Alteration processes of a thick basaltic lava flow of the Paraná Basin (Brazil): petrographic and mineralogical studies

Journal of South American Earth Sciences 16 (2003) 423–444 www.elsevier.com/locate/jsames

Alteration processes of a thick basaltic lava flow of the Parana´ Basin (Brazil): petrographic and mineralogical studies F. Schenatoa,*, M.L.L. Formosob, P. Dudoignonc, A. Meunierc, D. Proustc, A. Masc a Universidade Luterana do Brasil—ULBRA, Av. Miguel Tostes, 101 Bairro Sa˜o Luı´s-Pre´dio 1, 92420-280 Canoas, RS, Brazil Centro de Estudos em Petrologia e Geoquı´mica (CPGq)-I.G.-UFRGS, Universidade Federal do Rio Grande do Sul, Av. Bento Gonc¸alves, 9500. Porto Alegre, Brazil c Hydrasa UMR 6532 CNRS—Universite´ de Poitiers-40 Avenue du Recteur Pineau, 86022 Poitiers Cedex, France

b

Received 1 May 2002; accepted 1 April 2003

Abstract Petrographic and mineralogical studies of a 45 m thick basaltic lava flow in southern Parana´ Basin (Rio Grande do Sul, Brazil) enables a detailed description of three structural levels: the lower vesicular zone (LVZ), inner massive zone (IMZ), and upper vesicular zone (UVZ). The three levels, inherited from cooling stages, are characterized by vertical zonation of the petrographic features and associated secondary mineral assemblages, such as clay minerals and zeolites. Zeolite crystallization is limited to the vesicle infilling and partial replacement of albitized plagioclases. The clay mineral sequence observed in vesicle infilling is a celadonite, saponite, chlorite/saponite mixed layer. The mesostasis of the three levels, which constitutes reduced sites of clay mineral crystallization from the peripheral levels (top and base) to the inner, massive, and vesicle-free part of the flow, presents a saponite to C/S mixed layer sequence. Petrographic and chemical observations support three steps for the alteration mechanisms. The earliest alteration stages are related to postmagmatic mechanisms. They are marked by earliest celadonite precipitation in the oxidative condition of the highly permeable UVZ, saponite with homogeneous compositions in reducing conditions, and C/S mixed layer conversion in the inner part of the flow, where temperature gradients have been preserved during the final stages of cooling. The albitization of plagioclase associated with zeolite crystallization and the compositional changes in clay mineralogy should be attributed to high water/basalt alterations during the low-grade burial metamorphic conditions in more permeable vesicular levels of the flow. q 2003 Elsevier Ltd. All rights reserved. Keywords: Basaltic flow; Clay minerals; Cooling process; Low-grade metamorphism; Parana´ Basin; Postmagmatic; Zeolite assemblage

Resumen O estudo petrogra´fico e mineralo´gico de um derrame basa´ltico de 45m de espessura, na porc¸a˜o sudeste da Bacia do Parana´ (Rio Grande do Sul, Brasil), segue uma descric¸a˜o detalhada de treˆs nı´veis estruturais: nı´vel vesicular inferior (LVZ), nı´vel macic¸o intermedia´rio (IMZ) e nı´vel vesicular superior (UVZ). Estes nı´veis estruturais, resultado dos processos de resfriamento do derrame, sa˜o caracterizados pela zonac¸a˜o vertical das feic¸o˜es petrogra´ficas e dos minerais secunda´rios associados, tais como argilominerais e zeolitas. As zeolitas ocorrem em vesı´culas e como fases de substituic¸a˜o dos plagiocla´sios albitizados. Os argilominerais ocorrem preenchendo vesı´culas seguindo a sequeˆncia, celadonita, saponita, interestratificado C/S. A meso´stase da rocha, que constitui sı´tios reduzidos de cristalizac¸a˜o de argilominerais, apresenta, das zonas perife´ricas (topo e base) para as zonas internas do derrame, a sequeˆncia saponita-interestratificado C/S. As observac¸o˜es petrogra´ficas e quı´micas sugerem treˆs esta´gios de alterac¸a˜o. Os primeiros esta´gios esta˜o relacionados com os mecanismos de alterac¸a˜o po´s-magma´tica. Estes esta´gios sa˜o marcados pela precipitac¸a˜o precoce de celadonita nos nı´veis mais permea´veis (UVZ), em condic¸o˜es oxidantes, formac¸a˜o de saponita com composic¸a˜o relativamente homogeˆnea, em condic¸o˜es redutoras, e pela conversa˜o para interestratificado C/S, nas partes internas do derrame, cujo gradiente de temperatura tem sido preservado por mais tempo, durante os esta´gios finais de resfriamento. Por outro lado, a albitizac¸a˜o dos plagiocla´sios, associada a` cristalizac¸a˜o de zeolitas, e as mudanc¸as na composic¸a˜o dos argilominerais, devem estar associados aos

* Corresponding author. Address: Instituto de Geocieˆncias, Centro de Estudos em Petrologia e Geoquı´mica (CPGq)-I.G., Universidade Federal do Rio Grande do Sul—UFRGS, Av. Bento Gonc¸alves 9500, Porto Alegre, Brazil. Fax: þ55-513-316-7302. E-mail address: [email protected] (F. Schenato). 0895-9811/$ - see front matter q 2003 Elsevier Ltd. All rights reserved. doi:10.1016/S0895-9811(03)00098-1

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mecanismos de alterac¸a˜o, nos nı´veis mais permea´veis, com altas razo˜es a´gua/rocha, em condic¸o˜es de metamorfismo de baixo grau. q 2003 Elsevier Ltd. All rights reserved. Palabras clave: Bacia do Parana´; Derrame basa´ltico; Processos de resfriamento; Alterac¸a˜o po´s-magma´tica; Metamorfismo de baixo-grau; Argilominerais e zeolitas

1. Introduction Clay minerals and zeolites in thick piles of basalt flows are commonly considered the result of hydrothermal alteration events (Juteau et al., 1979; Kristmannsdottir, 1979; Andrews, 1980; Alt et al., 1996) or low-grade metamorphic reactions during the burial stage (Levi et al., 1982; Schmidt and Robinson, 1997; Alt, 1999; Neuhoff et al., 1999). In particular, clay minerals in basalt alteration systems were believed to have formed from the interaction of residual fluids during immediate posteruptive cooling (deuteric alteration) and hydrothermal or metamorphic fluids with primary phases and mesostasis glass (Bo¨hlke et al., 1980; Degraff et al., 1989; McMillan et al., 1989; Gonc¸alvez et al., 1990; Marescotti et al., 2000). In these alteration mechanisms, the more reactive phases are olivine and glass. A possible first alteration stage of olivine is the precipitation of iddingsite along microcracks during posteruptive deuteric alteration (Shelley, 1993). Iddingisite develops coherently within olivine, and its pseudomorphs can be recognized by the distinctive shapes of the original crystals. The features of glass alteration in low-temperature seawater/basalt interaction mechanisms or subaerial geothermal systems have been described as palagonitization (Honnorez et al., 1979; Mevel, 1980; Zhou and Fyfe, 1989). Glass is observed in chilled margins of thick lava flows induced by quenching. However, in the inner, massive parts of thick lava flows, the low cooling rates induce holocrystalline textures (Scott and Hajash, 1976; McPhie et al., 1993). The dark cryptocrystalline or ‘glasslike’ aspect of the interstitial material in mesostasis sites usually poses difficulties for the accurate identification of the nature of the constituents (mineral or glass) (Anderson et al., 1984; Long and Wood, 1986; Bates and Jackson, 1987; Goff, 1996). This paper is a detailed petrographic and mineral chemical study of a thick lava flow from the Parana´ Basin of southern Brazil (Estaˆncia Velha region, Rio Grande do Sul). It attempts to identify the clay minerals and associated zeolites formed during successive alteration stages—from postmagmatic to late, low-grade metamorphic events—in the different levels of the flow. The term ‘postmagmatic’ as used herein refers to immediate posteruptive alteration during cooling of the flow. The petrographic study was used to understand the cooling and degassing processes recorded in the structural vertical pattern of the basaltic flow. The petrographic results are compared with characteristics of a basalt flood reported for volcanic sequences of

the Columbia River province (USA) (Long and Wood, 1986; McMillan et al., 1989; Mangan et al., 1993) and in active basalt flows (Peck, 1978; Cashman et al., 1994). The clay mineral and zeolite assemblages of the flow are compared with sequences described and interpreted as typical of hydrothermal or low-grade metamorphic conditions (Hulen and Nielson, 1986; Lonker et al., 1993; Dudoignon et al., 1997; Neuhoff et al., 1999).

