Urumieh-Dokhtar magmatic belt trending NW-SE parallel to the SSZ (50 to 80 km wide) extends from south to northwest of Iran, situated next to Tabriz fault in northeast of this zone. There are some volcanic rocks cropping out in the south of Sarab. Magmatic activities in this belt started in the Eocene and continued until Quaternary (Omrani, 2008; Azizi, 2009). Magmatic activities are divided into two phases: a) volcanic activities in Eocene followed by

Mesozoic-Paleogene magmatism in the SSZ.

given in Ma, (Ghalamghash, 2009).

**2.2 Urumieh-Dokhtar magmatic belt(UDMB)**

There are scattered exposures of the CIB, overlain by the Cenozoic volcanic rocks in northeast of UDMB and Alborz-Azerbaijan mountain ranges. These rocks, cropping out in the study area, are pre-Triassic sedimentary rocks that are similar to those of the south Armenia and Tauride blocks in Turkey. According to Sosson et al. (2005) and Dilek et al. (2010), Central Iran, Armenia and Tauride blocks were formed the northern margin of the Gondwana land prior to Triassic rifting. These three blocks include Precambrian crystalline basement overlain by Paleozoic–Mesozoic sedimentary sequences (Dilek, 2009; Roland, 2009; Sosson, 2005), separated from Gondwana land in the Late Triassic and joined the southern margin of the Eurasian plate during Late Cretaceous–early Paleocene (Dilek, 2009; Roland, 2009). Paleomagnetic investigations on Middle Triassic igneous rocks in Sumkhit- Gharebagh (Azerbaijan) and south Armenia indicate that the latitude was at about 22 N for the time of their formation (Sosson, 2005). The Paleo-Tethys Ocean started closing as the Neo-Tethys Ocean opened during separation of Central Iran, Armenia and Tauride blocks from the Gondwana land moving northwards. According to Bazhenov et al. (1996), the site of paleo-Tethys ocean enclosure is the present place of Sumkhit-Gharabagh-Black Sea area.

#### **2.4 Alborz-Arasbaran-Lesser Caucasus Belt (AALCB)**

The Transcaucasian Massif includes Pan-African orogenic crust intruded by latest Proterozoic to Palaeozoic granitoids, experienced multiple deformation and migmatitization, and the Jurassic to Early Cretaceous plutons representing a magmatic arc (Zakaridze et al., 2007). This arc continued into the Eastern Pontide blocks in the west.

A Cretaceous island arc complex with calc-alkaline to alkaline extrusive rocks, and pyroclastic deposits, flysch units and marl-limestone rocks occurs north of the Sevan-Akera suture zone (Dilek et al., 2010).

Cenozoic magmatic rocks of the AALCB are mainly acidic to basic volcanic rocks and acidic to intermediate intrusive bodies, with calc-alkaline to alkaline affinities trending NW-SE and located in the northwestern Iran and Caucasus mountain ranges (Moayed, 2001; Jamali et al., 2010). Similar to UDMB, magmatic activities in this belt have been started in the Eocene and continued to Quaternary.

Azizi et al. (2009) and Alavi (2007) believe that the scattered pieces of ophiolites around Tabriz fault are the remnants of UDMB back-arc oceanic crust (basin) and its subduction was the cause of the formation of Alborz-Arasbaran magmatic belt. There are scattered exposures of ophiolites and metamorphic rocks from Anzali to Sevan- Akera at northern margin of the AALCB (Fig. 1) (Galoyan et al., 2009; Berberian et al., 1981; Jamali et al., 2010). Based on isotopic data, the formation of Sevan-Akera ophiolites have started in the Middle-Late Jurassic and continued to Early Cretaceous (Galoyan et al., 2009). According to Sosson (2005), based on interlayering of radiolarites and pillow basalts, the age of ophiolites in the northern Armenia is proposed to be Late Jurassic. The exposure sites of these ophiolites could be the site of the Mesozoic ocean that was closed in Late Cretaceous-Early Paleocene causing the AALCB magmatisms. Jurassic to Cretaceous magmatic rocks with alkaline to calc-alkaline affinities in the north Lesser Caucasus indicates arc magmatism characteristics and continued to the west joining the Pontide magmatic belt in northeastern Turkey (Dilek, 2010; Zakaridze, 2007). Intense volcanic activities with acidic to basic composition were started in the AALCB during Eocene. Eocene volcanic rocks with alkaline to calc-alkaline nature (Figs. 3 and 4) indicate continental margin magmatism. Distribution patterns of the

Fig. 3. K2O (wt%) versus SiO2 (wt%) diagram for clasification of NW Iran volcanic rocks (Peccarillo and Taylor, 1976).  Similar to UDMB, magmatic activities in this belt have been started in the Eocene

Azizi et al. (2009) and Alavi (2007) believe that the scattered pieces of ophiolites around Tabriz fault are the remnants of UDMB back-arc oceanic crust (basin) and its subduction was the cause of the formation of Alborz-Arasbaran magmatic belt. There are scattered exposures of ophiolites and metamorphic rocks from Anzali to Sevan- Akera at northern margin of the AALCB (Fig. 1) (Galoyan et al., 2009; Berberian et al., 1981; Jamali et al., 2010). Based on isotopic data, the formation of Sevan-Akera ophiolites have started in the Middle-Late Jurassic and continued to Early Cretaceous (Galoyan et al., 2009). According to Sosson (2005), based on interlayering of radiolarites and pillow basalts, the age of ophiolites in the northern Armenia is proposed to be Late Jurassic. The exposure sites of these ophiolites could be the site of the Mesozoic ocean that was closed in Late Cretaceous-Early Paleocene causing the AALCB magmatisms. Jurassic to Cretaceous magmatic rocks with alkaline to calc-alkaline affinities in the north Lesser Caucasus indicates arc magmatism characteristics and continued to the west joining the Pontide magmatic belt in northeastern Turkey (Dilek, 2010; Zakaridze, 2007). Intense volcanic activities with acidic to basic composition were started in the AALCB during Eocene. Eocene volcanic rocks with alkaline to calc-alkaline nature (Figs. 3 and 4) indicate continental margin magmatism. Distribution patterns of the

Fig. 3. K2O (wt%) versus SiO2 (wt%) diagram for clasification of NW Iran volcanic rocks (Peccarillo and Taylor, 1976). The abbreviations are as follow: M-Qt: Mio-Pliocene to Quaternary(Kheirkhah et al., 2009), Qt: Quaternary (Dilek et al., 2010), E: Eocene (Dilek et al., 2010), M-Pl: Mio-Pliocene (Dilek et al., 2010), Qt: Quaternary (Ahmadzadeh et al., 2010),

Cret: Cretaceous (Azizi et al., 2009), E: Eocene (Azizi et al., 2009), Qt: Quaternary

and continued to Quaternary.

(Ahmadzadeh et al., 2010).

REEs (weak or absent Eu negative anomaly and more differentiated pattern of REEs) also are indicative of continental subduction zones (Fig. 5). Most of the samples (except some Cretaceous volcanics in SSZ) plotted in (Nb/Zr)n-Zr - Zn diagram clustered in the collision related and within plate fields (Fig. 6). In Rb - Y+Nb and Nb - Y diagram (Pearce, 1996, 1984) they mainly plotted in the field of post collision granitoids (Fig. 7). In the Hf, Nb, Ta and Rb diagrams of Harris (1986), most of the samples plotted at the border of VAG, WPG and Post-Collision fields (Fig. 8).

Following to the Eocene intense volcanic activities, Oligocene magmatic activities occurred in the form of intrusive bodies, in the AALCB. The composition of Oligocene intrusions is mainly acidic to intermediate with lesser amounts of mafic rocks (Fig. 9). They include granite, granodiorite, quartz monzonite, monzonite, syenite, syeno-diorite and gabbro which are intruded by lamprophyric to dacitic dykes. Based on the geochemical studies, the intrusive rocks are of different origins. Monzodiorites, gabbros and syenites show alkaline and shoshonitic nature and probably were originated from partial melting of lithospheric mantle, while the acidic rocks (granites, granodiorites and quartz monzonites) with high-K, calc-alkaline affinities (Fig. 8) show C type adakitic magma characteristics. They also show high amounts of Ba and Sr and originated from partial melting of lower mafic and potassic crust due to increase of the crust's thickness. Finally, the lamprophyrs are alkaline type and formed from the melting of OIB type metasomatized lithospheric mantle (Aghazadeh, 2010).

Fig. 4. Total alkali (wt%) versus SiO2 (wt%) clasification of volcanic rocks (after Le Bas et al., 1986). (symbols as in Fig.3)

Fig. 5. C1 Chondrite normalized REE patterns for Meso-Cenozoic volcanic rocks in NW Iran. Mean values of REE used for each rock group. (symboles as in Fig. 3).

Fig. 6. (Nb/Zr)n versus Zr (ppm) diagram (Thieblemont and Tegyey, 1994) for volcanic rocks. Most of the samples plotted in the collision related field. (Abbreviations as in Fig. 3).

**8 10**

**La**

**Ce**

**Pr**

Mean values of REE used for each rock group. (symboles as in Fig. 3).

**Nd Sm**

**Eu**

Fig. 5. C1 Chondrite normalized REE patterns for Meso-Cenozoic volcanic rocks in NW Iran.

Fig. 6. (Nb/Zr)n versus Zr (ppm) diagram (Thieblemont and Tegyey, 1994) for volcanic rocks. Most of the samples plotted in the collision related field. (Abbreviations as in Fig. 3).

**Gd**

**Tb**

**Dy**

**Ho**

**Er**

**Tm**

**Yb**

**Lu**

**100**

**Sample/C1 Chondrite**

**300**

Fig. 7. Tectonic discrimination diagram of Rb versus Yb + Nb and Nb versus Y (after Pearce et al., 1984 and 1996). Most of the samples plotted in the post collision field (blue circles). (symbols as in Fig.3)

Fig. 8. Triangular diagrams of Harris et al. (1986), for various tectonic environments in collision zones. Most of the samples plotted at the border of VAG, WPG and Post-Collision field, which is characteristics of the post-collision magmatism. (symbols as in Fig. 3)

Field observations indicate that the acidic calc-alkaline rocks are older than intermediate to basic shoshonitic rocks. These observations are confirmed by zircon age dating of intrusive bodies of north and east of Ahar given by Aghazadeh (2010). The results of these dating indicate that the age of adakitic granodiorite is 31.8 Ma., gabbro-monzonite is 28 Ma and syenite and melasyenite is 24-26 Ma. In K2O-SiO2 diagram, lamprophyrs and synenites are plotted in shoshonite field and gabbros, monzonites and granites (resulted from their differentiation) belong to the high potassium series. Granites, granodiorites and quartz monzonites are situated in medium to high potassium calc-alkaline field (Fig. 10). In diagram of trace elements of Pearce (1984, 1996) and Thieblemont and Tegyey (1994), most of the samples belong to post collision environments and they have no direct relationships with subduction (Figs. 11 and 12).

Fig. 9. Geochemical classification of the AALC plutons. (A) Using the total alkali versus SiO2 (Wilson, 1989); the blue line separates alkaline rocks from sub-alkaline. (B) Zr/TiO2 versus SiO2 diagram of Winchester and Floyd (1977).

differentiation) belong to the high potassium series. Granites, granodiorites and quartz monzonites are situated in medium to high potassium calc-alkaline field (Fig. 10). In diagram of trace elements of Pearce (1984, 1996) and Thieblemont and Tegyey (1994), most of the samples belong to post collision environments and they have no direct relationships

Fig. 9. Geochemical classification of the AALC plutons. (A) Using the total alkali versus SiO2 (Wilson, 1989); the blue line separates alkaline rocks from sub-alkaline. (B) Zr/TiO2 versus

SiO2 diagram of Winchester and Floyd (1977).

with subduction (Figs. 11 and 12).

Fig. 10. K2O (wt%) versus SiO2 (wt%) diagram for clasification of the Arasbaran Oligicene plutonic rocks (Le Maitre et al.,1989; Rickwood, P.C., (1989). (symboles as in Fig.9).

Fig. 11. Tectonic discrimination diagram of Rb versus Yb + Nb and Nb versus Y (after Pearce et al., 1984; Pearce, 1996). Most of the samples plotted in post collision field (blue circle). (symboles as in Fig.9).

Fig. 12. (Nb/Zr)n versus Zr (ppm) diagram (Thieblemont and Tegyey, 1994) for the volcanic rocks. Most of the samples plotted in the collision related field.

