II International Coloquium Vendian-Cambrian of W-Gondwana

The Puncoviscana Formation is represented by a thick and heterogeneous succession of sedimentary rocks highly deformed and slightly metamorphosed, cropping out mainly along the Eastern Cordillera between 23º south (Tarija Department, southern Bolivia), and 27º south (Province of Tucumán, northern Argentina). They were originally described as part of the “basal Precambrian shield” of northwestern Argentina, and later reconsidered by means of the finding of several trace fossils localities including the Neoproterozoic-Lower Cambrian transition.

Widespread siliciclastic rocks characterize the succession (sandstones, shales and few conglomerates), but limestones, volcaniclastics and lava flows are also mentioned. The first mentioned lithology display poor sorted clasts with medium to fine grains, mono to polycrystalline cuartz, albite plagioclase, scarce ortose and some sedimentary and low grade metamorphics, as well as few chert. A compositional analysis corroborates that the Puncoviscana Formation was originated on the continental block (marginal plain / back-arc), while sedimentary and low grade metamorphic series were also developed, as shallow tectonic stages of a reworked multi-phase orogen (Jezek, 1990). Paleocurrents display a general SSE-NNW flow direction, while the conglomerates are described as intra-formational type, being placed on the western border. Several localities with vulcanites are also mentioned, whose sedimentary pattern probably allows us to refer them to submarine fan related type of sedimentation. Abundant red shales, by sectors of the basin, represent hemipelagic conditions with some oxygen content within the sediments, while the greenish shales are mentioned as deposition of deep ocean contour flow. The blackish and brownish limestones are interpreted to be deposited over dorsals or some structural highs within the sea floor.

The ichnological distribution display a remarkable alignment as belts, probably resulting on the morphology of the Puncoviscana basin, as well as a chronological record of the ichnofaunas. Three ichnoassociations are recognized: eastwards the Beltanelliformis, the Nereites is placed in the middle, while Oldhamia is located westwards (Aceñolaza & Alonso, 2000 with references). Several aspects have been discussed on this distribution: they may represent different facies on a same event, or different temporal levels on the evolution of the Puncoviscana basin. Due to the morphological characters of the basin, we believe as a more viable possibility the second one, and consider Oldhamia as a totally different ichnoassociation that will represents the “agronomic revolution” of Seilacher & Pflüger (1994). Oldhamia displays a Tommotian-Atdabannian age, while the other ichnassociations represent older elements, reaching down to the Ediacaran.

In this moment, discussions are focused on the way the Puncoviscana basin was formed and developed. The general idea is that the basin had an origin on a marginal aulacogenic “triple point” type of structure, placed in the western border of Gondwana. The southwestern branch of this aulacogenic structure is proposed as the responsible for the development of the Puncoviscana Basin.

We cannot recognize the units that represent the base and the source of materials for the basin, so we suppose that they belong to an older tectono-orogenic cycle as Sunsas (1100-900 Ma, sensu Litherland & Bloomfield (1981). Once this cycle stabilizes, a new process begins with the formation of the earlier mentioned “Bolivian Triple Point” (Suárez Soruco, 2000), separating westwards the Arequipa-Antofalla terrane, to the north-northeastern the Guaporé craton and to the south-southeast the Rio de la Plata craton. The identity of the Arequipa terrane as a Gondwanan bock is strongly supported by the data given by Luccassen et al. (2000). This opening of the Puncoviscana basin follows a north/south-southwest direction, isolating the Gondwanan border while receiving the sediments from the surrounding structural highs as Arequipa, Guaporé and the Rio de La Plata terranes and cratons, that will lastly end up filling the basin.

References

Aceñolaza, F.G. & Alonso, R. 2000. Icnoasociaciones de la transición Precàmbrico-Càmbrico en el noroeste de Argentina.Cuadernos de Geologìa Ibèrica (in press).Madrid.

Jezek, P. 1990. Análisis sedimentológico de la Formación Puncoviscana entre Tucumán y Salta. In Aceñolaza, F., Miller, H. and Toselli, A. (eds). El Ciclo Pampeano en el noroeste argentino. Correlación Geológica 4, 9-36. Tucumán.

Litherland, M. and Bloomfield, K. 1981. The Proterozoic history of eastern Bolivia. Precambrian Research. 15: 157-179.

Lucassen, F., Becchio, R., Wilke, H.G., Franz, G., Thirwall, M.F., Viramonte, J. & Wemmer, K. 2000. Proterozoic- Paleozoic development of the basement of the Central Andes (18º-26º S); a mobile belt of the South American Craton. Journal of South America Earth Sciences 13: 697-715.

Seilacher, A. & Pflüger, F. 1994. From biomats to benthic agriculture: a biohistoric revolution. In: Krumbein, W.E., Paterson, D.M. and Stal, L.J. (Eds.), Biostabilization of sediments. B.I.S. Oldemburg: 97-105.

Suarez Soruco, R. 2000. Compendio de geología de Bolivia. Revista Técnica de Yacimientos Petrolíferos Fiscales 18(1-2) 1-144. La Paz.

Marcante característica do registro sedimentar neoproterozóico é a abundância relativa de dolomitos (Tucker 1982). Tal relação demonstra que no Neoproterozóico as condições para dolomitização foram mais favoráveis, inclusive com possibilidade de precipitação direta (Grotzinger & Knoll 1995), o que também seria sugerido pela ocorrência de dolomitos com textura primária totalmente preservada (Tucker & Wright 1990).

Kazmierczak et al. (1985) explicaram que a abundância de dolomitos no Neoproterozóico, em comparação à ocorrência de calcários calcíticos, seria devido à existência de oceano pobre em Ca²⁺, tendo interpretado que o teor de Ca²⁺nos oceanos foi crescente, atingindo seu valor máximo no limite neoproterozóico/cambriano, quando a quantidade de Ca²⁺ foi tal que a biocalcificação teria sido uma forma encontrada pelos organismos de se desintoxicarem da elevada concentração de cálcio. Outra hipótese é a de que teria havido condições excepcionais de bombeamento da água do mar através dos sedimentos.

A constatação de precipitação de dolomita através de atividade microbiana (bactérias redutoras de sulfatos) em lagunas costeiras de Lagoa Vermelha - Rio de Janeiro (Vasconcelos et al. 1995) demonstra a importância destes microorganismos como centros de nucleação da dolomita, constituindo, também, uma das possíveis causas da relativa abundância de dolomitos no Proterozóico (Wright 1997), quando a atividade microbiana teria atingido seu apogeu, atestado pela abundância de estromatólitos.

No domínio da Faixa de Dobramentos Paraguai, duas unidades dolomíticas são encontradas: a Formação Bocaina e o Membro Superior da Formação Araras. A primeira unidade insere-se na base do Grupo Corumbá, aflorante na Serra da Bodoquena e a outra no Grupo Alto Paraguai com exposições na Serra das Araras (figura 1).

A Formação Bocaina, com espessuras entre de 30 e 80 m, é caracterizada por dolomitos puros, em determinadas localidades, apenas o mineral dolomita foi identificado sob difratometria de raios X, com gradativa predominância de silexitos no topo. Nela também ocorrem rochas fosfáticas e abundantes estruturas estromatolíticas.

O contato inferior dos dolomitos da Formação Bocaina é erosivo e estes ocorrem sobrepostos diretamente sobre o embasamento gnáissico-granítico. Apresentam-se também posicionados diretamente sobre os diamictitos, provavelmente glácio-marinhos, da Formação Puga, através de contato erosivo. O contato erosivo com o embasamento constitui extensa superfície de aplainamento denominada Superfície de Aplainamento Pedra Branca (Boggiani & Coimbra 1998).

Para a Formação Bocaina, a origem dos dolomitos é interpretada como relacionada à ocorrência de condições excepcionais de bombeamento da água do mar através dos sedimentos. Estas condições poderiam estar relacionadas à rápida transgressão marinha pós-glaciação associada à fragmentação de supercontinente.

Fig. 1: Exposições do Grupo Corumbá e da Formação Araras no domínio da Faixa Paraguai e sobre o Cráton Amazônico.

Outra característica da transgressão marinha da Formação Bocaina é a sua relativa pouca amplitude em termos de elevação do nível do mar, porém com avanço do mar por uma extensa área sobre o continente, previamente aplainado. Tal conformação geomorfológica teria possibilitado o desenvolvimento de extensas plataformas epineríticas sob condições de águas rasas e bem iluminadas, propícias à proliferação de atividade microbiana bentônica e formação dos estromatólitos, os quais teriam favorecido ainda mais a dolomitização.

A continuidade lateral por centenas de quilômetros e relativa homogeneidade das fácies sedimentares da Formação Araras são indícios de que estas se depositaram em área extensa e plana. Estas condições poderiam estar presentes em mar epicontinental ou mesmo em extensa rampa carbonática, onde a sedimentação teria se dado ao longo de planície de maré. Neste contexto, a Associação Faciológica Calcítica e Terrígena inferior teria se depositado em ambientes de inframaré e a Associação Faciológica Dolomítica, estromatolítica, em ambientes de inter e supramaré.

A sedimentação da base da Formação Araras teria ocorrido em ambiente de inframaré, sob condições euxínicas, com sedimentação carbonática calcítica sujeita a aporte periódico de terrígenos, resultantes de fatores provavelmente climáticos. Com a sedimentação, e conseqüente diminuição da profundidade da bacia, ocorreram condições de dolomitização dos sedimentos em ambiente de infra e supramaré, sob climas áridos, onde o aporte terrígeno passou a ser praticamente nulo.

A metade inferior da Formação Araras não sofreu dolomitização devido às intercalações de camadas pelíticas que teriam impedido a circulação dos fluidos dolomitizantes.

Os dados obtidos de isótopos de C e O (Tabela 1) são coerentes com estas interpretações. Na Formação Bocaina, apresentam d¹³C próximos de zero a ligeiramente negativos e os valores de d¹⁸O ao redor de -3 ⁰/₀₀PDB, sem grande dispersão dos valores, já na Formação Araras, os valores demonstram enriquecimento em ¹³C e ¹⁸O e são mais dispersos, o que corrobora a hipótese de ambiente relativamente mais restrito e condições evaporíticas.

Following the discovery of new rich fossil-„Lagerstätten“ in the last decade the problem of „cause and course“ of the seemingly explosive appearance of many metazoan groups during the Lower Cambrian has moved into the centre of international research subjects. Worldwide, new finds and occurrences of Ediacaran-type fossil communities have been found and are still being controversly discussed with respect to their constructional and systematic nature. In the field of trace fossil research new occurrences along with better age determinations give a complementary insight into the development of organismic complexity from 1000 Million years onward. Well known fossil-Lagerstätten like Chengjiang (Yangtze platform, China) and Sirius Passet (Greenland), yielding soft-bodied faunas, provide much information on the stages of the evolution of macroscopic metazoans from sponges to chordates at Lower Cambrian times. The systematic variety as well as the complexity of body construction plans already at the time of the first Trilobite-biozones in the Lower Cambrian strongly indicate a longer evolutive history in the Precambrian, results, which are supported by results coming from molecular biology.