2. Geological settings The Parana´ continental basaltic flood province covers an area of approximately 1.2 million km2 and represents a very important volcanic province related to the Parana´-Etendeka event, with 0.8 £ 106 km3 of lava flows (Fig. 1). It formed during the opening of the south Atlantic Ocean in the Early Cretaceous. The lava flows consist of predominantly tholeiitic basalts and andesitic basalts (90%), as well as minor amounts of andesites, lati-andesites, latites, rhyolites, and rhyodacites near the top of the sequence (Bellieni et al., 1986, 1988; Piccirillo et al., 1988). The total thickness of the volcanic pile is approximately 1600 m in the center of the basin. Two magmatic groups have been identified, according to their chemical characteristics: (1) a low-Ti group known as the Gramado, the Esmeralda, and the Ribeira magmatic series and (2) a high-Ti group named the Pitanga, the Paranapanema, and the Urubici (Peate et al., 1990, 1992; Hawkesworth et al., 1992; Peate and Hawkesworth, 1996). Major and trace element characteristics suggest different sources for the magma that generated the volcanic sequence. Detailed descriptions of these magmatic groups and their distribution in the Parana´ volcanic sequence are widespread (Peate et al., 1990, 1992; Hawkesworth et al., 1992; Peate and Hawkesworth, 1996). The studied basaltic flow belongs to the tholeiitic basalt sequence of the low-Ti group in the southern Parana´ Basin. It is 45 m thick and occurs in the Estaˆncia Velha region (Brazil); the first flow of the volcanic pile directly overlies the eolian sandstones of the Botucatu´ Formation. The total thickness of the volcanic pile above the flow is approximately 350 m. This flow shows an andesi-basalt composition with low Ti (TiO2 , 1.4%) and high FeO (mean value FeOT ¼ 10.24%) contents. Bulk rock chemical compositions show that it has typical tholeiitic affinities with normative quartz (Table 1). The trace element pattern normalized to the composition of the primitive mantle shows enrichment in

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Fig. 1. Map of the distribution of the volcanic rocks types in the Parana´ Basin (Peate et al., 1992). The studied flood is the first of the volcanic pile on the southern side of the Parana´ Basin, south of Brazil (Estaˆncia Velha region).

lithophile components and depletion in Nb, Ta, and Sr in comparison with trace element patterns of the Gramado and Esmeralda basalts (Peate and Hawkesworth, 1996). The Ti/Zr (, 60), Zr/Y (5.0 – 6.5), and Ti/Y (, 300) ratios coincide with the Gramado magmatic type, which

is predominant in southern Parana´ Basin (Figs. 2 and 3). The trace element pattern of the Gramado magmatic type, including the lava flow studied, indicates a significant crustal contribution (Peate et al., 1992; Peate and Hawkesworth, 1996).

47 200 27 158 19 306 98 137 271 4 0.9 3 ,1 21.4 44 22 4.4 1.2 0.9 2.6 0.36

Rb Sr Y Zr Nb Ba Ni Cr V Th U Hf Ta La Ce Nd Sm Eu Tb Yb Lu

b

n.a. 210 26 132 10 366 85 220 n.a. n.a. n.a. n.a. 20 n.a. 39 n.a. n.a. n.a. n.a. n.a. n.a.

53.99 14.8 10.97 5.5 0.16 1.26 9.09 2.15 0.97 0.19 1.6 100.68

BP-02 IMZ

61 201 27 161 14 311 101 129 273 5 1.2 3.9 1 21.2 45 21 4.4 1.3 0.9 2.5 0.37

53.73 14.38 11.62 4.96 0.18 1.27 9.03 2.29 1.66 0.19 1.28 100.59

BP-15 IMZ

42 201 27 157 13 307 97 140 276 4.9 0.9 3.6 ,1 21.1 46 23 4.5 1.2 1 2.6 0.35

53.77 14.47 11.55 4.94 0.16 1.27 9.18 2.33 1.61 0.19 0.98 100.45

BP-16 IMZ

48 199 26 153 13 302 101 125 265 4.6 0.9 3.4 ,1 20.4 44 18 4.2 1.2 0.9 2.4 0.34

53.23 14.49 9.96 4.93 0.16 1.24 9.14 2.28 1.54 0.16 0.97 98.1

BP-17 IMZ

44 200 27 164 15 303 99 131 269 4.7 0.7 4.5 ,1 20.9 45 18 4.3 1.3 0.8 2.5 0.35

53.66 14.61 11.52 4.95 0.16 1.24 9.24 2.32 1.57 0.19 1.03 100.49

BP-18 IMZ

67 191 27 158 14 312 100 139 276 5.2 0.9 4.2 ,1 21.4 47 19 4.4 1.3 0.6 2.7 0.34

53.4 14.46 11.71 5.04 0.17 1.26 8.96 2.41 1.68 0.19 0.82 100.1

BP-19 IMZ

59 191 26 155 13 328 103 132 270 4.6 0.9 3.9 ,1 21 42 18 4.3 1.2 0.8 2.5 0.36

53.59 14.76 11.14 4.93 0.16 1.25 8.76 2.58 1.79 0.19 0.53 99.68

BP-20 IMZ

Mean chemical composition of 120 samples of the Gramado magma type (Peate et al., 1992). Fe2O3 ¼ FeOT.

53.73 14.57 10.65 4.89 0.16 1.25 9.2 2.32 1.56 0.19 1.12 99.64

SiO2 Al2O3 Fe2O3b MgO MnO TiO2 CaO Na2O K2O P2O5 LOI Total

a

BP-01 IMZ

Samples

Table 1 Bulk rock chemical composition of the studied lava flow (Estaˆncia Velha quarry samples)

53 190 28 179 16 358 89 112 314 5.3 0.9 3.5 1 22.2 45 19 4.6 1.4 0.8 2.5 0.39

52.39 14.94 11.89 4.43 0.16 1.32 8.19 3.05 2.34 0.2 1.46 100.37

BP-21 UVZ

71 182 28 168 17 340 89 114 336 5.4 1 4.4 ,1 23.1 53 22 5 1.4 1 2.9 0.37

49.5 14.77 11.68 4.69 0.16 1.32 8.22 2.79 2.3 0.2 4.89 100.52

BP-22 UVZ

68 208 29 174 16 314 92 117 401 6.1 ,0.5 5 ,1 24.7 57 26 5.3 1.5 1.1 3.1 0.43

47.16 14.62 12.5 4.78 0.2 1.4 8.18 2.94 2.5 0.21 5.18 99.67

BP-24 UVZ

52 213 28 172 17 353 92 108 346 5.5 1.2 4 ,1 23.2 50 21 4.7 1.4 1.4 2.6 0.38

49.72 15.05 11.87 5.09 0.18 1.33 8.42 2.89 1.57 0.2 4.57 100.89

BP-25 UVZ

n.a. 205 27 144 10 376 92 180 n.a. n.a. n.a. n.a. 20 n.a. 48 n.a. n.a. n.a. n.a. n.a. n.a.

53.97 14.5 10.99 5.25 0.15 1.29 8.51 2.44 1.09 0.2 1.2 99.59

BP-05 UVZ

n.a. 181 28 156 10 356 88 170 n.a. n.a. n.a. n.a. 20 n.a. 52 n.a. n.a. n.a. n.a. n.a. n.a.

51.26 13.3 12.1 5.83 0.17 1.39 8.53 2.31 0.87 0.23 3.2 99.19

BP-06 UVZ

n.a. 175 29 155 10 409 94 180 n.a. n.a. n.a. n.a. 20 n.a. 46 n.a. n.a. n.a. n.a. n.a. n.a.

50.99 14.28 11.39 5.49 0.2 1.3 8.09 2.22 2.14 0.24 3.8 100.14

BP-07 UVZ

n.a. 204 27 147 10 394 80 170 n.a. n.a. n.a. n.a. 23 n.a. 46 n.a. n.a. n.a. n.a. n.a. n.a.

53.66 14.69 11.1 5.47 0.17 1.38 8.58 2.65 1.11 0.2 1.5 100.51

BP-12 UVZ

35 167 32 199 18 265 82 140 267 5.8 1.1 4.6 ,1 28.6 60 24 5.5 1.6 0.9 2.9 0.43

45.88 14.75 16.12 3.82 0.15 1.53 7.29 2.79 1.64 0.24 6.43 100.64

BP-29 LVZ

45 238 33 166 14 388 42 – – – – – – – – – – – – – –

53.68 14.26 12.64 4.9 0.19 1.43 8.64 2.68 1.33 0.21 – 99.96

Gramadoa

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Fig. 2. Whole-rock spider diagram of the studied basaltic flood in comparison with selected data of Gramado and Esmeralda magma types from Peate and Hawkesworth (1996). Analyses normalized to values of the primitive mantle. Key: B studied rock, A Gramado magma type, W Esmeralda magma type.