Oligocene intrusive bodies could be divided into two groups based on their structure, texture and the depth of emplacement: (1) Large intrusive bodies (batholites) with medium to coarse grained texture that formed at 3 to 4 km depth such as Sheivar Dagh, Khankandi, Youseflu and Kaleibar are poorly mineralized, and mineralization associated with them mainly occurs in the contact zone with their host rocks or within the country rocks. These intrusive bodies show high K calc-alkaline to shoshonitic affinity. (2) Sub-volcanic bodies with porphyritic textur and calc-alkaline character formed at 1-2 km depth and are associated with remarkable mineralization such as Sungun, Mivehrud, Masjed Daghi and Haftcheshmeh deposits. Most of the mineralization occurs in quartz monzonite-monzonite. Microscopic investigations of rock samples of quartz monzonite- monzonite indicate that amphibole and biotite phenocrysts are paragentically earlier than other minerals. This indicates that the amount of water in the original magma was more than 3% in the crystallization stage (Whitney, 1975; Burham, 1979; Whitney and Stromer, 1985). The lack of Eu negative anomaly may be assumed to be the result of high oxygen fugacity in the original magma, because the high oxygen fugacity oxidizes Eu2+ to Eu3+ and so it can not enter into the plagioclase structure (Hezarkhani, 2006). In addition to bearing water and oxidizing state of the primary magma, replacement in the lower depth caused water saturation and fluid separation in primary stages of crystallization of original magma followed by mineralization.

After Oligocene intrusion phase in the Arasbaran area the next magmatic activities occurred in the Late Miocene and continued to Quaternary. The Late Miocene-Pliocene magmatic rocks are mostly acidic in composition (granite–granodiorite) and they are in the form of sub-volcanic bodies and/or volcanic domes (such as Ghalehlar and Bohlul Daghi), that could be observed in the southwestern Arasbaran area. These magmatic rocks have alkaline adakitic nature and they are probably related to post collision processes such as slab break off or delamination (Jahangiry, 2007; Karimzade, 2004). Adakitic magmatism in the Plio-Quaternary was followed by basic alkaline volcanism. At first, the alkali basic magma was sodic in composition changing upward to more potassic in composition, and bears leucite phenocrysts. The Plio-Quaternary volcanic rocks include andesite, basalt-andesite, tephritebasalt, leucite basalt and tephrite. The amounts of K2O and MgO and the K2O/Na2O ratio in these rocks are high enough to classify them as high or even ultra potassic rocks (Fig. 3). They characterized with significant enrichment in LILE and LREE and depletion in highfield-strength elements. So, the metasomatized lithospheric mantle composed of garnet lherzulite and high fugacity of CO2/H2O ratio played a significant role in their genesis (Ahmadzadeh, 2010 and Khezerlu, 2008).

#### **2.5 Discussion and conclusions**

50 Petrology – New Perspectives and Applications

Fig. 12. (Nb/Zr)n versus Zr (ppm) diagram (Thieblemont and Tegyey, 1994) for the volcanic

Oligocene intrusive bodies could be divided into two groups based on their structure, texture and the depth of emplacement: (1) Large intrusive bodies (batholites) with medium to coarse grained texture that formed at 3 to 4 km depth such as Sheivar Dagh, Khankandi, Youseflu and Kaleibar are poorly mineralized, and mineralization associated with them mainly occurs in the contact zone with their host rocks or within the country rocks. These intrusive bodies show high K calc-alkaline to shoshonitic affinity. (2) Sub-volcanic bodies with porphyritic textur and calc-alkaline character formed at 1-2 km depth and are associated with remarkable mineralization such as Sungun, Mivehrud, Masjed Daghi and Haftcheshmeh deposits. Most of the mineralization occurs in quartz monzonite-monzonite. Microscopic investigations of rock samples of quartz monzonite- monzonite indicate that amphibole and biotite phenocrysts are paragentically earlier than other minerals. This indicates that the amount of water in the original magma was more than 3% in the crystallization stage (Whitney, 1975; Burham, 1979; Whitney and Stromer, 1985). The lack of Eu negative anomaly may be assumed to be the result of high oxygen fugacity in the original magma, because the high oxygen fugacity oxidizes Eu2+ to Eu3+ and so it can not enter into the plagioclase structure (Hezarkhani, 2006). In addition to bearing water and oxidizing state of the primary magma, replacement in the lower depth caused water saturation and fluid separation in primary stages of crystallization of original magma

After Oligocene intrusion phase in the Arasbaran area the next magmatic activities occurred in the Late Miocene and continued to Quaternary. The Late Miocene-Pliocene magmatic rocks are mostly acidic in composition (granite–granodiorite) and they are in the form of sub-volcanic bodies and/or volcanic domes (such as Ghalehlar and Bohlul Daghi), that could be observed in the southwestern Arasbaran area. These magmatic rocks have alkaline

rocks. Most of the samples plotted in the collision related field.

followed by mineralization.

Two oceanic island arcs magmatism, one in SSZ (Songhor-Baneh) and the other in the north Lesser Caucasus are distinguished. Except of these two oceanic arcs, all of the magmatic rocks in the SSZ, UDMB and AALCB are related to continental margins. As going toward north and south from Tabriz fault, the magmatism becomes older and its alkalinity decreases.

SSZ magmatic rocks that are mostly intrusive type were primarily related to an active subduction margin and later on about 80 million years ago reveals post–collision magmatic characteristics. There is no unanimity for the time of collision between the Arabian plate and SSZ. Some geologists believe that the collision had been occurred in the Late Cretaceous– Early Paleocene (Moaiyed, 2001; Ghalamghash, 2009; Mazhari, 2009), while some others think of the collision had been occurred during Middle–Late Miocene (Omrani, 2009; Azizi, 2009). However, the rare earth and trace elements diagrams indicate that the major plutonic and volcanic rocks of the UDMB and SSZ belong to the post-collision environment. So, it may be concluded that the time of collision is more likely Late Cretaceous–Early Paleocene, and the following magmatisms are not directly related to subduction. On the other hand they may be related to the post collisional magmatism such as slab break off and delamination.

From Tabriz fault toward north to Lesser Caucasus, a reverse zonation is observed, that means the age of magmatic rocks increases toward north. In the vicinity of Tabriz fault the magmatic rocks are of Miocene–Quaternary age, while in the Arasbaran region and southern part of Lesser Caucasus, the magmatism belongs to Eocene- Oligocene, and from Sevan-Akera suture zone toward north, the magmatic rocks are related to Jurassic-Late Cretaceous events. However, in the study area the older rocks (Late Mesozoic) show magmatic arc characteristics, while the younger rocks (Cenozoic) are indicative of postcollision environments.

In each side of the Tabriz fault, the zoning is reversed relative to the other side and it could not be justified to be related to one subduction zone. It is more likely to belong to two separate subduction zones. It seems that in the Zagros region the subduction relates to southern Neo-Tethys, while in Lesser Caucasus-Arasbaran, this event relates to the northern Neo-Tethys, and these two events justify the magmatism and geotectonic evolution of the region. On the other hand, a northern ocean was existed between the north of Central Iran, south Armenia and Tauride blocks with Eurasian plate, and a southern ocean was between the above blocks with Arabian plate. The oceans were closed during convergence of the Arabian plate moving northward towards Eurasian plate. The dip direction of subduction in southern ocean is supposed to be northeast, while there are many controversies about the dip direction of the subducting slab of the northern ocean, although magmatic and petrogenetic characteristics of the area and the distribution of the magmatic rocks in time and space are indicative of probable southwest dip direction for subducting slab of the northern ocean.

Gold Mineralization in the SSZ were previously supposed to be related to orogenic events, while recently it is believed that they are intrusion related type mineralization that belongs to the Eocene post collision events. Recent investigations have also indicated the probable existence of porphyry type mineralization in the Hashtrud, a part of northwestern UDMB.

Carlin and epithermal type gold mineralizations are associated with Miocene adakitic magmatism, i.e. in Ghorveh and Takab region. Oligocene magmatism in the Arasbaran zone is associated with remarkable porphyry, skarn and vein type Cu-Mo-Au mineralization. Epithermal gold mineralization is accompanying both Oligocene and Miocene magmatisms.

In general, the northern belt including the AALCB is highly mineralized compare to the southern belt including the UDMB and SSZ that have relatively poor mineralization.

#### **3. Acknowledgements**

This study was supported by the Geological Survey of Iran (GSI). We acknowledge the faculty research grants from Miami University in support of fourth author work in eastern Turkey, Azerbaijan, and Northern Iran. Special thanks to Mr M.B. Dorri, N. Abedian, and B. Borna (GSI) for logistical support.

#### **4. References**

the above blocks with Arabian plate. The oceans were closed during convergence of the Arabian plate moving northward towards Eurasian plate. The dip direction of subduction in southern ocean is supposed to be northeast, while there are many controversies about the dip direction of the subducting slab of the northern ocean, although magmatic and petrogenetic characteristics of the area and the distribution of the magmatic rocks in time and space are indicative of probable southwest dip direction for subducting slab of the

Gold Mineralization in the SSZ were previously supposed to be related to orogenic events, while recently it is believed that they are intrusion related type mineralization that belongs to the Eocene post collision events. Recent investigations have also indicated the probable existence of porphyry type mineralization in the Hashtrud, a part of

Carlin and epithermal type gold mineralizations are associated with Miocene adakitic magmatism, i.e. in Ghorveh and Takab region. Oligocene magmatism in the Arasbaran zone is associated with remarkable porphyry, skarn and vein type Cu-Mo-Au mineralization. Epithermal gold mineralization is accompanying both Oligocene and

In general, the northern belt including the AALCB is highly mineralized compare to the

This study was supported by the Geological Survey of Iran (GSI). We acknowledge the faculty research grants from Miami University in support of fourth author work in eastern Turkey, Azerbaijan, and Northern Iran. Special thanks to Mr M.B. Dorri, N. Abedian, and B.

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### **Petrologic Study of Explosive Pyroclastic Eruption Stage in Shirataka Volcano, NE Japan: Synchronized Eruption of Multiple Magma Chambers**

Masao Ban1, Shiho Hirotani2, Osamu Ishizuka2 and Naoyoshi Iwata1 *1Department of Earth and Environmental Sciences, Yamagata University, 2Geological Survey of Japan/AIST Japan*

#### **1. Introduction**

56 Petrology – New Perspectives and Applications

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series and their differentiation products using immobile elements. Chemical

and Solov'eva, N. (2007) Geochemistry and geochronology of the Neoproterozoic Pan-African basement of the Transcaucasian Massif (Republic of Georgia) and implications for island arc evolution in the Late Precambrian Arabian-Nubian

> Some of detailed petrologic studies on rock samples of middle to large sized explosive pyroclastic eruptions recently revealed that the eruptions were caused by simultaneous eruption of multiple distinct magma chambers beneath the volcanoes (e.g., Nakagawa et al. 2003: Shane et al. 2007). It is very important to examine the genetic relationships among the magmas to understand the magma feeding system which caused such explosive eruptions. The explosive pyroclastic eruption stage in Shirataka volcano, NE Japan (Fig. 1) is one of potential candidates for such kind of researches. The aim of this study is to reveal the magma feeding system beneath Shirataka volcano in the explosive pyroclastic eruption stage and examine the genetic relationships among magmas involved in the explosive eruption.

#### **2. Geologic outline of the Shirataka volcano**

The northeast Japan arc (Fig. 1) is one of the representative island arcs. The volume distribution of volcanic products, excluding caldera-related felsic rocks, clearly reveals the existence of two volcanic chains: the frontal row (Nasu volcanic zone) and the back arc row (Chokai volcanic zone) (e.g., Kawano et al. 1961; Tatsumi and Eggins 1995). The Shirataka volcano is situated in the back arc row (Chokai volcanic zone) (Fig. 1). The geologic studies of the Shirataka volcanoes have been performed by many researchers (e.g., Mimura & Kanno, 2000; Ishii & Saito, 1997). Based on the results of these studies, the activity of the Shirataka volcano (0.9-0.7 Ma) is summarized to three stages; strato-cone building, explosive pyroclastic eruption, and lava-dome + block-and-ash flow forming stages (Table 1). The Kokuzo lava, the Numata pumice flow deposit, and Hagino blockand-ash flow deposit + four lava domes (Shiratakayama, Kitsunegoe, Nishikuromoriyama lava dome group + Higashikuromoriyama lava dome) were formed in each stage. Subsequently, the Shirataka volcano collapsed, creating a horseshoe-shaped caldera of ca. 4 km diameter (Yagi et al., 2005).