Microscopically sized remains, gathered by special search- and processing techniques from Late Precambrian through Cambrian rocks are increasingly important in adding to theories and knowledge on the basis of macroscopic finds. Besides diverse „small shelly“ faunas, these small and important fossils now include mainly embryonic stages and developmental series as well as components of the ecologically very important meiofauna. Above all the „Orsten type“ phosphatic preservation may preserve details down to the cellular and even histological level and may be treated and analysed like recent biological material. In addition „cuticular preservation“ may add well preserved material in certain, clay- and silt-rich rocks, which is, however, somewhat more difficult to assign systematically. Other preservation stages, i.e. early mineralisation in iron-carbonate (Siderite), clay minerals or silicification are likely to yield more well preserved material upon further directed search.

Multiple techniques, applied to different lithologies and preservation stages must be used and transferred from other paleontological fields to be able to isolate and investigate these valuable remains adequately. The lecture gives an overview of examples, techniques and possible future perspectives of microscopic scale paleontological investigations to be used in further research in series around the Precambrian-Cambrian boundary:

Whole rock samples

Sectioning techniques. Phosphorites, being common and widespread worldwide as rock series, redeposited rock fragments or thin crusts (often hardgrounds) in rocks of this age interval, are often extremely fossiliferous, but only in some cases attracted major paleontological attention up to now. No chemical technique being available to extract phosphatic fossils from a phosphatic matrix without the danger of significant loss, sectioning techniques are still widely used to investigate the fossil content of these rocks. Many phosphorites however become barely transparent and do not allow for fine structure investigations under high magnificatins, as long as standard section thicknesses are applied. The peel technique, being most commonly in use in limestones, is a promising, fast, and very material-saving technique to be applied also in phosphorite-investigations. Peels gathered from phosphorites section planes are no pure surface replicas, but, due to organic and microcrystalline apatitic substance still adhering to the film, a kind of transfer preparations, which may be investigated with view on fine-structural as well as mineralogical investigations using highly magnifying oil-immersion objectives. Besides, individual sections of the peel may be sputter coated and investigated using the SEM.

Sectioning techniques in limestones, cherts etc. are applied mainly to organic micro- and meiofossils, not extractable from a limestone matrix by the use of acids. They may be investigated and documented in polished or varnish-covered sections using oblique incident light, since contrast and colour is better visible than in normal thin sections in transmitted light. Such sections of cherts, phosphorites, but also many other rocks including highly metamorphosed ones, also allow for investigations in incident light under high magnifications with the aid of oil immersion objectives. By this technique, extremely small sized tests and skeletons of mineralic as well as organic composition may be detected.

Etching techniques: Limestones, in which the fossils have not been subjected to early diagenetic phosphatization, may still be etched using careful etching techniques and concentrations (diluted acetic acid) to recover their organic fossil content. Investigation and sorting of the residue must take place without drying of the residue in a petri dish under shallow water cover in order to avoid incrustation and deformation of the thin and flexible walls.

Palynological processing techniques remain unsuccessful, if the palynomorphs are, by heat or small scale tectonic deformation, desintegrated into small pieces, held together only by the surrounding rock matrix. Palynomorphs of this type can still be prepared as whole pieces from the residue, if the Si-flouride flakes, often forming during HF action, are not dissolved but retained in the washing screen and embedded in glycerine-jelly on microscopic slide preparations. In this case the flouride flakes are taking over the protective and supporting function of the rock matrix and the organic bodies can be investigated in their original context through the almost transparent whitish matrix of the flouride flake.

Macroscopic fossils

Cuticular transfer techniques: Such transfer techniques, being applied in cuticular analysis in Paleobotany, may also be used during investigations of macroscopic paleozoological findings, if at least some of the original cuticular substance is preserved. Transferred organic cuticular substance may be inverstigated in great detail microscopically, including sensory organs, setation etc. Pyrite-impregnation, common in macrofossils from fresh, unweathered rocks, may be x-rayed and reconstructed three-dimensionally by the recently invented micro-computer-tomography technique, provided that the fossils are not completely flattened. If the pyritization is fine and clear enough, serial sectioning or serial grinding can be applied here as well.

Isolated micro- and meiofossils

The SEM-standard technique stays indispensable and most important for external documentation of phosphatized fossils. Other techniques must be applied, if information about internal structures is desired (soft integument preservation in closed carapaces or shells, eggs with embryonic structures). Thin sectioning of individual preselected small phosphatic bodies allows for their internal investigation within a reasonably short time. Such sections are more easily sectioned down to a thickness below 25 µm (“ultrathin sections”), allowing for a more detailed microscopic investigation compared to whole rock thin sections of average thickness.

The only (but major) disadvantage is the total loss of material outside the section plane. Possibilities to avoid this is a laborious “small scale” serial sectioning technique, accompanied by sequential documentation of the section planes or, again, micro-computertomography allowing for destruction-free single or sequential investigations.

Complex techniques following the similarly complex initial treatment of fossils and whole rock samples have been published recently in Paleolontologica electronica, vol. 4 (2001) by Sutton, Briggs, Siveter & Siveter, and will be the future for this kind of research: This is the combination of manual microscopic techniques including serial grinding with digital photography and graphic processing, resulting in 3-D-images, that can be focused through and rotated by video in electronic publications. This techniques eliminate the necessity of mechanical and/or chemical isolation of organic or mineralized remains from the rock, as long as the optical contrast is large enough for the fossil outlines to be traced on the digital images. If serial sectioning is used along with the peel technique, this allows for an even better “storage” and documentation of material during this destructive method.

The importance of microscopic studies

Microscopic remains, gathered by the above mentioned techniques, currently quickly gain importance in paleontological prospection and research, as they are allowing for a very close cooperation between neontologists working on the respective fields (embryology, construction morphology, phylogenetic systematics) and paleontologists on the basis of exceptionally well preserved material. The important processes of early diagenetic (“Tapho”)-mineralisation especially by phosphate, yielding excellent soft-integument or even soft-part and cellular preservation, are known to occur predominantly in smaller sized bodies.

Once again paleontology might be able to yield evidences for theoretically predicted “ground-plan” and transition stages of major groups. It is the idea of the author, brought foreward to the research group on construction morphology at Senckenberg-Museum/Frankfurt and currently jointly discussed and prepared by several authors, that the multitude of spherical small fossils occurring in phosphoritic Precambrian/Early Cambrian series represent not “only” embryos of various kinds, but also fossil evidence (in the Early Cambrian like a kind of “living fossils” by that time) of the published spherical to elliptical early metazoan ground-plan reconstructions (“gallertoids”), expected (and now found to be) only about one millimeter in size. Fossils of such organisms cannot be expected to be easily visible on rock surfaces, but must be isolated and investigated following complex chemical and/or petrographical treatment of the rock samples.

This review involves experiences and results from many sides. Detailed references will be given in a forthcoming more detailed publication.

La Formación Cerros de Aguirre (Vendiano superior): geocronología, características estructurales y litológicas

Néstor Campal*, Fernando Gancio*, Alejandro Schipilov* & Jorge Bossi*

*Cátedra de Geología, Facultad de Agronomía. Garzón 780, Montevideo-Uruguay

“A tropical paradise”?

Neoproterozoic glaciations from the southern Brazilian perspective

Toni T. Eerola*

*Department of Geology, Federal University of Santa Catarina, Brazil. E-mail: toni_eerola@hotmail.com

Introduction

The Rio de la Plata Craton (southern Brazil and Uruguay) collided with the Kalahari Craton during the Neoproterozoic-Cambrian times (Fernandes et al. 1992, Chemale Junior et al. 1995, Frimmel et al. 1996). This period records Sturtian (750-700 Ma) and Varangerian (620-580 Ma) glacial deposits in Namibia, Argentina, Brazil, Laurentia and other continents (Eerola 2001 and references therein), when the whole Earth seems to has been covered by glaciers (Hoffman et al. 1998). A discrete Cambrian glacial period occurred again in Namibia, western Africa and China (Germs 1995, Hoffman et al. 1998).

There are also some suggestions on Neoproterozoic glacial influence in southern Brazil (e.g. Carvalho & Pinto 1938, Bocchi 1970, Eerola 1995), but those were not confirmed or rejected by subsequent published studies. As the problem is mainly ignored in the literature on the region, the Neoproterozoic glacial record has been considered as absent. This is intriguing because glacial deposits occur in Namibia (Germs 1995, Hoffman et al. 1998) and other regions close to southern Brazil (Eerola 2001), where sedimentary basins were coeval with glaciations (Eerola 1994, Hartmann et al. 2000). This paradoxal situation was well illustrated by the map on ancient glaciations on the Gondwana continents by Crowell’s (1983) figure 2. It was also questioned by Eerola (1994).

The aim of this paper is to rediscuss this paradox and Neoproterozoic-Cambrian glaciations from the southern Brazilian perspective.

Importance of the presence or absence of glacial record in southern Brazil

The finding of Neoproterozoic glacial diamictites and dropstones in southern Brazil has important implications. It would give information on prevailed climate in the region during that time and would permit consistent chronostratigraphic correlations with Namibian basins, as well as with other coeval deposits in the world (Eerola 1994, 1995).

Their prooved absence is also significant. As southern Brazil was surrounded by regions affected by glacial influence (see Eerola 2001), the supposed lack of coeval glacial record in the middle of those is intriguing. There should be important reasons for that. Those could reflect tectonics, climate, paleolatitude, sedimentary processes and problems in preservation, recognition and interpretation of their record.

Reasons for the apparent absence of Neoproterozoic glacial record in southern Brazil

If glaciers cover large areas during glaciations, why such record has not been confirmed in southern Brazil, while this is found in Namibia and elswhere? There should be good reasons for this paradoxal situation. Some of those are suggested and discussed here:

(1) misinterpretation of sedimentary deposits as glacial in Namibia and elsewhere;

(2) no glaciers in southern Brazil, controlled by climate and/or topography;

(3) nonpreservation of glacial record in southern Brazil or

(4) glacially infuenced deposits have not been found or recognised in southern Brazil.

Discussion

As the Neoproterozoic glacial record is widespread and largely accepted in Namibia (e.g. Germs 1995, Frimmel et al. 1996, Hoffman et al. 1998), its misinterpretation there could be discarded.