3. Structure of the lava flow On the basis of the distribution pattern of the vesicles, three successive structural zones can be distinguished from the base to the top of the studied basaltic flow: a lower vesicular zone (LVZ) (, 0.1 m), an inner massive zone (IMZ) (30 m), and an upper vesicular zone (UVZ) (15 m). The LVZ represents the basal part of the lava flow that is in contact with the sedimentary rocks of the Parana´ Basin. It consists of a thin vesicular zone (5 – 20 cm thick), with less than 5 vol% of vesicles with diameters between 0.2 and 2.0 cm. This zone grades into the massive zone. The IMZ forms the central and main part of the flow. It consists of a dense unit in which vesicles are absent, and it is cross-cut by

427

multiple joint sets that define narrow columns. The UVZ occurs above the massive zone. It is characterized by a regular pattern of vesicle distribution. The quantity of vesicles induces a significant porosity and an important concentration of secondary minerals. In the UVZ, the size and abundance of vesicles change according to the depth (Fig. 4). The increasing abundance of vesicles, counted as equivalent porosity (vesicularity %), along the 15 m of the UVZ is inversely proportional to the average size of the vesicles. In other words, the vesicles increase in number and decrease in size from the base to the top of the UVZ. At the top of the flow (40 –44 m), vesicularity is approximately 20%. This level is made of vesicles with diameters ranging from 0.11 to 1.08 cm. At the base of the UVZ (30 – 35 m), the diameters of vesicles increase and range from 0.26 to 3.3 cm. The associated porosity is , 5%. The volume and diameter distribution patterns of the vesicles in the UVZ result from vesiculation processes and degassing related to the initial stages of cooling and crystallization of the flow. This pattern indicates mechanisms of bubble coalescence in the lowermost portion of the UVZ. It is similar to the patterns observed by Peck (1978) in the vesicular top zones of Alae Lava Lake (Hawaii) and by Aubele et al. (1988) in basaltic lava flows in New Mexico (USA). The vertical structure of the flow, with an inner massive and peripheral vesicular level, is a consequence of the progressive cooling of the unit. The vertical distribution of vesicles shows the progression of the upper cooling front into the inner part of the flow in the 1100 –1150 8C range (Aubele et al., 1988).

4. Analytical methods Chemical analyses of primary and secondary phases were performed using an CAMECA SX50 electron microprobe equipped with wavelength dispersive spectrometers

Fig. 3. Variation diagrams representing the low-Ti magmatic groups from Peate et al. (1992). (a) Ti/Zr versus Zr/Y; (b) Ti/Y versus Sr. Key: B studied basaltic flood, A mean chemical composition of Gramado type.

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Chemical analyses of bulk samples were performed by Activation Laboratories, Ontario, Canada. Major and trace elements (Rb, Sr, Y, Zr, Nb, Ba, Ni, Cr, V, Th, U, Hf, Ta) and REE were determined by inductively coupled plasma. Total Fe is indicated as Fe2O3. The image analysis was used to quantify (1) modal proportions of the primary minerals and the mesostasis (interstitial residuum) of the rock in thin sections and (2) volume, distribution, and size of the vesicles in the lava flow. The images were digitized using a Macintosh Squares 650, equipped with a Neotech card and OPTILABq and OPTISCANq software.

5. Petrography and mineral chemistry

Fig. 4. Vesicle distribution and diameters of vesicles voids in 30–45 m interval in the UVZ. (a) Vesicularity (porosity %) versus depth; (b) average vesicle size (cm) versus depth. Volume and size of the vesicles were quantified by image analysis on photographs and polished thin sections.

(WDS system). Accelerating voltage of 15 kV, sample current of 10 nA, and a beam with a diameter of 3– 5 mm were used. Counting time was 10 s per element, and the data were calibrated with synthetic oxides and natural silicate standards. Polished thin sections of the selected samples were examined with a JEOL JSM 6400 SEM operated at 15 kV, coupled with a AN1000 energy dispersive spectrometer (EDS) for the identification and quantification of, especially, the mineral phases present in small sites as the interstitial residuum of the rock (mesostasis), vesicles, and microvesicles, as well as primary phase alteration. Secondary minerals were examined by X-ray diffraction (XRD) on random powder and oriented preparation. The data were obtained through a Philips PW 1730 diffractometer equipped with Fe-filtered Co Ka radiation, 40 kV, and 40 nA at the University of Poitiers, France. Bulk samples were first gently ground and separated into , 2 mm fractions by centrifugation and sedimentation. Clay minerals were identified from the , 2 mm clay fraction. Air drying (AD), ethylene glycol solvates (EG), and heat treatments (300 8C) were carried out for each sample. Dioctahedral and trioctahedral phyllosilicates were distinguished on the basis of the (06,33) reflections recorded in the 59 –648 2u range with a step size of 0.0058 2u and a counting time of 10 s per step. Complex experimental diffraction patterns were decomposed by deconvolutions based on least-square fitting procedures (Gaussian and Lorentzian line-shapes) using DECOMPXRq program software (Lanson and Besson, 1992). Calculated XRD patterns were achieved using NEWMODq software (Reynolds, 1985). Zeolite minerals were hand picked from bulk samples and analyzed from random powder preparations.

The studied basaltic flow shows porphyritic to glomeroporphyritic texture containing euhedral to subhedral olivine and plagioclase phenocrysts. The abundant groundmass is composed of plagioclase (31 – 40 vol%), pyroxene (28 – 30 vol%), Fe – Ti oxide minerals (2 – 3 vol%), and interstitial residuum (mesostasis). The UVZ and LVZ of the flow show fine-grained intersertal texture, with 10 vol% of plagioclase phenocrysts and feathery groundmass crystals, indicative of fast cooling. These zones have up to 30 vol% of unglassy mesostasis. In contrast, the IMZ shows coarse-grained intergranular texture, with up to 10 vol% of phenocrysts and 15 vol% of unglassy mesostasis. Textural features, such as the skeletal mineral morphology and major mesostasis abundance in the UVZ and LVZ, accord with the faster cooling rate at the top and bottom of the flow (Lofgren, 1980; Long and Wood, 1986). 5.1. Primary minerals In the IMZ, plagioclase compositions range from An42 to An70, and the most calcic compositions occur in phenocrysts. In the UVZ and LVZ, groundmass and phenocryst plagioclase have a composition of An35 – 70, but the majority of the grains are albitized or replaced by an association of secondary albite þ zeolite. The clinopyroxenes in the three zones are augite and pigeonite. Augite compositions range from En35 – 50 in the IMZ to En36 – 40 in both the UVZ and the LVZ. The calcium content in augite of all zones is relatively low, but groundmass augites show lower Ca and higher Fs contents in comparison with augite phenocrysts. Pigeonite compositions range from En30 – 65 in the IMZ to En62 – 69 in the UVZ. A trend of decreasing Mg can be observed from phenocryst to groundmass pigeonites. Pigeonite does not occur in the LVZ. Ti – Fe oxide minerals are mainly magnetite and titanomagnetite in the LVZ and UVZ and titanomagnetite associated with ilmenite in the IMZ. Olivine pseudomorphs occur only at the peripheral vesicular levels. The olivine is completely replaced by secondary clay minerals.

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5.2. Phases of mesostasis Mesostasis in the three structural levels of the flow has cryptocrystalline aspects. It is formed in residual groundmass microsites usually smaller than 0.5 mm. In the LVZ, the basaltic texture is characterized by grain size evolution. From the lava flow contact to the inner parts,

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the mesostasis phases evolve from very fine grains (1 – 5 mm) to larger grains (50 mm) in a few centimeters, as is clearly displayed in backscattered images (Fig. 5a – c). In coarser grained textures, the individual microcrystal constituents of the mesostasis are observable, particularly K-feldspars closely associated with albite. The residual areas disseminated between K-feldspars and

Fig. 5. SEM images in backscattered electron mode of the mesostasis in the LVZ, IMZ, and UVZ. (a –d) In the LVZ, the texture is characterized by grain size evolution through the few centimeters away from the lava flow contact. Mesostasis sites show microlites of K-feldspars (F) and plagioclase phenocrysts (Pl) partially albitized. White points are titanomagnetite (Ti-mt) microcrystals. The saponite (Sa) has precipitated between feldspar grains. (e) In the IMZ, a random chlorite/saponite mixed layer (chr% , 20%) has crystallized between quartz (Qz) and feldspar (F) grains. Ap ¼ acicular apatite. (f) In the UVZ, the clay mineral (Sa) has precipitated in residual areas between feldspar laths. Mv ¼ microvesicle.

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albite microlites are filled by a clayey, cryptocrystalline matrix and disseminated Ti – Fe oxides (Fig. 5d). These grain size variations, observed in only one petrographic thin section, evince the evolution of the cooling rates at the lower contact of the flow. In the IMZ, the micro to cryptocrystalline mesostasis contains mainly interstitial quartz and feldspar (50 – 100 mm) associated with acicular apatite grains (10 – 30 mm), titanomagnetite dendrites, and aggregates of clay minerals (20 –50 mm) (Fig. 5e). These phases are clearly identified in backscattered images and SEM/EDS analyses. In the UVZ, the interstitial mesostasis is composed of an intimate mixture of cryptocrystalline clay minerals, Timagnetite microcrystals, and mainly acicular K-feldspar microcrystals (10 – 20 mm) (Fig. 5f). The mesostasis aspect in the three zones suggests an overgrowth of clay minerals in residual areas between feldspar laths, quartz, or apatites. There is no evidence of glass relicts in mesostasis sites, in either the inner massive or peripheral vesicular zones. 5.3. Secondary minerals Secondary minerals are mainly clay and zeolites observed in the three zones of the flow (Fig. 6). In the LVZ, the clay minerals are located in the rock groundmass mesostasis and olivine replacement; in the IMZ, only in the mesostasis sites; and in the UVZ, the clay mineral assemblage is distributed in groundmass mesostasis, olivine

replacement, and pore infilling (vesicles and microvesicles). In the UVZ, the zeolites have precipitated with albite to replace the plagioclase, as in the inner part of the vesicles. In LVZ samples, only one smectite phase was identified ˚ in AD and by the 001 reflections, indexed at 15.70 A ˚ ˚ 16.97 A (rounded to 17.00 A for a better mathematical fit) after EG treatment (Fig. 7a). In IMZ samples, only one random chlorite/smectite mixed layer was characterized by 001 reflections at 15.13 ˚ in AD and after EG treatment, respectively and 16.69 A (Fig. 7b). The percentage of the chlorite layer, using Reynolds’s (1985a,b) software, is approximately 20%. In UVZ samples, three types of clay minerals were identified after mathematical decompositions: smectite, random chlorite-smectite (C/S) mixed layer, and celadonite. The XRD pattern shows 001 reflections at 15.10 and ˚ in AD, which correspond to expansive phases and 10.07 A celadonite, respectively. After EG treatment, the solvated ˚ reflection decomposed in two bands samples show 16.50 A ˚ at 16.3 and 17.2 A, which correspond to the C/S mixed layer ˚ 001 and smectite, respectively (Fig. 7c). The 10.07 A reflection of celadonite is maintained after EG saturation. The zeolite was the last phase to crystallize in microsites of the UVZ after the precipitation of clay minerals. Typical vesicle infillings consist of heulandite, mesolite/scolecite, stilbite, thomsonite, laumontite, and mordenite, in decreasing order of abundance. Frequently, only one or two of these minerals is present in the same void. Laumontite and mordenite are abundant along the contact with the overlying