Fig. 1. Location of the Shirataka volcano in the northeast Japan arc. Circles represent Quaternary stratovolcanoes in the northeast Japan arc. The distribution of the volcanoes is from Nakagawa et al. (1988). The distribution of the volcanoes are divided into four volcanic zones: Aoso-Osore (green circle), Sekiryo (orange circle), Moriyoshi (yellow circle), and Chokai (purple circle) zones from trench to rear arc side (Nakagawa et al., 1988).

Table 1. Simplified stratigraphic relationships and rock series & types of the eruptive products form the Shirataka volcano. \*Four lava domes, the Shiratakayama, Kitsunegoe, Nishikuromoriyama lava domes group + Higashikuromoriyama lava dome

#### **3. Petrologic outline of eruptive products from the Shirataka volcano**

#### **3.1 Petrography**

58 Petrology – New Perspectives and Applications

Chokai zone Nasu zone

Fig. 1. Location of the Shirataka volcano in the northeast Japan arc. Circles represent Quaternary stratovolcanoes in the northeast Japan arc. The distribution of the volcanoes is from Nakagawa et al. (1988). The distribution of the volcanoes are divided into four volcanic zones: Aoso-Osore (green circle), Sekiryo (orange circle), Moriyoshi (yellow circle), and Chokai (purple circle) zones from trench to rear arc side (Nakagawa et al., 1988).

140 E 142 E

0 100km

*Pacific Ocean*

Shirataka

volcano

*Sea of Japan*

The eruptive products of the Shirataka volcano are andesitic to dacitic lavas in the first and third stages, but the products are mainly pumice with minor amount of scoria in the second explosive pyroclastic eruption stage. The banded pumice, composed of white pumice and black scoria parts, is rarely observed in the second stage. Mafic magmatic inclusions (~30 cm) are observed in all lavas (Hirotani & Ban, 2006; Hirotani et al., 2009). These are thought to be quenched products of the mafic magma (e.g., Eichelberger, 1980; Bacon, 1986). According to Hirotani & Ban (2006) and Hirotani et al. (2009), the lavas, pumice, and scoria are porphyritic (total phenocryst: ca. 23-40 vol.%) andesitic to dacitic (57.2-65.7 wt% SiO2), with orthopyroxene, clinopyroxene, plagioclase, and Fe-Ti oxide, with or without hornblende, quartz and olivine phenocrysts. The mafic inclusions are light to dark gray colored, usually rounded in form, and moderately vesiculated, are porphyritic basalt to andesite (48.3-58.5 wt% SiO2). The total volume of phenocrysts is <ca. 5 vol.% in basalt, and ca. 10 vol.% in basaltic-andesite to andesite, with orthopyroxene, clinopyroxene, and plagioclase, and with or without olivine and quartz phenocrysts. The groundmass has quenched diktytaxitic texture (Bacon, 1986).

#### **3.2 Petrographic evidence for magma mixing and variety of the phenocrysts**

The mafic inclusions and disequilibrium phenocryst assemblages (Sakuyama, 1981) such as Mg-rich olivine co-existing with quartz and/or magnesio-hornblende are observed in most of the samples. These features suggest that most of the rocks from Shirataka volcano were formed by magma mixing/mingling (Hirotani & Ban, 2006; Hirotani et al., 2009). Correspondingly, the chemical compositions of phenocrystic minerals show variability. Based on the chemical compositions with textural features, Hirotani & Ban (2006) grouped phenocrystic minerals into several groups. The phenocrysts form different groups cannot coexist in equilibrium each other by the point of view of their chemical compositions. In the cases of the stages 1 and 3 products, the groups are named to population I, II, and III by Hirotani et al., (2009). Here we refer "population" as "group". Group I includes An-rich plagioclase, Mgrich olivine and Mg-rich clinopyroxene. Most of these phenocrysts show compositional normal zoning. Group II includes An-poor plagioclase, Mg-poor orthopyroxene, Mg-poor clinopyroxene, hornblende, quartz, titanomagnetite, and ilmenite. Most of the phenocrysts in this population except for Fe-Ti oxides, show compositional reverse zoning and dissolution textures, such as sieved textures (Tsuchiyama, 1985) or embayed form. Group III is made of smaller phenocrysts of plagioclase, olivine, orthopyroxene, and clinopyroxene. These are interpreted to reflect relict cores from a mafic magma, relict cores from a silicic magma, and minerals growing in the hybrid magma caused by mixing the two magmas precipitated the crystal cores of the groups I and II (Hirotani & Ban 2006).

#### **3.3 Rock series and type of eruptive products of the Shirataka volcano**

Hirotani & Ban (2006) presented major and trace elements (Ba, Rb, Sr, Zr, Nb, Y, V, Cr, and Ni) compositional data for the studied samples from all stages of the Shirataka volcano. All samples of the Shirataka volcano plot within calc-alkaline series, and most of them belong to medium-K by the definition of Gill (1981), but the lavas and mafic inclusions from the earlier phase of the cone building stage belong to low-K series (Hirotani and Ban, 2006; Hirotani et al., 2009) (Fig. 2a). All of samples from stage 2 belong to medium-K series, but the pumice has slightly lower K2O contents than the scoria.

Fig. 2. K2O vs. SiO2 (a), in the left figure Cr vs. SiO2 (b) in the right figure variation diagrams of the eruptive products from the Shirataka volcano (modified after Hirotani et al., 2009). Data are from Hirotani & Ban (2006). The boundary lines between low-K and medium-K, and medium-K and high-K are from Gill (1981). Filled symbols are mafic inclusions and open ones are host rocks. \*Four domes consist of the Shiratakayama, Kitsunegoe, Nishikuromoriyama lava dome group + Higashikuromoriyama lava dome; \*\*Hagino, the Hagino block-and-ash flow deposit

of the stages 1 and 3 products, the groups are named to population I, II, and III by Hirotani et al., (2009). Here we refer "population" as "group". Group I includes An-rich plagioclase, Mgrich olivine and Mg-rich clinopyroxene. Most of these phenocrysts show compositional normal zoning. Group II includes An-poor plagioclase, Mg-poor orthopyroxene, Mg-poor clinopyroxene, hornblende, quartz, titanomagnetite, and ilmenite. Most of the phenocrysts in this population except for Fe-Ti oxides, show compositional reverse zoning and dissolution textures, such as sieved textures (Tsuchiyama, 1985) or embayed form. Group III is made of smaller phenocrysts of plagioclase, olivine, orthopyroxene, and clinopyroxene. These are interpreted to reflect relict cores from a mafic magma, relict cores from a silicic magma, and minerals growing in the hybrid magma caused by mixing the two magmas precipitated the

Hirotani & Ban (2006) presented major and trace elements (Ba, Rb, Sr, Zr, Nb, Y, V, Cr, and Ni) compositional data for the studied samples from all stages of the Shirataka volcano. All samples of the Shirataka volcano plot within calc-alkaline series, and most of them belong to medium-K by the definition of Gill (1981), but the lavas and mafic inclusions from the earlier phase of the cone building stage belong to low-K series (Hirotani and Ban, 2006; Hirotani et al., 2009) (Fig. 2a). All of samples from stage 2 belong to medium-K series, but

100

Cr, ppm

0

SiO2, wt% SiO2, wt%

Fig. 2. K2O vs. SiO2 (a), in the left figure Cr vs. SiO2 (b) in the right figure variation diagrams of the eruptive products from the Shirataka volcano (modified after Hirotani et al., 2009). Data are from Hirotani & Ban (2006). The boundary lines between low-K and medium-K, and medium-K and high-K are from Gill (1981). Filled symbols are mafic inclusions and open ones are host rocks. \*Four domes consist of the Shiratakayama, Kitsunegoe,

Nishikuromoriyama lava dome group + Higashikuromoriyama lava dome; \*\*Hagino, the

45 50 55 60 65 70

, stage 1: low-K , stage 1: medium-K

high-Cr type

, stage 2: pumice , stage 2: scoria

, stage 3: Hagino\*\* , stage 3: Four domes\*

20

low-Cr type

40

60

80

crystal cores of the groups I and II (Hirotani & Ban 2006).

the pumice has slightly lower K2O contents than the scoria.

(a) (b)

medium-K

low-K

(a) (b)

high-K

45 50 55 60 65 70

Hagino block-and-ash flow deposit

1

0

2

K2O, wt%

**3.3 Rock series and type of eruptive products of the Shirataka volcano**

Furthermore, Hirotani et al. (2009) defined the high- and low-Cr types for the rocks from the Shirataka volcano. Cr and Ni contents of the former type are >20 ppm and >5 ppm, 48-55 wt% of SiO2. Whereas, those of the latter type are ca. 15 ppm and ca. 3 ppm respectively for a given SiO2 range of 48-55 wt%. Both types are observed in the low and medium-K series. The host lavas having high-Cr type mafic inclusions show higher Cr and Ni contents than those with low-Cr compositions. Consequently these lavas are referred to as high- and low-Cr type lavas, respectively (Fig. 2b). All the samples from stage 2 are low-Cr type.

In addition, the high-Cr rocks always have lower Sr isotope ratio than the low-Cr type rocks for a given range of silica contents of a same geologic or petrologic unit in the cases of stages 1 and 3 (Hirotani et al., 2009). All the samples from stage 2 show similar Sr isotope ratio to the medium-K, low-Cr type rocks in the first stage or those in the early part of the third stage.

#### **4. Uranium, Thorium, and Hafnium contents of the eruptive products from the Shirataka volcano**

Hirotani & Ban (2006) and Hirotani et al. (2009) pointed out that the compositions of the samples from the explosive pyroclastic eruption stage show similar features as the medium-K series, low-Cr type rocks of the other stages (Fig. 2), except for the following features. The pumices have extremely higher Zr contents than any other samples from the Shirataka volcano (Fig. 3), and these show slightly lower K2O contents than the other medium-K lavas from the other stages (Fig. 2a). The higher Zr contents could be attributed to the presence of accessory minerals such as zircon and this would be confirmed by analysing Hf, Th, and U contents. (e.g., Rowe et al. 2007).

Fig. 3. Zr vs. SiO2 diagram of the eruptive products from the Shirataka volcano (modified after Hirotani & Ban, 2006). Data are from Hirotani & Ban (2006). Symbols are the same as in Fig. 2.

#### **4.1 Analytical method**

Concentrations of Hf, Th and U were determined by ICP-MS on a VG Platform instrument at the Geological Survey of Japan/AIST following the method of Ishizuka et al. (2010). About 100 mg of sample powder was dissolved in a HF-HNO3 mixture (5:1) using screw-top Teflon beakers on hotplate, followed by fusion with Na2CO3 (0.5 g) at 1,050℃. After evaporation to dryness, the residues were re-dissolved in 2% HNO3 prior to analysis. In and Re were used as internal standards, while JB2 and JB1a solutions diluted by the same method to the samples were used as external standards during ICP-MS measurements. Instrument calibration was performed using 5-6 solutions made from international rock standard materials. Reproducibility is generally better than ±6% (2 s.d.).

#### **4.2 The analytical results**

The representative data are listed in Table 2 and the data are plotted as the SiO2 variation diagrams in Fig. 4.

Fig. 4. Hf, Th, and U contents against SiO2 contents of the samples from the Shirataka volcano. Symbols are the same as in Fig. 2.

Re were used as internal standards, while JB2 and JB1a solutions diluted by the same method to the samples were used as external standards during ICP-MS measurements. Instrument calibration was performed using 5-6 solutions made from international rock

The representative data are listed in Table 2 and the data are plotted as the SiO2 variation

45 50 55 60 65 70

Fig. 4. Hf, Th, and U contents against SiO2 contents of the samples from the Shirataka

SiO2 wt%

standard materials. Reproducibility is generally better than ±6% (2 s.d.).

**4.2 The analytical results**

1.5

0

2

4

6

1

3

5

7

Th, ppm

Hf, ppm

U, ppm

1.0

0.5

0.0

volcano. Symbols are the same as in Fig. 2.

diagrams in Fig. 4.

Table 2. Hf, Th, and U contents of representative samples from the Shirataka volcano. Hagino\*, the Hagino block-and-ash flow deposit; Domes\*\*, the Shiratakayama, Kitsunegoe, Nishikuromoriyama lava dome group + Higashikuromoriyama lava dome

The Hf, Th, and U contents of rocks from stages 1 and 3 are plotted on the same compositional trends in the silica variation diagrams. These contents gradually increase with increasing SiO2 content (Fig. 4). The ranges of Hf, Th, and U contents of rocks from stages 1 and 3 are 2.1 to 4.4, 1.4 to 4.1, and 0.3 to 1.0, respectively (Table 2). The ranges of Hf, Th, and U contents of scoria samples from stage 2 fall in the ranges of those from stages 1 and 3. However, the pumice samples of stage 2 are plotted in the higher areas than the compositional trends defined by the samples from stages 1 and 3 in the Hf, Th, and U vs. SiO2 diagrams (Fig. 4). The contents of these elements in the pumice samples are extremely higher than the other ones (Table 2).

