Nós usamos tomografia sísmica de tempo de percurso para estudar o manto superior sob as regiões Sudeste e Centro-Oeste do Brasil. Este método baseia-se na inversão de resíduos relativos de tempo para as ondas P e S (VanDecar, 1991), que foram obtidos para mais de 80 estações em uma área de 20 x 20 graus. Mais de 11000 e 8000 resíduos de tempo foram obtidos para as ondas P (P direta e PKPdf) e S (S direta, ScS, SKS e SKKS), respectivamente, utilizando correlação cruzada de forma de ondas para até 12 estações operando simultaneamente. Para avaliar a robustez dos resultados com respeito aos dados, nós utilizamos o método estatístico de re-amostragem Jackknife, o qual inerentemente leva em conta a altamente variável cobertura dos raios e os erros das medidas, e pode fornecer limites de confiança para as anomalias. Inversões regionais foram realizadas para estudar a influência da parametrização nas anomalias sísmicas. Nossos resultados mostraram boa correlação das anomalias sísmicas com as principais estruturas tectônicas e revelaram novas anomalias que ainda não haviam sido observadas nos trabalhos anteriores. Anomalias de alta velocidade na porção oeste do Cráton do São Francisco apóiam a hipótese de que este cráton foi parte de uma placa Neoproterozóica maior. Anomalias de baixa velocidade sob a Província Tocantins (principalmente nas faixas móveis entre os Crátons Amazônico e do São Francisco) foram interpretadas como causadas por afinamento litosférico, consistente com a boa correlação entre a sismicidade intraplaca e as anomalias de baixa velocidade nesta região (Assumpção et al., 2004b). A melhora na resolução da anomalia de alta velocidade sob a Bacia do Paraná (~200 km) permitiu uma discussão sobre a geometria do núcleo cratônico desta Bacia. A subducção da Placa de Nazca foi observada como uma anomalia de alta velocidade sob a Bacia do Paraná (profundidades entre 700 e 1200 km). Nestas profundidades, uma grande anomalia de baixa velocidade aparece próxima da Placa de Nazca. Testes sintéticos mostraram que esta anomalia é um artefato da inversão gerado pela presença da Placa de Nazca. / We used travel time seismic tomography to study the upper mantle beneath SE and Central Brazil. This method is based on the inversion of P- and S-wave relative travel time residuals (VanDecar, 1991) obtained from more than 80 stations in an area of 20 x 20 degrees. More than 11000 P and PKP residuals, and more than 8000 S, ScS, SKS, and SKKS residuals were obtained from waveform cross-correlations for up to 12 simultaneous stations. To evaluate the robustness of the tomographic results with respect to the data, we use the Jackknife re-sampling method, which inherently take into account the highly variable ray coverage and measurement errors, and can provide confidence limits for the anomalies. Regional inversions were carried out to study influence of the parameterization on the seismic anomalies. Our results show correlations of seismic anomalies with the main tectonic structures and reveal new anomalies not yet observed in previous works. High velocity anomalies in the western portion of the São Francisco Craton support the hypothesis that this craton was part of a major Neoproterozoic plate. Low velocity anomalies beneath the Tocantins Province (mainly fold belts between the Amazon and São Francisco cratons) are interpreted as due to lithospheric thinning, consistent with a good correlation between intraplate seismicity and low velocity anomalies in this region (Assumpção et al., 2004b). The resolution improvement of the high velocity anomalies beneath the Parana Basin (at ~200 km) allows a discussion about the geometry of the cratonic nucleous of this Basin. The slab of the Nazca Plate is observed as a high velocity anomaly beneath the Parana basin (at 700-1200 km depth). At these depths, large low velocity anomaly appears next to the slab. Synthetic tests show that these anomalies are an artifact of the inversion generated by the presence of the slab.

A Província Aurífera de Alta Floresta insere-se na porção sul do Cráton Amazônico e constitui uma região alongada na direção WNW-ESSE onde se situam depósitos auríferos associados ao magmatismo plutonovulcânico Paleoproterozoico. A área pesquisada abrange o garimpo de ouro do Papagaio, situado em Paranaíta, Mato Grosso. Durante o mapeamento geológico, foram identificados corpos de rochas plutônicas a subvulcânicas de composições granodiorítica e granítica, além de rochas vulcânicas e piroclásticas dacíticas e riolíticas. Estes litotipos são pertencentes à série calcioalcalina de médio a alto potássio, meta- a peraluminosas correspondentes a granitos do tipo I de arco vulcânico em margem continental ativa. As idades U-Pb mostraram que o magmatismo na região ocorreu aproximadamente entre 1.80 - 1.78 Ga. com zircões herdados de até 2,1 Ga, mostrando evidências da existência de uma crosta mais profunda e antiga durante a subducção. Os dados de \'épsilon\'Nd(t) mostraram a existência de contribuição crustal e mantélica na fonte de idade TDM entre 2.15-2.02 Ga. Estes dados revelaram que o vulcanismo na área é pertencente a Suíte Colíder, correspondente de um magmatismo que ficou ativo por aproximadamente 16 Ma e que evoluiu de composição dacítica a riolítica. A lavra garimpeira ocorre sobre uma mineralização aurífera com cobre, zinco e baixos teores de chumbo, preferencialmente confinada em veios que cortam os halos de alteração potássica, sericítica e propilítica. Os veios existentes na região evoluem de forma sistematica em seis tipos: (I) Veio de quartzo estéril associado ao halo da alteração potássica; (II) Veio de quartzo com minerais de minério, por vezes, associado ao halo sericítico; (III) Veio sulfetado com halo de alteração sericítica bem desenvolvido nas salbandas; (IV) e (V) Veios de quartzo e carbonato tardios; (VI) sistema de veios de quartzo mais novo e sem relação com a evolução do depósito. Os três primeiros tipos são semelhantes ao sistema A-B-D descrito em depósitos do tipo pórfiro, enquanto que os mais novos apresentam características de um ambiente mais raso e epitermal. O ouro ocorre principalmente na zona central dos veios do tipo II, em paragênese com a calcopirita + esfalerita + pirita ± galena ± magnetita. O quartzo dos veios sofreu diferentes intensidades de recristalização decorrente de subsequentes pulsos hidrotermais e as imagens de catodoluminescência mostraram que as inclusões fluidas aquocarbônicas representam fluidos primários com importante participação no transporte de conteúdo metalífero. Os resultados de isótopos estáveis de D, O e S em quartzo, sericita e pirita hidrotemais mostram que os fluidos são magmáticos com contribuições de fluidos meteóricos. Neste contexto, o Garimpo do Papagaio se desenvolveu sobre um arco magmático continental, a partir da intrusão de corpos graníticos hidratados e oxidados que marcam o evento magmático final da Suíte Colíder. Representa um depósito do tipo pórfiro que foi sobreposto por características epitermais intermediate-sulfidation, como resultado da telescopagem hidrotermal consequente do rebaixamento da câmara magmática. / The Alta Floresta Gold Province is situated on the southern portion of Amazonian Craton and forms an elongate region with WNW-ESSE direction where auriferous deposits are associated with Paleoproterozoic plutonovolcanic magmatism. The area of research covers the Papagaio artisanal gold mining, located in Paranaíta, Mato Grosso. During geological mapping, plutonic to subvolcanic rocks bodies of granodioritic and granitic compositions were identified, as well as dacitic and rhyolitic volcanic and pyroclastic rocks. These lithotypes belong to the calc-alkaline series of medium to high potassium, meta- to peraluminous corresponding to type I granites of volcanic arc in an active continental margin. U-Pb ages showed that magmatism in the region has crystallized in the range of 1.80-1.78 Ga with inherited zircons up to 2.1 Ga, showing evidence of a deeper and older crust during subduction. \'épsilon\'Nd (t) data showed the contribution of crustal and mantle material in the source with TDM age between 2.15- 2.02 Ga. These data revealed that volcanism in the area belongs to the Colíder Suite, corresponding to a magmatism that was active for approximately 16 Ma and evolved from dacitic to a rhyolitic composition. The mining prospect occurs on gold mineralization with copper, zinc and low levels of lead, preferably confined in veins that cut potassic, seritic and propylitic alteration halos. The existing veins in the region evolve systematically into six types: (I) Barren quartz vein associated with potassic alteration halo; (II) Quartz vein with ore minerals, sometimes associated with sericitic halo; (III) Sulphide vein with a well developed sericitic alteration halo in the salbands; (IV) and (V) Late quartz and carbonate veins; (VI) system of quartz veins newer and unrelated to deposit evolution. The first three types are similar to A-B-D system described in porphyry deposits, while the newer ones have characteristics of a shallower and epithermal environment. Gold occurs mainly in the central zone of the type II veins, in paragenesis with chalcopyrite + sphalerite + pyrite ± galena ± magnetite. The quartz of the veins underwent different intensities of recrystallization due to subsequent hydrothermal pulses and cathodoluminescence images showed that aquocarbonic fluid inclusions represent primary fluids with important participation in metalliferous content transportation. The results of D, O and S stable isotope in hydrothermal quartz, sericite and pyrite show that fluids are magmatic with meteoric contributions. In this context, the Papagaio artisanal mining developed on a continental magmatic arc, from the intrusion of hydrated and oxidized granite bodies that mark the final magmatic event of the Colíder Suite. It represents a porphyry deposit that was superimposed by epithermal intermediate-sulfidation characteristics as a result of hydrothermal telescoping, resulting from the lowering of the magma chamber.