Fig. 6. Sketch of the clay mineral zonation through the UVZ, IMZ, and IVZ. Ce ¼ celadonite, Sa ¼ saponite, C/S ¼ chlorite/saponite mixed layer. The percentages in rectangles indicate the chlorite layer % in the mixed layers.

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Fig. 7. XRD patterns of clay minerals (,2 mm fraction) crystallized in the (a) LVZ, (b) IMZ, and (c) UVZ. The XRD patterns correspond to air-dried ˚ ; (b) random chlorite/smectite mixed preparations (AD) and ethylene glycol treatment (EG). (a) Only smectite phase indexed at 15.70 (AD) and 16.97 (EG) A ˚ ; and (c) three types of clay minerals: smectite, random S/C mixed layer, and celadonite, for which layer (C/S , 20%)(indexed at 15.13 (AD) and 16.69 (EG) A ˚ (AD) correspond to expansive phases and celadonite, respectively. The solvated sample (EG) shows 001 reflection at the 001 reflections at 15.10 and 10.07 A ˚ , which decomposes into two bands at 16.5 and 17.2 A ˚ , corresponding to a C/S mixed layer and smectite, respectively. 16.50 A

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Table 2 Representative compositions of zeolites from vesicles in the UVZ Anal. no.

Heulantite 38

39

Stilbite 68

146

147

148

151

162

Mesolite 160

57

50

87

88

89

SiO2 Al2O3 FeO MgO TiO2 MnO CaO Na2O K2 O Total

60.76 60.48 60.82 60.57 59.43 59.81 60.64 62.68 61.90 62.12 61.61 46.02 46.41 46.95 17.55 17.60 17.95 17.67 17.75 17.49 18.04 18.54 18.24 16.50 17.79 25.96 26.06 26.30 0.06 0.07 0.59 0.00 0.07 0.02 0.00 0.07 0.17 0.00 0.02 0.02 0.00 0.13 0.00 0.00 0.05 0.01 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.07 0.00 0.01 0.01 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.02 0.07 0.00 0.00 0.06 0.09 0.04 0.02 0.00 0.04 0.06 0.00 0.04 0.00 8.53 8.19 8.52 8.12 8.00 8.51 8.09 8.52 8.08 7.48 8.32 11.43 11.74 11.64 0.92 1.08 0.82 1.08 1.09 1.07 1.02 1.17 1.14 0.74 1.55 3.12 2.97 3.15 0.08 0.07 0.10 0.07 0.04 0.00 0.23 0.11 0.04 0.61 0.17 0.02 0.05 0.13 87.92 87.56 88.92 87.52 86.43 87.00 88.07 91.12 89.58 87.49 89.52 86.59 87.27 88.29

Formula unit composition Si Al Fe2þ Mg Ti Mn Ca Na K

26.89 26.87 26.70 26.89 26.74 26.78 26.78 26.78 26.84 27.51 26.84 9.15 9.22 9.29 9.25 9.41 9.23 9.39 9.34 9.32 8.61 9.13 0.02 0.03 0.22 0.00 0.03 0.01 0.00 0.03 0.07 0.00 0.01 0.00 0.00 0.03 0.01 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.03 0.00 0.00 0.02 0.03 0.01 0.01 0.00 0.02 0.02 4.04 3.90 4.01 3.86 3.86 4.08 3.83 3.90 3.76 3.55 3.88 0.79 0.93 0.70 0.93 0.95 0.93 0.87 0.97 0.97 0.63 1.31 0.05 0.04 0.06 0.04 0.02 0.00 0.13 0.06 0.02 0.34 0.10

Si/Al (Na þ K)/(Na þ K þ Ca þ Mg)

2.94 0.17

2.92 0.20

2.87 0.16

2.91 0.20

2.84 0.20

2.90 0.19

2.85 0.21

2.87 0.21

2.88 0.21

3.19 0.22

2.94 0.27

132

134

46.89 24.86 0.08 0.00 0.06 0.22 11.44 2.57 0.00 86.11

47.72 25.65 0.00 0.00 0.09 0.03 11.97 2.60 0.01 88.07

9.01 5.99 0.00 0.00 0.00 0.00 2.40 1.18 0.01

9.02 5.97 0.00 0.00 0.00 0.01 2.44 1.12 0.01

9.02 5.96 0.02 0.00 0.00 0.00 2.40 1.17 0.03

9.20 5.75 0.01 0.00 0.01 0.04 2.41 0.98 0.00

9.16 5.80 0.00 0.00 0.01 0.00 2.46 0.97 0.00

1.50 0.33

1.51 0.32

1.51 0.33

1.60 0.29

1.58 0.28

FeOT ¼ FeO.

lava flow. Heulandite and stilbite are ubiquitous in the UVZ. Thomsonite has restricted occurrence along the borders of the vesicles, in that it is locally precipitated between the C/S mixed layer lamellae and the preceding mesolite. Representative compositions of more abundant phases as heulandite, stilbite, and mesolite are given in Table 2. Exchangeable cation sites in heulandite and stilbite are dominantly occupied by Ca2þ, Naþ, and Kþ, which constitute only 20 –30% of the exchangeable ions. 5.4. Chemical composition of secondary minerals In the LVZ, smectite has homogeneous compositions in olivine replacement and mesostasis. This phase plots in the trioctahedral smectite domain (saponite) in the Mþ –3R2 – 2R3 diagram (Meunier et al., 1991; Fig. 8a). Calculated on the basis of O10(OH)2, the structural formula has an octahedral sum closer to 2.7– 3.0 (Table 3). Caþ2 is the dominant interlayer cation, though some analyses exhibit relatively high Kþ and Naþ contents in exchangeable positions. In the vesicles, the chemical compositions of the brown saponite are characterized by high Mg content (17 – 20%). The sum of the interlamellar cations (Ca þ Na þ K) ranges from 0.33 to 0.50. In the IMZ, the C/S mixed layer compositions, present in mesostasis sites, plot between the trioctahedral smectite

(saponite) and the chlorite compositional fields, closer to the saponite domain according to the low percentage of the chloritic layer (Fig. 8a). The structural formula calculation of the C/S mixed layer, on the basis of O10(OH)8, gives an octahedral sum between 5.14 and 5.8 (Table 3). In the UVZ, smectite compositions in the mesostasis sites plot in the trioctahedral smectite domain (saponite) (Meunier et al., 1991; Fig. 8a, Table 3). Olivine phenocrysts are completely replaced by celadonite. The celadonite composition is closer to the K(Mg,Fe2þ)(Al,Fe 3þ )Si 4O 10(OH) 2 end-member composition, as defined by Wise and Eugster (1964), with a silicon content between 3.95 and 4.00 and a potassium content (0.73 –0.83) lower than ideal (Table 4, Fig. 8b). In the vesicles, the clay mineral sequence from the walls to the center of the void is celadonite, saponite, and C/S mixed layer (Fig. 9). The representative compositions of this celadonite differ from celadonite of olivine replacement in their slightly lower silicon (3.82 –3.93) contents (Table 5, Fig. 8c). The deviation from an ideal composition domain in the Mþ –4Si – 3R2 diagram may be due to all Fe being counted as Fe3þ in the structural formula. In the same vesicles, smectite is brown to brownish-green and radiates fine lamella lining the wall voids. Its compositions plot in the trioctahedral smectite domain of saponite (Fig. 8c). The calculation of structural formula,

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433

Fig. 8. Clay minerals compositions in the Mþ –4Si–3R2 diagram (Meunier et al., 1991) (Mþ ¼ 2Ca þ Na þ K; 4Si ¼ Si/4; 3R2þ ¼ (Fe2þ þ Mg þ Mn)/3). sap ¼ saponite; chl ¼ chlorite; cel ¼ celadonite. (a) Clay minerals of the mesostasis sites in the UVZ (þ), IMZ (W), and LVZ (X) zones plot in saponite and C/S mixed layer fields; (b) celadonite composition replaces olivine (FeOT ¼ Fe2O3); (c) clay minerals of the vesicles infilling in the UVZ plot in celadonite, saponite, and C/S mixed layer domains.

on the basis of O10(OH)2, gives an octahedral sum closer to 2.8 – 3.0. The magnesium content in saponite is relatively high (15.5%; Table 5). The C/S mixed layer occurs as large, radiating bundles crystallized in the center of the voids (Fig. 9), locally associated with

zeolite crystals. Their compositions fall between the compositional fields of saponite and chlorite (Fig. 8c). The structural formula, calculated on the O10(OH)8 base, has an octahedral sum between 5.5 and 5.9 (Table 5). The chlorite interlayer percentage, estimated after

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Table 3 Representative clay mineral analyses from mesostasis sites of the UVZ, IMZ, and LVZ Anal no.