### **5. Discussion**

Previous petrologic studies (Hirotani & Ban, 2006; Hirotani et al., 2009) on samples from the first and third stages revealed that (1) all products were formed by magma mixing between mafic and felsic end-member magmas, (2) both end-member magmas changed their compositions with time, (3) the felsic end-member magma was formed by the melt extraction or re-melting of partially or fully crystallized mafic end-member magma of the same activity. The evidence for (1) includes the existence of mafic inclusions in all lava samples, disequilibrium phenocryst assemblages (Sakuyama, 1981) such as Mg-rich olivine and quartz, and wide compositional ranges in plagioclase phenocryst in all samples. The estimated end-member magmas for geologic or petrologic units of stages 1 and 3 are presented in Table 3. As can be observed in the table, the compositions of both the two endmembers change temporally.

Table 3. Temporal change of the mafic and felsic end-members compositions from stages 1 and 3 in the Shirataka volcano. Data are from Hirotani et al. (2009). \*Four lava domes, the Shiratakayama, Kitsunegoe, Nishikuromoriyama lava domes group + Higashikuromoriyama lava dome

Similarity of Sr and Nd isotope ratios within the same Cr-type host rocks and mafic inclusions of the same geologic unit shows the consanguinity of the mafic and felsic endmembers for each case. The trace element calculations showed that the felsic end-member

The Hf, Th, and U contents of rocks from stages 1 and 3 are plotted on the same compositional trends in the silica variation diagrams. These contents gradually increase with increasing SiO2 content (Fig. 4). The ranges of Hf, Th, and U contents of rocks from stages 1 and 3 are 2.1 to 4.4, 1.4 to 4.1, and 0.3 to 1.0, respectively (Table 2). The ranges of Hf, Th, and U contents of scoria samples from stage 2 fall in the ranges of those from stages 1 and 3. However, the pumice samples of stage 2 are plotted in the higher areas than the compositional trends defined by the samples from stages 1 and 3 in the Hf, Th, and U vs. SiO2 diagrams (Fig. 4). The contents of these elements in the pumice samples are extremely

Previous petrologic studies (Hirotani & Ban, 2006; Hirotani et al., 2009) on samples from the first and third stages revealed that (1) all products were formed by magma mixing between mafic and felsic end-member magmas, (2) both end-member magmas changed their compositions with time, (3) the felsic end-member magma was formed by the melt extraction or re-melting of partially or fully crystallized mafic end-member magma of the same activity. The evidence for (1) includes the existence of mafic inclusions in all lava samples, disequilibrium phenocryst assemblages (Sakuyama, 1981) such as Mg-rich olivine and quartz, and wide compositional ranges in plagioclase phenocryst in all samples. The estimated end-member magmas for geologic or petrologic units of stages 1 and 3 are presented in Table 3. As can be observed in the table, the compositions of both the two end-

; 52.00% SiO2 and 0.60% K2O (mafic end-member),

70.00% SiO2 and 2.20% K2O (felsic end-member) ; 49.50% SiO2 and 0.64% K2O (mafic end-member), 66.50% SiO2 and 1.75% K2O (felsic end-member) ; 50.50% SiO2 and 0.80% K2O (mafic end-member), 66.00% SiO2 and 1.65% K2O (felsic end-member)

; 51.00% SiO2 and 0.50% K2O (mafic end-member), 66.50% SiO2 and 1.85% K2O (felsic end-member) ; 52.00% SiO2 and 0.85% K2O (mafic end-member), 67.50% SiO2 and 1.70% K2O (felsic end-member)

; 48.00% SiO2 and 0.20% K2O (mafic end-member), 64.50% SiO2 and 0.85% K2O (felsic end-member) ; 49.00% SiO2 and 0.21% K2O (mafic end-member), 64.00% SiO2 and 0.75% K2O (felsic end-member)

higher than the other ones (Table 2).

members change temporally.

*stage 3: lava-dome & block-and-ash flow forming stage*

; medium-K, high-Cr type

; medium-K, low-Cr type

; medium-K, high-Cr type

; medium-K, low-Cr type

; low-K, low-Cr type

Shiratakayama, Kitsunegoe, Nishikuromoriyama lava domes group +

Table 3. Temporal change of the mafic and felsic end-members compositions from stages 1 and 3 in the Shirataka volcano. Data are from Hirotani et al. (2009). \*Four lava domes, the

Similarity of Sr and Nd isotope ratios within the same Cr-type host rocks and mafic inclusions of the same geologic unit shows the consanguinity of the mafic and felsic endmembers for each case.  Discussion**

cannot be explained by the fractional crystallization process but by the remelting process of the solidified mafic end-member leaving behind gabbroic residue (Hirotani et al., 2009). The genesis of felsic magmas have been explained by similar process in many cases (e.g. Feeley et al, 1998; Hansen et al., 2002, Smith et al., 2006; Vogel et al, 2006; Ban et al., 2007).

The pumice and scoria in the stage 2 possess evidence of mama mixing, such as dissolution textures in plagioclase phenocrysts, disequilibrium phenocryst assemblages, and wide compositional variation in plagioclase phenocrysts. Thus the pumice and scoria could also be formed by magma mixing (Hirotani & Ban, 2006). The compositions of the end-member magmas for the products of stages 1 and 3 have been well discussed (Hirotani et al., 2009), because the compositional ranges of the samples from stages 1 and 3 are wide and thus it was possible to determine the compositions of end-member magmas using the magma mixing lines in the compositional variation diagrams. However, pumice and scoria samples from the stage 2 display very narrow compositional ranges. It was impossible to determine the end-member compositions in these cases using the same method applied for stages 1 and 3 samples. Thus the compositions of the endmember magmas still remain in unknown. In the following sections, we will define these end-member compositions using another method.

#### **5.1 Estimation of the compositions of the end-member magmas for the pumice and scoria in the explosive pyroclastic eruption stage**

The juvenile fragments of the explosive pyroclastic eruption stage are mainly pumice with minor amount of scoria, and these also possess petrologic features of mixing of two endmembers (Hirotani & Ban, 2006). However there is very little clue to estimate the endmember compositions because of the very narrow compositional ranges of the pumice and the scoria. Thus, here we assume that the end-member compositions of the pumice and scoria are similar to one of the estimated end-member compositions for stages 1 and 3 eruptive products. The chemical compositions of the scoria fall in the ranges of the products of the early part of the third stage (Figs. 2, 3 & 4). Thus the end-member compositions of the scoria would be the same as in the early part of the third stage. On the other hand, the pumice shows higher Zr, Hf, Th, and U contents than any other samples in the Shirataka volcano.

In Fig. 5, we compare the compositions of glass inclusions in pyroxene phenocrysts of pumice to those from stages 1 and 3. Representative compositional data of the glass inclusions are presented in Hirotani & Ban (2006). The compositional range of the glass compositions from pumice is similar to those from the medium-K products of stage 1 in Fig. 5. Therefore, it is reasonable to consider that the felsic end-member compositions of the pumice are also similar to the above mentioned ones. In this case, the mafic end-member compositions should plot on the silica poorer extensions of the mixing lines tying the felsic end-member and pumice compositions (Fig. 6). In this figure, the mafic end-member of the low-K samples of stage 1 is on the extension line (orange colored dotted line in Fig. 6). Here, we assume that the compositions of the mafic end-member are the same as those of low-K products of stage 1. In the next step, we determined the felsic end-member compositions of the pumice, using the tie lines of the assumed mafic end-member and pumice compositions on the SiO2 variation diagrams, assuming the silica content of the felsic end-member is the same as that of the medium-K products of stage 1. The obtained chemical compositions of the felsic end-member are similar to those of the medium-K products of stage 1, except for higher Zr, Hf, Th, and U contents.

Fig. 5. K2O-SiO2 diagram showing the compositions of glass inclusions in pyroxene phenocrysts from each stage in the Shirataka volcano. \*Four domes, the Shiratakayama, Kitsunegoe, Nishikuromoriyama lava dome group + Higashikuromoriyama lava dome; \*\*Hagino, the Hagino block-and-ash flow deposit

Fig. 6. K2O-SiO2 diagram showing the end-member compositions for the pumice and scoria from stage 2 in the Shirataka volcano. The boundary lines of low-K and medium-K, and medium-K and high-K are from Gill (1981). The dotted line shows the mixing line for the pumice. \*Four domes, the Shiratakayama, Kitsunegoe, Nishikuromoriyama lava dome group + Higashikuromoriyama lava dome; \*\*Hagino, the Hagino block-and-ash flow deposit

stage 3, Four domes

stage 2, scoria Hagino

stage 2, pumice stage 1, medium-K stage 3,

glasses in pyroxene phenocrysts

stage 1, low-K

, stage 1: medium-K

.

, stage 2: pumice , stage 2: scoria

felsic end-members

, stage 3: Hagino\*\* , stage 3: Four domes\*

SiO2, wt% 50 60 70 80

45 50 55 60 65 70 , stage 1: low-K

, mafic & felsic end-members of stage 1: low-K , mafic & felsic end-members of stage 1: medium-K , mafic & felsic end-members of stage 3: Hagino\*\* , mafic & felsic end-members of stage 3: Four domes\*

Fig. 6. K2O-SiO2 diagram showing the end-member compositions for the pumice and scoria from stage 2 in the Shirataka volcano. The boundary lines of low-K and medium-K, and medium-K and high-K are from Gill (1981). The dotted line shows the mixing line for the pumice. \*Four domes, the Shiratakayama, Kitsunegoe, Nishikuromoriyama lava dome group + Higashikuromoriyama lava dome; \*\*Hagino, the Hagino block-and-ash flow deposit

SiO2, wt%

Fig. 5. K2O-SiO2 diagram showing the compositions of glass inclusions in pyroxene phenocrysts from each stage in the Shirataka volcano. \*Four domes, the Shiratakayama, Kitsunegoe, Nishikuromoriyama lava dome group + Higashikuromoriyama lava dome;

pumice

5

K2O, wt%

\*\*Hagino, the Hagino block-and-ash flow deposit

K2O, wt%

4

3

2

scoria

whole rocks

stage 1, low-K stage 1, medium-K stage 2, pumice stage 2, scoria stage 3, Hagino\*\* stage 3, Four domes\*

1

0

0

&

& &

&

1

end-members

2

mafic

3

#### **5.2 Melting of accessory zircon with thorite inclusions in producing the felsic endmember magmas for the pumice in stage 3 in the Shirataka volcano**

During dormancy between stages 1 and 2, the magma chamber filled with the felsic endmember of the medium-K samples in the first stage would become a mush chamber (e.g., Huber et al., 2010) by the crystallization. In the second stage, new magma injection from deeper levels would re-activate the mushy chamber. This re-activated magma would become to felsic end-member magma for the pumice in the second stage. Such re-activation process of the old magma chambers is thought to be common in producing felsic magmas (e.g., Tamura & Tatsumi, 2002; Girard & Stix, 2009; McCurry & Rodgers, 2009). The higher Zr, Hf, Th, and U contents of the felsic end-member magma can be explained by the melting of the accessory zircon minerals during the re-activation (Wolff et al., 2006). The estimated Zr content of the felsic magma can be obtained by melting of zircon-bearing rock (0.04% of the melt should be from zircon melting). Accordingly, the Hf, Th, and U contents in the zircon are calculated to be ca. 11,500, 10,000, and 1,100 ppm, respectively (Fig. 7). The Hf, Th, and U contents in zircon minerals were reported by many researches (e.g., Rubatto & Hermann, 2007). The calculated Hf content of ca. 11,500 ppm is in the range of reported Hf contents in zircon. On the other hand, the calculated Th and U contents of ca. 10,000 and 1,100 ppm are exceed the ranges of Th and U contents in zircon in equilibrium with the felsic end-member (e.g., Rubatto & Hermann, 2007). Recently, Paquette & Mergoil-Daniel (2009) reported that zircon crystals sometimes include tiny thorite inclusions. The Th and U contents in thorite are much higher than in zircon. Thus, the high Th and U contents can be ascribed to the melting of thorite-bearing components during the re-activation.

Fig. 7. Plots of Zr, Hf, Th, and U vs. SiO2 showing the estimated felsic end-member magma compositions for the studied pumice. Symbols, other than the end-members are the same as in Fig. 2.