Glacial influence might has been avoided in southern Brazil if the climate was warm, as has been proposed (e.g. Fragoso Cesar et al. 1984). Then the Neoproterozoic glaciations were not global, but some areas were preserved from those (Hyde et al. 2000). However, this seems to be highly unlike, as southern Brazil was surrounded by regions with glacial influence (Eerola 2001) and the Neoproterozoic glacial record is worldwide (Hoffman et al. 1998). Moreover, southern Brazil has been located near the South Pole during the Neoproterozoic (e.g. Grunow et al. 1996), when glaciers should have been naturally expected there. However, the coeval glacial record seems to be absent from Uruguay (Gaucher 2000).

Eerola (1994) suggested that the relief was higher in Namibia and alpine glaciers were generated in SW Africa, but not in southern Brazil. However, as the widespread Dom Feliciano Belt developed in southern Brazil (Fernandes et al. 1992, Chemale Junior et al. 1995, Frimmel et al. 1996), mountain building and uplift should have been more vigorous in southern Brazil than in Namibia. It is then more probable that alpine glaciers might have been generated in southern Brazil rather than in Namibia.

However, if the glacial record is lacking from southern Brazil, its possible nonpreservation in continental or alpine conditions should be considered. Possible glacial record may have been also deformed, reworked or it may be unexposed, so that its finding and recognition is difficult. However, suggestions on glacial influence should be re-examined.

If the glacial record exists in the region, it should be looked in the metamorphosed and deformed Vacacaí Supergroup (750-700 Ma, Hartmann et al. 2000), where it was suggested by Bocchi (1970). This glacial record may be correlative with the Açungui Group in Paraná State, southern Brazil, where glacials were described by Perdoncini & Soares (1992). Varangerian glacial record may be found at the bases of the Camaquã and Itajaí groups (~600 Ma, Hartmann et al. 2000), where glacial influence was suggested by Eerola (1995) and Carvalho & Pinto (1938).

Conclusions

The existence of the Neoproterozoic glacial record in southern Brazil remains open. This is an intriguing mystery, considering that coeval glacial deposits are so widespread in closely located Namibia and elsewhere. Although there are some suggestions on glacial influence in the region, those are still unconfirmed by subsequent published studies.

Four reasons were suggested here for this apparent absence of glacial deposits in southern Brazil: (1) misinterpretation of deposits in Namibia; (2) no glaciers in southern Brazil, controlled by a) climate or b) topography; (3) nonpreservation of the glacial record or (4) such record has not been found or recognised. It seems that the last two are the most reasonable explanations for such absence.

To obtain answere for this and other related questions, the Neoproterozoic-Cambrian diamictite and lonestone ocurrences should be assessed, mapped and studied in detail in southern Brazil, associated with geochronological, paleontological, paleomagnetic and isotopic investigations. The published and informal suggestions (see Eerola 1995) on glacial influence should be checked at the first place. Further, intensive research on such occurrences in southern Brazil will probably reveal were there Neoproterozoic glaciers or not, and which of the presented reasons are valid, or are there other hypothesis not considered here.

References

Bocchi, P.R. 1970 Geologia da folha Caçapava do Sul. Boletim da Divisão de Geologia e Mineralogia, Rio de Janeiro, 245: 1-83, 1 mapa escala 1:50.000.

Carvalho, P.F. & Pinto, E.A. 1938 Reconhecimento geológico no Estado de Santa Catarina. Boletim do Serviço Geológico e Minaralógico do Brasil, 92, 30 p.

Chemale Júnior, F., Hartmann, L.A. & Da Silva, L.C. 1995 Stratigraphy and tectonism of the Brasiliano Cycle in southern Brazil. Communications of the Geological Survey of Namibia 10: 151-166.

Eerola, T.T. 1994 The Proterozoic to Eopaleozoic supracrustal sequences of the Rio Grande do Sul Shield and the contemporaneous glaciogenic record of the SW Africa: possible tectonic, paleogeographic and paleoclimatic implications. Boletim de Resumos Expandidos, 38^o Congresso Brasileiro de Geologia, Camboriú, SBG, v. 1: 624-625.

Eerola, T.T. 1995 From ophiolites to glaciers? Review on geology of the Neoproterozoic-Cambrian Lavras do Sul region, southern Brazil. In: Autio, S. (ed.) Geological Survey of Finland Current Research 1993-1994. Geological Survey of Finland, Special Paper 20: 5-16.

Eerola, T.T. 2001 Climate changes at the Neoproterozoic-Cambrian transition. In: Zhuravlev, A. & Riding, R. (eds.) The Ecology of the Cambrian Radiation. New York, Columbia University Press, 90-106.

Fernandes, L.A.D., Tommasi, A. & Porcher, C.C. 1992 Deformation patterns in the southern Brazilian branch of the Dom Feliciano Belt, a reappraisal. Journal of South American Earth Sciences 5: 77-96.

Fragoso Cesar, A.R.S., Lavina, E.L., Paim, P.S.G. & Faccini, U.F. 1984 A antefossa molássica do Cinturão Dom Feliciano no Escudo do Rio Grande do Sul (in Portuguese, with English summary). In: Congresso Brasileiro de Geologia, 33, Rio de Janeiro. Anais, SBG, v. 7: 3272-3283.

Frimmel, H.E., Hartnady, C.J.H. & Koller, F. 1996 Geochemistry and tectonic setting of magmatic units in the Pan-African Gariep Belt, Namibia. Chemical Geology, 130: 101-121.

Gaucher, C. 2000 Sedimentology, paleontology, and stratigraphy of the Arroyo del Soldado Group (Vendian to Cambrian, Uruguay. Beringeria 26, 120 p.

Germs, G.J.B. 1995 The Neoproterozoic of southwestern Africa, with emphasis on platform stratigraphy and paleontology. Precambrian Research 73: 137-151.

Grunow, A., Hanson, R. & Wilson, T. 1996 Were aspects of Pan-African deformation linked to Iapetus opening? Geology, 24 (12): 1063-1066.

Hartmann, L.A., Leite, J.A.D., Da Silva, L.C., Remus, M.V.D., McNaughton, N.J., Groves,

D.I., Fletcher, R., Santos, J.O.S. & Vasconcellos, M..A.Z. 2000 Advances in SHRIMP geochronology and their impact on understanding the tectonic and metallogenic evolution of southern Brazil. Australian Journal of Earth Sciences 47 (5): 829-844.

Hoffman, P.F., Kaufman, A.J., Halverson, G.P. & Schrag, D.P. 1998 A Neoproterozoic snow-ball Earth. Science 281: 1342-1346.

Hyde, W.T., Crowley, T.J., Baum, S.K. & Peltier, W.R. 2000 Neoproterozoic “snowball Earth” simulations with acoupled climate/ice-sheet model. Nature 405: 425-429.

Perdoncini, L. C. & Soares, P.C. 1992 Depósitos glaciogênicos do Proterozóico superior no leste do Paraná. In: Congresso Brasileiro de Geologia, 37, São Paulo, 1992, Boletim de Resumos Expandidos. São Paulo, SBG, v. 2: 452-453.

Chemostratigraphy of the lower Arroyo del Soldado Group (Vendian, Uruguay): a potential reference section for W-Gondwana and its implications

Claudio Gaucher*; Alcides Nóbrega Sial**; Gonzalo Blanco* & Peter Sprechmann*

*Departamento de Paleontología, Facultad de Ciencias, Iguá 4225, 11400 Montevideo, Uruguay. e-mail: gaucher@chasque.apc.org

**Departamento de Geología, Centro de Tecnologia e Geociências, Universidade Federal de Pernambuco

(UFPE), Caixa Postal 7852, Recife PE-50732-970, Brasil.

Introduction

The Arroyo del Soldado Group (ASG) is a 5 km-thick platform sequence that overlies with angular and erosional unconformity an Archean to Mesoproterozoic basement of the Nico Pérez Terrane of Uruguay (Gaucher et al. 1996, 1998, Gaucher 2000). It is notable for its extensive and very thick carbonate deposits, namely (a) the Polanco Formation, with maximum thickness of more than 900 m, (b) the Cerro Victoria Formation, composed of up to 150 m of stromatolitic and oolitic limestones, and (c) various isolated dolomite and limestone strata in the lower Cerro Espuelitas Formation and upper Yerbal Formation (Fig. 1). These are excellent prerequisites for chemostratigraphic studies in a time of extreme biogeochemical oscillations, environmental instability and biological innovations. Carbon and oxygen isotopic data from the ASG have been reported by Boggiani (1998), Kawashita et al. (1999), Gaucher (2000) and Gaucher et al. (2001). Kawashita el al. (1999) present four Sr-isotopic determinations from the lowermost Polanco Formation. We report here the preliminary results of a series of 50 C- and O-isotopic determinations from the Yerbal, Polanco, lower Barriga Negra and lower Cerro Espuelitas formations, and combine them with available data, as well as with reported Sr-isotopic data.

Character of isotopic signatures

The carbonates in the studied sections meet the criteria proposed by Jacobsen & Kaufman (1999) to distinguish between primary and diagenetic signatures, namely: (1) the sampled micro-domains in the carbonates were chosen after careful petrographic examination, only from primary mineral phases; (2) d¹⁸O values >-10 $ PDB in ca. 90% of the analyzed samples; and (3) Mn/Sr relations range between 0.03 and 0.62 (50 analyses), and are thus clearly below the boundary of 2 accepted for non-altered carbonates. Therefore, the primary nature of the C- and Sr-isotopic data is corroborated by the above mentioned facts.

Chemostratigraphy

The sections studied are: Arroyo Campanero (Gaucher 2000: 25) for the Yerbal Formation; Arroyo Yerbalito Syncline (Gaucher 2000: 23), Calera de Recalde Syncline (Fig. 1) and Cerro Espuelitas Syncline for the Polanco Formation; the stratotype and parastratotype (Gaucher 2000) of the Barriga Negra Formation and the Tapes Grande Syncline (Gaucher 2000) for both the Polanco and Cerro Espuelitas Formation.