Nós usamos tomografia sísmica de tempo de percurso para estudar o manto superior sob as regiões Sudeste e Centro-Oeste do Brasil. Este método baseia-se na inversão de resíduos relativos de tempo para as ondas P e S (VanDecar, 1991), que foram obtidos para mais de 80 estações em uma área de 20 x 20 graus. Mais de 11000 e 8000 resíduos de tempo foram obtidos para as ondas P (P direta e PKPdf) e S (S direta, ScS, SKS e SKKS), respectivamente, utilizando correlação cruzada de forma de ondas para até 12 estações operando simultaneamente. Para avaliar a robustez dos resultados com respeito aos dados, nós utilizamos o método estatístico de re-amostragem Jackknife, o qual inerentemente leva em conta a altamente variável cobertura dos raios e os erros das medidas, e pode fornecer limites de confiança para as anomalias. Inversões regionais foram realizadas para estudar a influência da parametrização nas anomalias sísmicas. Nossos resultados mostraram boa correlação das anomalias sísmicas com as principais estruturas tectônicas e revelaram novas anomalias que ainda não haviam sido observadas nos trabalhos anteriores. Anomalias de alta velocidade na porção oeste do Cráton do São Francisco apóiam a hipótese de que este cráton foi parte de uma placa Neoproterozóica maior. Anomalias de baixa velocidade sob a Província Tocantins (principalmente nas faixas móveis entre os Crátons Amazônico e do São Francisco) foram interpretadas como causadas por afinamento litosférico, consistente com a boa correlação entre a sismicidade intraplaca e as anomalias de baixa velocidade nesta região (Assumpção et al., 2004b). A melhora na resolução da anomalia de alta velocidade sob a Bacia do Paraná (~200 km) permitiu uma discussão sobre a geometria do núcleo cratônico desta Bacia. A subducção da Placa de Nazca foi observada como uma anomalia de alta velocidade sob a Bacia do Paraná (profundidades entre 700 e 1200 km). Nestas profundidades, uma grande anomalia de baixa velocidade aparece próxima da Placa de Nazca. Testes sintéticos mostraram que esta anomalia é um artefato da inversão gerado pela presença da Placa de Nazca. / We used travel time seismic tomography to study the upper mantle beneath SE and Central Brazil. This method is based on the inversion of P- and S-wave relative travel time residuals (VanDecar, 1991) obtained from more than 80 stations in an area of 20 x 20 degrees. More than 11000 P and PKP residuals, and more than 8000 S, ScS, SKS, and SKKS residuals were obtained from waveform cross-correlations for up to 12 simultaneous stations. To evaluate the robustness of the tomographic results with respect to the data, we use the Jackknife re-sampling method, which inherently take into account the highly variable ray coverage and measurement errors, and can provide confidence limits for the anomalies. Regional inversions were carried out to study influence of the parameterization on the seismic anomalies. Our results show correlations of seismic anomalies with the main tectonic structures and reveal new anomalies not yet observed in previous works. High velocity anomalies in the western portion of the São Francisco Craton support the hypothesis that this craton was part of a major Neoproterozoic plate. Low velocity anomalies beneath the Tocantins Province (mainly fold belts between the Amazon and São Francisco cratons) are interpreted as due to lithospheric thinning, consistent with a good correlation between intraplate seismicity and low velocity anomalies in this region (Assumpção et al., 2004b). The resolution improvement of the high velocity anomalies beneath the Parana Basin (at ~200 km) allows a discussion about the geometry of the cratonic nucleous of this Basin. The slab of the Nazca Plate is observed as a high velocity anomaly beneath the Parana basin (at 700-1200 km depth). At these depths, large low velocity anomaly appears next to the slab. Synthetic tests show that these anomalies are an artifact of the inversion generated by the presence of the slab.

Le refroidissement seculaire du manteau terrestre et de la lithosphere continentale se traduit par l'augmentation de la surface de terres emergees. L'augmentation corollaire des processus d'alteration et d'erosion des silicates a des consequences majeures pour les enveloppes externes. Nous avons developpe un modele physique qui permet d'evaluer la surface de terres emergees en fonction de la temperature du manteau, de la surface totale de continents, et de la distribution des altitudes continentales. Nos resultats numeriques montrent qu'a la fin de l'Archeen, moinsde 15% de la surface terrestre etaient emergee, en accord avec nombre d'observations geologiques et geochimiques. Pour estimer le refroidissement seculaire de la lithosphere continentale, nous avons combine des modeles thermo-mecaniques avec des observations de terrain. Nos resultats, contraints par des donnees geologiques, suggerent que la temperature au Moho a diminue de ~ 200ºC en 2,7 Ga dans le craton des Pilbaras. Pour evaluer l'effet de la croissance continentale sur l'evolution de la surface de terres emergees, nous avons developpe un modele base sur un modele d'evolution thermique publie. Nos resultats suggerent que la surface emergee, de moins de 5% de la surface terrestre a l'Archeen, depend peu de la croissance continentale. Ceci permet de reconcilier l'evolution du 87Sr/86Sr oceanique avec une croissance continentale precoce. Les continents sont enrichis en phosphate, element essentiel a la biosphere. Leur emergence aurait donc engendre une augmentation de la production d'oxygene par des micro-organismes photosynthetiques, contribuant ainsi a l'oxydation de l'atmosphere il y a 2,4 Ga.

科学研究費補助金 研究種目:基盤研究(C)(2) 課題番号:07640638 研究代表者:星野 光雄 研究期間:1995-196年度

Na América do Sul, o Cráton Amazônico representa um componente essencial nas reconstruções de supercontinentes, entretanto, há uma grande escassez de dados paleomagnéticos de qualidade para esta unidade geotectônica, principalmente, para o Proterozóico. Com o intuito de esclarecer a participação do Cráton Amazônico na evolução do ciclo continental, este trabalho apresenta um estudo paleomagnético realizado em quatro unidades geológicas Paleo- a Mesoproterozóicas pertencentes ao Cráton Amazônico. As unidades escolhidas para este estudo foram às rochas vulcânicas do Grupo Surumu (1980-1960 Ma, U-Pb), as soleiras máficas Avanavero (ca. 1780 Ma, U-Pb) ambas situadas no norte do Estado de Roraima (Escudo das Guianas), os enxames de diques Nova Guarita e a intrusiva máfica Guadalupe ambas localizadas no norte do Estado do Mato Grosso (Escudo Brasil-Central). Determinações 40Ar/39Ar realizadas em biotitas de quatro diques de Nova Guarita mostraram resultados coerentes, fornecendo uma idade média de 1418,5 ± 3,5 Ma para a época de intrusão dos diques. Idades U-Pb obtidas em rochas da intrusiva máfica Guadalupe indicam uma idade mínima de 1530 Ma para estas amostras. As análises paleomagnéticas realizadas em mais de 1100 espécimes de rocha através dos tratamentos térmicos e por campos magnéticos alternados revelaram direções características coerentes para as quatro unidades de rochas estudadas: (1) as rochas do Grupo Surumu apresentaram direções noroeste com inclinações positivas. Foi calculada uma direção média Dm = 298,6°, Im = 39,4° (N = 20, alfa95 = 10,1°, K = 11,4), a qual foi interpretada como sendo de origem primária; (2) as rochas máficas Avanavero apresentaram direções sudeste com inclinações positivas/negativas baixas, sendo determinada uma direção média Dm = 135,6°, Im = -2,1° (N = 10, alfa95 = 15,9°, K = 10,2°). Um teste de contato cozido realizado para um dos sítios amostrados atesta o caráter primário da magnetização remanente isolada, a qual foi adquirida pelas rochas há ca.1780 Ma atrás; (3) os diques máficos Nova Guarita apresentaram polaridades reversas e normais, tendo sido isoladas direções sul/sudoeste com inclinações positivas e nordeste com inclinações negativas. Um teste de contato cozido positivo foi obtido para um dique que intrude o Granito Matupá, o qual confirma que a magnetização remanente (Dm = 220,5°, Im = 45,9°, N = 19, alfa95 = 6,5°, K = 27,7) isolada para estas rochas corresponde a uma magnetização termorremanente adquirida durante a formação da rocha há ca. 1419 Ma atrás; (4) rochas pertencentes a Intrusiva Máfica Guadalupe também apresentaram polaridades reversas e normais. Direções noroeste/nordeste com inclinações positivas ou sul/sudeste com inclinações negativas foram isoladas para estas rochas, para as quais foi calculada a direção média Dm = 356,6°, Im = 59,4°, (N = 10, alfa95 = 10,2°, K = 23,2). A idade desta componente, entretanto, ainda não está bem estabelecida, podendo representar uma remagnetização adquirida durante o evento Brasiliano, já que ela é similar às magnetizações adquiridas há 520 Ma, presentes em formações geológicas do Cráton Amazônico e do Cráton do São Francisco. A caracterização da mineralogia magnética de todas as amostras investigadas foi obtida através de curvas termomagnéticas, curvas de histerese e curvas de magnetização remanente isotérmica. Quatro pólos paleomagnéticos para o Cráton Amazônico foram determinados para estas componentes, os quais estão localizados em 234,8° E, 27,4°N (A95=9,8°) (pólo GS, Grupo Surumu), 27,5°E, -45,8°N (A95=11,5°) (pólo AV, Avanavero), 245,9°E, -47,9°N (A95=7,0°) (pólo NG, Nova Guarita) e 306,2°E, 38,9°N (A95=13,7°) (pólo GUA, Guadalupe). Os resultados paleomagnéticos obtidos para as rochas Surumu (pólo GS) contribuíram para um melhor ajuste da curva de deriva polar aparente (CDPA) para o Escudo das Guianas durante o Paleoproterozóico (2070-1960 Ma). A comparação desta CDPA com a construída para o Cráton Oeste-África para o mesmo período de tempo sugere que estes blocos cratônicos estavam unidos há 1970-2000 Ma atrás, em uma paleogeografia em que as zonas de cisalhamento Guri, no Escudo das Guianas, e Sassandra, no Cráton Oeste-África estavam alinhadas como sugerido em modelos anteriores. O pólo Avanavero de 1780 Ma é consistente com a paleogeografia do supercontinente Columbia em que o proto-Cráton Amazônico e a Báltica estavam unidos como no modelo SAMBA (South America-Baltica) proposto anteriormente com base em evidências geológicas. No cenário proposto aqui para o Supercontinente Columbia há 1780 Ma atrás, o Cráton Oeste-África estava unido ao proto-Cráton Amazônico na mesma configuração sugerida pelos dados paleomagnéticos de 1790-2000 Ma. O atual lado leste da Laurentia estava unido ao norte (atual) da Báltica. A Sibéria estava unida com a atual costa Ártica da Laurentia e a proto-Austrália, com a atual costa oeste da Laurentia, em posição similar ao modelo SWEAT. Embora os dados paleomagnéticos disponíveis para o Cráton Norte da China e Índia indiquem paleolatitudes equatorias para estes dois blocos, nesta época, suas posições no supercontinente Columbia são ainda incertas. No modelo do Columbia apresentado neste trabalho, o Norte da China foi colocado ao lado da Sibéria e a Índia, ao lado da proto-Austrália, em decorrência de evidências geológicas. Outros blocos cratônicos, tais como, Congo-São Francisco, Kalahari e Rio de La Plata não foram incluídos, pela ausência de pólos paleomagnéticos desta idade. Os dados paleomagnéticos atualmente existentes para a Báltica e a Laurentia mostram que estes dois blocos continentais permaneceram unidos desde 1830 Ma até, pelo menos, 1270 Ma atrás. Já o pólo paleomagnético obtido para os diques Nova Guarita de 1419 Ma e o pólo de mesma idade, recentemente obtido para a Intrusiva Indiavaí, quando comparados com pólos de mesma idade da Báltica e da Laurentia, sugerem que o proto-Craton Amazônico já havia iniciado sua ruptura no Supercontinente Columbia nessa época. De modo alternativo, porém, essa diferença na posição dos pólos do proto-Cráton Amazônico e da Báltica/Laurentia, pode ser explicada por movimentos transcorrentes dextrais que teriam ocorrido entre o Escudo das Guianas e a parte sul do Cráton Amazônico em tempos posteriores a 1420 Ma. Neste caso, esta grande massa continental do Supercontinente Columbia, composta pelo proto-Cráton Amazônico, Báltica e Laurentia, pode ter permanecida unida por, pelo menos, 400 Ma. / The Amazonian Craton is an important component in Paleoproterozoic reconstructions, however, paleomagnetic data for this craton are yet scarce. Aiming to decipher the involvement of the Amazonian Craton in the Contiental cycle evolution, paleomagnetic studies were carried out in four Paleo- to Mesoproterozoic geological units. The chosen units are the volcanic rocks from the Surumu Group (1,980-1,960 Ma, U-Pb), the Avanavero mafic sills (ca. 1,780 Ma, U-Pb), both from the northern Roraima State (Guyana Shield), and the Nova Guarita dyke swarm and Guadalupe mafic intrusive, both from the northern Mato Grosso State (Central- Brazil Shield). 