consistent with the sinistral sense of shear of main structures parallel to the margin of South America (Cunningham, 1993; Olivares et al., 2003). These interpretations are coherent with the migration of the Antarctic Peninsula towards the south starting in the latest Jurassic (Hervé et al., 2006; König and Jokat, 2006; Miller, 2007), which was situated parallel to the west of Patagonia at that time (Miller, 2007, and references therein). Moreover, it is suggested that the late Triassic deformation in northern Antarctic Peninsula (Península Orogeny), which affeets the Trinity Península Group accre-tionary complex (TPG, Hyden and Tanner, 1981), is associated with sinistral strike-slip movements, while dextral strike-slip is mainly a Cretaceous phenomenon in the Antarctic Peninsula (written communication, A. Vaughan, October 2006).

The deposition of sediments of the DYC in high southern latitudes (Lacassie, 2003; Lacassie et al., 2006), contrasts with the second scenario, which involves deposition of the DYC in lower and warmer latitudes, perhaps associated with subsequent sinistral strike-slip movements of the entire MDAC along the Panthalassan margin of Gondwana.

3.      Sampling and Methods

Palynological data were acquired from one sample of limestone of the TL and seven samples of metasediments of the DYC: five from Madre de Dios Archipelago and two from Diego de Almagro Archipelago (Table 1; Fig. 1). All samples were processed by standard palynological methods. All but one of the samples (MD05-20 from Guarello Island) yielded poorly preserved palynomorphs. The study and the description of the specimens were made with an optical microscope. The slides are housed at the Laboratory of Paleopalynology of the Departamento de Ciencias de la Tierra, Universidad de Concepción under codes 1396 to 1401.


Two metasedimentary samples (FO04-21 and FO04-22) from the units of the MDAC were collected for U-Pb zircon dating by SHRIMP RG (sensitive high resolution ion microprobe, reverse geometry) at the Research School of Earth Sciences, The Australian National University. Zircon grains were separated from total rock samples using standard crushing, washing, heavy liquid and paramagnetic procedures. The zircon-rich heavy mineral concentrates were poured onto double-sided tape, mounted in epoxy together with chips of the reference zircons (FCl and SL13), sectioned approximately in half, and polished. Reflected and transmitted light photomicrographs were prepared for all zircons. Cathodolumines-cence (CL) Scanning Electron Microscope (SEM) images were prepared for all zircon grains. The CL images were used to decipher the internal structures of the sectioned grains and to ensure that the ~20 um SHRIMP spot was wholly within a single age component (usually the youngest) within the sectioned grains.

The U-Th-Pb analyses were made using SHRIMP RG. The zircon grains were analyzed sequentially and randomly. Each analysis consisted of 4 scans through the mass range, with a reference zircon analyzed for every five unknown zircon analyses; SHRIMP analytical method follows Williams (1998, and references therein). The data have been reduced using the SQUID Excel Macro of Ludwig (2001). The U-Pb ratios have been normalized relative to a valué of 0.01859 for the FCl reference zircon, equivalent to an age of 1,099 Ma (Paces and Miller, 1993). Uncertainties given for individual analyses (ratios and ages) are at the one sigma level (Tables 2 and 3). Tera-Wasserburg concordia plots, probability density plots with stacked histograms and weighted mean 206Pb/238U age calculations were carried out using ISOPLOT/EX (Ludwig, 2003). The 'Mixture Mo-delling' algorithm of Sambridge and Compston (1994), via ISOPLOT/EX, was used to un-mix statistical age populations or groupings; from these groups weighted mean 206Pb/238U ages were calculated and the uncertainties are reported as 95% confidence limits.

An estimate for the maximum age for the deposition of the sediment sample may be de-termined from the weighted mean age of the youngest peak in these distributions, where >3 analyses are within analytical uncertainty. Such an age grouping has taken into account isolated cases of inferred radiogenic Pb-loss, which can produce minor scatter to younger ages. Ages for individual grains are reported at the 68% confidence level, and Geological Time Scale referred throughout the text is that of Gradstein et al. (2004).