#### **5.3 The magma feeding system for the explosive pyroclastic eruption stage 5.3.1 Formation of mixed magma for the pumice**

As discussed above, the felsic end-member of medium-K samples of the first stage would be re-activated at the second stage and became the felsic end-member magma for the pumice. The end-member magma chamber would be stored in shallower part of the crust. The reactivation would be induced as a result of the ascent of the mafic end-member of the low-K products of the first stage. This magma would be still active also in the second stage (Fig. 8). The mafic magma from deeper levels is characterized by high-Cr type. This type mafic magma differentiate into low-Cr type through the AFC process as described for the formation of the low-Cr mafic magmas of stages 1 and 3 by Hirotani et al. (2009). The low-Cr type mafic magma would ascent further. When the mafic magma reached shallow crustal level where the remnant of stage 1 medium-K felsic magma chamber locates, the mafic magma would re-activate this felsic chamber by its heat. The higher Zr, Hf, U, and Th contents of the felsic end-member magma can be explained by the melting of the accessory zircon minerals with thorite inclusions during the re-activation. The re-activated felsic magma would mix with the low-Cr type mafic magma in the chamber. The schematic representation of the magma feeding system is presented in Fig. 8.

Fig. 8. Schematic representation of the magma feeding system of the stage 2 in the Shirataka volcano.

#### **5.3.2 Formation of mixed magma for the scoria**

The third stage mafic end-member magma would antecedently appear as the mafic endmember for scoria at the second stage. The felsic end-member magma for the scoria was generated by the melt extraction or re-melting process, and subsequently mixed with the mafic end-member in a crustal magma chamber. The Mg# and the estimated precipitated temperature of two-pyroxenes in the scoria are higher than those in the other felsic end-members (Hirotani & Ban, 2006). Furthermore the size of the pyroxene phenocrysts is smaller than those in the other felsic end-members, but larger than the groundmass minerals. These features indicate that the pyroxenes of the scoria are grouped to group III (population III of Hirotani et al. (2009)), and precipitated from the mixed magma for the scoria in the chamber, suggesting that the mixed magma was stored in the crust for a while.

#### **5.3.3 Mingling between two mixed magmas**

Consequently, two kinds of mixed magma were stored at shallow crust level. One is for the pumice and the other is for the scoria. Further ascent of the mafic end-member triggered the ascent of the mixed magma of the scoria. This mixed magma would tap the shallow chamber filled with the other mixed magma of the pumice. Finally, these two mixed magmas erupted synchronously, which made the eruption explosive. In addition, rarely observed banded pumice would be formed probably during the ascent through the conduit.

#### **6. Conclusions**

68 Petrology – New Perspectives and Applications

As discussed above, the felsic end-member of medium-K samples of the first stage would be re-activated at the second stage and became the felsic end-member magma for the pumice. The end-member magma chamber would be stored in shallower part of the crust. The reactivation would be induced as a result of the ascent of the mafic end-member of the low-K products of the first stage. This magma would be still active also in the second stage (Fig. 8). The mafic magma from deeper levels is characterized by high-Cr type. This type mafic magma differentiate into low-Cr type through the AFC process as described for the formation of the low-Cr mafic magmas of stages 1 and 3 by Hirotani et al. (2009). The low-Cr type mafic magma would ascent further. When the mafic magma reached shallow crustal level where the remnant of stage 1 medium-K felsic magma chamber locates, the mafic magma would re-activate this felsic chamber by its heat. The higher Zr, Hf, U, and Th contents of the felsic end-member magma can be explained by the melting of the accessory zircon minerals with thorite inclusions during the re-activation. The re-activated felsic magma would mix with the low-Cr type mafic magma in the chamber. The schematic

AFC; formation of low-Cr mafic magma

medium-K mafic magma (high-Cr),

medium-K felsic magma (low-Cr), similar to the early part of the stage 3

formation of mixed magma for the

mingling between the two mixed magmas; magmas for the pumice

mixing between medium-K mafic

scoria in stage 2;

and the scoria

and felsic magmas

medium-K mafic magma (low-Cr), similar to the early part of the stage 3

**5.3 The magma feeding system for the explosive pyroclastic eruption stage**

**5.3.1 Formation of mixed magma for the pumice**

representation of the magma feeding system is presented in Fig. 8.

precursory activity of stage 3 low-K mafic magma (high-Cr),

Fig. 8. Schematic representation of the magma feeding system of the stage 2 in the Shirataka

low-K mafic magma (low-Cr), similar to that formed in stage 1

similar to that of stage 1

volcano.

solidified magmatic bodies in stage 1

remnant of medium-K felsic magma (low-Cr) formed in stage 1

formation of mixed magma for the pumice in stage 2; mixing between reactivated (including melting of accessory minerals)

medium-K felsic magma and the low-K mafic magma

Detailed petrologic study on the explosive pyroclastic eruption stage of the Shirataka volcano, NE Japan, produced the following results. The eruptive products are mostly pumice with minor amount of scoria. Banded pumice is rarely observed.

#### **7. Acknowledgment**

We greatly appreciate anonymous reviewer for many constructive review comments on the early version of the manuscript. The manuscript was greatly improved by the comments. We are grateful for Dr. S. Nakano for supporting this study. We also thank K. Yamanobe for assisting with the ICP-MS measurements.

#### **8. References**

We greatly appreciate anonymous reviewer for many constructive review comments on the early version of the manuscript. The manuscript was greatly improved by the comments. We are grateful for Dr. S. Nakano for supporting this study. We also thank K. Yamanobe for

Bacon, C. R. (1986) Magmatic inclusions in silicic and intermediate volcanic rocks. *J. Geophys.*

Ban, M., Hirotani, S., Wako, A., Suga, T., Iai, Y., Kagashima, S., Shuto, K. & Kagami, H.

Eichelberger, J. C. (1980) Vesiculation of mafic magma during replenishment of silicic

Feeley, T. C., Dungan, M. A. & Frey, F. A. (1998) Geochemical constraints on the origin of

Girard, G. & Stix, J. (2009) Magma recharge and crystal mush rejuvenation associated with

Hansen, J., Skjerlie, K. P., Pederson, R. B. & Rosa, J. D. L. (2002) Crustal melting in the lower

Hirotani, S. & Ban, M. (2006) Origin of silicic magma and magma feeding system of Shirataka volcano, NE Japan. *J. Volcanol. Geotherm. Res.,* 156, 229-251. Hirotani, S., Ban, M. & Nakagawa, M. (2009) Petrogenesis of mafic and associated silicic

Huber, C., Bachmann, O. & Dufek, J. (2010) The limitations of melting on the reactivation of

Ishii, M. & Saito, K. (1997) A K-Ar age study on Shirataka volcano, Yamagata Prefecture.

Ishizuka, O., Yuasa, M., Tamura, Y., Shukuno, H., Stern, R. J., Naka, J., Joshima, M. & Taylor,

Kawano, Y., Yagi, K. & Aoki, K. (1961) Petrography and petrochemistry of the volcanic

Mimura, K. & Kanno, K. (2000) Stratigraphy and history of Shirataka volcano, NE Japan.

oceanic arc rift propagation. *Earth Planet. Sci. Lett.,* 294, 111–122.

Gill, J. B. (1981) *Orogenic andesites and plate tectonics,* Springer-Verlag, Berlin. pp 385.

Norwegian Caledonides. *Contrib. Mineral. Petrol.,* 143, 316-335.

(2007). Origin of silicic magmas in a large-caldera-related stratovolcano in the central part of NE Japan - Petrogenesis of the Takamatsu volcano -. *J. Volcanol.*

mafic and silicic magmas at Cordon El Guadal, Tatara-San Pedro Complex, central

early post-collapse upper basin member rhyolites, Yellowstone Caldera, Wyoming.

parts of island arcs: an example from the Bremanger Granitoid Complex, west

end-member magmas for calc-alkaline mixed rocks in the Shirataka volcano, NE

R. N. (2010) Migrating shoshonitic magmatism tracks Izu–Bonin–Mariana intra-

rocks of Quaternary volcanoes of northeastern Japan. *Sci. Rep. Tohoku Univ. Ser*

**7. Acknowledgment**

**8. References**

assisting with the ICP-MS measurements.

*Geotherm. Res.,* 167, 100-118.

*J. Petrol.,* 50, 2095-2125.

*III,* 7, 1-46.

magma reservoirs. *Nature*, 288, 446-450.

Chile. *Contrib. Mineral. Petrol.,* 131, 393-411.

Japan. *Contrib. Mineral. Petrol.,* 157, 709-734.

*Bull. Yamagata Univ., Nat. Sci.*, 14, 99-108.

*Bull. Volcanol. Soc. Jpn*., 45, 13-23.

silicic mushes. *J. Volcanol. Geotherm. Res.,* 195, 97-105.

*Res,* 91, 6091-6112.

Yagi, H., Soda, T, Inokuchi, T., Haraguchi, T. & Ban, M. (2005) Catastrophic collapse of Mt. Zao and Mt. Shirataka and their chronological timing. *The Quaternay Res.,* 44, 263- 272.

### **Late to Post-Orogenic Brasiliano-Pan-African Volcano-Sedimentary Basins in the Dom Feliciano Belt, Southernmost Brazil**

Delia del Pilar Montecinos de Almeida1, Farid Chemale Jr.2 and Adriane Machado3 *1Universidade Federal do Pampa (UNIPAMPA) 2Universidade de Brasília (UnB) 3Centro de Geofísica da Universidade de Coimbra (CGUC) 1,2Brazil 3Portugal*

#### **1. Introduction**

72 Petrology – New Perspectives and Applications

Yagi, H., Soda, T, Inokuchi, T., Haraguchi, T. & Ban, M. (2005) Catastrophic collapse of Mt.

272.

Zao and Mt. Shirataka and their chronological timing. *The Quaternay Res.,* 44, 263-

The Neoproterozoic/Early Paleozoic transition in Southern Brazil was marked by the last phases of the Brasiliano/Pan-African Cycle (900-540 Ma, Chemale Jr., 2000). These include the development of a trough series related to the different Brasiliano Orogenesis phases, which occupied one preferential subsiding locus called Camaquã Basin (Paim et al., 2000). This basin is located in the Sul-Riograndense Shield (Rio Grande do Sul State, Southernmost Brazil, Fig. 2). The Camaquã Basin is a northeast-southwest elongated basin positioned along the Dom Feliciano Belt limits (to the east) and the Rio de la Plata Craton (to the west). The basin was filled with a thick (ca. 11 km) sedimentary rocks package deposited in marine (at the base) and continental (at the top) environments, and a volcanic rocks series interbedded with the sediments. At least three volcanic events were recognized and related to the Hilário and Acampamento Velho Formations, and Rodeio Velho Member, from the base to the top (Fig. 1).

Paim et al. (2000) divided the Camaquã Basin into four sub-basins named as Taquarembó-Ramada, Santa Bárbara, Guaritas and Piquiri-Arroio Boici, from West to East (Fig.1). Each subbasin showed a partial independent evolution, and consequently, the stratigraphic framework shows some differences in a regional scale. Santa Bárbara and Guaritas sub-basins show the most complete stratigraphic section. Paim et al. (2000) defined five major depositional cycles from tectonic origin separated by unconformities. Every cycle corresponds to a unit ranking as group or allogroup. From the base to the top, there are the following groups: Maricá, Bom Jardim, Cerro do Bugio, Santa Bárbara and Guaritas (Table 1).

This chapter aims to make the characterization of the two last volcanic events associated to the filling of Camaquã Basin, named as Acampamento Velho Formation and Rodeio Velho Member in Rio Grande do Sul, Brazil. The book chapter will be composed of an introduction. After the introduction, it will be presented the Camaquã Basin stratigraphy and evolution compilation, aiming to introduce the reader to the volcanic events characterization. We also present a compilation of geochronological data from different units that fill the Camaquã Basin as well as the U-Pb ages obtained by us for Rodeio Velho, Acampamento Velho, Hilário volcanic events and Maricá Formation (in sandstone). The mineralogy, petrography, geochemistry (major, trace and rare-earth elements) and isotope data will be presented to the Acampamento Velho Formation and Rodeio Velho Member.