LVZ 33

LVZ 47

LVZ 51

LVZ 66

LVZ 6

LVZ 7

SiO2 Al2O3 FeO MgO TiO2 MnO CaO Na2O K2 O Total

35.63 44.20 40.04 41.21 42.27 45.93 33.90 6.84 6.59 6.18 6.39 7.35 6.73 10.00 8.78 5.44 7.63 8.64 8.08 9.43 19.61 15.93 17.04 16.60 13.04 15.34 14.75 12.22 0.14 0.15 0.28 0.36 0.11 0.17 0.07 0.29 0.19 0.23 0.08 0.13 0.10 0.15 1.94 2.30 2.15 2.06 0.81 2.41 1.95 0.67 0.34 0.40 0.73 1.47 0.29 0.28 0.66 0.33 0.09 0.30 1.19 0.32 0.51 70.87 76.59 73.59 72.80 76.75 80.10 78.69

Si AlIV P T

Calculated based on 11 3.32 3.64 3.51 0.68 0.36 0.49 4.00 4.00 4.00

oxygen 3.65 3.56 0.35 0.44 4.00 4.00

3.69 0.31 4.00

IMZ 90

IMZ 91

IMZ 97

IMZ 107

IMZ 109

UVZ 119

UVZ 121

UVZ 127

32.38 9.96 19.77 13.04 0.00 0.04 1.57 0.33 0.47 77.55

35.11 10.56 19.23 12.91 0.06 0.02 1.72 0.18 0.34 80.12

36.79 9.85 20.91 12.16 0.04 0.09 1.50 0.24 0.49 82.08

29.95 38.90 42.27 37.86 25.54 27.70 29.97 11.28 8.22 7.64 7.92 7.78 8.26 8.15 18.00 8.50 8.39 7.72 11.47 14.72 15.91 12.22 14.70 16.26 16.33 7.03 7.29 7.70 0.04 0.02 0.07 0.00 0.00 0.00 0.06 0.18 0.04 0.16 0.37 0.00 0.13 0.21 1.28 3.10 2.64 2.71 2.19 2.59 2.58 0.43 0.11 0.12 0.07 0.02 0.01 0.13 0.71 0.20 0.92 0.31 0.09 0.08 0.04 74.08 73.78 78.46 73.29 54.11 60.77 64.75

Calculated based on 14 oxygen Calculated based on 11 3.86 3.76 3.88 4.00 3.63 3.42 3.49 3.36 0.14 0.24 0.12 0.00 0.37 0.58 0.51 0.64 4.00 4.00 4.00 4.00 4.00 4.00 4.00 4.00

UVZ 2

UVZ 3

UVZ 19

oxygen 3.22 3.17 0.78 0.83 4.00 4.00

3.22 0.78 4.00

AlVI Fe2þ Mg Ti Mn P oct

0.07 0.68 2.21 0.01 0.02 3.00

0.29 0.38 2.09 0.01 0.01 2.78

0.15 0.56 2.17 0.02 0.02 2.91

0.32 0.64 1.72 0.02 0.01 2.71

0.29 0.57 1.93 0.01 0.01 2.81

0.32 0.63 1.76 0.01 0.01 2.74

1.20 1.87 2.07 0.01 0.01 5.16

1.12 1.92 2.26 0.00 0.00 5.30

1.26 1.78 2.13 0.00 0.00 5.18

1.26 1.90 1.97 0.00 0.01 5.14

1.24 1.82 2.21 0.00 0.02 5.29

0.28 0.63 1.93 0.00 0.00 2.84

0.24 0.58 2.00 0.00 0.01 2.84

0.19 0.57 2.16 0.00 0.03 2.95

0.37 1.21 1.32 0.00 0.00 2.90

0.28 1.41 1.24 0.00 0.01 2.95

0.25 1.43 1.23 0.00 0.02 2.94

Mg Ca Na K

0.00 0.19 0.12 0.08

0.00 0.20 0.05 0.04

0.00 0.20 0.07 0.01

0.00 0.20 0.13 0.03

0.00 0.07 0.24 0.13

0.00 0.21 0.05 0.03

0.00 0.24 0.06 0.07

0.00 0.19 0.07 0.07

0.00 0.20 0.04 0.05

0.00 0.17 0.05 0.07

0.00 0.17 0.10 0.11

0.00 0.29 0.02 0.02

0.00 0.23 0.02 0.10

0.00 0.26 0.01 0.04

0.00 0.30 0.00 0.01

0.00 0.32 0.00 0.01

0.00 0.30 0.03 0.01

Mþ 4Si 3R2þ

0.58 0.83 0.97

0.49 0.91 0.83

0.48 0.88 0.91

0.55 0.91 0.79

0.51 0.89 0.84

0.49 0.92 0.8

0.61 0.96 1.32

0.53 0.94 1.39

0.49 0.97 1.30

0.47 1.00 1.29

0.54 0.91 1.35

0.62 0.86 0.85

0.58 0.87 0.86

0.57 0.84 0.92

0.61 0.81 0.84

0.65 0.79 0.89

0.64 0.81 0.89

FeOt/(MgO þ FeOt)

0.36

0.24

0.31

0.40

0.35

0.39

0.62

0.60

0.60

0.63

0.60

0.37

0.34

0.32

0.62

0.67

0.67

FeOT ¼ FeO.

Reynolds’s (1985a,b) simulation, increases with depth in the UVZ from , 10% at the top to , 40% near the bottom (Fig. 6).

6. Discussion 6.1. Bulk rock chemistry variation The behavior of major elements, such as SiO2, Al2O3, MgO, and K2O measured in bulk rock, is relatively homogeneous throughout the massive, 30 m thick basalt. Some variations of these elements are verified only in the UVZ; a decrease of SiO2 and CaO and a significant enrichment in K2O, Na2O, and LOI is observed from 30 m to the top of the flow. TiO2 and P2O5 remain almost unchangeable throughout the 45 m of the flow (Table 1, Fig. 10). The samples present relatively homogeneous compositions of trace elements such as Zr and Sr in all three structural zones. The Ba content increases weakly in the vesicular zone, probably due to an increase of K2O. Remarkable variations in the chemical composition of thick lava flows (. 30 m thickness) are usually attributed to

in situ differentiation or alteration processes (Levi et al., 1982; Bailey, 1989; Schmidt and Robinson, 1997; Robinson and Bevins, 1999). The observed variations, mainly in SiO2, Na2O, CaO, and K2O contents, and the significant LOI increase from 30 to 45 m are probably associated with the abundance of hydrated minerals (i.e. clay minerals and zeolites) and thus with an alteration process. The concentration of Na2O with the increase of LOI in the vesicular zone probably reflects localized secondary albite (plus zeolite) replacing the primary calcic plagioclase. 6.2. Observed clay minerals/zeolite sequence The texture of the basaltic rock and the nature of the mesostasis evolve through the thickness of the lava flow, following the three zones formed by variations in the cooling rates from the peripheries to the inner part of the flow. The composition of secondary clay minerals also evolves according to the structural zonation of the flow. The upward-progressing cooling front is evident in the narrow LVZ, developed at the bottom of the lava flow. The basaltic and mesostasis textures are characterized by evolution of the very fine grains into larger grains in a few

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Table 4 Representative celadonite analyses of olivine replacement in the UVZ

Anal. no.