40Ar/39Ar determinations on biotites from samples belonging to four Nova Guarita dykes yielded well-defined plateau ages whose mean 1,418.5 ± 3.5 Ma is interpreted as the age of dyke intrusion. U-Pb (SHRIMP) determinations on rocks from the Guadalupe mafic Intrusive indicate a minimum age of 1,530 Ma for this unit. Paleomagnetic analysis performed on more than 1,100 specimens by thermal and alternating magnetic field (AF) treatments revealed stable characteristic remanent magnetizions (ChRM) for all geological units: (1) northwestern directions with positive inclinations were isolated for samples from the Surumu Group (mean: Dm = 298.6°, Im = 39.4°, N = 20, alpha95 = 10.1°, K = 11.4), which were interpreted to be primary. (2) Southeastern directions with low downward/upward inclinations were isolated for the Avanavero rocks, for which a mean direction was calculated: Dm=135.6°, Im = -2.1° (N=10, alpha95 = 15.9°, K = 10.2°). A positive baked contact test attests for the primary origin of this ChRM direction, which was probably acquired at about 1,780 Ma ago; (3) both south/southwestern directions with downward inclinations or northeastern directions with upward inclinations were isolated for the Nova Guarita dykes. A positive baked contact test attests for the primary nature of the ChRM directions (Dm = 220.5°, Im = 45.9°, N=19, alpha95=6.5°, K = 27.7) which most probably correspond to a termo-remanent magnetization (TRM) acquired at ca. 1,419 Ma ago; 10 (4) both northwest/northeastern directions with downward inclinations or outhsoutheastern directions with upward inclinations were isolated for rocks from the Guadalupe intrusive, whose mean direction is: Dm=356.6°, Im=59.4°, (N =10, alpha95=10.2°, K = 23.2). The age of this component is yet uncertain. U-Pb geochronology suggests an age of (or older than) 1,530 Ma for these rocks, however, a remagnetization effect at Cambrian times (520 Ma) cannot be rolled out as these directions are very similar to those found for younger geological units in the Amazonian Craton and Sao Francisco Craton. Four new paleomagnetic poles for the Amazonian Craton were obtained from these magnetic components, which are located at: 234.8°E, 27.4°N (A95=9.8°) (GS pole, Surumu Group), 27.5°E, 45.8°S (A95=11.5°) (AV pole, Avanavero), 245.9°E, 47.9°S (A95=7.0°) (NG pole, Nova Guarita) and 306.2°E, 38.9°N (A95 = 13.7°) (GUA pole, Guadalupe). The 1,960 Ma Surumu pole contributes to better define the APW path traced for the Guyana Shield in the time interval between 2,070 Ma and 1,960 Ma. Comparison of this APW path with that traced for West-Africa Craton for the same time interval suggests that these two cratonic blocks were linked together, in a paleogeography where the Guri (Guyana Shield) and Sassandra (West-Africa Craton) shear zones are aligned, as suggested by previous models. The Avanavero pole is consistent with the proto-Amazonian Craton and Baltica link as in the SAMBA (South America-Baltica) model at ca. 1,780 Ma ago, as previously proposed based on geological evidence. In the scenario proposed here for the Columbia Supercontinent at 1,780 Ma ago, the West-Africa Craton was linked to the proto-Amazonian Craton in the same configuration as suggested by Paleoproterozoic (1,960-2,000 Ma) paleomagnetic data (see above). Actual eastern Laurentia was linked to northern Baltica. Siberia was located at the actual Arctic Coast of Laurentia, and proto-Australia at the western coast of Laurentia, in a position similar to that of SWEAT model. Although available 1,780 Ma paleomagnetic data from North China and India indicate low paleolatitudes for these two blocks, their positions in the supercontinent Columbia are yet uncertain. In our model, North China is located beside Siberia, and India beside proto-Australia, based on geological evidences. Other cratonic blocks, such as Congo-Sao Francisco, Kalahari and Rio de la Plata were not included as no 1,780 Ma paleomagnetic poles are presently available for them. The paleomagnetic poles presently available for Baltica and Laurentia, show that these two blocks remained as a single continental mass since 1,830 Ma up to at least 1,270 Ma. However, the 1,419 Ma Nova Guarita pole and the recently published 1,416 Ma Indiavai pole from the Amazonian Craton, when compared with poles of similar age from Baltica and Laurentia suggest that the proto-Amazonian Craton had already broke-up from the Columbia Supercontinent at that time. Alternatively, the difference in the position of the 1,420 Ma poles from the proto-Amazonian Craton and those from Baltica/Laurentia, may be explained by dextral transcurrent movements between the Guyana Shield and the southern part of the Amazonian Craton at times later than 1,420 Ma. If so, this great continental mass, formed by proto-Amazonian Craton, Baltica and Laurentia may have remained as a single continental block for at least 400 Ma.

Lithium Cesium Tantalum (LCT) pegmatites are important resources for rare metals like Cesium, Lithium or Tantalum, whose demand increased markedly during the past decade. At present, Cs is known to occur in economic quantities only from the two LCT pegmatite deposits at Bikita located in Zimbabwe and Tanco in Canada. Host for this Cs mineralisation is the extreme rare zeolite group mineral pollucite. However, at Bikita and Tanco, pollucite forms huge massive, lensoid shaped and almost monomineralic pollucite mineralisations that occur within the upper portions of the pegmatite. In addition, both pegmatite deposits have a comparable regional geological background as they are hosted within greenstone belts and yield a Neoarchean age of about 2,600 Ma. Furthermore, at present the genesis of these massive pollucite mineralisations was not yet investigated in detail. Major portions of Western Australia consist of Meso- to Neoarchean crustal units (e.g., Yilgarn Craton, Pilbara Craton) that are known to host a large number of LCT pegmatite systems. Among them are the LCT pegmatite deposits Greenbushes (Li, Ta) and Wodgina (Ta, Sn). In addition, small amounts of pollucite were recovered from one single diamond drill core at the Londonderry pegmatite field. Despite that, no systematic investigations and/or exploration studies were conducted for the mode of occurrence of Cs and especially that of pollucite in Western Australia. In the course of the present study nineteen individual pegmatites and pegmatite fields located on the Yilgarn Craton, Pilbara Craton and Kimberley province have been visited and inspected for the occurrence of the Cs mineral pollucite. However, no pollucite could be detected in any of the investigated pegmatites. Four of the inspected LCT-pegmatite systems, namely the Londonderry pegmatite field, the Mount Deans pegmatite field, the Cattlin Creek LCT pegmatite deposit (Yilgarn Craton) and the Wodgina LCT pegmatite deposit (Pilbara Craton) was sampled and investigated in detail. In addition, samples from the Bikita pegmatite field (Zimbabwe Craton) were included into the present study in order to compare the Western Australian pegmatites with a massive pollucite mineralisation bearing LCT pegmatite system. This thesis presents new petrographical, mineralogical, mineralchemical, geochemical, geochronological, fluid inclusion and stable and radiogenic isotope data. The careful interpretation of this data enhances the understanding of the LCT pegmatite systems in Western Australia and Zimbabwe. All of the four investigated LCT pegmatite systems in Western Australia, crop out in similar geological settings, exhibit comparable internal structures, geochemistry and mineralogy to that of the Bikita pegmatite field in Zimbabwe. Furthermore, in all LCT pegmatite systems evidences for late stage hydrothermal processes (e.g., replacement of feldspars) and associated Cs enrichment (e.g., Cs enriched rims on mica, beryl and tourmaline) is documented. With the exception of the Wodgina LCT pegmatite deposit, that yield a Mesoarchean crystallisation age (approx. 2,850 Ma), all other LCT pegmatite systems gave comparable Neoarchean ages of 2,630 Ma to 2,600 Ma. The almost identical ages of the LCT pegmatite systems of the Yilgarn and Zimbabwe cratons suggests, that the process of LCT pegmatite formation at the end of the Neoarchean was active worldwide. Nevertheless, essential distinguishing feature of the Bikita pegmatite field is the presence of massive pollucite mineralisations that resulted from a process that is not part of the general development of LCT pegmatites and is associated with the extreme enrichment of Cs. The new findings of the present study obtained from the Bikita pegmatite field and the Western Australian LCT pegmatite systems significantly improve the knowledge of Cs behaviour in LCT pegmatite systems. Therefore, it is now possible to suggest a genetical model for the formation of massive pollucite mineralisations within LCT pegmatite systems. LCT pegmatites are generally granitic in composition and are interpreted to represent highly fractionated and geochemically specialised derivates from granitic melts. Massive pollucite mineralisation bearing LCT pegmatites evolve from large and voluminous pegmatite melts that intrude as single body along structures within an extensional tectonic setting. After emplacement, initial crystallisation will develop the border and wall zone of the pegmatites, while due to fractionated crystallisation immobile elements (i.e., Cs, Rb) become enriched within the remaining melt and associated hydrothermal fluids. Following this initial crystallisation, a relatively small portion (0.5–1 vol.%) of immiscible melt or fluid will separate during cooling. This immiscible partial melt/fluid is enriched in Al2O3 and Na2O, as well as depleted in SiO2 and will crystallise as analcime. In addition, this melt might allready contains up to 1–2 wt.