4.      Results

4.1.      Palynology

The palynological analysis revealed a palynoflora composed predominantly by Gymnospermopsida monosaccate pollen grains, although Gymnospermopsida bisaccate pollen grains were also observed. Selected species are illustrated in figure 3. The samples show a very low frequency of palynomorphs, and in most cases an exact identification of the species is impossible because of the bad preservation state of the palynomorphs.The palynomorphs detected within the TL are Punctatisporites punctatus (Ibrahim) Ibrahim, a Pteridophyta known from the Carboniferous to Triassic in New Zealand, Australia, Asia, Europe and South America (Alpern and Doubinger, 1973; Owens et al., 2002; Pérez Loinaze, 2008).


FIG. 3. Selected pollen grains, spores and algae from the studied samples. Black line represents 10 µm. A. Granatisporítes sp.; B. Protohaploxypinus sp.; C. Cannanoropollis sp.; D. Botryococcus brauii; E. Punctatisporites punctatus; F. Plicatipollenites sp.

The palynological association observed in the metasediments of the DYC (sampled in the Madre de Dios Archipelago and in the Diego de Almagro Island) is characterized by Gymnospermopsida pollen. In addition, Pteridophyta spores as well as rare green algae (Botryococcus brauii Kützing, Lower Carboniferous to Recent) and epiphyllous fungal spores (Granatisporites and Multicellaesporites spp.) have been observed. The Gymnospermopsida include Coniferales and Cordaitales. The more frequent monosaccate pollens are Plicatipollenites and Cannanoropollis spp., which are also represented in the Carboniferous-Permian of Gondwana (e.g., Vergel, 2008; Di Pasquo, 2009). The bisaccate pollen grains of Gymnospermopsida are assigned to the Protohaploxypinus sp., also known in middle Carboniferous successions of Argentina, butrecognized also in the Permian of Brazil, South Africa, India, Antarctica, Australia and North America (Césari and Gutiérrez, 2000, and reference therein). Remnants of polyplicate and monocolpate pollen grains, of 80-95 um, assigned to Praecolpatites sinuosus (Balme and Hennelly) Bharadwaj and Srivastava, have been observed (sample FO04-22, Fig. 4). They have broad distribution in the Permian of Argentina, Brazil, Australia, Africa, Antarctica and New Zealand (e.g., Lindström, 1995). Therefore, a Permian age for the deposition of the sediments of the DYC is inferred.


FIG. 4. A. Outcrops at Seno Contreras, where the stratigraphical contact between deformed banded cherts of the DC (above the white line) and shales and sandstones of the DYC (below) is observed; B. Stratigraphic column showing the disposition of the samples analyzed by U-Pb SHRIMP and by palynological methods. 1. Sandstones; 2. Banded cherts; 3. Shales; 4. Tuffaceous layer.

In addition, a humid environment of conifer and/ or cordaitales forests with an undergrowth of fems (probably developed on wet shaded slopes) is pro-posed from the palynological association recorded in the DYC.

4.2.      Petrography

The metasedimentary samples collected for U-Pb zircon dating were obtained from an outcrop where the DC and the DYC are in conformable and inter-calated stratigraphic contact (Fig. 4). Significantly, this site corresponds to one of only two localities were this type of contact between these complexes is recorded. The samples were spatially associated, stratigraphically separated by ca. 10 m. The first sample (FO04-21) comes from a deformed (folded) metasedimentary horizon (0.04 to 0.06 m thick) of tuffaceous character, interbedded in metacherts of the DC. The microscopic petrographic description of this sample shows that it is mainly composed by very angular fragments (0.01-0.2 mm) of quartz (55%), altered feldspars (30%) and biotite flakes (15%) in a cryptocrystalline siliceous matrix. Accessory minerals include zircon, garnet, sphene, apatite and Fe-oxides. Scarce small and highly altered shard fragments were also observed. The biotites are oriented parallel to the contacts with the underlying radiolarian chert. This last feature coupled with the normal grading observed in this bed agrees with subaquatic conditions of deposition. The second sample (FO04-22) is a quartz rich metasandstone of the DYC previously analyzed by palynological methods. The sample was extracted from a massive and structureless sandstone bed (20 m of minimum thickness) nearly 3 m above the contact with the banded cherts of the DC. It is a feldspathic arenite formed by well to moderately sorted subangular and highly spherical fragments, with sizes between 0.02 and 1.2 mm (0.3 mm in average). Main fragments are quartz (60%), feldspars (30%), biotite (8%) and white mica (2%). Accessories include zircon, apatite, lithic fragments (basalts and rhyolites), garnet, pyrite and Fe-oxides.