UVZ 55

UVZ 124

UVZ 36

UVZ 37

UVZ 127

UVZ 128

UVZ 137

UVZ 138

UVZ 144

SiO2 Al2O3 FeO MgO TiO2 MnO CaO Na2O K2O Total

52.52 6.48 20.54 4.12 0.12 0.00 0.42 0.09 8.02 92.30

52.90 5.26 24.73 3.51 0.10 0.00 0.24 0.05 7.78 94.57

52.52 4.00 24.49 3.90 0.03 0.05 0.47 0.13 8.56 94.15

52.40 5.36 22.91 3.54 0.16 0.00 0.61 0.17 8.01 93.17

49.08 6.72 18.14 3.53 0.07 0.20 0.54 0.10 7.20 85.57

48.72 6.62 21.14 3.43 0.30 0.00 0.34 0.03 7.36 87.93

52.63 6.34 19.65 4.04 0.14 0.02 0.70 0.09 7.50 91.10

51.18 6.13 21.03 4.04 0.08 0.03 0.61 0.05 7.68 90.82

52.48 5.96 20.70 3.81 0.11 0.00 0.57 0.07 8.09 91.78

Calculated based on 11 oxygen Si 3.97 Al 0.03 P T 4.00

3.97 0.03 4.00

4.00 0.00 4.00

3.98 0.02 4.00

3.97 0.03 4.00

3.90 0.10 4.00

4.00 0.00 4.00

3.95 0.05 4.00

4.00 0.00 4.00

Al Fe2þ Mg Ti Mn P oct

0.54 1.30 0.46 0.01 0.00 2.31

0.44 1.55 0.39 0.01 0.00 2.39

0.35 1.56 0.44 0.00 0.00 2.36

0.46 1.45 0.40 0.01 0.00 2.32

0.61 1.23 0.43 0.00 0.01 2.28

0.52 1.41 0.41 0.02 0.00 2.36

0.57 1.25 0.46 0.01 0.00 2.28

0.51 1.36 0.46 0.00 0.00 2.33

0.53 1.32 0.43 0.01 0.00 2.29

Mg Ca Na K

0.00 0.03 0.01 0.77

0.00 0.02 0.01 0.75

0.00 0.04 0.02 0.83

0.00 0.05 0.03 0.78

0.00 0.05 0.02 0.74

0.00 0.03 0.00 0.75

0.00 0.06 0.01 0.73

0.00 0.05 0.01 0.76

0.00 0.05 0.01 0.79

Mþ 4Si 3R2þ

0.85 0.99 0.26

0.79 0.99 0.26

0.93 1.00 0.27

0.90 0.99 0.25

0.85 0.99 0.25

0.81 0.97 0.25

0.85 1.00 0.25

0.87 0.99 0.26

0.89 1.00 0.25

FeOT ¼ Fe2O3.

centimeters from the bottom of the flow. The progressing cooling fronts trap a quantity of vesicles in the lower and upper peripheral zones of the flow, thereby inducing a significant increase of porosity and the crystallization of important amounts of secondary minerals. Although great amounts of secondary minerals occur mainly in vesicular

zones as vesicle infilling and replacements of primary minerals, the areas of mesostasis are restricted microsites of clay mineral crystallization widespread in the three zones. The nature of clay minerals precipitated in the mesostasis changes from the LVZ and UVZ to the IMZ: saponite in the LVZ, random chlorite/saponite mixed layer in the IMZ, and

Fig. 9. Microphotographs of clay-bearing microsites of the UVZ. (a) SEM backscattered images of the zoned sequence in microvesicles. The mineral clay sequence from wall-rock to the center of the void is celadonite (Ce), saponite (Sa), and chlorite/smectite mixed layer (C/S); (b) detailed SEM image of the large radiating bundles of C/S crystallized in the center of the voids (5– 30 mm).

0.91 0.96 0.24

Mþ 4Si 3R2þ

FeOT ¼ FeO.

0.00 0.05 0.03 0.78

Mg Ca Na K

0.94 0.97 0.26

0.00 0.06 0.04 0.78

0.93 0.98 0.25

0.00 0.04 0.03 0.82

0.65 1.04 0.49 0.03 0.00 2.21

0.87 0.98 0.24

0.00 0.04 0.03 0.76 0.82 0.98 0.25

0.00 0.04 0.01 0.73

0.81 0.82 0.53 0.03 0.00 2.19

0.78 0.93 0.48 0.02 0.01 2.21

Al Fe2þ Mg Ti Mn P oct

0.79 0.89 0.50 0.02 0.00 2.20

3.92 0.08 4.00

Calculated based on 11 oxygen 3.85 3.86 3.93 3.91 0.15 0.14 0.07 0.09 4.00 4.00 4.00 4.00

Si Al P T

0.61 1.11 0.50 0.03 0.00 2.26

49.05 9.38 12.27 4.46 0.47 0.05 0.51 0.04 7.11 83.34

49.04 9.34 13.34 4.23 0.32 0.00 0.42 0.17 7.47 84.33

49.37 7.63 15.67 4.11 0.48 0.00 0.51 0.16 8.11 86.04

48.45 8.03 16.74 4.23 0.56 0.00 0.66 0.27 7.69 86.62

UVZ 57

49.85 10.22 14.34 4.19 0.27 0.20 0.64 0.21 7.97 87.87

UVZ 56

SiO2 Al2O3 FeO MgO TiO2 MnO CaO Na2O K2O Total

UVZ 55

UVZ 54

UVZ 82

Anal. no.

0.84 0.98 0.26

0.00 0.04 0.01 0.75

0.74 0.93 0.54 0.02 0.00 2.23

3.92 0.08 4.00

49.98 8.92 14.17 4.62 0.34 0.00 0.44 0.09 7.52 86.07

UVZ 58

0.56 0.84 0.91

0.00 0.25 0.01 0.04

0.21 0.66 2.05 0.00 0.01 2.93

3.36 0.64 4.00

39.57 8.48 9.36 16.24 0.05 0.08 2.77 0.08 0.39 77.01

UVZ 34

0.53 0.85 0.88

0.00 0.26 0.02 0.01

0.27 0.69 1.93 0.00 0.01 2.90

3.39 0.61 4.00

37.83 8.33 9.18 14.46 0.05 0.14 2.66 0.09 0.06 72.80

UVZ 36

Table 5 Representative clay mineral analyses of vesicle and microvesicle infilling in the UVZ

0.54 0.86 0.90

0.00 0.25 0.02 0.02

0.21 0.62 2.08 0.00 0.00 2.91

3.43 0.57 4.00

37.89 7.34 8.22 15.42 0.00 0.01 2.55 0.13 0.18 71.75

UVZ 40

0.52 0.84 0.91

0.00 0.25 0.01 0.02

0.22 0.67 2.03 0.00 0.02 2.94

3.36 0.64 4.00

37.28 8.08 8.93 15.11 0.07 0.21 2.59 0.05 0.14 72.45

UVZ 41

0.55 0.82 0.93

0.00 0.26 0.01 0.01

0.20 0.77 2.02 0.00 0.00 2.99

3.28 0.72 4.00

35.91 8.57 10.04 14.84 0.05 0.04 2.68 0.08 0.09 72.29

UVZ 42

0.59 0.83 0.91

0.00 0.28 0.01 0.02

0.21 0.67 2.05 0.00 0.00 2.93

3.33 0.67 4.00

36.41 8.11 8.76 15.04 0.00 0.00 2.88 0.05 0.15 71.39

UVZ 51

0.50 0.85 0.90

0.00 0.23 0.01 0.03

0.22 0.60 2.10 0.00 0.01 2.93

3.41 0.59 4.00

39.22 7.85 8.24 16.19 0.04 0.10 2.49 0.04 0.26 74.43

UVZ 52 33.12 14.71 11.68 22.23 0.02 0.50 0.94 0.04 0.04 83.26

UVZ 91 30.94 15.06 22.34 15.01 0.00 0.41 0.93 0.07 0.03 84.78

UVZ 100 29.87 13.92 21.90 15.85 0.00 0.49 0.83 0.05 0.04 82.95

UVZ 101

0.19 0.85 1.49

0.00 0.09 0.00 0.01

1.17 0.78 3.63 0.00 0.05 5.62

0.21 0.85 1.48

0.00 0.10 0.01 0.00

1.17 1.00 3.39 0.00 0.04 5.61

0.23 0.83 1.47

0.00 0.11 0.01 0.00

1.20 2.00 2.39 0.00 0.04 5.63

0.21 0.82 1.55

0.00 0.10 0.01 0.01

1.07 2.01 2.59 0.00 0.05 5.72

Calculated based on 14 oxygen 3.41 3.39 3.31 3.28 0.59 0.61 0.69 0.72 4.00 4.00 4.00 4.00

33.69 14.74 9.16 24.06 0.00 0.54 0.80 0.01 0.11 83.10

UVZ 90

0.04 0.85 1.58

0.00 0.02 0.00 0.01

1.03 2.05 2.69 0.00 0.00 5.77

3.40 0.60 4.00

32.05 13.05 23.09 17.05 0.01 0.00 0.14 0.00 0.05 85.43

UVZ 136

0.05 0.81 1.67

0.00 0.02 0.00 0.01

0.89 2.18 2.81 0.00 0.03 5.91

3.23 0.77 4.00

30.43 13.32 24.58 17.80 0.04 0.33 0.16 0.02 0.06 86.74

UVZ 4

0.27 0.86 1.48

0.00 0.12 0.02 0.00

1.14 1.10 3.30 0.00 0.04 5.57

3.44 0.56 4.00

38.13 15.90 14.57 24.48 0.06 0.47 1.25 0.12 0.04 95.03

UVZ 6

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Fig. 10. Variation diagrams of some major and trace elements of the studied basaltic flood as a function of the 45 m depth. Black squares represent the location of the analyzed samples.

a saponite to chlorite/saponite mixed layer sequence in the UVZ. Other secondary phases, such as celadonite and zeolite, are limited to the UVZ. Celadonite replaces olivine and lining vesicles. Zeolite minerals (heulandite, mesolite/scolecite, stilbite, thomsonite, laumontite, and mordenite) are associated with the pervasive albitization of calcic plagioclase and vesicle infilling. Laumontite and mordenite are limited to areas along the contact with the upper lava flow. This distribution of alteration features and secondary mineralogy is generally described as a low-grade

metamorphic product (Alt et al., 1996; Schmidt and Robinson, 1997; Alt, 1999; Neuhoff et al., 1999; Robinson and Bevins, 1999) whose main characteristics are (1) the secondary paragenetic sequence following the porosity grade of the basaltic rock, namely, UVZ and LVZ levels and the IMZ level (Schmidt and Robinson, 1997) and (2) the crystallochemistry of clay minerals evolving from a C/S mixed layer with low chlorite content to C/S with higher chlorite content to pure chlorite. Minerals of the celadonite group are systematically present in continental and oceanic basaltic flows. Most of these celadonitic minerals are