Dolostones of the upper Yerbal Formation underlying Cloudina, Waltheria, Titanotheca, Bavlinella and Soldadophycus-bearing siltstones yielded positive d¹³C values of up to +2.15 $ PDB, and increasing up-section. This trend continues into limestone/dolostone rhytmites of the lowermost Polanco Formation, reaching peak values (P₁) of +5.3 $ PDB 180 m above the contact with the Yerbal Formation (Fig. 1). Basal rhytmites of the Polanco Formation yielded ⁸⁷Sr/⁸⁶Sr values of 0.7078 according to Kawashita et al. (1999). An impressive negative d¹³C-excursion to –3.3 $ PDB follows (N₂), and corresponds to an interval dominated by calcareous tempestites (Fig. 1). Up section, a positive d¹³C-peak of +2.1 $ PDB occurs (P₂), comprising black, carbonaceous and laminated calcisiltites (Fig. 1).The return to slightly negative values (-0.70 $ PDB) is recorded in calcareous tempestites 650 m above the base of the Polanco Formation in the Calera de Recalde Syncline (N₃). At that section, uppermost tempestitic calcarenites and interbedded rhytmites show positive d¹³C-values of up to +2.8 $ PDB (P₃, Fig. 1). The contact between the Polanco and Barriga Negra formations is characterized by a negative d¹³C-excursion to –1.9 $ PDB, which continues into the lowermost Barriga Negra Formation (N₄). Carbonates interbedded in the lower-middle Cerro Espuelitas Formation at its stratotype yielded values of +2.4 $ PDB, probably indicating a new positive excursion.

The isotopic signal at the stratotype of the Polanco Formation in the Tapes Grande Syncline shows approximately the same peaks described above, but shifted to negative values. This can be interpreted as a deep-water signal (Calver 2000), as also suggested by sedimentary structures and facies (Gaucher 2000). Therefore, it is here proposed to locate the chemostratotype of the Polanco Formation in the Calera de Recalde Syncline (Fig. 1), because of the unequivocal preservation and interpretation of primary isotopic signatures.

Interpretation

The value of 0.70780 is considered by Kawashita et al. (1999) as the most confident estimate of the primary ⁸⁷Sr/⁸⁶Sr-ratio. This value corresponds according to Jacobsen & Kaufman (1999) and Walter et al. (2000: 391) to an age of 580 Ma, which is taken here as the age of the base of the Polanco Formation. This corroborates the biostratigraphic data (Cloudina and Bavlinella-Soldadophycus assemblage), which indicate a post-Varangerian/Marinoan (600-590 Ma) age for the whole lower ASG. The upper Yerbal-lower Polanco positive excursion (+5.3 $ PDB) corresponds to a time interval between 585 and 574 Ma in the global d 13C-curves of Jacobsen & Kaufman (1999: 48) and Walter et al. (2000: 414), corroborating the Sr-isotopic data mentioned above. The N₂ peak of the middle Polanco Formation (-3.3 $ PDB) matches well a negative excursion at 573 Ma in the global curve (Jacobsen & Kaufman 1999, Walter et al. 2000). Peaks P₂to N₄ of the upper Polanco Formation and lowermost Barriga Negra-Cerro Espuelitas formations, probably correspond to the 570-550 Ma interval in the curve of Walter et al. (2000), although the upper date is still uncertain. It is noteworthy that the curve here reported for the Polanco Formation shows more details than most of the curves reported for the period, and could be due to the great thickness and almost pure carbonates typical of the unit.

From the sedimentlogical point of view, the relationship between C-isotopic composition and facies of the Polanco Formation is evident. The finer, more dolomitic and organic-carbon rich lithologies show predominantly positive values, while the intervals dominated by tempestites are characterized by ¹³C-depleted carbonates. This is interpreted here as a result of climatic-driven, eustatic sea-level oscillations. It is well known (Frimmel et al., in press) that carbonates associated with Neoproterozoic glacigenic rocks are strongly ¹³C-depleted. The predominance of tempestites in the negative d 13C intervals of the Polanco Formation points to a moderate lowering of palaeobathymetry (above storm wave-

Fig. 1: Stratigraphic column and chemostratigraphy of the Polanco Formation in the Calera de Recalde Syncline.

The Tandilia System is an orographic belt located in the Buenos Aires Province, between latitudes 36º 30´ - 38º 10´ South and longitudes 57º 30´ - 61º West. Its maximum length is 350 km in the NW-SE direction. The hills are composed of an igneous metamorphic basement and a Precambrian and Lower Palaeozoic sedimentary cover, which shows a horizontal to subhorizontal arrangement (<12º).

The Precambrian sedimentary succession lies in the north-western part, around Olavarría and Barker-San Manuel region, whereas the Lower Palaeozoic crops out in two areas: a western belt of the range and mainly towards the south-east (Balcarce-Mar del Plata area). These deposits lie on a crystalline basement (Buenos Aires Complex), which is over 2,000 Ma old and is composed of granitoids, migmatites, ectinites, milonites, anphibolites and basic dykes.

The stratigraphic scheme shows a sequential arrangement composed of three Riphean sequences, a Riphean-Vendian sequence and a final Ordovician sequence. From the lithostratigraphic standpoint the Precambrian sedimentary successions comprise: a) the Villa Mónica Formation, b) the Cerro Largo Formation, and c) the Loma Negra Formation, all of these being constituents of the Sierras Bayas Group, and d) the Cerro Negro Formation and its equivalent Las Aguilas Formation. The Lower Palaeozoic succession is known as the Balcarce Formation.

These lithostratigraphic units were grouped into five depositional sequences: the Tofoletti (I), Malegni (II) and Villa Fortabat (III) sequences are Riphean, the La Providencia sequence (IV) is Riphean-Vendian and the Batán sequence (V) is Cambrian-Ordovician.

Between the crystalline basement and the sedimentary cover, arkosic and quartz-kaolinitic saprolites indicate palaeoweathering surfaces. The presence of diamictites between the crystalline basement and the Balcarce Formation is a peculiar feature reported in the El Volcán Hill.

All these units being unfossiliferous, they carry biogenic sedimentary structures (trace fossils and stromatolites) and acritarchs as the only evidence of biocoenosis in the Precambrian and Lower Palaeozoic seas of this region.

The focus of this contribution is to give an overview of the main sedimentary features of the Precambrian units in the Olavarría region and talk about their biological evidences (stromatolites, trace fossils and microfossils.

UPPER PRECAMBRIAN SEDIMENTARY ROCKS

The Upper Precambrian sedimentary cover of Tandilia in Sierras Bayas (close to Olavarría) is a 167 m thick succession (Sierras Bayas Group) composed of three depositional sequences separated by regional unconformities.

The oldest depositional sequence (Tofoletti, 52 m thick) shows two sedimentary facies associations: a) quartz-arkosic arenites to the base and b) dolostones and shales to the top. The former is composed of shallow marine siliciclastic rocks (conglomerates, quartz and arkosic sandstones, diamictites and shales), and the latter is characterised by shallow marine stromatolitic dolostones and shales. This sequence has been dated by stromatolites and Rb/Sr ages in 800-900 Ma.

The second depositional sequence (Malegni, 75 m thick) consists of a basal succession composed of chert breccia, fine-stratified glauconitic shales and fine-grained sandstones, followed by cross-bedded quartz arenites which are in turn covered by siltstones and claystones. This sequence represents a shallowing upward succession from subtidal nearshore to intertidal flat deposits. An age between 700-800 Ma has been defined from Rb/Sr dating.

The youngest Riphean depositional sequence (Villa Fortabat) is a 40 m thick unit composed almost exclusively of red and black micritic limestones, originated by suspension fall-out in open marine ramp and lagoonal environments.

On top of the Sierras Bayas Group a regional unconformity is recognised. This surface has been related with sea-level drop. Meteoric dissolution of the carbonatic sediments is interpreted as a karstic surface on which residual clays and brecciated chert accumulated.

The Vendian Cerro Negro Formation (La Providencia Depositional Sequence) appears on top of the above described unconformity. It is a more than 100 m thick unit characterised by claystones and heterolithic, fine-grained sandstone–claystone interbeds, mainly formed in upper to lower intertidal flats. Through radiometric dating (680Ma) and the presence of acritarchs the underlying Cerro Negro Formation has been dated in the Vendian.

STROMATOLITES

The dolostones of the Villa Mónica Formation support a very good assemblage of stromatolites, which is composed of Colonella fm., Conophyton ?ressootti, Conophyton fm., Cryptozoon fm., Gongylina fm., Gymnosolem fm., Inzeria fm., Jacutophyton fm., Jurusonia nisvensis, Katavia fm., Kotuikania fm., Kussiella fm., Minjaria fm., Parmites fm., Parmites cf. cocrescens and Stratifera fm. (Poiré, 1987, 1989, 1993).

These stromatolites have been studied on four different scales (Poiré, 1987, 1989): megascale, mesoscale, small scale and microscale. The megascale shows that the lower and the upper part of the dolostones are mainly composed of 0.5-1m thick domal biostromes and 0.05-0.30m thick interbiostromal green shales. A few bioherms are present in the top of the unit. In the middle part of the dolostones they are absent, with only one level of Cryptozoon fm. In the mesoscale, two kind of patterns have been distinguished dependent of the vertical and lateral arrangement of the stromatolites inside the biostromes. In one case, when the stromatolites do not present any variation along the bed, these bioconstructions have been named “monostromatolite layers”. In the other case, “stromatolites cycles” are defined when there are changes in stromatolite morphology. Thus, three “monostromatolite layers” and six “stromatolite cycles” have been identified in the Villa Mónica Formation. The small scale has been used to look at the gross morphology and to classify stromotolites (Groups) and the microscale for the microstructure (forms).

Sedimentologically, Poiré (1987, 1990) considered that each stromatolite cycle records changes in the hydrodynamic conditions during its growth, due to sea level fluctuations. Almost all of them are interpreted as shallowing-upward sequences, while others are considered as originating by deepening processes whenever the group Conophyton is deeper than convex-columnar stromatolites.

TRACE FOSSILS

References about the Precambrian trace fossils of Tandilia are scarce. Palaeophycus isp. and Didymaulichnus isp. have been described in the Cerro Largo Formation by Poiré et al. (1984), while Helminthopsis isp. and probable medusa resting traces have been mentioned for the Loma Negra Formation. Skolithos isp. have recently been registered in the lower part of the Cerro Negro Formation.

In the Precambrian units, trace fossils are scarce and show a poor ichnodiversity. Palaeophycus isp. and Didymaulichnus isp. have been described in the Cerro Largo Formation (Poiré et al., 1984), while Helminthopsis isp. and probable medusa resting traces have been found in the Loma Negra Formation. Skolithos isp. has recently been registered in the lower part of the Cerro Negro Formation.

On the other hand, trace fossils from the Lower Paleozoic Balcarce Formation are very abundant and show a high ichnodiversity: The following updated list of trace fossils was recognised: Ancorichnus ancorichnus, Arthrophycus alleghaniensis, Arthrophycus isp., Bergaueria isp., Cochlichnus isp., Conostichus isp., Cruziana furcifera, Cruziana isp., Daedalus labeckei, Didymaulichnus lyelli, Didymaulichnus isp., Diplichnites isp., Diplocraterion isp., Herradurichnus scagliai, ?Monocreterion isp., Monomorphichnus isp., Palaeophycus alternatus, Palaeophycus tubularis, Palaeophycus isp., Phycodes aff. pedum, Phycodes isp., Plagiogmus isp., Planolites isp., Rusophycus isp., Scolicia isp. and Teichichnus isp.