4.3.      U-Pb SHRIMP ages of detrital zircons

The Tera-Wasserburg diagrams plot the total ratios, uncorrected for common Pb, and show that the data generally plot close to Concordia (Fig. 5). Relative probability spectra of the detrital zircon ages are presented in figure 5. For sample FO04-21 54 grains were analyzed, whereas 42 grains were examined for sample FO04-22.


  FIG. 5. Tera-Wasseburg diagrams for zircon U-Pb data. Analyses are plotted as total ratios calibrated for U-Pb, but un-corrected for common Pb. The error ellipses are 68.3% confidence limits. The dotted arrow in FO04-21 shows the direction of common Pb at 270 Ma. Alignment along this arrow suggests a real inferred age on Concordia with variable degrees of incorporated common Pb at the time of crystallization.

FO04-21. The zircons of this sample are prismatic and euhedral crystals, with zoned magmatic intemal structures as seen under CL imaging (Fig. 6). This is compatible with its textural and mineralogical characteristics, which are indicative of the tuffaceous character of the metasediment. Although some of the youngest individual ages involve significant common Pb correction (Fig. 5); a correction has been applied to derive the radiogenic ratios and age of these analyses (Table 2). The age spectrum shows a narrow range of provenance ages with a major peak in the early middle Permian, representing ca. 76% of the analyses. Minor peaks are observed in the Carboniferous, Devonian, Ordovician, Cambrian and Neoproterozoic, each one equivalent to less than 8% of the total analyzed grains. The Permian analyses yield a weighted mean 206VbP-3S\J age of 270.4±2.7 Ma (MSWD=1.2), interpreted as the maximum possible depositional age of the analyzed metasediment.


FIG. 6. U-Pb zireon age provenance pat-terns (age versus relative proba-bility) of the analyzed samples. Insets show the entire population of zireon ages. Montages ofrepresentative portions of the cathode-luminescence images for each sample, with individual spot ages.

FO04-22. The zircon grains show zoned internal structures (Fig. 6), and subrounded to subangu-lar shapes with high sphericity are predominant, although prismatic grains are observed as well. The grains analyzed from this sample are very low in common Pb. The relative probability plots of the detrital zircon ages display a prominent component in the early Permian (ca. 40% of the analyses), with other subordínate peaks in the Carboniferous and Devonian (ca. 17% of the analyses each one). There are scattered older ages ranging from Early Paleozoic to Neo and Mesoproterozoic aged noise and one Paleoproterozoic aged zircon. A weighted mean 206PB238U age of 289.7±2.1 Ma (MSWD=1.3) place a constraint on the maximum age of deposition of this metasediment.

5.           Discussion

5.1.           Regional paleoenvironmental conditions

As pointed out by Cúneo (1996) and Buatois et al. (2006), it seems that the end of the late Paleozoic Ice Age in Gondwana is diachronous, waning first in South America, as revealed by litho- and bio-stratigraphic records (ca. 280 Ma; Iannuzzi et al., 2007), and then in Australia (ca. 260 Ma; Fielding et al., 2008). This diachronism has been habitually attributed to the Gondwana drift across the South Pole (López-Gamundí et al., 1994; Visser, 1996), but the possibility of more than one glacial event cannot be ruled out (Limarino et al., 2006). Roscher and Schneider (2006) show that there is a general trend of aridization in the Permo-Carboniferous interrupted by wet phases related to the waxing and waning of the Gondwana icecap. Lindström and McLoughlin (2007) indicate that during the middle to late Permian a gradual warming trend is evident from the western to the eastern parts of Gondwana. Furthermore, semiarid and arid climatic conditions in middle Permian times appear like a common feature in many western Gondwana basins (López-Gamundí et al., 1992; Limarino et al., 2006; Spalletti and Limarino, 2006; Souza et al., 2007), while the start of arid conditions in western Gondwana has been situated towards the end of the early Permian by Césari et al. (2007).