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deficient in potassium compared with the end member (Kaleda and Cherkes, 1991; Alt, 1999). Moreover, the vesicular levels frequently show albitization (Schmidt and Robinson, 1997; Alt, 1999; Neuhoff et al., 1999; Robinson and Bevins, 1999). The vesicles are filled by a C/S mixed layer in which the increase in chlorite content is explained by the progressive increase in temperature and time during the lava burial (Neuhoff et al., 1999). After flowing in a subaerial or submarine environment, successive events can affect the primary rock. They can operate early during the cooling or seawater/hydrothermal alteration processes or during the late burial, low-grade metamorphism. These successive events give nearly identical paragenesis. Moreover, the largely developed events, such as low-grade metamorphism and fossil geothermal fields, display a paragenesis that follows the temperature gradient at the hectometric to kilometric scale and preserve the structure of each lava flow unit (Harrison and Merriman, 1984; Hulen and Nielson, 1986; Lonker et al., 1993; Dudoignon et al., 1997). The composition of clay minerals evolves through three structural zones (Fig. 11). In the LVZ, secondary clays are saponite, which occurs as a mesostasis component and in olivine replacement, with an MgO/(MgO þ FeO) ratio close to 0.66%. Saponite that fills the vesicles shows a MgO/(MgO þ FeO) ratio ranging from 0.66 to 0.80%. In the IMZ, clay minerals are restricted to mesostasis sites and constitute random C/S mixed layers. They are approximately 20% chlorite layers and characterized by an MgO/ (MgO þ FeO) ratio close to 0.40%. In the UVZ, the distribution of secondary minerals is more complex (Fig. 11):

† Celadonite as a pseudomorph on olivine and precipitated along vesicle rims bears MgO/(MgO þ FeO) ratios close to 0.15 and 0.23%, respectively; † Saponite as a mesostasis component displays a large range of MgO/(MgO þ FeO) ratios, from 0.66 to 0.40%; † Saponite to C/S mixed layer clay sequence as a vesicle infilling has MgO/(MgO þ FeO) ratios decreasing from 0.66% in saponite to 0.42 – 0.37% in the C/S mixed layer; and † Pervasive albitization of plagioclases yields zeolites crystallized into vesicles. 6.3. Temperature and time conditions during the cooling stage The gradual temperature decrease in lava flows occurs in two stages. The first is dominated by degassing and fractional crystallization, with temperatures between the liquidus (1200 8C) and the solidus (980 8C). The second stage cools below the solidus (980 8C to ambient temperature). During the first stage, solidification fronts, represented by a partially liquid and partially crystalline portion of the magma, advance from the top and base of the flow to the center (Fig. 12). The concomitant degassing/vesiculation process and the progression of the two solidification fronts build the vertical zonation of the flow (i.e. UVZ, IMZ, and LVZ). Simple calculations based on Jaeger’s (1961, 1968) equations, which only take conduction into account, show that isotherms of the solidus 980 8C intercept one another at a depth of 29.40 m. This distance corresponds to approximately two-thirds from the top in a 45 m thick flow. Thus, the top of the flow, in contact with the atmosphere, cooled

Fig. 11. Sketch of the clay mineral zonation in the UVZ, IMZ, and LVZ on the basis of their MgO/MgO þ FeO ratios. (0) corresponds to a postmagmatic stage in oxidative conditions of initial high porosity due to degassing in the UVZ; (1) corresponds to a postmagmatic stage in reducing conditions; and (2) corresponds to a late low-grade metamorphism stage. Sap ¼ saponite, C/S , 20% ¼ random chlorite/saponite mixed layer with less than 20% chloritic layers, and cel ¼ celadonite. The value of the MgO/MgO þ FeO ratio is 0.66.

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Fig. 12. Progressing curves of the upper and lower solidification fronts (1150 and 980 8C isotherms) of inward lava flow (Jaeger, 1961, 1968).

more rapidly than the base, which was in contact with the substrate. The time estimated for complete solidification of the 45 m thick flood basalt is on the order of 35 years (Schenato, 1997). Cooling proceeds when the solidus temperature drops to ambient temperature after the flow is degassed and crystallized (Jaeger, 1961, 1968; McMillan et al., 1987). During cooling, the temperature varies within the flow. Cooling occurs faster at the top of the flow, whereas high temperatures remain for long periods in the center and the base. Cooling time for the entire lava flow from 980 8C to ambient temperature, following Jaeger’s (1961, 1968) model calculations, is estimated at 560 years (Fig. 13) (Schenato, 1997). The composition of clay minerals precipitated in different volcanic environments varies according to the nature of the reacting fluids, the water/rock ratio, and the temperature gradient. The dominant clay minerals newly formed in basaltic rocks are trioctahedral smectites (saponite), which may have different origins, including seawater alteration, hydrothermal activity, diagenesis, and

low-grade metamorphism. Whatever the natural environment, the temperature stability for smectite crystallization is , 200 8C (Pytte and Reynolds, 1989; Tucker, 1991; Dudoignon et al., 1997). The conversion of saponite to chlorite has been studied in diagenetic systems (April, 1981; Velde, 1985; April and Keller, 1992) and hydrothermal alterations (Tomasson and Kristmannsdottir, 1972; Kristmannsdottir, 1979; Stackes and O’Neil, 1982; Inoue et al., 1984; Inoue, 1985; Meunier et al., 1991; Dudoignon et al., 1997). This conversion involves a C/S mixed layer formation stage. In general, the temperature stability of saponite is 100– 200 8C, C/S mixed layer ranges from 200 to 230 8C, and chlorite forms at temperatures . 240 8C (Evards and Schiffman, 1983; Bettison and Schiffman, 1988; Bettison et al., 1991). The percentage of chlorite in the mixed layers gradually increases as a function of temperature and time. In basaltic environments, large domains of temperature stability have been considered for saponite formation from 200 – 240 8C (Andrews, 1980) to 100– 300 8C (Banks and Melson, 1966). Furthermore, saponite was observed as

Fig. 13. (a) Progressing of cooling (under solidus) through the lava flow during the first 100 years. The curve asymmetry reflects the difference in surface temperatures of the flow and the base. (b) Progressing of cooling through the lava flow between 100 and 560 years. Calculated curves are based on Jaeger’s (1961, 1968) equations.

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a metastable phase formed at high temperatures in the geothermal fields of Milos (Greece) and Chipilapa (El Salvador) (Papanagiotou et al., 1995; Patrier et al., 1996). Studies of active geothermal fields (Beaufort et al., 1996; Patrier et al., 1996) attest to the main role of kinetics in the conversion of saponite to C/S in the 200– 300 8C temperature range. They demonstrate the increase of percentage chlorite layers in C/S mixed clays as a function of time when temperature is maintained. In the geothermal fields, the temperature for crystallization of saponite, calculated from d 18O isotopic data, ranges between 130 and 240 8C (Kristmannsdottir, 1979; Stackes and O’Neil, 1982; Dudoignon et al., 1997). Chlorite/saponite mixed layers and chlorite present crystallization temperatures that may reach 350 8C (Lonker et al., 1993; Dudoignon et al., 1997). The temperature for saponite crystallization in the quenched peripheries of submarine and subaerial lava flows identified as synchronous to the lava cooling were calculated from d 18O data as equal to 80 –100 8C (Dudoignon et al., 1989, 1992). Moreover, experimental syntheses have shown that trioctahedral, magnesium-rich clays (saponite) can be the first minerals to crystallize at high temperature conditions up to 600 8C in the SiO2 – Al2O3 –MgO – H2O system (Iijima and Roy, 1963; Harder, 1971; Whitney, 1983; Kloprogge et al., 1999). In both continental and oceanic basaltic flows, celadonite is associated with low temperature events, often , 40 8C (Alt, 1999). Nevertheless, synthesis experiments show that celadonite is stable in elevated temperature/low pressure conditions (Wise and Eugster, 1964; Velde, 1972). The celadonite – biotite joint is reached at approximately 400 8C. In addition to reports of these phyllosilicates in hydrothermal or hydrothermal-metamorphic alteration (Alt et al., 1996; Alt, 1999) on basaltic lava bodies, they have been attributed to the final products of fluids during immediate posteruptive (deuteric or postmagmatic) cooling in basaltic lavas (Scheidegger and Stakes, 1977; Walker and Ineson, 1983; Shayan et al., 1990; Shelley, 1993). The 300– 80 8C temperature range recorded in the inner part of the 45 m thick flow is approximately 150 years (Jaeger, 1961, 1968; Schenato, 1997). This must have been the time available for the crystallization of secondary minerals during a postmagmatic stage, mainly in the residuum mesostasis sites. The rate of this process remained constant in the massive central part of the flow for long periods, whereas the temperature decreased quickly in the upper part. In these conditions, following magnesium-rich saponite crystallization, the C/S mixed layer in the inner part of the flow may be explained by a saponite– chlorite conversion process that produced a C/S series (Meunier et al., 2000). 6.4. The source of H2O During degassing, considerable quantities of fluids initially dissolved in the basaltic melt escape into