% Cs2O. However, due to the effects of fluxing components (e.g., H2O, F, B) this analcime melt becomes undercooled which prevents crystallisation of the analcime as intergranular grains. Since this analcime melt exhibits a lower relative gravity when compared to the remaining pegmatite melt the less dense analcime melt will start to ascent gravitationally and accumulate within the upper portion of the pegmatite sheet. At the same time, the remaining melt will start to crystallise separately and form the inner portions of the pegmatite. This crystallisation is characterised by still ongoing fractionation and enrichment of incompatible elements (i.e., Cs, Rb) within the last crystallising minerals (e.g., lepidolite) or concentration of these incompatible elements within exsolving hydrothermal fluids. As analcime and pollucite form a continuous solid solution series, the analcime melt is able to incorporate any available Cs from the melt and/or associated hydrothermal fluids and crystallise as Cs-analcime in the upper portion of the pegmatite sheet. Continuing hydrothermal activity and ongoing substitution of Cs will then start to shift the composition from Cs-analcime composition towards Na-pollucite composition. In addition, if analcime is cooled below 400 °C it is subjected to a negative thermal expansion of about 1 vol.%. This contraction results in the formation of a prominent network of cracks that is filled by late stage minerals (e.g., lepidolite, quartz, feldspar and petalite). Certainly, prior to filling, this network of cracks enhances the available conduits for late stage hydrothermal fluids and the Cs substitution mechanism within the massive pollucite mineralisation. Furthermore, during cooling of the pegmatite, prominent late stage mineral replacement reactions (e.g., replacement of K-feldspar by lepidolite, cleavelandite, and quartz) as well as subsolidus self organisation processes in feldspars take place. These processes are suggested to release additional incompatible elements (e.g., Cs, Rb) into late stage hydrothermal fluids. As feldspar forms large portions of pegmatite a considerable amount of Cs is released and transported via the hydrothermal fluids towards the massive pollucite mineralisation in the upper portion of the pegmatite. Consequently, the initial analcime can accumulate enough Cs in order to shift its composition from the Cs-analcime member (>2 wt.% Cs2O) towards the Na-pollucite member (23–43 wt.% Cs2O) of the solid solution series. The timing of this late stage Cs enrichment is interpreted to be quasi contemporaneous or immediately after the complete crystallisation of the pegmatite melt. However, much younger hydrothermal events that overprint the pegmatite are also interpreted to cause similar results. Hence, it has been demonstrated that the combination of this magmatic and hydrothermal processes is capable to generate an extreme enrichment in Cs in order to explain the formation of massive pollucite mineralisations within LCT pegmatite systems. This genetic model can now be applied to evaluate the potential for occurrences of massive pollucite mineralisations within LCT pegmatite systems in Western Australia and worldwide. / Lithium-Caesium-Tantal-(LCT) Pegmatite repräsentieren eine bedeutende Quelle für seltene Metalle, deren Bedarf im letzten Jahrzehnt beträchtlich angestiegen ist. Im Falle von Caesium sind zurzeit weltweit nur zwei LCT-Pegmatitlagerstätten bekannt, die abbauwürdige Vorräte an Cs enthalten. Dies sind die LCT-Pegmatitlagerstätten Bikita in Simbabwe und Tanco in Kanada. Das Wirtsmineral für diese Cs-Mineralisation ist das extrem selten auftretende Zeolith-Gruppen-Mineral Pollucit. In den Lagerstätten Bikita und Tanco bildet Pollucit dagegen massive, linsenförmige und fast monomineralische Pollucitmineralisationen, die in den oberen Bereichen der Pegmatitkörper anstehen. Zusätzlich befinden sich beide Lagerstätten in geologisch vergleichbaren Einheiten. Die Nebengesteine sind Grünsteingürtel die ein neoarchaisches Alter von ca. 2,600 Ma aufweisen. Die Bildung derartiger massiver Pollucitmineralisationen ist bis jetzt noch nicht detailliert untersucht worden. Große Bereiche von Westaustralien werden von meso- bis neoarchaischen Krusteneinheiten (z.B. Yilgarn Kraton, Pilbara Kraton) aufgebaut, von denen auch eine große Anzahl an LCT-Pegmatitsystemen bekannt sind. Darunter befinden sich unter anderem die LCT-Pegmatitlagerstätten Greenbushes (Li, Ta) und Wodgina (Ta, Sn). Zusätzlich wurden kleine Mengen an Pollucit in einer einzigen Kernbohrung im Londonderry Pegmatitfeld angetroffen. Ungeachtet dessen, wurden in Westaustralien bis jetzt keine systematischen Untersuchungen und/oder Explorationskampagnen auf Vorkommen von Cs und speziell der von Pollucit durchgeführt. Im Verlauf dieser Studie wurden insgesamt neunzehn verschiedene Pegmatitvorkommen und Pegmatitfelder des Yilgarn Kratons, Pilbara Kratons und der Kimberley Provinz auf das Vorkommen des Minerals Pollucit untersucht. Allerdings konnte in keinem der untersuchten LCT-Pegmatitsystemen Pollucit nachgewiesen werden. Von vier der untersuchten LCT-Pegmatitsystemen, dem Londonderry Pegmatitfeld, dem Mount Deans Pegmatitfeld, der Cattlin Creek LCT-Pegmatitlagerstätte (Yilgarn Kraton) und der Wodgina LCT-Pegmatitlagerstätte (Pilbara Kraton) wurden detailliert Proben entnommen und weitergehend untersucht. Zusätzlich wurden die massiven Pollucitmineralisationen im Bikita Pegmatitfeld beprobt und in die detailierten Untersuchungen einbezogen. Der Probensatz aus dem Bikita Pegmatitfeld dient als Referenzmaterial mit dem die Pegmatitproben aus Westaustralien verglichen werden. Die vorliegende Arbeit fasst die wesentlichen Ergebnisse der petrographischen, mineralogischen, mineralchemischen, geochemischen und geochronologischen Untersuchungen sowie der Flüssigkeitseinschlussuntersuchungen und stabilen und radiogenen Isotopenzusammensetzungen zusammen. Alle vier der in Westaustralien untersuchten LCT-Pegmatitsysteme kommen in geologisch ähnlichen Rahmengesteinen vor, weisen einen vergleichbaren internen Aufbau, geochemische Zusammensetzung und Mineralogie zu dem des Bikita Pegmatitfeldes in Simbabwe auf. Weiterhin konnten in allen LCT-Pegmatitsystemen Hinweise für späte hydrothermale Prozesse (z.B. Verdrängung von Feldspat) nachgewiesen werden, die einhergehend mit einer Anreicherung von Cs verbunden sind (z.B. Cs-angereicherte Säume um Glimmer, Beryll und Turmalin). Mit der Ausnahme der Wodgina LCT-Pegmatitlagerstätte, in der ein mesoarchaisches Kristallisationsalter (ca. 2,850 Ma) nachgewiesen wurde, lieferten die Altersdatierungen in den anderen LCT-Pegmatitsystemen übereinstimmende neoarchaische Alter von 2,630 Ma bis 2,600 Ma. Diese fast identischen Alter der LCT-Pegmatitsysteme des Yilgarn und Zimbabwe Kratons suggerieren, dass die Prozesse, die zur LCT-Pegmatitbildung am Ende des Neoarchaikums führten, weltweit aktiv waren. Ungeachtet dessen stellt das Vorhandensein von massiver Pollucitmineralisation das Alleinstellungsmerkmal des Bikita Pegmatitfeldes dar, welche sich infolge eines Prozesses gebildet haben der nicht Bestandteil der üblichen LCT-Pegmatitentwicklung ist und sich durch eine extreme Anreicherung an Cs unterscheidet. Die neuen Ergebnisse die in dieser Studie von den Bikita Pegmatitfeld und den Westaustralischen LCT-Pegmatitsystemen gewonnen wurden, verbessern das Verständnis des Verhaltens von Cs in LCT-Pegmatitsystemen deutlich. Somit ist es nun möglich, ein genetisches Modell für die Bildung von massiven Pollucitmineralisationen in LCT-Pegmatitsystemen vorzustellen. LCT-Pegmatite weisen im Allgemeinen eine granitische Zusammensetzung auf und werden als Kristallisat von hoch fraktionierten und geochemisch spezialisierten granitischen Restschmelzen interpretiert. Die Bildung von massiven Pollucitmineralisationen ist nur aus großen und voluminösen Pegmatitschmelzen, die als einzelner Körper entlang von Störungen in extensionalen Stressregimen intrudieren möglich. Nach Platznahme der Schmelze bildet die beginnende Kristallisation zunächst die Kontakt- und Randzone des Pegmatits, wobei infolge von fraktionierter Kristallisation die immobilen Elemente (v.a. Cs, Rb) in der verbleibenden Restschmelze angereichert werden. Im Anschluss an diese erste Kristallisation entmischt sich nach Abkühlung eine sehr kleine Menge (0.5–1 vol.%) Schmelze und/oder Fluid von der Restschmelze. Diese nicht mischbare Teilschmelze/-fluid ist angereichert an Al2O3 und Na2O sowie verarmt an SiO2 und kristallisiert als Analcim. Zusätzlich kann diese Schmelze bereits mit 1–2 wt.% Cs2O angereichert sein. Aufgrund der Auswirkung von Flussmitteln (z.B. H2O, F, B) wird allerdings der Schmelzpunkt dieser Analcimschmelze herabgesetzt und so die Kristallisation des Analcims als intergranulare Körner verhindert. Da diese Analcimschmelze im Vergleich zu der restlichen Schmelze eine geringere relative Dichte besitzt, beginnt sie gravitativ aufzusteigen und sich in den oberen Bereichen des Pegmatitkörpers zu akkumulieren. Währenddessen beginnt die restliche Schmelze separat zu kristallisieren und die inneren Bereiche des Pegmatits zu bilden. Diese Kristallisation ist einhergehend mit fortschreitender Fraktionierung und der Anreicherung von inkompatiblen Elementen (v.a. Cs, Rb) in den sich als letztes bildenden Mineralphasen (z.B. Lepidolit) oder der Konzentration der inkompatiblen Element in die sich entmischenden hydrothermalen Fluiden. Da Analcim und Pollucit eine lückenlose Mischungsreihe bilden, ist die Analcimschmelze in der Lage, alles verfügbare Cs von der Restschmelze und/oder assoziierten hydrothermalen Fluiden an sich zu binden und als Cs-Analcim im oberen Bereich des Pegmatitkörpers zu kristallisieren. Fortschreitende hydrothermale Aktivität und Substitution von Cs verschiebt dann die Zusammensetzung des Analcims von der Cs-Analcim- zu Na-Pollucitzusammensetzung. Zusätzlich erfährt der Analcim bei Abkühlung unter 400 °C eine negative thermische Expansion von ca. 1 vol.%. Diese Kontraktion führt zu der Bildung des markanten Rissnetzwerkes das durch späte Mineralphasen (z.B. Lepidolit, Quarz, Feldspat und Petalit) gefüllt wird. Vor der Mineralisation allerdings, erhöht dieses Netzwerk an Rissen die verfügbaren Wegsamkeiten für die späten hydrothermalen Fluide und begünstigt somit den Cs-Substitutionsmechanismus in der massiven Pollucitmineralisation. Weiterhin kommt es bei der Abkühlung des Pegmatits zu späten Mineralverdrängungsreaktionen (z.B. Verdrängung von K-Feldspat durch Lepidolit, Cleavelandit und Quarz), sowie zu Subsolidus-Selbstordnungsprozessen in Feldspäten. Diese Prozesse werden weiterhin interpretiert inkompatible Elemente (z.B. Cs, Rb) in die späten hydrothermalen Fluide freizusetzen. Da Feldspäte große Teile der Pegmatite bilden, kann somit eine beträchtliche Menge an Cs freigeben werden und durch die späten hydrothermalen Fluide in die massive Pollucitmineralisation in den oberen Bereichen des Pegmatitkörpers transportiert werden. Infolgedessen ist es möglich, dass genügend Cs frei gesetzt werden kann, um die Zusammensetzung innerhalb der Mischkristallreihe von Cs-Analcim (>2 wt.% Cs2O) zu Na-Pollucit (23–43 wt.% Cs2O) zu verschieben. Die zeitliche Einordnung dieser späten Cs-Anreicherung wird als quasi zeitgleich oder im direkten Anschluss an die vollständige Kristallisation der Pegmatitschmelze interpretiert. Es kann allerdings nicht vernachlässigt werden, dass auch jüngere hydrothermale Ereignisse, die den Pegmatitkörper nachträglich überprägen, ähnliche hydrothermale Prozesse hervorrufen können. Somit konnte gezeigt werden, dass es durch Kombination dieser magmatischen und hydrothermalen Prozessen möglich ist, genügend Cs anzureichern, um die Bildung von massiven Pollucitmineralisationen in LCT-Pegmatitsystemen zu ermöglichen. Dieses genetische Modell kann nun dazu genutzt werden, um das Potential von Vorkommen von massiven Pollucitmineralisationen in LCT-Pegmatitsystemen in Westaustralien und weltweit besser einzuschätzen.