Palynomorphs documented in this study are comparable to those recorded in Carboniferous and Permian strata of other late Paleozoic Gondwanan basins, such as those exposed in the Chaco-Paraná Basin in Argentina and Uruguay (Césari et al., 1995; Beri et al., 2006), and at Rio Grande do Sul in Paraná Basin, Brazil (Souza and Marques-Toigo, 2005). It is noteworthy to emphasize that the only palynomor-ph with exclusive Permian record (Praecolpatites sinuosus) is in one of the samples with a Permian maximum depositional age (Fig. 4). Regarding the paleoenvironmental background mentioned above and the Permian age obtained in the DYC, the proposed paleoclimatic setting for the deposition of the late Paleozoic basins of southwestern Gondwana, in particular those situated in Patagonia, helps to establish regional correlations and some paleoclimatic inferences. Archangelsky et al. (1996) and Limarino et al. (1996) suggested a specially humid and températe climate, even subtropical, for the deposition of the La Golondrina Formation (the lower member in the La Golondrina Basin, Patagonia), whose age would be restricted between the Sakmarian and the Kungurian (Limarino and Spalletti, 2006). The age and paleoclimatic features of this formation are remarkable, particularly because it crops out in a relatively close position (ca. 500 km apart) to the present-day position of the outerops of the DYC (Fig. 7). Another late Paleozoic Patagonian basin that shares similar paleoclimatic characteristics and also an adjacent location to the deposits of the DYC is the Tepuel-Genoa Basin (López-Gamundí and Limarino, 1984; Andreis et al. 1987) (Fig. 7), likely related to metamorphic rocks outcropping in the Coastal Range of Chile (Hervé, 1988; Duhart et al., 2001). The early Permian component of this basin (Sakmarian-Artinskian, according to Césari et al., 2007) is represented by the Río Ge-noa Formation, whose sediments have also been interpreted as deposited in a humid and subtropical climate (Archangelsky et al., 1996; Limarino et al., 1996). However, it is not possible to determine a more precise paleoclimatic connection with the La Golondrina Formation, mostly because late early Permian (Kungurian) sedimentary rocks have not been identified in the Tepuel-Genoa Basin (Limarino and Spalletti, 2006). If it is assumed a fixed position of the DYC with respect to Patagonia since its deposition, and considering the previous examples and the age obtained for the DYC, it would be probable that warm paleoclimatic conditions were recorded in the metasediments of this complex. This was, however, impossible to register in this study, mainly due to the very low proportion of palynological material in the samples.


FIG. 7. Paleogeographic reconstruction for South America in the early Permian, based on the pole of Rapalini et al. (2006). Positions of the late Paleozoic basins (from Limarino and Spalleti, 2006), Choiyoi deposits (modified from Kay et al., 1989; Ramos, 2000), and the currcnt location of the outcrops of the Madre de Dios Accretionary Complex (MDAC) and the Trinity Peninsula Group (TPG) are shown, as well as the probable position of the Antarctic Peninsula (AP). The southern extension of the Rio Chico-Punta Dúngenes high is taken from Ramos (2008), and the distribution of Permian igneous rocks of the North Patagonian Massif is from Pankhurst et al. (2006).