the atmosphere; only some remain in the frozen magma. Experimental data show that water solubility in a basaltic melt depends on lithostatic pressure (Burnham, 1979). At a confining pressure of 13.6 bar (45 m deep) at the flow base, the fraction of water mass dissolved in the lava at the time of extrusion is estimated to be , 0.28%. This corresponds to the calculated minimum quantity of water dissolved in the melt at the time of extrusion. Authigenic water is exsolved into vesicles, though some may remain dissolved in the magma, depending on the lithostatic pressure. Degassing and bubbling stop when the viscosity increases drastically (. 105 poise) in the 1150– 1000 8C temperature range (Aubele et al., 1988; Goff, 1996). Thus, the isothermal line at 1150 8C is considered the bubble ‘freezing’ line. During the first stage, the processes of nucleation, growth, and rise of bubbles concur with the progression of the upper and lower freezing lines. Following a decrease in lithostatic pressure from the base to the top of the flow, the mechanisms can be simplified as follows: † Instantaneous trapping of early nucleated and small vesicles at the base of the flow (LVZ formation); † Nucleation, growth, and rise of vesicles through the inner and unfrozen part of the flow (IMZ vesicle-free zone); and † Progressive trapping and coalescence of ascending vesicles at the upper freezing line contact, which progress down toward the inner part of the flow (formation of the UVZ). Thus, the fractional crystallization of the hydrated melt (in equilibrium with the lithostatic pressure) produces early formed anhydrous minerals during the cooling stage of the flow and a small quantity of remaining fluids with H2O and volatiles that may be trapped and concentrated into the interstitial residuum of rock (Walker, 1989). These residual and differentiated hot fluids may account for the first crystallization of clay minerals in both massive and vesicular levels in mesostasis sites. In a second step, the high permeability of the vesicular level probably enables the interaction between circulating hot aqueous fluids and the development of large alteration processes. According to the petrographic features, the chemical composition, and the mineralogy of the clay minerals, three steps may be involved in the formation of secondary phases in the lava unit: 1. Celadonite precipitation in oxidative conditions according to the high porosity and associated high permeability of the upper vesicular level (UVZ). The celadonite precipitate in the most sensitive microsites: olivine crystals and the wall rock periphery of the vesicle. The low calculated MgO/(MgO þ FeO) ratio corresponds to the high amount of Fe trapped as Fe3þ in the celadonitic clay mineral.

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2. Saponite crystallization, with a MgO/(MgO þ FeO) ratio close to 0.66%, trapped throughout the lava flow. Saponite is located in mesostasis sites (infilling microsites between K-feldspar grains) in the LVZ and UVZ but is absent in the IMZ. Saponite progressively evolves to a C/S mixed layer (MgO/(MgO þ FeO) ¼ 0.40%) with a low percentage chlorite layer (, 20%) during the long cooling time and temperature stability of the IMZ. 3. The high water/basalt interaction in the outer permeable level, which results in progressive crystallization of C/S with an increased chlorite layer percentage and decreased MgO/(MgO þ FeO) ratio (, 0.40%). The saponite to C/S evolution is evident from the outer rims to the center of the vesicles. The richer chloritic layer C/S precipitates in the inner parts of the vesicles. The high W/R ratio also causes chemical changes in the initial composition of the saponite in mesostasis—its MgO/(MgO þ FeO) ratio, initially close to 0.60%, shifts to 0.40%. In addition, albitization of plagioclases occurs and secondary albite and zeolite paragenesis is formed. The intensity of albitization depends on the permeability of the medium. It is more discrete in the LVZ, where plagioclases are partially albitized without any associated zeolites. In the UVZ, the pervasive albitization of calcic plagioclases induces large compositional variations in the zeolites crystallized in vesicles. The great variety of zeolites is usually distributed according to the depth and T gradient in low-metamorphic and hydrothermal piles. In this case, zeolites are concentrated in a 10 –15 m level. This late crystallization stage related to the zeolite minerals seems synchronous with late C/S formation in the vesicles.

7. Conclusions The vertical zonation classically observed throughout a lava flow body is the result of the cooling stages of the unit. These induce the peripheral vesicular zones, the IMZ and associated fracture network, and characteristics of the entablature or columnar levels. The clay minerals that have crystallized in different volcanic piles generally have been studied in vesicle or columnar joint infillings or as primary phase replacements. These clay minerals are commonly associated with zeolite crystallization and explained as a result of hydrothermal or low-grade metamorphic alteration processes. In these alteration mechanisms, the primary phases most sensitive to dissolution and clay mineral replacement are olivine and glass. Albitization of plagioclases is generally observed in more permeable parts. The detailed petrography of a 45 m thick basaltic lava flow shows variations in the alteration features and associated clay minerals. These variations are supported by the vertical structural differences of the lava flow and the porosity of the three levels. Larger secondary mineral

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varieties are observable in the UVZ: a celadonite/saponite to C/S sequence in vesicles; celadonite replaces olivines, and secondary albite plus zeolite replaces plagioclases. In contrast, the mesostasis areas are restricted microsites for the crystallization of clay mineral in all zones. In mesostasis, only saponite is observable in the vesicular levels and C/S mixed clays (with a chloritic layer of approximately 20%) in the inner massive part, which suggests a saponite to C/S conversion in the slowly cooling inner part. The vertical clay mineralogy is usually reported as due to low-grade metamorphism during burial, whereas the variation in mineralogy is attributed to differences in porosity (Schmidt and Robinson, 1997; Neuhoff, 1999). The high porosity grade supports the alteration intensity. However, pressure and temperature conditions during the cooling steps are not homogeneous in the 45 m thick flow units, according to the vertical curves of the cooling rate in a single lava flow. The clay mineralogy shows that the first alteration stage depends on the primary porosity of the flow. It was marked by celadonite precipitation in oxidative conditions, according to the initially permeable system and the associated high porosity of the UVZ. It suggests the earliest precipitation of celadonite in the most sensitive microsites: olivine crystals and the wall rock periphery of the vesicles. The progressive porosity decreases, induced by the first clay mineral precipitation, enable secondary clay mineral precipitation in reduced conditions (saponite and C/S). Thus, homogenous saponite precipitates in reduced conditions and then is converted to saponite-C/S in the inner part of the flow, where temperature gradients are preserved for a longer time during the final cooling stages. Petrographic and chemical observations support three steps in the alteration mechanisms: (1) earliest celadonite and/or saponite crystallization, (2) saponite to C/S conversion during the cooling stages, and (3) pervasive alteration in the more permeable vesicular levels, probably related to low-grade metamorphic alterations. This chronology suggests a high temperature of celadonite crystallization. The saponite-C/S mixed layer sequence identified in the mesostasis is commonly associated with a crystallization temperature range of 80 –300 8C (Schenato, 1997). This temperature range persisted for approximately 150 years in the inner part of the flow. After the early precipitation of anhydrous phases due to fractional crystallization, the H2O-rich residuum promoted the precipitation of clay minerals of deuteric or postmagmatic origin. Postmagmatic crystallization of saponite to a C/S mixed layer in mesostasis sites implies a glass problem. Glass is undoubtedly observed along the chilled margins of flows or in the outer skins of pillow lava, though its presence is questionable in the massive parts, where crystallization of the lava body occurs at a lower cooling rate. Petrographic features at a crystal scale show an

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evolution of grain size from the outer contact of the flow to the inner parts, which suggests different cooling rates and undermines the hypothesis of glass formation or devitrification. Moreover, the interstitial material usually described as glass in basaltic systems has been identified as a granitic-like residuum (acid composition with high silicon content of 60 – 75%) that is due to microlocal fractionation (Lambert et al., 1989; Hoover and Murphy, 1989; Shelley, 1993). In tholeiitic basalts, granophyric material is often observed in the interstices following primary phase crystallization of olivine, pyroxene, and plagioclase (Shelley, 1993). This accords with the presence of K-feldspar and quartz in mesostasis sites. The fluids remaining after fractional crystallization processes have concentrations of H2O and volatile species, such as Na2O, which were probably abundant enough to decrease the effective viscosity of the residual liquid and prohibit glass formation in undercooled conditions (Burnham, 1979; Cox et al., 1979; Jaupart and Tait, 1990). It is often easier to invoke glass devitrification or late glass alteration to explain the temperature gap between primary phases, such as plagioclase, K-feldspar, and quartz, and clay minerals in interstitial sites as mesostasis. Direct precipitation of clay minerals from deuteric fluids during the immediate postmagmatic stages could be envisaged instead of glass alteration, because the following favorable conditions occur during the cooling period of a solidified lava flow: (1) high temperatures are preserved for a longer time in the center of lava bodies because of the low thermal conductivity of silicate rocks and (2) residual fluids are enriched in volatile (H2O) and incompatible components, and thus, hydrated silicates (i.e. clay minerals) can be formed.

Acknowledgements This research was supported by Project Capes/Cofecub no. 185/96 and Fapergs (RS, Brazil).

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