Lithium Cesium Tantalum (LCT) pegmatites are important resources for rare metals like Cesium, Lithium or Tantalum, whose demand increased markedly during the past decade. At present, Cs is known to occur in economic quantities only from the two LCT pegmatite deposits at Bikita located in Zimbabwe and Tanco in Canada. Host for this Cs mineralisation is the extreme rare zeolite group mineral pollucite. However, at Bikita and Tanco, pollucite forms huge massive, lensoid shaped and almost monomineralic pollucite mineralisations that occur within the upper portions of the pegmatite. In addition, both pegmatite deposits have a comparable regional geological background as they are hosted within greenstone belts and yield a Neoarchean age of about 2,600 Ma. Furthermore, at present the genesis of these massive pollucite mineralisations was not yet investigated in detail. Major portions of Western Australia consist of Meso- to Neoarchean crustal units (e.g., Yilgarn Craton, Pilbara Craton) that are known to host a large number of LCT pegmatite systems. Among them are the LCT pegmatite deposits Greenbushes (Li, Ta) and Wodgina (Ta, Sn). In addition, small amounts of pollucite were recovered from one single diamond drill core at the Londonderry pegmatite field. Despite that, no systematic investigations and/or exploration studies were conducted for the mode of occurrence of Cs and especially that of pollucite in Western Australia. In the course of the present study nineteen individual pegmatites and pegmatite fields located on the Yilgarn Craton, Pilbara Craton and Kimberley province have been visited and inspected for the occurrence of the Cs mineral pollucite. However, no pollucite could be detected in any of the investigated pegmatites. Four of the inspected LCT-pegmatite systems, namely the Londonderry pegmatite field, the Mount Deans pegmatite field, the Cattlin Creek LCT pegmatite deposit (Yilgarn Craton) and the Wodgina LCT pegmatite deposit (Pilbara Craton) was sampled and investigated in detail. In addition, samples from the Bikita pegmatite field (Zimbabwe Craton) were included into the present study in order to compare the Western Australian pegmatites with a massive pollucite mineralisation bearing LCT pegmatite system. This thesis presents new petrographical, mineralogical, mineralchemical, geochemical, geochronological, fluid inclusion and stable and radiogenic isotope data. The careful interpretation of this data enhances the understanding of the LCT pegmatite systems in Western Australia and Zimbabwe. All of the four investigated LCT pegmatite systems in Western Australia, crop out in similar geological settings, exhibit comparable internal structures, geochemistry and mineralogy to that of the Bikita pegmatite field in Zimbabwe. Furthermore, in all LCT pegmatite systems evidences for late stage hydrothermal processes (e.g., replacement of feldspars) and associated Cs enrichment (e.g., Cs enriched rims on mica, beryl and tourmaline) is documented. With the exception of the Wodgina LCT pegmatite deposit, that yield a Mesoarchean crystallisation age (approx. 2,850 Ma), all other LCT pegmatite systems gave comparable Neoarchean ages of 2,630 Ma to 2,600 Ma. The almost identical ages of the LCT pegmatite systems of the Yilgarn and Zimbabwe cratons suggests, that the process of LCT pegmatite formation at the end of the Neoarchean was active worldwide. Nevertheless, essential distinguishing feature of the Bikita pegmatite field is the presence of massive pollucite mineralisations that resulted from a process that is not part of the general development of LCT pegmatites and is associated with the extreme enrichment of Cs. The new findings of the present study obtained from the Bikita pegmatite field and the Western Australian LCT pegmatite systems significantly improve the knowledge of Cs behaviour in LCT pegmatite systems. Therefore, it is now possible to suggest a genetical model for the formation of massive pollucite mineralisations within LCT pegmatite systems. LCT pegmatites are generally granitic in composition and are interpreted to represent highly fractionated and geochemically specialised derivates from granitic melts. Massive pollucite mineralisation bearing LCT pegmatites evolve from large and voluminous pegmatite melts that intrude as single body along structures within an extensional tectonic setting. After emplacement, initial crystallisation will develop the border and wall zone of the pegmatites, while due to fractionated crystallisation immobile elements (i.e., Cs, Rb) become enriched within the remaining melt and associated hydrothermal fluids. Following this initial crystallisation, a relatively small portion (0.5–1 vol.%) of immiscible melt or fluid will separate during cooling. This immiscible partial melt/fluid is enriched in Al2O3 and Na2O, as well as depleted in SiO2 and will crystallise as analcime. In addition, this melt might allready contains up to 1–2 wt.% Cs2O. However, due to the effects of fluxing components (e.g., H2O, F, B) this analcime melt becomes undercooled which prevents crystallisation of the analcime as intergranular grains. Since this analcime melt exhibits a lower relative gravity when compared to the remaining pegmatite melt the less dense analcime melt will start to ascent gravitationally and accumulate within the upper portion of the pegmatite sheet. At the same time, the remaining melt will start to crystallise separately and form the inner portions of the pegmatite. This crystallisation is characterised by still ongoing fractionation and enrichment of incompatible elements (i.e., Cs, Rb) within the last crystallising minerals (e.g., lepidolite) or concentration of these incompatible elements within exsolving hydrothermal fluids. As analcime and pollucite form a continuous solid solution series, the analcime melt is able to incorporate any available Cs from the melt and/or associated hydrothermal fluids and crystallise as Cs-analcime in the upper portion of the pegmatite sheet. Continuing hydrothermal activity and ongoing substitution of Cs will then start to shift the composition from Cs-analcime composition towards Na-pollucite composition. In addition, if analcime is cooled below 400 °C it is subjected to a negative thermal expansion of about 1 vol.%. This contraction results in the formation of a prominent network of cracks that is filled by late stage minerals (e.g., lepidolite, quartz, feldspar and petalite). Certainly, prior to filling, this network of cracks enhances the available conduits for late stage hydrothermal fluids and the Cs substitution mechanism within the massive pollucite mineralisation. Furthermore, during cooling of the pegmatite, prominent late stage mineral replacement reactions (e.g., replacement of K-feldspar by lepidolite, cleavelandite, and quartz) as well as subsolidus self organisation processes in feldspars take place. These processes are suggested to release additional incompatible elements (e.g., Cs, Rb) into late stage hydrothermal fluids. As feldspar forms large portions of pegmatite a considerable amount of Cs is released and transported via the hydrothermal fluids towards the massive pollucite mineralisation in the upper portion of the pegmatite. Consequently, the initial analcime can accumulate enough Cs in order to shift its composition from the Cs-analcime member (>2 wt.% Cs2O) towards the Na-pollucite member (23–43 wt.% Cs2O) of the solid solution series. The timing of this late stage Cs enrichment is interpreted to be quasi contemporaneous or immediately after the complete crystallisation of the pegmatite melt. However, much younger hydrothermal events that overprint the pegmatite are also interpreted to cause similar results. Hence, it has been demonstrated that the combination of this magmatic and hydrothermal processes is capable to generate an extreme enrichment in Cs in order to explain the formation of massive pollucite mineralisations within LCT pegmatite systems. This genetic model can now be applied to evaluate the potential for occurrences of massive pollucite mineralisations within LCT pegmatite systems in Western Australia and worldwide.:Contents Abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . iii Zusammenfassung . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v Versicherung . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . vii Acknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ix Contents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xi 1. Introduction 1 1.1. Motivation and Scope of the Thesis . . . . . . . . . . . . . . . . . . . . . . . . 1 1.2. Structure of the Thesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 2. Fundamentals 7 2.1. The Alkali Metal Cesium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 2.1.1. Distribution of Cesium . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 2.1.2. Mineralogy of Cesium . . . . . . . . . . . . . . . . . . . . . . . . . . . 8 2.1.3. Geochemical Behaviour of Cesium . . . . . . . . . . . . . . . . . . . . 13 2.1.4. Economy of Cesium . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15 2.2. Pollucite – (Cs,Na)2Al2Si4O12×H2O . . . . . . . . . . . . . . . . . . . . . . . . 16 2.2.1. Crystal Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 2.2.2. Analcime–Pollucite–Series . . . . . . . . . . . . . . . . . . . . . . . . . 17 2.2.3. Formation of Pollucite . . . . . . . . . . . . . . . . . . . . . . . . . . . 17 2.2.4. Pollucite Occurences . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 2.3. Pegmatites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34 2.3.1. General Characteristics of Pegmatites . . . . . . . . . . . . . . . . . . 34 2.3.2. Controls on Pegmatite Formation and Evolution . . . . . . . . . . . . . 40 2.3.3. Pegmatite Age Distribution and Continental Crust Formation . . . . . . 43 3. Geological Settings of Archean Cratons 47 3.1. Zimbabwe Craton . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47 3.1.1. Tectonostratigraphic Subdivision . . . . . . . . . . . . . . . . . . . . . 48 3.1.2. Tectonometamorphic Evolution of the Northern Limpopo Thrust Zone . 49 3.1.3. Pegmatites within the Zimbabwe Craton . . . . . . . . . . . . . . . . . 52 3.1.4. Masvingo Greenstone Belt . . . . . . . . . . . . . . . . . . . . . . . . . 53 3.1.5. Geological Setting of the Bikita Pegmatite District . . . . . . . . . . . . 58 3.2. Yilgarn Craton . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 3.2.1. Tectonostratigraphic Framework and Geological Development . . . . . 62 3.2.2. Tectonic Models for the Development . . . . . . . . . . . . . . . . . . . 70 3.2.3. Pegmatites within the Yilgarn Craton . . . . . . . . . . . . . . . . . . . 76 3.2.4. Geological setting of the Londonderry Pegmatite Field . . . . . . . . . . 76 3.2.5. Geological Setting of the Mount Deans Pegmatite Field . . . . . . . . . 85 3.2.6. Geological Setting of the Cattlin Creek Pegmatite Deposit . . . . . . . . 