MICROFOSSILS

A microfossil assemblage was described by Pothe de Baldis et al. (1983) from the upper shales of the Cerro Largo Formation. This are composed Paleorivularia ontarica, Chuaria olavarriensis and Leiosphaeridia sp.

Acritarchs were also recorded by Cingolani et al. (1991) from the lower part of Cerro Negro Formation, which consists of simple forms of Sphaeromorphitae: Synphaeridium sp., Trachysphaeridium sp. and Leiosphaeridia sp. By K/Ar analysis, these authors have obtained an age younger than 680 Ma. for the microfossil- bearing sedimentary rocks.

AGE OF THE SEDIMENTARY COVER

There are few elements to assign a precise age to the Precambrian units. The dolostones of Villa Mónica Formation have been dated by stromatolites and Rb/Sr ages in 800-900 Ma. An age between 700-800 Ma has been defined from Rb/Sr dating for Cerro Largo Formation, while through radiometric dating (680Ma) and the presence of acritarchs the Cerro Negro Formation has been dated in the Vendian (Cingolani et al., 1991).

The precise age of the Balcarce Formation is difficult to determine as well. The underlying init (Cerro Negro) is Vendian and the upper limit of the Balcarce Formation is sustained by an intrusive diabase body dated around 450 Ma and 498 Ma (Rapela et al., 1974). Consequently the unfossiliferous Balcarce Formation would be assigned to the Cambrian-Ordovician lapse. During the seventies and eighties, Cruziana has been considered a useful biostratigraphic indicator. Based on this concept, the presence of Cruziana furcifera has been one of the most substantial elements to accept a Lower Ordovician age for the Balcarce Formation. In the nineties C. furcifera was found in the Lower Cambrian of Alberta, Canada. Therefore, the range of this ichnospecies could be extended at least from the Lower Cambrian up to the Lower-Middle Ordovician. On the othar hand, the presence of Plagiogmus isp. would strongly indicate a Cambrian age.

REFERENCES

BONHOMME, M. and CINGOLANI, C., 1980. Mineralogía y geocronología Rb-Sr y K-Ar de fracciones finas de la “Formación La Tinta”, provincia de Buenos Aires. Asociación Geológica Argentina Revista 35: 519-538. Buenos Aires.

CINGOLANI, C.A., RAUSCHER, R. and BONHOMME, M., 1991. Grupo La Tinta (Precámbrico y Paleozoico inferior) provincia de Buenos Aires, República Argentina. Nuevos datos geocronológicos y micropaleontológicos en las sedimentitas de Villa Cacique, partido de Juarez. Revista Técnica de YPFB 12(2): 177-191, Santa Cruz.

IÑIGUEZ, A.M., DEL VALLE, A., POIRÉ, D.G., SPALLETTI, L.A. and ZALBA, P.E., 1989. Cuenca Precámbrica / Paleozoica inferior de Tandilia, Provincia de Buenos Aires. In: Chebli, G. and Spalletti, L. (Eds.) Cuencas Sedimentarias Argentinas. Serie Correlación Geológica 6 (Universidad Nacional de Tucumán): 245-263. S. M. de Tucumán.

POIRÉ, D.G., 1987. Mineralogía y sedimentología de la Formación Sierras Bayas en el núcleo septentrional de las sierras homónimas, Partido de Olavarría, Provincia de Buenos Aires. PhD Thesis, Facultad de Ciencias Naturales y Museo, Universidad Nacional de La Plata: 494: 271 pp. La Plata (unpublished).

POIRE, D.G. Stromatolites of the Sierras Bayas Group, Upper proterozoic of Olavarría, Sierras Septentrionales, Argentina. Stromatolite Newsletter 14:58-61.

POIRÉ, D.G., 1993. Estratigrafía del Precámbrico sedimentario de Olavarría, Sierras Bayas, provincia de Buenos Aires, Argentina. XII Congreso Geológico Argentino and II Congreso Exploración de Hidrocarburos Actas II: 1-11. Mendoza.

POTHE de BALDIS, E.D., BALDIS, B.A. and CUOMO, J. Los fósiles precámbricos de la Formación Sierras Bayas (Olavarría) y su importancia intercontinental. Revista de la Asociación Geológica Argentina 38(1): 73-83, Buenos Aires.

RAPELA, C.W., DALLA SALDA, L. and CINGOLANI, C., 1974. Un intrusivo básico ordovícico en la “Formación La Tinta” (Sa. De los Barrientos, Buenos Aires). Revista de la Asociación Geológica Argentina 29(3):319-331, Buenos Aires.

There is a fairly high preservation potential for microbial mat signatures in Neoproterozoic peritidal siliciclastic sediments. The good potential for preservation of such structures is physico-biologically controlled, since microbial mats colonizing siliciclastic sediment surfaces lead to ‘biostabilization’ of the surface layer and thus lesser erodability of the sediment. The probability of preservation is increased if the mats are buried underneath land-directed storm deposits (tempestites), as documented from the Recent by Reineck & Gerdes (1997). In analogy to modern siliciclastic peritidal domains where microbial mats preferably grow in the (upper) intertidal to supratidal zones including coastal lagoons (e.g. Gunatilaka, 1975; Horodyski & Bloeser, 1977; Park, 1977; Gerdes & Krumbein, 1987; Gerdes et al., 2000), it is envisaged that Neoproterozoic mats also colonized these zones of dominantly low-energy hydrodynamic conditions. It may furthermore be expected that, in the absence of destructing organisms, Neoproterozoic microbial mats will extensively have grown in these zones thus providing, quite numerically, favourable prerequisites for the preservation of related structures. Viceversa, if mat-related structures are preserved and identified, they may serve as sensitive facies indicators in peritidal siliciclastic settings.

The fundamental process for the formation and preservation of ‘microbially induced sedimentary structures’ (‘MISS’; Noffke et al., 1996; 2001a, b) is ‘biostabilization’ of the siliciclastic sediment surface, which means a fixation of the sediment grains by cyanobacterial filaments and extracellular mucilages of the microbes (e.g. Paterson & Black, 2000). The process is related to the development of microbial films (‘biofilms’) which adhere to the grains and transform the upper few millimetres of the sediment into a cohesive clastic layer. At suitable conditions, biofilms may evolve into microbial mats forming gelatinous to leathery organic layers on the sediment surface (e.g. Krumbein, 1983; Stolz, 2000).

Biostabilization is thus partly an early diagenetic process of ‘bio-agglutination’ acting on siliciclastic sediment surfaces. It changes the physical properties of the sediment surface which becomes resistant against ‘normal’ intertidal wave action and is not disintegrated into individual sedimentary grains. Resistance increases with mat growth but may, at any stage, be overcome by sufficiently high current velocities, e.g. during spring tides or storms. In such cases, the biofilm or mat may be deformed and partly destroyed. Resulting structures are certain types of ‘wrinkle marks’ (Fig. 1a, b) as well as ‘erosion marks’ (Fig. 1c), ‘microbial sand chips’ (Fig. 1e), and ‘microbial mat chips’ (e.g. Hagadorn & Bottjer, 1997, 1999; Gerdes & Krumbein, 1987; Pflüger & Gresse, 1996).

The cohesive layer or mat atop the sediment is also able to trap ascending gasses, mainly methane, released from buried decaying organic material. Resulting structures are ‘gas domes’ and specific types of wrinkle structures (e.g. Reineck et al., 1990; Gerdes et al., 1993; Pflüger, 1999). Likewise, the mat forms a barrier for ascending water released from the underlying sediments and particularly from buried mats during compaction. The resulting increase of the water/sediment ratio may lead to destabilization and increased tendency for liquefaction of the sediment immediately underlying the mat. This may play an important role in the preservation of wrinkle structures (Fig. 1c).

A critical situation affecting microbial mats is their subaerial exposure which is most likely to occur in the upper intertidal and supratidal zones. Desiccation of subaerially exposed mats induces shrinkage of the water-rich organic layer and may produce cracks. Strongly desiccated mats develop large polygonal patterns of wide desiccation cracks (e.g. Gerdes et al., 2000). However, as capillary water is ascending and moistening the mat from below during the initial stages of subaerial exposure, desiccation and shrinkage will proceed slowly in the beginning. However, at a critical stage cracking of the mat surface will occur, frequently starting with triple junctions or isolated sigmoidal fractures (e.g. Porada & Löffler, 2000). With proceeding desiccation and shrinkage, new cracks are continuously forming, whereas propagation of existing cracks, typically following curvilinear paths, may lead to more or less evolved networks (Fig. 1f, g). If the next flooding event occurs at this stage of still incomplete desiccation, the cracks will be filled with sediment whereas the mat will recover and overgrow the crack fillings. In this way, the trapped crack fillings will be preserved as Manchuriophycus type ‘microbial shrinkage cracks’, which are characterized by their independence and easy separability from the beds below and above.

Wrinkle structures and Manchuriophycus type cracks have been described from shallow marine siliciclastic sediments since the Paleoproterozoic and up to the Jurassic (e.g. Frarey & McLaren, 1963; Bloos, 1976). However, there appears to be a maximum of occurrence in the Neoproterozoic, between ca 900 and 545 Ma, and a rapid decrease during the Cambrian related to the ‘Cambrian substrate revolution (Bottjer et, al., 2000). The Neoproterozoic era includes two major, possibly worldwide glaciations, the Sturtian-Rapitan glacial interval (ca. 750-720 Ma) and the Marinoan-Varanger glacial interval (ca. 610-600 Ma). These will probably have widely restricted the distribution of microbial mats in peritidal settings and may have had an impact on the biological evolution in general. Under such aspects it might be of interest to check if mat-related signatures of pre-Sturtian age (900-750 Ma) differ from those of Vendian age (600-545 Ma).

References:

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Bottjer, D.J., Hagadorn, J.W., Dormbos, S.Q. 2000. The Cambrian substrate revolution. GSA Today, 10.

Frarey, M.J. and McLaren, D.J. 1963. Possible metazoans from the Early Proterozoic of the Canadian Shield. Nature, 200, 461-462.

Gerdes, G., Krumbein, W.E. 1987. Biolaminated Deposits. In: Bhattacharya, G.M., Friedmann, G.M., Neugebauer, H.J., Seilacher, A. (Eds.). Lecture Notes in Earth Sciences 9, Springer-Verlag, Berlin, pp. 1-183.