Even though we cannot precise the paleoclimatic conditions during the deposition of the DYC, some observations can be done. The types of deposits accumulated as a consequence of ice activity are very varied, and several non-glacially related mechanisms can produce similar deposits. The recognition of these accumulations is more difficult if we consider that glacial deposits are frequently reworked in outwashes or by mass fiow. For these reasons, the assignment of a glacial origin to any deposit, or its refutation, needs the combination of the properties of the deposit itself, the adjacent rocks units, as well as climatic and paleogeographic conditions at the time of deposition (e.g., Charrier, 1986). Additionally, to establish that a sedimentary succession, or part of it, preserves a record of glacial, proglacial or periglacial depositional environments, multiple facies criteria are needed (diamictites, chaotic fabrics, rhythmites, laminated mudrocks with outsized dispersed clasts (lonestones), striatedpavement, and faceted, bullet-shaped and striated clasts, among others; Miller, 1996), or it is required a single criterion that unam-biguously appears to indicate glacial influence by virtue of its occurrence out of context with enclosing facies (e.g., Fielding et al., 2008). In this context, even though Hiere have been records of diamictites in the DYC (Cecioni, 1956; Forsythe and Mpodozis, 1983), this does not give conclusive evidence of a glacial environment of deposition. Moreover, there are only few localities in the whole extension of this accretionary complex (more than 1,000 km2) where diamictites have been identified, but the existence of glacial characteristics sensu stricto like lonestones (dropstones), striated pavements or faceted clasts have never been reported.

The large extension and volume of the DYC could be attributed to the great volume of fresh water produced by ice-melting during the waning of the glaciation in Gondwana (Buatois et al., 2006). This huge volume of water probably reworked large amounts of sediments that were subsequently deposited at the margins of the continent. The ichnofaunas registered in DYC (Scalarituba isp., Chondrites isp., Planolites isp., Palaeophycus isp. and Ancorichnus isp.; in Lacassie, 2003), are distinctive of marine environments (written communication, L. Buatois, January 2008), and the low content of palynological material in the samples suggests a marine offshore environment of deposition of the host rocks. Both proposals match with the interpretation of the DYC rocks as turbidites, and would exelude a fjord-like setting for the origin of this succession and indicate that the place of deposition of the DYC was located far away from the direct influence of fresh water formed during the Gondwanan deglaciation.

5.2.      Age and sources of metasediments

The detrital zircon data for sample FO04-21 of the DC reveal a maximum possible depositional age of this tuff-rich sediment of 270.4±2.7 Ma, roughly in the limit between the early Permian (Kungurian) and the middle Permian (Roadian). This result is identical to the youngest predominant detrital zircon U-Pb SHRIMP age component previously reported for samples of the DYC (ca. 270 Ma, Hervé et al., 2003; Hervé et al., 2006), and suggests that this peak is probably linked to a significant contribution of volcanic (tuffaceous) material deposited near the continental margin. This is also in agreement with the age of widespread ash fall deposits and tuffaceous horizons present in basins of the west Gondwana (Turner, 1999; Stollhofen et al., 2000; López-Gamundí, 2006; Santos et al., 2006, Tohver et al., 2007), commonly correlated with the peak of the Choiyoi silicic volcanism during the late early Permian and middle Permian along the Andean Cordillera and its equivalents in Patagonia (López-Gamundí, 2006).

The data obtained for sample FO04-22 of the DYC (289.7±2.1 Ma) indicate an early Permian (Sakmarian) maximum possible depositional age. This is nearly 20 m.y. older than the youngest U-Pb SHRIMP ages component recorded for detrital zircons in the DYC (Hervé et al., 2003). Nonetheless, the former authors have recognized this Sakmarian peak (ca. 290 Ma) as an individual Permian population in the age spectrum of the detrital zircons in the DYC. The apparent inverted stratigraphical position of the analyzed samples (Fig. 4) can be explained as FO04-21 being the airborne volcanic material deposited near the continental margin of Gondwana, and FO04-22 as resedimented detritus, formerly deposited somewhere between its local source area and the final depositional site, and then redeposited as turbidite fiows above the cherts which include the tuffaceous layer represented by FO04-21.