91 3.3. Pilbara Craton . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 3.3.1. Tectonostratigraphic Framework and Geological Development . . . . . 99 3.3.2. Tectonic Model for the Development . . . . . . . . . . . . . . . . . . . 101 3.3.3. Pegmatites within the Pilbara Craton . . . . . . . . . . . . . . . . . . . 105 3.3.4. Geological Setting of the Wodgina Pegmatite District . . . . . . . . . . 106 4. Fieldwork and Sampling of Selected Pegmatites and Pegmatite Fields 115 4.1. Bikita Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 115 4.2. Londonderry Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . . . . . . 115 4.2.1. Londonderry Feldspar Quarry Pegmatite . . . . . . . . . . . . . . . . . 115 4.2.2. Lepidolite Hill Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . 117 4.2.3. Tantalite Hill Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . . 118 4.3. Mount Deans Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . . . . . 118 4.3.1. Type I – Flat Lying Pegmatites . . . . . . . . . . . . . . . . . . . . . . . 118 4.3.2. Type II – Steeply Dipping Pegmatites . . . . . . . . . . . . . . . . . . . 120 4.4. Cattlin Creek Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 120 4.5. Wodgina LCT-Pegmatite Deposit . . . . . . . . . . . . . . . . . . . . . . . . . . 121 4.5.1. Mount Tinstone Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . 123 4.5.2. Mount Cassiterite Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . 123 5. Petrography and Mineralogy 139 5.1. Quantitative Mineralogy by Means of Mineral Liberation Analysis . . . . . . . . 141 5.2. Mineralogical and Petrographical Characteristics of Individual Mineral Groups . 141 5.2.1. Feldspar . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 141 5.2.2. Quartz . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 144 5.2.3. Mica . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 5.2.4. Pollucite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147 5.2.5. Petalite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149 5.2.6. Spodumene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149 5.2.7. Beryl . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 154 5.2.8. Tourmaline . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 155 5.2.9. Apatite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 5.2.10. Ta-, Nb- and Sn-oxides . . . . . . . . . . . . . . . . . . . . . . . . . . 157 5.3. Reconstruction of the General Crystallisation Sequence . . . . . . . . . . . . . 162 6. Geochemistry 165 6.1. Major Elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 165 6.2. Selected Minor and Trace Elements . . . . . . . . . . . . . . . . . . . . . . . . 174 6.3. Fractionation Indicators . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 175 6.4. Rare Earth Elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 185 7. Geochronology 193 7.1. 40Ar/39Ar-Method on Mica . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 193 7.1.1. Bikita Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . . . . . 194 7.1.2. Mount Deans Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . 195 7.1.3. Londonderry Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . 195 7.1.4. Cattlin Creek Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . 195 7.1.5. Wodgina Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . . . 198 7.2. Th-U-Total Pb Monazite Dating . . . . . . . . . . . . . . . . . . . . . . . . . . . 201 7.2.1. Monazite Ages . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 201 7.3. U/Pb Dating of Selected Ta-, Nb- and Sn-Oxide Minerals . . . . . . . . . . . . 203 7.3.1. Bikita Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . . . . . 203 7.3.2. Londonderry Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . 203 7.3.3. Mount Deans Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . 206 7.3.4. Cattlin Creek Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . 206 7.3.5. Wodgina Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . . . 207 8. Fluid Inclusion Study 211 8.1. Bikita Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 211 8.2. Wodgina Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 211 8.3. Carbon Isotope Analysis on Fluid Inclusion Gas of Selected Mineral Phases . . 212 9. Stable and Radiogenic Isotopes 217 9.1. Whole Rock Sm/Nd-Isotopy . . . . . . . . . . . . . . . . . . . . . . . . . . . . 217 9.1.1. New Whole Rock Sm/Nd Data . . . . . . . . . . . . . . . . . . . . . . 217 9.2. Lithium Isotope Analysis on Selected Mineral Phases . . . . . . . . . . . . . . . 220 9.2.1. New Lithium Isotope Data . . . . . . . . . . . . . . . . . . . . . . . . . 220 10.Discussion 227 10.1. Regional Geological and Tectonomagmatic Development . . . . . . . . . . . . 227 10.1.1. Constraints from Field Evidence . . . . . . . . . . . . . . . . . . . . . . 227 10.1.2. Petrographical and Mineralogical Constraints . . . . . . . . . . . . . . 229 10.1.3. Geochemical Constraints . . . . . . . . . . . . . . . . . . . . . . . . . 230 10.1.4. Isotopic Constraints . . . . . . . . . . . . . . . . . . . . . . . . . . . . 232 10.1.5. Constraints from Fluid Inclusion Data . . . . . . . . . . . . . . . . . . . 233 10.1.6. Geochronological Constrains . . . . . . . . . . . . . . . . . . . . . . . 233 10.2. Massive Pollucite Mineralisations . . . . . . . . . . . . . . . . . . . . . . . . . . 243 10.2.1. Unique Characteristics of Massive Pollucite Mineralisations . . . . . . . 243 10.2.2. New Concepts for the Formation of Massive Pollucite Mineralisations . . 252 10.3. Genetic Model for the Formation of Massive Pollucite Mineralisations within LCT Pegmatite Systems . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 264 11.Summary and Conclusions 267 References 273 Lists of Abbreviations 309 General Abbreviations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 309 Mineral Abbreviations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 310 List of Figures 311 List of Tables 315 Appendix 317 A. Legend for Topographic Maps 319 B. Sample List 323 C. Methodology 331 C.1. Quantitative Mineralogy by Means of Mineral Liberation Analysis . . . . . . . . 331 C.2. Geochemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 331 C.3. 40Ar/39Ar-Method on Mica . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 335 C.4. Th-U-Total Pb Monazite Dating . . . . . . . . . . . . . . . . . . . . . . . . . . . 335 C.5. U/Pb Dating of Selected Ta-, Nb- and Sn-Oxide Minerals . . . . . . . . . . . . 336 C.6. Fluid Inclusion Study . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 337 C.7. Whole Rock Sm/Nd-Isotopy . . . . . . . . . . . . . . . . . . . . . . . . . . . . 338 C.8. Lithium Isotope Analysis on Selected Mineral Phases . . . . . . . . . . . . . . . 338 D. Data – Mineral Liberation Analysis 341 E. Data – Geochemistry 345 F. Data – Geochronology 349 G. Data – Stable and Radiogenic Isotopes 353 / Lithium-Caesium-Tantal-(LCT) Pegmatite repräsentieren eine bedeutende Quelle für seltene Metalle, deren Bedarf im letzten Jahrzehnt beträchtlich angestiegen ist. Im Falle von Caesium sind zurzeit weltweit nur zwei LCT-Pegmatitlagerstätten bekannt, die abbauwürdige Vorräte an Cs enthalten. Dies sind die LCT-Pegmatitlagerstätten Bikita in Simbabwe und Tanco in Kanada. Das Wirtsmineral für diese Cs-Mineralisation ist das extrem selten auftretende Zeolith-Gruppen-Mineral Pollucit. In den Lagerstätten Bikita und Tanco bildet Pollucit dagegen massive, linsenförmige und fast monomineralische Pollucitmineralisationen, die in den oberen Bereichen der Pegmatitkörper anstehen. Zusätzlich befinden sich beide Lagerstätten in geologisch vergleichbaren Einheiten. Die Nebengesteine sind Grünsteingürtel die ein neoarchaisches Alter von ca. 2,600 Ma aufweisen. Die Bildung derartiger massiver Pollucitmineralisationen ist bis jetzt noch nicht detailliert untersucht worden. Große Bereiche von Westaustralien werden von meso- bis neoarchaischen Krusteneinheiten (z.B. Yilgarn Kraton, Pilbara Kraton) aufgebaut, von denen auch eine große Anzahl an LCT-Pegmatitsystemen bekannt sind. Darunter befinden sich unter anderem die LCT-Pegmatitlagerstätten Greenbushes (Li, Ta) und Wodgina (Ta, Sn). Zusätzlich wurden kleine Mengen an Pollucit in einer einzigen Kernbohrung im Londonderry Pegmatitfeld angetroffen. Ungeachtet dessen, wurden in Westaustralien bis jetzt keine systematischen Untersuchungen und/oder Explorationskampagnen auf Vorkommen von Cs und speziell der von Pollucit durchgeführt. Im Verlauf dieser Studie wurden insgesamt neunzehn verschiedene Pegmatitvorkommen und Pegmatitfelder des Yilgarn Kratons, Pilbara Kratons und der Kimberley Provinz auf das Vorkommen des Minerals Pollucit untersucht. Allerdings konnte in keinem der untersuchten LCT-Pegmatitsystemen Pollucit nachgewiesen werden. Von vier der untersuchten LCT-Pegmatitsystemen, dem Londonderry Pegmatitfeld, dem Mount Deans Pegmatitfeld, der Cattlin Creek LCT-Pegmatitlagerstätte (Yilgarn Kraton) und der Wodgina LCT-Pegmatitlagerstätte (Pilbara Kraton) wurden detailliert Proben entnommen und weitergehend untersucht. Zusätzlich wurden die massiven Pollucitmineralisationen im Bikita Pegmatitfeld beprobt und in die detailierten Untersuchungen einbezogen. Der Probensatz aus dem Bikita Pegmatitfeld dient als Referenzmaterial mit dem die Pegmatitproben aus Westaustralien verglichen werden. Die vorliegende Arbeit fasst die wesentlichen Ergebnisse der petrographischen, mineralogischen, mineralchemischen, geochemischen und geochronologischen Untersuchungen sowie der Flüssigkeitseinschlussuntersuchungen und stabilen und radiogenen Isotopenzusammensetzungen zusammen. Alle vier der in Westaustralien untersuchten LCT-Pegmatitsysteme kommen in geologisch ähnlichen Rahmengesteinen vor, weisen einen vergleichbaren internen Aufbau, geochemische Zusammensetzung und Mineralogie zu dem des Bikita Pegmatitfeldes in Simbabwe auf. Weiterhin konnten in allen LCT-Pegmatitsystemen Hinweise für späte hydrothermale Prozesse (z.B. Verdrängung von Feldspat) nachgewiesen werden, die einhergehend mit einer Anreicherung von Cs verbunden sind (z.B. Cs-angereicherte Säume um Glimmer, Beryll und Turmalin). Mit der Ausnahme der Wodgina LCT-Pegmatitlagerstätte, in der ein mesoarchaisches Kristallisationsalter (ca. 2,850 Ma) nachgewiesen wurde, lieferten die Altersdatierungen in den anderen LCT-Pegmatitsystemen übereinstimmende neoarchaische Alter von 2,630 Ma bis 2,600 Ma. Diese fast identischen Alter der LCT-Pegmatitsysteme des Yilgarn und Zimbabwe Kratons suggerieren, dass die Prozesse, die zur LCT-Pegmatitbildung am Ende des Neoarchaikums führten, weltweit aktiv waren. Ungeachtet dessen stellt das Vorhandensein von massiver Pollucitmineralisation das Alleinstellungsmerkmal des Bikita Pegmatitfeldes dar, welche sich infolge eines Prozesses gebildet haben der nicht Bestandteil der üblichen LCT-Pegmatitentwicklung ist und sich durch eine extreme Anreicherung an Cs unterscheidet. Die neuen Ergebnisse die in dieser Studie von den Bikita Pegmatitfeld und den Westaustralischen LCT-Pegmatitsystemen gewonnen wurden, verbessern das Verständnis des Verhaltens von Cs in LCT-Pegmatitsystemen deutlich. Somit ist es nun möglich, ein genetisches Modell für die Bildung von massiven Pollucitmineralisationen in LCT-Pegmatitsystemen vorzustellen. LCT-Pegmatite weisen im Allgemeinen eine granitische Zusammensetzung auf und werden als Kristallisat von hoch fraktionierten und geochemisch spezialisierten granitischen Restschmelzen interpretiert. Die Bildung von massiven Pollucitmineralisationen ist nur aus großen und voluminösen Pegmatitschmelzen, die als einzelner Körper entlang von Störungen in extensionalen Stressregimen intrudieren möglich. Nach Platznahme der Schmelze bildet die beginnende Kristallisation zunächst die Kontakt- und Randzone des Pegmatits, wobei infolge von fraktionierter Kristallisation die immobilen Elemente (v.a. Cs, Rb) in der verbleibenden Restschmelze angereichert werden. Im Anschluss an diese erste Kristallisation entmischt sich nach Abkühlung eine sehr kleine Menge (0.5–1 vol.