Gerdes, G., Claes, M., Dunajtschik-Piewak, K., Riege, H., Krumbein, W.E. and Reineck, H.-E. 1993. Contribution of microbial mats to sedimentary surface structures. Facies 29, 61-74.

Gerdes, G., Klenke, Th., Noffke, N. 2000. Microbial signatures in peritidal siliciclastic sediments: a catalogue. Sedimentology 47, 279-308.

Gunatilaka, A. 1975. Some aspects of the biology and sedimentology of laminated algal mats from Mannar Lagoon, northwest Ceylon. Sedimentary Geol. 14, 275-300.

Hagadorn, J.W., Bottjer, D.J. 1997. Wrinkle structures: Microbially mediated sedimentary structures common in subtidal siliciclastic settings at the Proterozoic-Phanerozoic transition. Geology 25, 1047-1050.

Hagadorn, J.W., Bottjer, D.J. 1999. Restriction of a late Neoproterozoic biotope: suspect-microbial structures and trace fossils at the Vendian-Cambrian transition. Palaios 14, 73-85.

Horodyski, R.J., Bloeser, B. 1977. Laminated algal mats from a coastal lagoon, Laguna Mormona, Baja California, Mexico. J. Sed. Petrol. 47, 680-696.

Krumbein, W.E. 1983. Stromatolites –The challenge of a term in time and space. Precambrian Res., 20, 493-531.

Noffke, N., Gerdes, G., Klenke, T., Krumbein, W.E. 1996. Microbially induced sedimentary structures – examples from modern sediments of siliciclastic tidal flats. Zentralbl. Geol. Paläont. I, 1995 (1/2), 307-316.

Noffke, N., Gerdes, G., Klenke, T., Krumbein, W.E. 2001a. Microbially induced sedimentary structures –a new category within the classification of primary sedimentary structures. J. Sedimentary Res. 71, 649-656.

Noffke, N., Gerdes, G., Klenke, T., Krumbein, W.E. 2001b. Microbially induced sedimentary structures indicating climatological, hydrological and depositional conditions within Recent and Pleistocene coastal facies zones (southern Tunesia). Facies 44, 23-30.

Park, R.K. 1977. The preservation potential of some recent stromatolites. Sedimentology 24, 485-506.

Paterson,D.M. and Black, K.S. 2000.Siliciclastic intertidal microbial sediments. In: Riding, R. and Awramik, S. (Eds.) Microbial Sediments. Springer-Verlag, Berlin, p. 217-225.

Pflüger, F. 1999. Matground structures and redox facies. Palaios 14, 25-39.

Pflüger, F., Gresse, P.G. 1996. Microbial sand chips - a non-actualistic sedimentary structure. Sedimentary Geol. 102, 263-274.

Porada, H., Löffler, T. 2000. Microbial shrinkage cracks in siliciclastic rocks of the Neoproterozoic Nosib Group (Damara Supergroup) of central Namibia. Communs. Geol. Surv. Namibia 12, 63-72.

Reineck, H.E., Gerdes, G. 1997. Tempestites in Recent shelf and tidal environments of the southern North Sea. Cour. Forsch.-Inst. Senckenberg 201, 361-370.

Reineck, H.E., Gerdes, G., Claes, M., Dunajtschik, K., Riege, H., Krumbein, W.E. 1990. Microbial modification of sedimentary surface structures. In: Heling, D., Rothe, P., Förstner, U., Stoffers, P. (Eds.). Sediments and Environmental Geochemistry. Springer-Verlag, Berlin, pp. 254-276.

Stolz, J.F. 2000. Structure of microbial mats and biofilms. In: Riding, R. and Awramik, S. (Eds.) Microbial Sediments. Springer-Verlag, Berlin, p. 1-8.

Introducción

El Terreno Punta del Este (TPE) esta constituido por una serie de gneises y migmatitas formadas en el intervalo de 1000 y 900 Ma, y que han sido intensamente retrabajadas durante la orogénesis Rio Doce (ca. 600-500 Ma). Este segmento cortical representa un terreno metamórfico de alto grado, correlacionable con complejos gnéisicos del Suroeste de África, en particular con el Cinturón Kibariáno - Namaqualano, reconocido en la porción sudoccidental del continente africano (Namibia). Edades U-Pb en circón de granitoides tonalíticos indican valores entre 1000 y 900 Ma y fueron interpretados como indicativas del momento de generación de estás rocas. Esta debió ser también la edad del metamorfismo de alto grado que afecto una gran parte de las rocas gnéisicas de la región. Por otra parte, los movilizados anatécticos relacionados con leucosomas de migmatitas dieron edades ca. 520-540 Ma, indicando que las condiciones metamórficas (superpuestas) durante la orogénesis Rio Doce alcanzaron por lo menos la facies anfibolita. La cobertura metasedimentaria del TPE ocurre en la proximidad de las ciudades de La Paloma y Rocha. Estas se encuentran representadas por una sucesión metasedimentaria siliciclástica representada por la Formación Rocha. A pesar de la deformación polifásica y del bajo grado de metamorfismo que afecta a esta formación, estructuras primarías son frecuentes tales como estratificación plano-paralela, estratificación cruzada, hummocky y niveles masivos con estratificación gradacional. Así como el basamento del TPE se correlaciona con el Complejo Gnéissico de Namaqua, la Formación Rocha podría correlacionarse con las supracorticales del Grupo Gariep. La Formación Cerro Aguirre representa una secuencia volcánica de composición intermedia a ácida (Campal & Gancio, 1993), que sufrió una compresión generando pliegues abiertos con orientación axial N30-40E, y desarrollo local de un clivaje de plano axial. Muchos granitoides isótropos circunscritos, representan la manifestáción magmática más importante que afecta al TPE. Esos granitoides de filiación alcalina están representados por los plutones de José Ignacio y Santa Teresa. Las edades Rb-Sr dadas por isócronas combinando minerales y datos de roca total, así como obtenidas exclusivamente en minerales presentan valores entre 611-590 Ma y 550-537 Ma, para Jose Ignacio y Santa Teresa respectivamente. La cuenca Gariep–Rocha no debió tener una apertura oceánica considerable. Por eso, la rama principal del océano Adamastor se desarrollo al Oeste del cinturón Gariep – Rocha y su consumo generó el cinturón granítico El cerramiento de la cuenca Gariep – Rocha y la deformación subsiguiente de la pila supracortical ocurrio ca. 545 - 570 Ma, como consecuencia del transporte hacia el este del cinturón Gariep sobre los terrenos africanos. Es asígnada a este periódo la obducción de las rocas volcánicas (corteza oceánica) del terreno Marmora sobre las unidades metasedimentarias de la zona de Port Nolloth. Es posible que la separación del TPE y su equivalente africano ocurriera únicamente durante el desmembramiento de la Pangea y la apertura del Océano Atlántico.

Geología del Terreno Punta del Este (TPE)

El Terreno Punta del Este (TPE) está separado de los granitoides neoproterozoicoos del Batolito de Aiguá, o granitoides centrales, por la Zona de Cizalla Alferez – Cordillera representada por una delgada faja de milonitas y de gneises miloníticos en condiciones de anatexis. Granitoides con tendencia subalcalina se encuentran a lo largo de ese lineamiento. La foliación milonítica presenta una dirección general entre N15E y N40E, con inclinación subvertical en dónde se observa el desarrollo de cuarzo en bandas dúctiles y moscovita. Las lineaciones de estiramiento observadas presentan inclinaciones de 20º a 30º para el SW. Los indicadores cinemáticos sugieren un movimiento oblicuo predominantemente dextrógiro.

Dominio de basamento

El basamento del TPE está constituido, esencialmente, por granitoides deformados con biotita y moscovita, granitoides oftalmíticos y ortogneisses en la porción central, mostrando una atenuación de esa deformación llegando a términos isótropos. Desde el punto de vista de la composición consisten en granitos a biotita porfiríticos con variaciones diversas, ocurriendo subordinadamente, términos dioríticos. Mineralógicamente, están constituidos por cuarzo con extinción ondulante, plagioclasa (oligoclasa a andesina), biotita distribuida por toda la matriz formando niveles asociados con hornblenda y más raramente, asociada con moscovita en crecimientos subparalelos. Los principales accesorios son granate, cordierita, titanita, apatito, circón, magnetita y epidota. Estos granitoides deformados ocurren en medio de gneises regionales con estructuración NE-SW. Esas rocas pueden ser agrupadas en tres unidades mayores: Unidad Cerro Olivo, caracterizado por presentar un predominio de leuco-gneises oftalmíticos, generalmente con biotita, moscovita y granate. Presentan abundante cantidad de cuarzo, microclina pertítica en cristales subautomorfos, generalmente sericitizados, oligoclasa subautomorfa con diferentes grados de alteración, biotita y moscovita en crecimiento epitáxico. La biotita suele presentar fenómenos de desferrificación y cloritización. Como principales accesorios se encuentran granate, titanita, magnetita, circón y epidota. La Unidad Cerro Centinela representada por gneises graníticos con estructuras bandeadas y segregación de máficos; y finalmente la Unidad Chafalote, incluye un conjunto de rocas máficas ricas en biotita con porcentaje variables de cuarzo, plagioclasa, granate, moscovita y ocasionalmente anfíbol. En general son rocas gruesas, granulares, foliadas que frecuentemente presentan enclaves de milonitas graníticas. De la misma manera que en los conjuntos anteriores las estructuras principales tienen dirección NE-SW. Paralela a esas direcciones son observados paragneises de alto grado metamórfico frecuentemente con almandino, cuarzo, cordierita y silimanita. Ocasionalmente pueden ser observados gneises calco-silicatados con bandas ricas en cuarzo y plagioclasa y niveles dónde predominan los anfíboles. Ortoanfibolitas hornblendíticas con andesina y subordinamente magnetita, zoisita y microclina, se desarrollan como cuerpos alargados en medio de los gneises regionales. En esas rocas son observados piroxenos parcialmente reabsorbidos por el anfíbol. En medio de las rocas gnéisicas descriptas ocurren áreas dónde predominan migmatitas de diversas estructuras. Predominan migmatitas biotíticas a veces con anfíbol y con estructuras dominantemente estromáticas a flebíticas y variaciones para los términos nebulíticos. El leucosoma está constituido por niveles félsicos dónde el cuarzo y la microclina-oligoclasa son los minerales predominantes con biotita pudiendo llegar hasta 25%. Las bandas máficas pueden contener moscovita, almandino, silimanita, circón, apatito y magnetita.