#### **2. Camaquã Basin stratigraphy and evolution**

#### **2.1 Camaquã Basin tectonic evolution**

Camaquã Basin is related to a system of late to post tectonic basins associated with the Pan African Brasiliano Orogeny. In the late stages of the Brasiliano Orogenic phase (700-500 Ma); the depressions formed received sediments from the erosion of mountain areas still undergoing active uplifting (Almeida, 1969). Almeida (1969) and Almeida et al. (1976, 1981) subdivided the deposits of these basins in lower and upper molasse. The lower molasse sequence is folded and associated with volcanism, whereas the upper molasse is horizontal or tilted, associated with magmatism. Almeida (1969) attributed the configuration of these molassic basins to reactivation processes of the deep faults. Wernick et al. (1978) related their genesis to strike-slip movements. Other authors suggested that these basins were formed in depressions generated by synforms developed during the Brasiliano Orogenic event (Fragoso Cesar et al., 1982). According to Fragoso Cesar et al. (2000), Camaquã Basin is a Neoproterozoic – Early Cambrian extensional, post orogenic and pre-cratonic rift system. Paim et al. (2000) interpreted this basin as a depositional locus in which different basins succeeded one another, each of them with their own lithological record and distinctive subsidence mechanisms. These authors admitted a connection between the evolution of the Camaquã Basin and the final tectonic stages of the Pan-African Brasiliano Orogeny in Southern Brazil. They also mentioned the continentalization that took place in later stages, even as the clear tectonic control in the stratigraphic arrangement of the allogroups, thus suggesting, a residual tendency of crustal uplifting during the final stages of the Pan-African-Brasiliano Orogeny.

According to Chemale Jr. (2000), the subsidence of the Neoproterozoic Rio de la Plata Plate beneath the Kalahari Plate led to the formation of oceanic crust around 700 Ma. Rifting of the Kalahari Plate resulted in the formation of the Encantadas Microcontinent (or microplate), with a converging margin to the west and a diverging one to the east. The proto-ocean Adamastor was formed between 700 and 650 Ma. Between 650 and 540 Ma, it started to subside to the west, beneath the Encantadas Microplate. The calc-alkaline orogenic magmatism occurred between 630-610 Ma, and the later to post orogenic around 600-540 Ma (Phase III). The closing of this proto-ocean was achieved initially through transgression, changing to transcurrence in the final stages. The last deformational pulses related to the Brasiliano Event (Dom Feliciano) took place around 540 Ma. The deposition in the Camaquã Basin formed in the retroarc region, initially preceded through the accumulation of sedimentary sequences and associated to the magmatism at 620 Ma. There is evidence for progressive continentalization and alkaline signatures toward the top of the rock record about 500 Ma, when the final accretion of west Gondwana took place. Thus, these events marked the beginning of the stabilization of Gondwana Supercontinent and the initiation of erosional processes in the Brasiliano mountain ranges. Subsidence of the Paraná Basin occurred from 500 Ma onwards, with the accumulation of basal sediments in a rift

characterization. We also present a compilation of geochronological data from different units that fill the Camaquã Basin as well as the U-Pb ages obtained by us for Rodeio Velho, Acampamento Velho, Hilário volcanic events and Maricá Formation (in sandstone). The mineralogy, petrography, geochemistry (major, trace and rare-earth elements) and isotope data will be presented to the Acampamento Velho Formation and Rodeio Velho Member.

Camaquã Basin is related to a system of late to post tectonic basins associated with the Pan African Brasiliano Orogeny. In the late stages of the Brasiliano Orogenic phase (700-500 Ma); the depressions formed received sediments from the erosion of mountain areas still undergoing active uplifting (Almeida, 1969). Almeida (1969) and Almeida et al. (1976, 1981) subdivided the deposits of these basins in lower and upper molasse. The lower molasse sequence is folded and associated with volcanism, whereas the upper molasse is horizontal or tilted, associated with magmatism. Almeida (1969) attributed the configuration of these molassic basins to reactivation processes of the deep faults. Wernick et al. (1978) related their genesis to strike-slip movements. Other authors suggested that these basins were formed in depressions generated by synforms developed during the Brasiliano Orogenic event (Fragoso Cesar et al., 1982). According to Fragoso Cesar et al. (2000), Camaquã Basin is a Neoproterozoic – Early Cambrian extensional, post orogenic and pre-cratonic rift system. Paim et al. (2000) interpreted this basin as a depositional locus in which different basins succeeded one another, each of them with their own lithological record and distinctive subsidence mechanisms. These authors admitted a connection between the evolution of the Camaquã Basin and the final tectonic stages of the Pan-African Brasiliano Orogeny in Southern Brazil. They also mentioned the continentalization that took place in later stages, even as the clear tectonic control in the stratigraphic arrangement of the allogroups, thus suggesting, a residual tendency of crustal uplifting during the final stages of the Pan-

According to Chemale Jr. (2000), the subsidence of the Neoproterozoic Rio de la Plata Plate beneath the Kalahari Plate led to the formation of oceanic crust around 700 Ma. Rifting of the Kalahari Plate resulted in the formation of the Encantadas Microcontinent (or microplate), with a converging margin to the west and a diverging one to the east. The proto-ocean Adamastor was formed between 700 and 650 Ma. Between 650 and 540 Ma, it started to subside to the west, beneath the Encantadas Microplate. The calc-alkaline orogenic magmatism occurred between 630-610 Ma, and the later to post orogenic around 600-540 Ma (Phase III). The closing of this proto-ocean was achieved initially through transgression, changing to transcurrence in the final stages. The last deformational pulses related to the Brasiliano Event (Dom Feliciano) took place around 540 Ma. The deposition in the Camaquã Basin formed in the retroarc region, initially preceded through the accumulation of sedimentary sequences and associated to the magmatism at 620 Ma. There is evidence for progressive continentalization and alkaline signatures toward the top of the rock record about 500 Ma, when the final accretion of west Gondwana took place. Thus, these events marked the beginning of the stabilization of Gondwana Supercontinent and the initiation of erosional processes in the Brasiliano mountain ranges. Subsidence of the Paraná Basin occurred from 500 Ma onwards, with the accumulation of basal sediments in a rift

**2. Camaquã Basin stratigraphy and evolution**

**2.1 Camaquã Basin tectonic evolution**

African-Brasiliano Orogeny.

environment. In the Sul-Riograndense Shield, this is probably represented by the Guaritas Group (Fig. 2).

Fig. 1. Geological map showing the main Camaquã Basin stratigraphic units (modified from Paim et al, 2000) and the localization of the main regions.

According to Basei et al. (2000), the magmatic, metamorphic structural and geotectonic features in Southeastern South America recorded the superposition of two Neoproterozoic-Eopaleozoic orogenies: Brasiliano (700-620 Ma) and Rio Doce (620-530 Ma) orogenies. The end of the extensional regime related to the Brasiliano cycle is marked by late magmatism (600±10 Ma), conspicuous in Southern Brazil. The foreland basins, with sedimentation around 560±20 Ma and deformation around 530 Ma, represent the main volcanicsedimentary record of Rio Doce Orogeny in Southern Brazil. The Camaquã Basin is considered to represent this type of basin. Differently from the current thoughts, Fragoso-Cesar et al. (2001) suggested an intraplate evolution to this basin, detached from the evolution of the Brasiliano Cycle. These authors related the development of the basin to the reactivation of the Brasiliano structures formed in the Neoproterozoic III to Eocambrian. They stated that the knowledge about the tectonic events and deformation of the sedimentary and volcanic rocks is still incipient, and there are no radiometric ages to position the deformation phases in time. They also ratify the connection between the Guaritas sedimentation and volcanics rocks from Rodeio Velho, and Paraná Basin. These units would be related to the development of SW-NE rifts during the initial stages of the evolution from Paraná Basin.

Notes: Includes, at the base, the Martins Andesite Member. At the base, contains Rodeio Velho Member. Composed by lower (Hilário) and upper (Acampamento Velho) members. Comprise the Vargas and Mangueirão members. Made up by Hilário and Arroio dos Nobres members. Constituted by the Rodeio Velho, at the base, Lanceiros, Varzinha and Guaritas members. Composed by Passo dos Bravos at the base, Vargas and Mangueirão members.

Table 1. Stratigraphic concepts for Camaquã Basin.

Silva et al. (2005) as cited in Borba et al. (2008) focusing on the tectonic evolution of Brasiliano events in the Mantiqueira Province (South and Central Brazil and Uruguay, Fig. 3) revealed three main orogenic systems: (a) Brasiliano I, with juvenile island-arc accretion and collisional climaxes at ca. 790 Ma (Embu Domain, Southeastern Brazil) and 730-700Ma (São Gabriel Orogeny, Southern Brazil); (b) Brasiliano II, with dominance of crustal

evolution of the Brasiliano Cycle. These authors related the development of the basin to the reactivation of the Brasiliano structures formed in the Neoproterozoic III to Eocambrian. They stated that the knowledge about the tectonic events and deformation of the sedimentary and volcanic rocks is still incipient, and there are no radiometric ages to position the deformation phases in time. They also ratify the connection between the Guaritas sedimentation and volcanics rocks from Rodeio Velho, and Paraná Basin. These units would be related to the development of SW-NE rifts during the initial stages of the

evolution from Paraná Basin.

Table 1. Stratigraphic concepts for Camaquã Basin.

Silva et al. (2005) as cited in Borba et al. (2008) focusing on the tectonic evolution of Brasiliano events in the Mantiqueira Province (South and Central Brazil and Uruguay, Fig. 3) revealed three main orogenic systems: (a) Brasiliano I, with juvenile island-arc accretion and collisional climaxes at ca. 790 Ma (Embu Domain, Southeastern Brazil) and 730-700Ma (São Gabriel Orogeny, Southern Brazil); (b) Brasiliano II, with dominance of crustal

and Mangueirão members.

Notes: Includes, at the base, the Martins Andesite Member. At the base, contains Rodeio Velho Member. Composed by lower (Hilário) and upper (Acampamento Velho) members. Comprise the Vargas and Mangueirão members. Made up by Hilário and Arroio dos Nobres members. Constituted by the Rodeio Velho, at the base, Lanceiros, Varzinha and Guaritas members. Composed by Passo dos Bravos at the base, Vargas

Fig. 2. Evolution model of the Sul-Riograndense Shield at the end of Neoproterozoic (Chemale Jr., 2000).

Fig. 3. Schematic geological map of the Southern portion of the Mantiqueira Province (b) located in Eastern South America (a); key to the terrains/blocks/belts/orogenies names: F= Florida; NP= Nico Perez; CD = Cuchilla Dionísio; Val= Valentines; R= Rivera; Taq=Taquarembó; SG= São Gabriel; SBV= Santana da BoaVista; P= Pelotas; FP= Florianópolis; BSQ = Brusque; LA= Luiz Alves; PI = Piên; PA= Paranapiacaba; unmetamorphosed volcanic-sedimentary successions are named as Camaquã (CQ ) and Itajaí (IT) (Silva et al., 2005, in Borba et al., 2008).

recycling and collisional climaxes at 640-620Ma (Dom Feliciano Orogeny, Southern Brazil and Uruguay) and 600Ma (Paranapiacaba and Rio Piên Orogenies, Southeastern Brazil); and (c) Brasiliano III, also characterized by crustal recycling and collisional climaxes between 590-560Ma (Araçuai Orogeny) and 520-500Ma (Búzios Orogeny), recorded mainly in Eastern Brazil. Foreland basins associated with the Brasiliano II orogenic system comprise portions of volcanic-sedimentary (metamorphosed or unmetamorphosed) successions along the 3000 km of length from Mantiqueira Province.

#### **2.2 Stratigraphy**

Thick accumulations of sedimentary and volcanogenic rocks are present in the Camaquã and Itajaí basins of Southern Brazil, considered to be the "molassic fore deep of the Dom Feliciano Belt" by Fragoso-Cesar et al. (1984). These sedimentary successions evolved from marine to alluvial-aeolian continental strata and have been considered to be foreland basins along with their African counterparts (Gresse et al., 1996). Recent research considers the Camaquã Basin to be a series of the different basins (foreland, strike-slip and rift) with a shared locus of subsidence (Paim et al., 2000). In terms of stratigraphy, five major unconformity bounded units (Allogroups) have been defined by Paim et al. (2000): the Maricá, Bom Jardim, Cerro do Bugio, Santa Bárbara and Guaritas Allogroups (Table 1). They reflect the progressive continentalization of the Camaquã Basin. These units can be interpreted as the evolution of temporally and spatially restricted basins, evolved from a foreland retroarc into a rift (or hemigraben) tectonic setting. In this more recent interpretation, only the Maricá Formation was deposited in a foreland basin.