Augustsson et al. (2006) suggest that the Permian sediments present in the metasedimentary complexes of Patagonia were probably largely supplied from local Patagonian and West Antarctic sources. Furthermore, Pankhurst et al. (2006) claim that the Permo-Triassic granites present in the North Patagonian Massif (extending between 41° and 44°S, approximately) can be identified as the most important source so far recognized for the provenance of detritus in the late Paleozoic metasedimentary rocks along the Pacific margin of Gondwana. This would imply transport by a system of long rivers over a wide and relatively fíat pre-Andean platform (Hervé et al., 2003), or even eolian transport over hundreds of kilometers (e.g., Dickinson and Gehrels, 2009).

The widespread abundance of radiometric data from ash fall deposits within the range of 280 to 260 Ma has been commonly attributed to a period of intense silicic volcanism along the continental margin of southwestern Gondwana that peaked around the 270 Ma (López-Gamundí, 2006). Besides the model of Pankhurst et al. (2006), until now no analogous magmatic process responsible for the ubiquitous Sakmarian peak (ca. 290 Ma) in the U-Pb spectrum from samples of the DYC has been clearly identified. Similar and equivalent radiometric ages have been obtained from adjacent late Paleozoic basins of southwestern Gondwana. In the southern-most Karoo Basin of South Africa, Bangert et al. (1999) reported 288.0±3.0 and 289.6±3.8 Ma from bentonitic tuff beds intercalated in sedimentary rocks of the Prince Albert Formation, lower Ecca Group. Rocha-Campos et al. (2006) and Guerra-Sommer et al. (2008) acquired comparable Permian ages from ash-fall deposits interbedded in coal successions of the southern Paraná Basin in Brazil. The last authors claim that this information supports the presence of an active and extensive volcanic event in western Gondwana around the Carboniferous-Permian boundary (ca. 299 Ma). It is, however, difficult to restrict the magmatic activity in southwestern Gondwana to a single instant in Permian times, since most of the limited available data are in the range of 270 to 290 Ma and usually they overlap within their analytical uncertainty. Instead, the entire early Permian could be regarded as a period of active and geographically widespread magmatism in this region of Gondwana. This scenario would also explain the strong coincidence of the main Permian peaks in the detrital zircon U-Pb spectrum of the metasediments of the DYC, the Rakaia Terrane (New Zealand) and the eastern Le May Group (Alexander Island, Antarctica) (Lacassie et al., 2006). Nevertheless, the cause of the petrographic and geochemical similarities between these three metasedimentary units is an issue not discussed here as it is beyond the scope of this work.

5.3.      Paleogeographic correlations

So far, there has been no consensus in the place of accretion of the TL and DC and the supposed subsequent sense of movement of the MDAC (as a coherent block) along the southwestern Gondwana margin. Lacassie et al. (2006) proposed accretion of the DC-TL assemblage at the Antarctic-Australian portion of the Gondwana margin, followed by dextral translation of the MDAC (and henee of the DYC) parallel to the margin. On the other hand, the possibility of a virtually fixed position for the DYC since its deposition (with Patagonia as reference) is hampered by the fact that there is no clear indication of a coeval Permian magmatic are at latitude similar to the current position of the MDAC in southern Patagonia (e.g., Hervé et al., 2006). This magmatic are could be represented either by the Choiyoi acid magmatic province (Kay et al., 1989; Mpodozis and Kay, 1990) or by the Permian igneous rocks in the North Patagonian Massif (Pankhurst et al., 2006), both sources currently located north of the 40° and 44°S, respectively However, the Choiyoi Formation presents ages of ca. 281 Ma near its base (Rocha-Campos et al., 2006; Suárez et al., 2009), and thus cannot account for the 290 Ma peak in the U-Pb age spectrum. Conversely, it is well known that the subduction of bathymetrically elevated oceanic features such as ridges or plateau (DC and TL in this case) can flatten the subducting slab and prevent the magmatic activity in the vicinity of the continental margin. According to the above-mentioned situation, the DYC would not have had displacement since its deposition and the remnants of the associated Permian magmatic are could still be hidden below the Mesozoic sedimentary cover somewhere in the southeastern Patagonia or even farther eastward. This option has been recently explored by Ramos (2008), who proposes a late Paleozoic magmatic are with a southern extension in the NNE trending Río Chico-Punta Dúngenes High (Fig. 7). This hy-pothesis would preclude defining the DYC as an allochthonous terrane. Aprecedent that supports this hypothesis is the presence of metasedimentary rocks with detrital zircons with U-Pb SHRIMP ages similar (ca. 290 Ma) to the ones recorded in the DYC to the east of the South Patagonian Batholith, at almost the same latitude of the MDAC (Augustsson et al., 2006). In brief, additional and more detailed paleomagnetic, geochronological and isotopic work is needed to give more convincing arguments supporting one of the proposed hypotheses, an autochthonous or an allochthonous origin of the DYC.