%) Schmelze und/oder Fluid von der Restschmelze. Diese nicht mischbare Teilschmelze/-fluid ist angereichert an Al2O3 und Na2O sowie verarmt an SiO2 und kristallisiert als Analcim. Zusätzlich kann diese Schmelze bereits mit 1–2 wt.% Cs2O angereichert sein. Aufgrund der Auswirkung von Flussmitteln (z.B. H2O, F, B) wird allerdings der Schmelzpunkt dieser Analcimschmelze herabgesetzt und so die Kristallisation des Analcims als intergranulare Körner verhindert. Da diese Analcimschmelze im Vergleich zu der restlichen Schmelze eine geringere relative Dichte besitzt, beginnt sie gravitativ aufzusteigen und sich in den oberen Bereichen des Pegmatitkörpers zu akkumulieren. Währenddessen beginnt die restliche Schmelze separat zu kristallisieren und die inneren Bereiche des Pegmatits zu bilden. Diese Kristallisation ist einhergehend mit fortschreitender Fraktionierung und der Anreicherung von inkompatiblen Elementen (v.a. Cs, Rb) in den sich als letztes bildenden Mineralphasen (z.B. Lepidolit) oder der Konzentration der inkompatiblen Element in die sich entmischenden hydrothermalen Fluiden. Da Analcim und Pollucit eine lückenlose Mischungsreihe bilden, ist die Analcimschmelze in der Lage, alles verfügbare Cs von der Restschmelze und/oder assoziierten hydrothermalen Fluiden an sich zu binden und als Cs-Analcim im oberen Bereich des Pegmatitkörpers zu kristallisieren. Fortschreitende hydrothermale Aktivität und Substitution von Cs verschiebt dann die Zusammensetzung des Analcims von der Cs-Analcim- zu Na-Pollucitzusammensetzung. Zusätzlich erfährt der Analcim bei Abkühlung unter 400 °C eine negative thermische Expansion von ca. 1 vol.%. Diese Kontraktion führt zu der Bildung des markanten Rissnetzwerkes das durch späte Mineralphasen (z.B. Lepidolit, Quarz, Feldspat und Petalit) gefüllt wird. Vor der Mineralisation allerdings, erhöht dieses Netzwerk an Rissen die verfügbaren Wegsamkeiten für die späten hydrothermalen Fluide und begünstigt somit den Cs-Substitutionsmechanismus in der massiven Pollucitmineralisation. Weiterhin kommt es bei der Abkühlung des Pegmatits zu späten Mineralverdrängungsreaktionen (z.B. Verdrängung von K-Feldspat durch Lepidolit, Cleavelandit und Quarz), sowie zu Subsolidus-Selbstordnungsprozessen in Feldspäten. Diese Prozesse werden weiterhin interpretiert inkompatible Elemente (z.B. Cs, Rb) in die späten hydrothermalen Fluide freizusetzen. Da Feldspäte große Teile der Pegmatite bilden, kann somit eine beträchtliche Menge an Cs freigeben werden und durch die späten hydrothermalen Fluide in die massive Pollucitmineralisation in den oberen Bereichen des Pegmatitkörpers transportiert werden. Infolgedessen ist es möglich, dass genügend Cs frei gesetzt werden kann, um die Zusammensetzung innerhalb der Mischkristallreihe von Cs-Analcim (>2 wt.% Cs2O) zu Na-Pollucit (23–43 wt.% Cs2O) zu verschieben. Die zeitliche Einordnung dieser späten Cs-Anreicherung wird als quasi zeitgleich oder im direkten Anschluss an die vollständige Kristallisation der Pegmatitschmelze interpretiert. Es kann allerdings nicht vernachlässigt werden, dass auch jüngere hydrothermale Ereignisse, die den Pegmatitkörper nachträglich überprägen, ähnliche hydrothermale Prozesse hervorrufen können. Somit konnte gezeigt werden, dass es durch Kombination dieser magmatischen und hydrothermalen Prozessen möglich ist, genügend Cs anzureichern, um die Bildung von massiven Pollucitmineralisationen in LCT-Pegmatitsystemen zu ermöglichen. Dieses genetische Modell kann nun dazu genutzt werden, um das Potential von Vorkommen von massiven Pollucitmineralisationen in LCT-Pegmatitsystemen in Westaustralien und weltweit besser einzuschätzen.:Contents Abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . iii Zusammenfassung . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v Versicherung . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . vii Acknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ix Contents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . xi 1. Introduction 1 1.1. Motivation and Scope of the Thesis . . . . . . . . . . . . . . . . . . . . . . . . 1 1.2. Structure of the Thesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 2. Fundamentals 7 2.1. The Alkali Metal Cesium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 2.1.1. Distribution of Cesium . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 2.1.2. Mineralogy of Cesium . . . . . . . . . . . . . . . . . . . . . . . . . . . 8 2.1.3. Geochemical Behaviour of Cesium . . . . . . . . . . . . . . . . . . . . 13 2.1.4. Economy of Cesium . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15 2.2. Pollucite – (Cs,Na)2Al2Si4O12×H2O . . . . . . . . . . . . . . . . . . . . . . . . 16 2.2.1. Crystal Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 2.2.2. Analcime–Pollucite–Series . . . . . . . . . . . . . . . . . . . . . . . . . 17 2.2.3. Formation of Pollucite . . . . . . . . . . . . . . . . . . . . . . . . . . . 17 2.2.4. Pollucite Occurences . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 2.3. Pegmatites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34 2.3.1. General Characteristics of Pegmatites . . . . . . . . . . . . . . . . . . 34 2.3.2. Controls on Pegmatite Formation and Evolution . . . . . . . . . . . . . 40 2.3.3. Pegmatite Age Distribution and Continental Crust Formation . . . . . . 43 3. Geological Settings of Archean Cratons 47 3.1. Zimbabwe Craton . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47 3.1.1. Tectonostratigraphic Subdivision . . . . . . . . . . . . . . . . . . . . . 48 3.1.2. Tectonometamorphic Evolution of the Northern Limpopo Thrust Zone . 49 3.1.3. Pegmatites within the Zimbabwe Craton . . . . . . . . . . . . . . . . . 52 3.1.4. Masvingo Greenstone Belt . . . . . . . . . . . . . . . . . . . . . . . . . 53 3.1.5. Geological Setting of the Bikita Pegmatite District . . . . . . . . . . . . 58 3.2. Yilgarn Craton . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 3.2.1. Tectonostratigraphic Framework and Geological Development . . . . . 62 3.2.2. Tectonic Models for the Development . . . . . . . . . . . . . . . . . . . 70 3.2.3. Pegmatites within the Yilgarn Craton . . . . . . . . . . . . . . . . . . . 76 3.2.4. Geological setting of the Londonderry Pegmatite Field . . . . . . . . . . 76 3.2.5. Geological Setting of the Mount Deans Pegmatite Field . . . . . . . . . 85 3.2.6. Geological Setting of the Cattlin Creek Pegmatite Deposit . . . . . . . . 91 3.3. Pilbara Craton . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99 3.3.1. Tectonostratigraphic Framework and Geological Development . . . . . 99 3.3.2. Tectonic Model for the Development . . . . . . . . . . . . . . . . . . . 101 3.3.3. Pegmatites within the Pilbara Craton . . . . . . . . . . . . . . . . . . . 105 3.3.4. Geological Setting of the Wodgina Pegmatite District . . . . . . . . . . 106 4. Fieldwork and Sampling of Selected Pegmatites and Pegmatite Fields 115 4.1. Bikita Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 115 4.2. Londonderry Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . . . . . . 115 4.2.1. Londonderry Feldspar Quarry Pegmatite . . . . . . . . . . . . . . . . . 115 4.2.2. Lepidolite Hill Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . 117 4.2.3. Tantalite Hill Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . . 118 4.3. Mount Deans Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . . . . . 118 4.3.1. Type I – Flat Lying Pegmatites . . . . . . . . . . . . . . . . . . . . . . . 118 4.3.2. Type II – Steeply Dipping Pegmatites . . . . . . . . . . . . . . . . . . . 120 4.4. Cattlin Creek Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 120 4.5. Wodgina LCT-Pegmatite Deposit . . . . . . . . . . . . . . . . . . . . . . . . . . 121 4.5.1. Mount Tinstone Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . 123 4.5.2. Mount Cassiterite Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . 123 5. Petrography and Mineralogy 139 5.1. Quantitative Mineralogy by Means of Mineral Liberation Analysis . . . . . . . . 141 5.2. Mineralogical and Petrographical Characteristics of Individual Mineral Groups . 141 5.2.1. Feldspar . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 141 5.2.2. Quartz . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 144 5.2.3. Mica . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146 5.2.4. Pollucite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147 5.2.5. Petalite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149 5.2.6. Spodumene . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149 5.2.7. Beryl . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 154 5.2.8. Tourmaline . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 155 5.2.9. Apatite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 5.2.10. Ta-, Nb- and Sn-oxides . . . . . . . . . . . . . . . . . . . . . . . . . . 157 5.3. Reconstruction of the General Crystallisation Sequence . . . . . . . . . . . . . 162 6. Geochemistry 165 6.1. Major Elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 165 6.2. Selected Minor and Trace Elements . . . . . . . . . . . . . . . . . . . . . . . . 174 6.3. Fractionation Indicators . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 175 6.4. Rare Earth Elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 185 7. Geochronology 193 7.1. 40Ar/39Ar-Method on Mica . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 193 7.1.1. Bikita Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . . . . . 194 7.1.2. Mount Deans Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . 195 7.1.3. Londonderry Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . 195 7.1.4. Cattlin Creek Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . 195 7.1.5. Wodgina Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . . . 198 7.2. Th-U-Total Pb Monazite Dating . . . . . . . . . . . . . . . . . . . . . . . . . . . 201 7.2.1. Monazite Ages . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 201 7.3. U/Pb Dating of Selected Ta-, Nb- and Sn-Oxide Minerals . . . . . . . . . . . . 203 7.3.1. Bikita Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . . . . . 203 7.3.2. Londonderry Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . 203 7.3.3. Mount Deans Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . 206 7.3.4. Cattlin Creek Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . 206 7.3.5. Wodgina Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . . . 207 8. Fluid Inclusion Study 211 8.1. Bikita Pegmatite Field . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 211 8.2. Wodgina Pegmatite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 211 8.3. Carbon Isotope Analysis on Fluid Inclusion Gas of Selected Mineral Phases . . 212 9. Stable and Radiogenic Isotopes 217 9.1. Whole Rock Sm/Nd-Isotopy . . . . . . . . . . . . . . . . . . . . . . . . . . . . 217 9.1.1. New Whole Rock Sm/Nd Data . . . . . . . . . . . . . . . . . . . . . . 217 9.2. Lithium Isotope Analysis on Selected Mineral Phases . . . . . . . . . . . . . . . 220 9.2.1. New Lithium Isotope Data . . . . . . . . . . . . . . . . . . . . . . . . . 220 10.Discussion 227 10.1. Regional Geological and Tectonomagmatic Development . . . . . . . . . . . . 227 10.1.1. Constraints from Field Evidence . . . . . . . . . . . . . . . . . . . . . . 227 10.1.2. Petrographical and Mineralogical Constraints . . . . . . . . . . . . . . 229 10.1.3. Geochemical Constraints . . . . . . . . . . . . . . . . . . . . . . . . . 230 10.1.4. Isotopic Constraints . . . . . . . . . . . . . . . . . . . . . . . . . . . . 232 10.1.5. Constraints from Fluid Inclusion Data . . . . . . . . . . . . . . . . . . . 233 10.1.6. Geochronological Constrains . . . . . . . . . . . . . . . . . . . . . . . 233 10.2. Massive Pollucite Mineralisations . . . . . . . . . . . . . . . . . . . . . . . . . . 243 10.2.1. Unique Characteristics of Massive Pollucite Mineralisations . . . . . . . 243 10.2.2. New Concepts for the Formation of Massive Pollucite Mineralisations . . 252 10.3. Genetic Model for the Formation of Massive Pollucite Mineralisations within LCT Pegmatite Systems . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 264 11.Summary and Conclusions 267 References 273 Lists of Abbreviations 309 General Abbreviations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 309 Mineral Abbreviations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 310 List of Figures 311 List of Tables 315 Appendix 317 A. Legend for Topographic Maps 319 B. Sample List 323 C. Methodology 331 C.1. Quantitative Mineralogy by Means of Mineral Liberation Analysis . . . . . . . . 331 C.2. Geochemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 331 C.3. 40Ar/39Ar-Method on Mica . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 335 C.4. Th-U-Total Pb Monazite Dating . . . . . . . . . . . . . . . . . . . . . . . . . . . 335 C.5. U/Pb Dating of Selected Ta-, Nb- and Sn-Oxide Minerals . . . . . . . . . . . . 336 C.6. Fluid Inclusion Study . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 337 C.7. Whole Rock Sm/Nd-Isotopy . . . . . . . . . . . . . . . . . . . . . . . . . . . . 338 C.8. Lithium Isotope Analysis on Selected Mineral Phases . . . . . . . . . . . . . . . 338 D. Data – Mineral Liberation Analysis 341 E. Data – Geochemistry 345 F. Data – Geochronology 349 G. Data – Stable and Radiogenic Isotopes 353

La période à 2.1 Ga du Craton d'Afrique de l'Ouest est une période de production massive de leucogranites. Le batholite de Ferkessédougou (Côte d'Ivoire) est, tant du point de vue de sa localisation, de son extension, que de son histoire dans la littérature géologique, un élément majeur de l'interprétation de la géodynamique du Paléoprotérozoïque. Tour à tour considéré comme un granite de ride augéosynclinale puis un granite typique de collision de type moderne, il n'en demeure pas moins le jalon d'un accident crustal NNE-SSW important du Paléoprotérozoïque : le linéament GFB (Greenville-Ferkessédougou-Bobodioulasso). Son étude et son interprétation s'inscrivent dans le débat sur la validité de la tectonique des plaques à cette période. Par analyse satellitale, nous montrons que ce batholite de 400 km de long sur une cinquantaine de large est en fait constitué d'une grande quantité de plutons adjacents relativement proches pétrographiquement. Le contexte général de son encaissant métavolcanique et/ou métasédimentaire birimien (2.2 Ga à 2.1 Ga) faiblement déformé et épizonal ainsi que l'absence de déformations tangentielles additionnelles traduisent la faiblesse de l'épaississement crustal synchrone de sa genèse. Sa fabrique déterminée par analyse structurale classique et par analyse de susceptibilité magnétique est double. La première "magmatique" est le plus souvent coaxiale, d'orientation et de pendages variables. Elle dessine des trajectoires courbes et traduit la mise en place des plutons. Elle est l'équivalent de ce qui a été rapporté à la déformation D1 autour de granites anciens. La seconde, magmatique à post-solidus, beaucoup mieux réglée, souvent constrictive ou non-coaxiale, est l'image d'un champ de contrainte régional transcurrent qui se rapporte à la déformation D2 classique dans la chronologie du craton birimien. Les plutons du batholite granitique de Ferké sont syntectoniques de cette grande phase de raccourcissement (2.1 Ga à 2.07 Ga) du birimien. Le passage en revue des différents modèles de plutons, l'analyse du secteur SW de la Côte d'Ivoire présentant des dykes granites à deux micas et des faciès métamorphiques de plus haut grade, un survol des formations de l'Imataca au Surinam, ainsi que le bilan de l'analyse structurale, nous amènent à proposer un modèle 3D pour le batholite de Ferké. Il correspondrait à un ensemble de petits plutons de forme laccolithique engagés dans la croûte superficielle entre les métasédiments épizonaux birimiens dits "de bassin", des gneiss et micaschistes probablement d'âge compris entre 2.2 et 2.1 Ga. Ces plutons seraient alimentés par un faisceau de dykes issus de la fracturation par ouverture en pull-apart du bassin sédimentaire SASCA-Bandama-Bobodioulasso. La fusion de la croûte essentiellement de type ceinture verte/TTG (absence de traces de matériel archéen), à la verticale du bassin subsident, aura incorporé du matériel sédimentaire de bassin. L'apport de chaleur nécessaire à la compensation de la faiblesse de l'épaississement aurait été fourni par le jeu décrochant des fractures majeures des pull-aparts. La place nécessaire à la mise en place des plutons exige une montée de ces derniers à la fin de la période extensive à la faveur de multiples cauldron-subsidences. Les conditions et régimes des deux déformations principales impliquent un lapse de temps très court entre la stabilisation de l'ouverture des bassins et leur fermeture. Par ailleurs, la variation du niveau structural du Nord-Est vers le Sud-Ouest, en particulier vers la zone probable de l'équivalent-racine au Vénézuéla-Surinam, montre un approfondissement et le rapprochement d'une croûte archéenne vers le SW alors qu'elle n'était jusqu'à présent reconnue que vers l'Ouest (Guinée, Dorsale de Man) ou soupçonnée vers l'Est au Ghana.

La composition chimique de la croûte continentale a significativement évolué à la transition Archéen-Protérozoïque (3000-2500 Ma), témoignant de changements géodynamiques majeurs à cette époque. Afin d'étudier l'expression et les origines de ces changements, qui sont encore mal contraints, j'ai étudié une diversité de granitoïdes qui se sont mis en place dans cette gamme d'âges à la marge Nord du craton du Kaapvaal, en Afrique du Sud. Ce travail a permis de préciser la typologie et l'origine des granitoïdes tardi-archéens ; ceux-ci peuvent être classés dans trois grands groupes : (1) Les sanukitoïdes, représentés en Afrique du Sud par le pluton de Bulai, sont des magmas dérivant de l'interaction entre une péridotite mantellique et un composant riche en éléments incompatibles (TTG, liquide issu de la fusion de sédiments, et, plus rarement, fluide aqueux). Les sanukitoïdes peuvent être classés en deux groupes distincts, selon les mécanismes de cette hybridation : les low-Ti sanukitoids proviennent d'une simple hybridation du liquide silicaté avec la péridotite, alors que les high-Ti sanukitoids sont issus de la fusion d'un assemblage métasomatique à amphibole et phlogopite, résultant de ces interactions. Enfin, les mécanismes de différenciation des suites sanukitoïdes au niveau de la croûte sont contrôlées par des mécanismes de cristallisation fractionnée ou (moins vraisemblablement) de fusion partielle. (2) Les sanukitoïdes " marginaux ", représentés dans le craton du Kaapvaal par les plutons de Mashashane, Matlala, Matok et Moletsi, sont des granitoïdes résultant de l'interaction entre des sanukitoïdes et des magmas provenant de la fusion de croûte préexistante. Etant donné la large gamme de sources possibles (TTG, métasédiments, roches mafiques) d'un craton à l'autre, ce groupe est extrêmement diversifié. Leurs mécanismes de différenciation sont contrôlés par la cristallisation fractionnée. (3) Certains granites, tels que le batholite de Turfloop en Afrique du Sud, sont directement issus de la fusion de lithologies crustales (TTG, métasédiments et amphibolites). Au sein du craton du Kaapvaal, l'évolution spatio-temporelle du magmatisme tardi-archéen suit un schéma très caractéristique : les TTG se mettent en place entre ~3300 et ~2800 Ma, puis laissent la place à la genèse de l'ensemble des granitoïdes présentés ci-dessus, qui se déroule entre 2780 et 2590 Ma. Cette séquence d'évènements est reproduite au sein de tous les cratons du monde à la fin de l'Archéen. Elle témoigne de l'avènement des processus de recyclage crustal, puisque, par opposition aux TTG archéennes qui dérivent de métabasaltes juvéniles, les magmas tardi-archéens sont issus à la fois de la différenciation intracrustale et de l'interaction entre une péridotite et du matériel continental introduit dans le manteau. Cette dualité de processus pétrogénétiques est aussi très typique des épisodes magmatiques qui ont lieu à la fin des cycles de subduction-collision post-archéens. Ainsi, l'évolution de la composition des granitoïdes entre 3000 et 2500 Ma traduit vraisemblablement l'initiation d'une forme de tectonique des plaques proche du régime actuel. Celle-ci serait liée au refroidissement planétaire global, qui a probablement entraîné un " effet de seuil " dans l'évolution de l'épaisseur de la croûte océanique ainsi que la rhéologie et le volume de la croûte continentale, permettant ainsi à la subduction et à la collision de ne devenir thermo-mécaniquement stables qu'à partir de la fin de l'Archéen.