The Maricá Formation was defined by Leinz et al. (1941) as a basal unit of Camaquã Basin. This Formation is a ±2500 m-thick sedimentary package with coastal deposits, including fluvial to shallow marine sediments with scarce volcanic contribution (Leites et al., 1990). The nature of the basal contact of this unit with the Neoproterozoic juvenile terrain is difficult to be defined, either an unconformity or a thrust plane, due to low relief, thick soil profiles and vegetation, and widespread younger intrusive bodies. According to Borba (2004), the Maricá formation can be divided into three major packages: the lower, intermediate and upper successions. The lower succession comprises conglomeratic and coarse-grained sandstones with very well rounded granite–gneiss pebbles and cobbles either marking the stratification planes or dispersed through the beds. The intermediate succession is made up of indurated siltstones and shales, either massive or showing planeparallel lamination, wave and climbing ripples, and lenticular lamination, as well as hummocky sandstones, suggesting a marine origin. The upper succession is very similar to the lower one, with the sandstones featuring large-scale troughcross and low-angle stratifications. Well-rounded granite–gneiss and rare volcanic, rhyolitic pebbles and cobbles are also present within the sandstone beds. The lower and upper successions correspond to braided fluvial deposits, while the intermediate package is related to a shallow marine depositional environment. The strata of the Maricá Formation were strongly affected by subsequent events of volcanism, intrusion, sedimentation and uplift. Volcanic rocks of the Bom Jardim Group (andesite Hilário) and the Acampamento Velho Formation (rhyolitic flow, rhyolitic pyroclastic rocks and basic flow) caused metamorphism contact in the Maricá Formation. The unit is unconformably overlain by the alluvial, fluvial and eolian deposits of the Santa Bárbara and Guaritas formations (Borba & Mizusaki, 2003). A syn-depositional volcanic event (Santos et al., 1978) has been recognized in the rock record of the Maricá Formation, evidenced by the occurrence of

recycling and collisional climaxes at 640-620Ma (Dom Feliciano Orogeny, Southern Brazil and Uruguay) and 600Ma (Paranapiacaba and Rio Piên Orogenies, Southeastern Brazil); and (c) Brasiliano III, also characterized by crustal recycling and collisional climaxes between 590-560Ma (Araçuai Orogeny) and 520-500Ma (Búzios Orogeny), recorded mainly in Eastern Brazil. Foreland basins associated with the Brasiliano II orogenic system comprise portions of volcanic-sedimentary (metamorphosed or unmetamorphosed) successions along the 3000

Thick accumulations of sedimentary and volcanogenic rocks are present in the Camaquã and Itajaí basins of Southern Brazil, considered to be the "molassic fore deep of the Dom Feliciano Belt" by Fragoso-Cesar et al. (1984). These sedimentary successions evolved from marine to alluvial-aeolian continental strata and have been considered to be foreland basins along with their African counterparts (Gresse et al., 1996). Recent research considers the Camaquã Basin to be a series of the different basins (foreland, strike-slip and rift) with a shared locus of subsidence (Paim et al., 2000). In terms of stratigraphy, five major unconformity bounded units (Allogroups) have been defined by Paim et al. (2000): the Maricá, Bom Jardim, Cerro do Bugio, Santa Bárbara and Guaritas Allogroups (Table 1). They reflect the progressive continentalization of the Camaquã Basin. These units can be interpreted as the evolution of temporally and spatially restricted basins, evolved from a foreland retroarc into a rift (or hemigraben) tectonic setting. In this more recent

The Maricá Formation was defined by Leinz et al. (1941) as a basal unit of Camaquã Basin. This Formation is a ±2500 m-thick sedimentary package with coastal deposits, including fluvial to shallow marine sediments with scarce volcanic contribution (Leites et al., 1990). The nature of the basal contact of this unit with the Neoproterozoic juvenile terrain is difficult to be defined, either an unconformity or a thrust plane, due to low relief, thick soil profiles and vegetation, and widespread younger intrusive bodies. According to Borba (2004), the Maricá formation can be divided into three major packages: the lower, intermediate and upper successions. The lower succession comprises conglomeratic and coarse-grained sandstones with very well rounded granite–gneiss pebbles and cobbles either marking the stratification planes or dispersed through the beds. The intermediate succession is made up of indurated siltstones and shales, either massive or showing planeparallel lamination, wave and climbing ripples, and lenticular lamination, as well as hummocky sandstones, suggesting a marine origin. The upper succession is very similar to the lower one, with the sandstones featuring large-scale troughcross and low-angle stratifications. Well-rounded granite–gneiss and rare volcanic, rhyolitic pebbles and cobbles are also present within the sandstone beds. The lower and upper successions correspond to braided fluvial deposits, while the intermediate package is related to a shallow marine depositional environment. The strata of the Maricá Formation were strongly affected by subsequent events of volcanism, intrusion, sedimentation and uplift. Volcanic rocks of the Bom Jardim Group (andesite Hilário) and the Acampamento Velho Formation (rhyolitic flow, rhyolitic pyroclastic rocks and basic flow) caused metamorphism contact in the Maricá Formation. The unit is unconformably overlain by the alluvial, fluvial and eolian deposits of the Santa Bárbara and Guaritas formations (Borba & Mizusaki, 2003). A syn-depositional volcanic event (Santos et al., 1978) has been recognized in the rock record of the Maricá Formation, evidenced by the occurrence of

interpretation, only the Maricá Formation was deposited in a foreland basin.

km of length from Mantiqueira Province.

**2.2 Stratigraphy**

volcanogenic beds interlayred with sedimentary rocks of the lower succession. Circular structures resembling degassing pipes also suggest the activity of volcanic hot springs (Borba et al., 2004). According to Paim et al (2000), in terms of deformation at the Taquarembó/Ramada sub-basin, are recognizes reverse ruptile faults and weak folds, while the Piquiri/Boici sub-basin is dominated by structures associated to the transcurrent processes in ruptile to ductile-ruptile structural level.

The Bom Jardim Group (Paim et al, 2000), is a volcano-sedimentary sequence that shows a lithostratigraphic complexity in its area-type and similarity lithological of some its units with others from Camaquã Supergroup units (Table 1). Bom Jardim Group, in its area-type (Janikian et al., 2003), shows approximately 4,000 thicknesses and constituted of andesitic volcanic rocks, pyroclastic rocks (lapilli tuffs, lithic tuffs and thick vitreous tuffs – Hilário Formation) and sedimentary rocks. These rocks were generated in an active tectonic lacustrine basin. Robertson (1966) made the first attempted correlation among the volcanosedimentary units from Bom Jardim area and others expositions from Camaquã Supergroup. Ribeiro et al. (1966) and Tessari & Picada (1966) correlated the outcropping units from Bom Jardim and Vale Piquiri (Piquiri Sub-basin) areas, covering the Arroio dos Nobres Formation.

The Hilário term was first used by Robertson (1966) to designate a variety of rock types composed of mainly andesitic to dacitic composition (flows, tuffs, volcanic conglomerates, mudflows, dykes and intrusive mafic rocks). Ribeiro et al. (1966) formalized this unit as a member of Crespos Formation - Bom Jardim Group. Santos et al. (1978) used the Hilário term to refer a member of Cerro dos Martins Formation, while Horbach et al. (1986) upgraded these rocks to the formation level. Several studies have been carried out Hilário Formation rocks. Nardi & Lima (1985), Porcher et al. (1995), Lima & Nardi (1998) works are especially relevant. Thus, Lima & Nardi (1998) defined the Lavras do Sul Shoshonitic Association, which comprises a wide compositional range of intrusive and extrusive basic to acid rocks. The extrusive rock includes the Hilário Formation and the spessartitic lamprophyres. In the Bom Jardim Group, the volcanic rocks are mostly in the lower stratigraphic levels, as lava flows or small shallow intrusion.

It the Santa Bárbara Sub-basin (Paim et al., 2000) toward the Guaritas Sub-basin, the volcanic rocks of this group are replaced gradational by alluvial conglomerates rich in volcanic clasts from basic to intermediate composition, and more distally, by sandy and silty related to the bottom flows (turbidites) , which are bounded to underwater deltaic systems. Next to the high topographic that today subdivides the Santa Bárbara and Guaritas sub-basins (e.g. Alto Caçapava), is common the occurrence of the erratic conglomerates containing local derivation clasts (shales and marble), indicating that the partitioning of the Camaquã Basin in several sub-basins occurred prior to the deposition of this unit.

Cerro do Bugio Allogroup (or Group) is composed of volcanic rocks at the base (Acampamento Velho Alloformation or Formation) and alluvial conglomerate deposits from Santa Fé Alloformation (ou Formation). According Paim (2000), Cerro do Bugio is limited by two angular unconformity that delineate its contact with the lower (Bom Jardim Allogroup) and upper (Santa Bárbara Allogroup) units. Cerro do Bugio Unit shows approximately 500 m thick and is composed of, from the base to the top, Acampamento Velho (volcanic) and Santa Fé (alluvial conglomerates, and secondarily, by sand-pelitic rhytmits and pelites), separated by an erosive unconformity. The Santa Fé Formation consists mainly of alluvial conglomerates rich in volcanic and plutonic clasts of the acid composition, which vary vertically to sand-pelitic rhytmits. These lithologies represent the alluvial systems interlaced, considering the transversal (deltaic fans, conglomeratic with paleocurrents for SE) and longitudinal (fans of the interlaced plain) characters. This allogroup was affected by dip-slip and transcurrent faults generated in crustal to ductile-ruptile level. Considering the field work observations, the Acampamento Velho Formation volcanic rocks are often interdigitated with rocks from Santa Fé Formation, suggesting the contemporaneity between both units.

Acampamento Velho Formation (sensu Cordani et al., 1974) is the basal unit of the Santa Bárbara Group and crops out mainly in the Santa Bárbara and Taquarembó-Ramada subbasins. The basal contact of the Acampamento Velho Formation is unconformable with the Maricá and Bom Jardim Groups. According to Almeida et al. (2002), in the Cerro do Bugio , Perau and Serra de Santa Bárbara (Fig. 1), the volcanic rocks (basaltic and rhyolitic compositions) upper contact is a disconformity by the alluvial conglomerates, and locally, by deltaic-lacustrine sandstones and mudstones from Santa Fé Formation.

The Santa Bárbara Group, defined by Robertson (1966), Ribeiro et al., (1966) and Ribeiro & Fantinel (1978) as Santa Bárbara Formation, comprises red-colored, sandstones and siltstones, and includes the Acampamento Velho and Santa Bárbara formations. It was named as Santa Bárbara Allogroup by Paim et al. (2000). This group is composed of Acampamento Velho, Santa Fé, Lanceiros and Segredo Formations, from the base to the top (Table 1). According to the same authors, the Lanceiros Formation is related to the progradation of an interlaced and sandy deltaic system of the longitudinal character, and the Pedra do Segredo Formation can be bounded to the progradation of an interlaced and sandy-conglomeratic of transversal nature. Borba & Mizusaki (2003) organized the Santa Bárbara Formation rocks in three depositional sequences named as, from the base to the top: Sequence I, Sequence II and Sequence III. The two basal sequences (I and II) represent a coherent depositional pattern, with axial fluvial and deltaic fan systems, which deposit North-Eastward with lateral contribution from alluvial fans. Such coarse deposits are composed mainly of metamorphic clasts derived from Alto Caçapava and Eastern steep Santa Bárbara Basin. The Sequence I shows a dip ranging from 32o to 40o NE in the Northern outcrops and 10o to 16o NE in the South, and the Sequence II from 28o to 30o SE. The Sequence III lies unconformable over basal subunits and reflects the axial system inversion, in that the gravel bed deposits paleocurrents systematically point South/South-Westward. The alluvial fan deposits of the Sequence III also suggest a tectonic rearrangement basin, with partial erosion from basal sequences, and the presence of granitic composition fragments, which reflects the deeper denudation stage of the Alto Caçapava and possibly a significant hiatus at the Sequence III base. The third dip sequence is between 23o to 26o ENE. The Acampamento Velho Formation is not included in the Santa Bárbara Formation (Borba & Misuzaki, 2003).