Little is known about the exact paleogeographic configuration of the Antarctic Peninsula during the late Paleozoic, though one of the most accepted current paleogeographic reconstructions lócate the Antarctic Peninsula lying west of Patagonia in the Middle Jurassic (Konig and Jokat, 2006). Therefore, the possibility of a geological correlation between the geologic units present in western Patagonia and those exposed in the Antarctic Peninsula must be considered. In this context, it has been suggested based on the similarity of lithology and detrital zircon age pattems, that the Trinity Península Group (TPG) is the equivalent counterpart of the DYC in the Antarctic Peninsula (Hervé et al., 2006) (Fig. 7). Willan (2003) assumed that the TPG could have been derived from a glaciated continental margin, though his result is based only on indirect evidence (geochemical weathering) and, at this time, there are no palynological reports on this unit. According to the paleogeographic reconstruc-tions for the late Paleozoic (e.g., Cawood and Buchan, 2007), the TPG was supposedly located in a higher paleolatitudinal position than the late Paleozoic Patagonian basins, and therefore the paleoclimatic conditions wouldhave been colder during its deposition. However, caution must be taken with this interpretation, mainly because the TPG is part of the Western Domain of the Antarctic Peninsula (Vaughan and Storey, 2000), which has been regarded as a suspect terrane and even as al-lochthonous to the rest of the terranes of the Antarctic Peninsula (Willan, 2003).

6.      Conclusions

This contribution presents the first palynological record for the late Paleozoic in Chile. The palynological assemblage recorded in the DYC is composed mainly of Gymnospermopsida pollen, with also Pte-ridophyta and fungal spores. The studied association indicates a humid environment of forests with an undergrowth of ferns.

The palynological data indicate a Permian age for the deposition of the DYC. This age is also sup-ported by new U-Pb SHRIMP detrital zircon ages, which constrain the maximum depositional age of the DYC to the limit between the early Permian and the middle Permian (ca. 270 Ma), confirming the maximum depositional age obtained by previous geochronological data (Hervé et al., 2003).

The available data indicate that the allochtho-nous hypothesis for the DYC is not completely proved, and an autochthonous tectonic setting (with respect to Patagonia) could also be a possible interpretation.

Acknowledgements

CONICYT doctoral grant to FAS, Fondecyt 1050431, Proyecto Anillo Antartico ARGT 04, Compañía de Aceros del Pacífico (CAP-Mina Guarello) and Expedición Última Patagonia 2006 supported this research. D. Quiroz (SERNAGEOMIN), A. Vidal and S. Martini are thanked for field support. We are indebted to Dr. C. Juliani (Universidade de Sao Paulo) for make possible sampling on Diego de Almagro Archipelago. Thanks also for Dr. A. Rapalini (Universidad de Buenos Aires) for bis paleogeographic reconstructions. Drs. J.P. Lacassie and P. Vásquez contributed with fruit-ful discussions. We also thank C. Limarino (Universidad de Buenos Aires), H. Bahlburg (WWU Münster) and an anonymous referee for their helpful comments on an early version of this manuscript, as well as the careful editorial handling by M. Suárez.

 

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Manuscript received: December 22, 2008; revised/accepted: January 13,2010.

©  2011  Servicio Nacional de Geología y Minería