Dominio de Cobertura

La principal cobertura metasedimentaría observada en el TPE está representada por la Formación Rocha (Sánchez Bettucci & Mezzano, 1993) que ocurre, en la región homónima, en una faja NE con 20 a 30 km de ancho y cerca de 120 km de extensión. Esta unidad está constituida por un conjunto de metamorfitas de bajo grado. Litológicamente se caracteriza por metasedimentitas siliciclásticas de granulometrías variada (metapelitas a metapsamitas) que sufrieron una evolución tectónica polifásica. Presenta un metamórfismo de bajo grado, que localmente alcanza el grado medio. En los términos menos metamórficos es frecuente la presencia de estructuras primarias tales como estratificación cruzada y plano-paralelas, cruzadas, hummocky y niveles de metareniscas con estratificación gradacional. Las estructuras primarias sugieren que el conjunto fue formado en ambiente de aguas rasas cuya y afectada por ondas de tempestad. Sirviéndose de informaciones sedimentarias Restos de la cobertura metasedimentaria son observados aisladamente. El caso más conspicuo esta en las proximidades de la ciudad de San Carlos y se encuentran intensamente afectadas por la zona de cizalla de Sierra Ballena. Estos niveles están constituidos por conglomerados polimícticos en la base que pasan a metarcosas con intercalaciones de esquistos sericíticos. La Formación Cerros Aguirre representa una cuenca volcano-sedimentaria. Consiste en un conjunto de rocas piroclásticas intermediarías a ácidas en dónde predominan ignimbritas de composición riolítica - dacítica y tobas de diferentes granulometrías y texturas. Ese conjunto, a pesar de ser posterior al desarrollo de la Formación Rocha exhibe un plegamiento con orientación axial N30-40E, asociado al cual ocurre el desarrollo de una esquistosidad de plano axial. El espesor máximo de este conjunto es de orden de 1200 m. Una fotointerpretación cuidadosa permite observar estructuras de interferencia de pliegues de Tipo III de Ramsay, indicando una evolución compleja de la deformación de éstas volcanitas, vinculadas con el movimiento de la zona de cizalla Laguna de Rocha. Numerosos granitos intrusivos son observados a lo largo de toda la extensión del TPE, ocurriendo como cuerpos circunscriptos de dimensiones batolíticas a pequeños stocks. Son cuerpos isótropos, en general leucocráticos, que representan la última actividad magmática importante observada en el TPE. Se encuentran intruyendo a las metasedimentitas de la Formación Rocha. Presentan edades entre 540-580 Ma y 530 a 680 Ma en intrusivos dentro del complejo gnéisico-migmatítico. Los principales cuerpos son Santa Teresa, José Ignacio, Rocha y Alferez..

Geocronologia

Resultados U-Pb, Rb-Sr y K-Ar

Los datos geocronológicos disponibles para el TPE se distribuyen entre 680 y 600 Ma para las rocas deformadas del basamento mientras que los valores para los granitoides tardios se ubican ca. 550 Ma El análisis de minerales aislados, indican para ambos grupos, valores algo más jóvenes. Es interesante notar que las edades alrededor de 1000 Ma, semejantes a las obtenidas por el método U-Pb, no fueron obtenidas utilizando la sistemática Rb-Sr. Las edades isocrónicas Rb-Sr de las rocas gnéisicas y migmatíticas del basamento del TPE deben de estar asociadas a la época de su deformación, indicando entonces, la importancia del evento Brasiliano / Panafricano en la deformación de las rocas ígneas pretéritas. Un conjunto muestras, todas del basamento del TPE, fueron analizadas por el método U-Pb en circones. Esas muestras están representadas por gneises oftalmíticos, en dónde los megacristales de feldespato potásico, bastante deformados, se destacan en medio de una matriz rica en biotita, plagioclasa y cuarzo. Paralelamente a ese bandeado ocurre la asociación de una intensa deformación cizallante. La selección de los circones a ser disueltos fue efectuada con lupa binocular después de una concentración de los mismos conseguida en un proceso combinando la utilización de la mesa separadora de Wilfley, líquidos densos y separador isomagnético. La concentración y purificación del U y del Pb, fueron efectuadas a través del pasaje de la solución, en columnas de intercambio iónico, según el método clásico presentado por Krough (1973). Los análisis espectrométricos fueron obtenidos en un aparato VG 354 con multi-colectores. Las edades fueron calculadas utilizándose el software Isoplot (Ludwig 1993) después de la corrección de las relaciones espectrométricas para el fraccionamiento isotópico, blanco de laboratorio y Pb inicial. En todas las muestras fueron reconocidos por lo menos dos tipos de circones. El tipo predominante, escogido para el análisis, está representado por circones prismáticos, generalmente 2 a 3x1, biterminados, sin inclusiones, poco fracturados, de coloración levemente rosada a transparente. En general presentan las aristas algo redondeadas. Mientras tanto, las edades obtenidas para cada roca individualmente son próximas al valor de 1040 +- 38 Ma obtenido pro el conjunto de muestras. El segundo tipo está caracterizado por circones redondeados, menos transparentes y algo fracturados, siendo analizada una única fracción constituida por este tipo de circones. La edad obtenida de 510 +- 135 Ma debe ser entendida como la de la época de formación del leucosoma. Mientras tanto, ese valor debe ser encarado como una edad mínima para ese proceso ya que pequeñas pérdidas de Pb, ocurridas después de la formación de la roca, mismo por difusión continua, pueden provocar modificaciones significativas en la edad obtenida en el intercepto inferior. El protolito a partir del cual el leucosoma fue generado tiene una posible edad mesoproterozoica conforme lo que indica el intercepto superior Por otra parte, los valores de las edades ²⁰⁷Pb/²⁰⁶Pb calculadas para cada fracción se sitúan entre 908 y 1172 Ma. La superposición tectono-metamórfica responsable por el rejuvenecimiento parcial del sistema U-Pb en circones es también confirmada, desde el punto de vista geocrónologico, por los datos Rb-Sr e K-Ar disponibles. Biotitas presentaron valores entre 656 y 515 Ma, demostrando que los gneises de Rocha alcanzaron, en el Neoproterozoico III, temperaturas del orden de 250ºC.

Siguiendo las características de las deformaciones observadas y las edades K-Ar y Rb-Sr disponibles, se sugiere que los eventos de edad neoproterozoica deben corresponder no solamente a un calentamiento regional sino también al desarrollo de la principal característica estructural observada en los gneises del TPE y el desarrollo de las grandes zonas de cizalla, tanto internas al TPE como a las que se producen en el contacto con el dominio de los granitoides centrales.

Isótopos de Nd y Sr

Fueron analizados gneises del basamento del TPE en roca total. Las edades modelo obtenidas para las rocas gnéisico–migmatíticas indicaron una amplia variación entre 2,4 e 1.8 Ga indicando una larga residencia cortical para estos materiales antes de que el mismo fuera colocado en forma de granitoides, en niveles corticales superiores, durante la orogénesis Namaqua y fueron afectados por procesos de deformación y anatexis en el neoproterozoico. Esa larga residencia crustal es corroborada por valores bastante negativos de e Nd entre –13 e –14.3. Dos análisis fueron efectuados en metasedimentitas de la Formación Rocha. Los valores obtenidos a pesar de que las muestras era similares entre si, fueron bastante diferentes, con la edad más antigua indicando 1.97 Ga y la más nueva un valor de 1.54 Ga. Es posible que los dos valores representen edades de mezcla entre materiales de origen diverso, sin que implique un significado geológico más preciso. En este caso la edad de la muestra URPR 35, más joven, representaría la época máxima para la depositación de la Formación Rocha. Nuevamente, los resultados bastante negativos para los valores de e Nd reforzaron la sugerencia de una larga residencia cortical para el protolito de esas rocas.

Yuxtaposición de las unidades diferentes del cinturón Dom Feliciano

Como se presento previamente los diferentes segmentos del Cinturón Dom Feliciano y los terrenos adyacentes tienen características geológicas, estructurales y geocronológicas diferentes La cintura metamórfica presenta una evolución deformacional polifásica asociada a repliegues de la foliación. La fase principal de deformación está caracterizada por pliegues con ejes NE y bajo buzamiento de los planos axiales al SE y desarrollo de una foliación de transposición S2. Este armazón se generó entre 750 y 640 Ma durante la orogénesis Brasiliana asociada con corrimientos con transporte al NW. La mayor parte de la evolución magmática del Cinturón Granítico tuvo lugar ca. 620 y 590 Ma. La generación de las suites calco-alcalinas más viejas está asociada con la fase de subducción de parte del océano de Adamastor y su deformación se generó con la colisión entre los Cinturones Granítico y Metamórfico que ocurrió ca. 600 Ma. Poco después de esta fase colisional se desarrolló la intrusión tardía de granitoides, isótropos, alcalinos. Reactivaciones de los principales cinturones miloníticos previos con características transcurrentes sinestrales y alta temperatura en la porción sur del Cinturón y dextrales y de baja temperatura en la porción norte tuvo lugar como resultado de un importante movimiento transpresional del Cinturón. La mejor estimación para está fase está representada a través de edades Ar-Ar en el entorno de 534 +-3 Ma para las zonas de cizallamiento de alto ángulo de las regiones de Canguçu y Pinheiro. Como resultado de la convergencia del Cinturón Metamórfico y el Granítico se generaron en la porción occidental cuencas tipo folerand cuyos rellenos diacrónicos tuvieron lugar en el Vendiano en Uruguay a ca. 560 Ma. La fase de deformación colisional que empezó alrededor de los 600 Ma en las regiones interiores del Cinturón Granítico causó en el cinturón metamórfico reactivaciones, cabalgamientos y replegamiento con planos axiales buzando al SE, sólo alcanzando a las cuencas foreland alrededor de los 535 Ma. En éstas cuencas, las deformaciones están caracterizadas por cabalgamientos y plegamiento suave con vergencia NW que localmente desarrolla un clivaje plano-axial en las condiciones de esquistos verdes. En resumen, la historia metamórfica y deformacional de los tres segmentos del CDF se sitúan en un intervalo entre 750 y 530 Ma. Inicialmente, no se estableció bien el contexto tectónico, la principal fase de deformación y metamorfismo del Cinturón Metamórfico tuvo lugar entre 750 Ma, y 640 Ma siendo transportada contra los dominios cratónicos occidentales (micro placa Luis Alves y Cratón del Rio de la Plata). Hacia los 620 Ma se produce subducción hacia el E de la corteza oceánica existente (Adamastor) y la generación del Cinturón Granítico tiene lugar en un contexto de arco magmático. Alrededor de 600 Ma tiene lugar la colisión entre el Cinturón Granítico y el Cinturón Metamórfico y las cuencas foreland empiezan a formarse. Alrededor de 570 Ma un importante transporte de masa ocurre paralelo a la dirección del cinturón y los últimos granitoides post-colisionales son generados. Cerca de 535 Ma la deformación del cinturón alcanzo las cuencas foreland y reactiva numerosas zonas de cizallamiento de alto ángulo..

Correlaciones

El interés de correlacionar terrenos a ambos lados del Océano Atlántico Sur refleja una curiosidad natural entre los investigadores que trabajan en la Sur-América oriental y de aquellos que estudian África del sudoeste El problema mayor de las reconstrucciones pasadas están en el hecho que algunos investigadores enfocaron la correlación directa entre los Cinturones metasedimentarios observados en ambos lados del Atlántico. Por consiguiente, varios propuestas, correlacionan directamente los Cinturones Kaoko/ Gariep de Africa del Sur con sus posibles equivalentes de Brusque/Poróngos/ Lavalleja, en Sudamerica. A pesar de presentar similares edades (Pan-Africano/Brasiliano), un examen más preciso y principalmente el reciente incremento de un gran número de datos radiométricos, éstos cinturones plegados tienen un significado completamente diferentes, dejando fuera los modelos tectónicos que comienzan con la formación de una extensa y única cuenca sedimentaría desde el comienzo de sus historias geológicas. Considerando que el Cinturón Granítico más jóven no pudo ser el arco magmático asociado con el desarrollo del Cinturón Dom Feliciano, los sedimentos de una probable cuenca asociada con este arco faltaría del lado Sudamericano. De una manera análoga, en el lado africano, las rocas representativas del arco magmático que debió existir durante la estructuración de las paleocuencas de Kaoko/Gariep está ausente El modelo tectónico presentado, basada en edades radimétricas y signaturas isotópicas, sugiere que el arco magmático, activo durante la deposición de las unidades del Kaoko (Damara Costero)/Gariep en una posible cuenca de retroarco, podría estar representada en el E del continente Sud-americano. Se propone aquí que este arco magmático desarrollado en el periodo 620-590 Ma, se generó por la subducción al E de corteza oceánica bajo los terrenos antiguos del lado africano, con el Cinturón Granítico del CDF representando sus raíces. El principal argumento para sugerir que la inmersión de la corteza oceánica debe de ser al E y no al W como fue señalado por trabajos anteriores, es la signatura Sm-Nd del Cinturón granítico completamente diferente del modelo observado en los otros dominios tectónicos Sud-americanos. Edades modelo (TDM) para el Cinturón Granítico son sistemáticamente más jóvenes (entre 1.2-1.7 Ga, concentrándose en el 1.4-1.6 intervalo de Ga) que aquellos para las unidades del Cinturón Metamórfico(rocas graníticas y metasedimentarias con valores alrededor de 2.0 Ga) y mucho más jóvenes que aquéllas encontradas en las unidades de basamento del Tipo Luis Alves-o Cratón de la Plata (> 2.1 y principalmente concentrándose alrededor de 2.7-3.2 Ga). La afinidad geoquímica del magmatismo del arco debió ser más alta en la margen activa dónde el arco fue instalado que el material de la placa opuesta (margen pasivo). Se sugiere aquí que ésta debe ser la explicación para la diferencia isotópicas entre el arco y los terrenos localizaron al oeste, y las similitudes entre el Cinturón Granítico y el magmatismo africano del SW. Por consiguiente, esto sugiere que la litosfera sobre la placa subductante, funde para la generación del Cinturón Granítico y difiere de aquella situada al oeste, una vez que los granitoides de edad similar tienen una signatura isotópica diferente. Por consiguiente, si sobre el lado Sudamericano las edades modelo Nd entre 1.2 y 1.6 Ga son características del Cinturón Granítico, valores en el mismo intervalo son comunes en varias regiones de la porción Sur-Africana. Edades Mesoproterozoicas predominan para las rocas del oeste del Damara (región entre el Walvis Bay/Karibib/Rio Huab) con granitoides calco-alcalinos Tipo Palmental y sienitas de intra-placa y granitos que presentan edades modelos entre 1.1-1.5 Ga. Un modelo muy similar también se encontró para granitoides metaaluminosos de Tipo A de la región central del Damara Los valores en el mismo intérvalo son igualmente comunes en el Cinturón Neoproterozoico Mozambique en Tanzania, dónde varias metapelitas y charonockitas localizadas en la porción oriental del cinturón Nd presenta edades modelo entre 1.1 y 1.5 Ga

Por otro lado, edades Arqueanas (Cinturón Limpopo , xenolitos en Kaapvaal y algunos núcleos de basamento dentro de los metasedimentitas de Damara) o Paleoproterozoico (Namaqua y xenolitos dentro de los granitoides deformados del Cinturón Namaqua son características de las rocas asociadas con el basamento de las coberturas neoproterozoicas Parte de las metasedimentitas del Damara (notablemente Rossing y Kuiseb) y Nama (principalmente Kuibis y Schwarzrand) muestran edades modelo Nd similares a aquéllas obtenidas para el Cinturón Granítico, sugiriendo que ésta debió haber sido una fuente importante para estos metasedimentitas , reforzando el modelo de una evolución en una cuenca de retroarco para las unidades de Kaoko y Gariep y por consiguiente la región interior para el Nama. La tectónica colisional hacia el oeste que involucra al Cinturón Granítico debió de desarrollarse en el lado Sur-americano alrededor de 600 Ma. Sólo después de 545 Ma los cabalgamientos hacia el este posicionan estos granitoides en las unidades supracorticales del lado africano de una manera sincrónica con su pico metamórfico y reactivando las estructuras del lado Sur-americano que deforman las cuencas Foreland. En el lado africano, las cuencas de edades similares (e.g. Nama) sólo presentan deformación después de 506 Ma.

Discusión

Como ha sido presentado,el modelo tectónico involucrando una subducción de corteza oceánica al NW produciendo el Cinturón Granítico y en condiciones de retro-arco el Cinturón Metamórfico no es aceptada. Tal modelo, a pesar de representar una lectura logica del zoneamiento Petrotectónico del CDF, cuando es examinado en detalle, presenta varios problemás que lo invalidan. Dentro del cuadro general, toda la parte occidental del sudeste de Brasil y Uruguay está caracterizada por dominios antiguos agupados en el Cratón del Rio de la Plata (RS y UY) y Microplaca Luis Alves (SC) sobre los cuales ocurrio el principal transporte de masa generando el CDF. En la región occidental, la Microplaca Luis Alves y el Terreno Piedra Altas ( parte uruguaya del Cratón del Rio de la Plata) fue preservada de la tectónica Neoproterozoica que afecto de una manera significante los otros dominios antiguos (Bloque Taquarembó , RS, y el Terreno Nico Pérez-Valentines, Uruguay), causando calentamiento general (rejuvenación de las edades K-Ar en biotitas y anfíboles) y generando varios cuerpos de granitoides. En el dominio del Cinturón Dom Feliciano , el Cinturón Metamórfico presenta valores metamórfico alrededor de 120 Ma, más viejo que aquellos obtenidos para la formación y emplazamiento de los granitoides más viejos reconocidos en el Cinturón Granítico adyacente. Por consiguiente, es necésario disociar la fase principal de metamórfismo del Cinturón Metamórfico del Cinturón Granítico. En el momento del comienzo de su generación (~620 Ma) las rocas supracrustales ya habian alcanzado su climax metamórfico e incluso sufrido la tectónica que los arrojo sobre las areas cratonicas occidentales por medio de nappes NW que caracteriza la deformación tardi-metamórfica del Cinturón Metamórfico (~640 Ma). Granitoides biotíticos intrusivos en las rocas supracrustales, a pesar de presentar edades comparables con el magmatismo del Cinturón Granítico,tienen características petrologicas, geoquimicas e isotópicas que los diferencian del último. El desarrollo de una intensa aureóla de metamórfica de contacto entre estos granitoides y la foliación principal de las rocas metamórficas del Cinturón Metamórfico es comun. Las incertidumbres geológicas son todavia grandes como para precisar el posicionamiento tectónico de este segmento del crustal. El desarrollo del magmatismo del Cinturón Granítico (entre 620 y 590 Ma) se asocia con la generación de un arco magmático maduro con participación crustal importante, generada por la subducción hacia el este, evolucionando en una posición geográfica distinta y de una manera disociada de los terrenos localizados al oeste La probable cuenca de retroarco de este Cinturón se sitúaría en el lado africano (parte de Los cinturones costeros Damara/Kaoko/Gariep). En este contexto la cuenca de Nama sería una cuenca interna. La distribución geometrica y gran parte de los rasgos estructurales sistematicamente observadas en las unidades que componen el Cinturón Dom Feliciano pudieron ser generadas en la transición Proterozoica-cámbrica, después de la fase colisional que yuxtapuso el Cinturón Granítico al Cinturón Metamórfico. Las cuencas Foreland (Itajaí, Camaquã, el del de Arroyo Soldado-Piriápolis) se producen en los dominios occidentales en respuestá a la apróximación del Cinturón Granítico (600-560 Ma) y no, generadas durante el tectónica deformacional más vieja (750-640 Ma) del Cinturón Metamórfico. Por consiguiente, estás cuencas deben entenderse como cuencas Foreland syn - tardi-colisionales de la Orogénesis de Rio Doce.

Una hipótesis especulativa para la formación del Gondwana

Augusto E. Rapalini* & Leda Sánchez Bettucci**

*Instituto de Geofísica Daniel Valencio (INGEODAV), Departamento de Ciencias Geológicas, Fac. Ciencias Exactas y Naturales, Universidad de Buenos Aires, CONICET, Pabellón 2, Ciudad Universitaria, C1428EHA, Buenos Aires, rapalini@gl.fcen.uba.ar

**Area Geofísica y Geotectónica, Instituto de Geología y Paleontología, Facultad de Ciencias, Univ. Nac. de la República, Montevideo, 11400, Uruguay, leda@fcien.edu.uy

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Formação Bocaina			Formação Araras
amostra	d¹³C ⁰/₀₀^PDB	d¹⁸O ⁰/₀₀^PDB	Amostra	d¹³C ⁰/₀₀^PDB	d¹⁸O ⁰/₀₀^PDB
cj-127-A	- 0,61	- 3,58	MT-03-D	-0,81	- 7,68
cj-127- B	- 1,03	- 3,93	MT-03-C	1,51	-1,08
cj-127- C	- 0,99	- 4,24	MT-03-B	0,50	-1,49
cj-127- D	- 0,41	- 4,13	MT-03-A	4,16	-0,09
cj-127- E	0,57	- 3,69
cj-127- F	0,14	- 4,74