The Guaritas was defined by Goñi et al. (1962), Ribeiro et al (1966), Robertson (1966) and Ribeiro & Fantinel (1978) and is lithologically similar to the Santa Bárbara Group. This unit represents the last great depositional event preserved in Camaquã Basin and is above unconformity of the previous units. It was denominated Guaritas Allogroup (Paim et al., 2000) and is composed of two alloformations or formations, from the base to the top: Pedra Pintada and Varzinha (Table 1). These units are bounded by an erosive unconformity. The Guaritas Group shows around 800 m thick and represents the last depositional episode preserved within of the Camaquã Basin. The Pedra Pintada Formation is temporal and spatially related to the volcanic rocks named as Rodeio Velho Member. It was originated in a desert environment with crescent eolian dunes, in an inter-dune area that describes interchange of dry and wet periods, and basal level plains that comprehend dry seasons during wetter periods (Paim et al., 2000). It is characterized by fine to medium sandstone, well selected, with large to very-large size cross-stratification and bulk sandstones, and occurs secondarily mudstones and fine to medium sandstones with crossed lamination by stream and waves (Paim et al., 2000). The rocks that belong to this formation are found in the Santa Bárbara and Guaritas sub-basins (Fig. 1). The Varzinha Alloformation is composed of sedimentary faces that represent a braided fluvial system at the West Guaritas Sub-basin portion, and a fanlike alluvial system at the East portion. According to Paim et al. (2000), analyzing the paleocurrents data of this unit, we detect a lateral association of two different alluvial systems: (1) interlaced fluvial system, West side of the Guaritas sub-basin; (2) alluvial fans represented by, at the least two lobes, at the East side of the Guaritas Sub-basin (Paim, 1995). The Varzinha Alloformation upper levels correspond to the several sets of the progradacional parasequences from deltaic source. They were developed de within a shallow lacustrine basin. The deltaic deposits were associated to a tributary deltaic fan system (lateral contribution) at the East edge of the Guaritas Sub-basin, and an interlaced plain delta system (longitudinal contribution), which shows progradation to the SW, at the Eest edge of the Guaritas Sub-basin. Almeida et al. (2003a) described the Varzinha Alloformation as a covering of the area that present four cones (Rodeio Velho Member) at Santa Bárbara Sub-basin (see iten 5). According to Takehara et al (2010), the Guaritas Group, which comprises the aeolian and alluvial plain as well as fan deposits, is covering the Santa Bárbara Group rocks in Santa Bárbara Sub-Basin.

Rodeio Velho Member (sensu Ribeiro et al., 1966) has been described as being at least three vesicular andesite flows, with thickness estimated at 100 m and no evidence of explosive activity. Silva Filho (1996) showed the intrusive character of the magmatism, rejecting to form ideas of an exclusively volcanic event. Fragoso Cesar et al. (2000) named these rocks as Rodeio Velho Intrusive Suite, which is represented by tabular intrusions within the subhorizontal continental deposits of the Guaritas Group. Almeida et al. (2000, 2003a) stated that the before mentioned event has a basaltic andesitic composition with alkaline affinity and the rocks cropped out as lava flows, pyroclastic deposits and shallow intrusions.

#### **3. Geochronological data**

80 Petrology – New Perspectives and Applications

interlaced, considering the transversal (deltaic fans, conglomeratic with paleocurrents for SE) and longitudinal (fans of the interlaced plain) characters. This allogroup was affected by dip-slip and transcurrent faults generated in crustal to ductile-ruptile level. Considering the field work observations, the Acampamento Velho Formation volcanic rocks are often interdigitated with rocks from Santa Fé Formation, suggesting the contemporaneity between

Acampamento Velho Formation (sensu Cordani et al., 1974) is the basal unit of the Santa Bárbara Group and crops out mainly in the Santa Bárbara and Taquarembó-Ramada subbasins. The basal contact of the Acampamento Velho Formation is unconformable with the Maricá and Bom Jardim Groups. According to Almeida et al. (2002), in the Cerro do Bugio , Perau and Serra de Santa Bárbara (Fig. 1), the volcanic rocks (basaltic and rhyolitic compositions) upper contact is a disconformity by the alluvial conglomerates, and locally,

The Santa Bárbara Group, defined by Robertson (1966), Ribeiro et al., (1966) and Ribeiro & Fantinel (1978) as Santa Bárbara Formation, comprises red-colored, sandstones and siltstones, and includes the Acampamento Velho and Santa Bárbara formations. It was named as Santa Bárbara Allogroup by Paim et al. (2000). This group is composed of Acampamento Velho, Santa Fé, Lanceiros and Segredo Formations, from the base to the top (Table 1). According to the same authors, the Lanceiros Formation is related to the progradation of an interlaced and sandy deltaic system of the longitudinal character, and the Pedra do Segredo Formation can be bounded to the progradation of an interlaced and sandy-conglomeratic of transversal nature. Borba & Mizusaki (2003) organized the Santa Bárbara Formation rocks in three depositional sequences named as, from the base to the top: Sequence I, Sequence II and Sequence III. The two basal sequences (I and II) represent a coherent depositional pattern, with axial fluvial and deltaic fan systems, which deposit North-Eastward with lateral contribution from alluvial fans. Such coarse deposits are composed mainly of metamorphic clasts derived from Alto Caçapava and Eastern steep Santa Bárbara Basin. The Sequence I shows a dip ranging from 32o to 40o NE in the Northern outcrops and 10o to 16o NE in the South, and the Sequence II from 28o to 30o SE. The Sequence III lies unconformable over basal subunits and reflects the axial system inversion, in that the gravel bed deposits paleocurrents systematically point South/South-Westward. The alluvial fan deposits of the Sequence III also suggest a tectonic rearrangement basin, with partial erosion from basal sequences, and the presence of granitic composition fragments, which reflects the deeper denudation stage of the Alto Caçapava and possibly a significant hiatus at the Sequence III base. The third dip sequence is between 23o to 26o ENE. The Acampamento Velho Formation is not included in the Santa Bárbara Formation (Borba

The Guaritas was defined by Goñi et al. (1962), Ribeiro et al (1966), Robertson (1966) and Ribeiro & Fantinel (1978) and is lithologically similar to the Santa Bárbara Group. This unit represents the last great depositional event preserved in Camaquã Basin and is above unconformity of the previous units. It was denominated Guaritas Allogroup (Paim et al., 2000) and is composed of two alloformations or formations, from the base to the top: Pedra Pintada and Varzinha (Table 1). These units are bounded by an erosive unconformity. The Guaritas Group shows around 800 m thick and represents the last depositional episode preserved within of the Camaquã Basin. The Pedra Pintada Formation is temporal and spatially related to the volcanic rocks named as Rodeio Velho Member. It was originated in a desert environment with crescent eolian dunes, in an inter-dune area that describes

by deltaic-lacustrine sandstones and mudstones from Santa Fé Formation.

both units.

& Misuzaki, 2003).

We present a synthesis of geochronological ages for the different lithologies mentioned in the literature as well as new U-Pb ages obtained by us in samples from Maricá Formation (sandstone), Hilário Formation (lamprophyre), Acampamento Velho base (andesitic basalt) and Rodeio Velho Member (alkaline basalt).

#### **3.1 Analytical procedures used for new U-Pb ages**

All zircons were mounted in epoxy with 2.5 cm diameter and polished until the zircons were just revealed. Images of zircons were obtained using the optical microscope (Leica MZ 125) and back-scatter electron microscope (Jeol JSM 5800). Zircon grains were dated with laser ablation microprobe (New Wave UP213) coupled to a MC-ICP-MS (Neptune) at the Isotope Geology Laboratory from UFRGS. Isotope data were acquired using static mode with spot size of 25 and 40 um. Laser-induced elemental fractional and instrumental mass discrimination were corrected by the reference zircon (GJ-1) (Simon et al., 2004), following the measurement of two GJ-1 analyses to every ten sample zircon spots. The external error is calculated after propagation error of the GJ-1 mean and the individual sample zircon (or spot).

The common 204Pb, after Hg correction based on 202Hg simultaneously measured, is insignificant in most situations. For instance, typical signal intensity due to 204Hg during a laser ablation on standard Zircon is 600-1000 cps range, while the calculated count rate for 204Pb is less than statistical error of ca. 25-33 cps. We assume that the 204Pb values obtained from zircons have common Pb composition, assuming concordant age of 206Pb/238Pb and 207Pb/206Pb (as estimated age). In this case, we estimate the radiogenic composition of 206Pb and Pb207 using the equation as fraction of non radiogenic 206Pb (Williams, 1998):

f206= [206Pb/204Pb]c/[206 Pb /204Pb]s

f207= [207Pb/204Pb]c/[207/204]s

For common lead isotope composition, we assumed an isotope composition evolution proposed by Stacey and Kramers (1975), which is required to attribute an initial estimated age.

#### **3.2 Geochronological data for different units**

Soliani Jr. et al. (1984) presented a synthesis of all the radio-chronological determinations obtained for the crystalline and sedimentary rocks from Sul-Riograndense Shield meridional portion, including part of the Santa Catarina, the whole Rio Grande do Sul and Uruguay, thus preserving the Cordani et al (1974) and Teixeira (1982) proposal, which says that the West Sul-Riograndense Shield region would have had an Transamazonic evolution, and the meridional portion would have been generated in the Brasiliano Cycle. Sartori & Kawashita (1985) obtained the age of 550 Ma for Caçapava do Sul Granitic Batholith using 21 Rb-Sr isotope analyses on total-rock.

According to Soliani et al. (2000), the first work of the regional integration and geochronological synthesis was made by Cordani et al (1974). The authors relate the K-Ar and Rb-Sr radiometric data to the Eopaleozoic rocks, which fill the Camaquã Basin. Thus, the Rb-Sr data of the Eastern portion with syntectonic characteristics (granites and migmatites) and post-tectonic (isotropic granites) show respectively, values from 650-600 Ma. K-Ar ages of the same region are in the same coherent interval characteristics of the Brasiliano Cycle. For the Midwest region of the shield were found radiometric values that suggest an older age, although a significant data percentage show Brasiliano ages for the granitic massive such as the Caçapava do Sul and Lavras do Sul among others. In the same way was confirmed that the volcanic rocks (Hilário and Acampamento Velho formations) were generated at the end of the Pre-Cambrian or at the beginning of the Phanerozoic, exactly as has been commented by Minioli & Kawashite (1971).

Leite et al. (1995, 1998) and Leite (1997) mentioned for the Rio Grande do Sul region, the first U-Pb data obtained in zircons using the sensitive high mass-resolution ion microprobe method, being the isotope data for the Caçapava do Sul granitic massives of the 561± 6 Ma and 541± 11 Ma. The first age is considered as corresponding to the igneous protolith that generated the Caçapava Granite. The second age is related to a magmatic event that generated the granite. Remus et al. (1996, 1997) used sensitive high mass-resolution ion microprobe method to obtain the U-Pb ages in zircons for the Lavras do Sul Granite core and edge, respectively: 592±5 and 580 ±5 Ma. The age of 565± 14 Ma was obtained for the Caçapava do Sul Granite.

#### **3.2.1 Maricá Formation**

82 Petrology – New Perspectives and Applications

The common 204Pb, after Hg correction based on 202Hg simultaneously measured, is insignificant in most situations. For instance, typical signal intensity due to 204Hg during a laser ablation on standard Zircon is 600-1000 cps range, while the calculated count rate for 204Pb is less than statistical error of ca. 25-33 cps. We assume that the 204Pb values obtained from zircons have common Pb composition, assuming concordant age of 206Pb/238Pb and 207Pb/206Pb (as estimated age). In this case, we estimate the radiogenic composition of 206Pb

For common lead isotope composition, we assumed an isotope composition evolution proposed by Stacey and Kramers (1975), which is required to attribute an initial estimated

Soliani Jr. et al. (1984) presented a synthesis of all the radio-chronological determinations obtained for the crystalline and sedimentary rocks from Sul-Riograndense Shield meridional portion, including part of the Santa Catarina, the whole Rio Grande do Sul and Uruguay, thus preserving the Cordani et al (1974) and Teixeira (1982) proposal, which says that the West Sul-Riograndense Shield region would have had an Transamazonic evolution, and the meridional portion would have been generated in the Brasiliano Cycle. Sartori & Kawashita (1985) obtained the age of 550 Ma for Caçapava do Sul Granitic Batholith using 21 Rb-Sr

According to Soliani et al. (2000), the first work of the regional integration and geochronological synthesis was made by Cordani et al (1974). The authors relate the K-Ar and Rb-Sr radiometric data to the Eopaleozoic rocks, which fill the Camaquã Basin. Thus, the Rb-Sr data of the Eastern portion with syntectonic characteristics (granites and migmatites) and post-tectonic (isotropic granites) show respectively, values from 650-600 Ma. K-Ar ages of the same region are in the same coherent interval characteristics of the Brasiliano Cycle. For the Midwest region of the shield were found radiometric values that suggest an older age, although a significant data percentage show Brasiliano ages for the granitic massive such as the Caçapava do Sul and Lavras do Sul among others. In the same way was confirmed that the volcanic rocks (Hilário and Acampamento Velho formations) were generated at the end of the Pre-Cambrian or at the beginning of the Phanerozoic,

Leite et al. (1995, 1998) and Leite (1997) mentioned for the Rio Grande do Sul region, the first U-Pb data obtained in zircons using the sensitive high mass-resolution ion microprobe

and Pb207 using the equation as fraction of non radiogenic 206Pb (Williams, 1998):

Cup configuration:

Coolant flow (Ar) 15 l/min Auxiliary flow (Ar) 0.8 l/min

Acquisition 50 cycles of 1.048 s

Gas input:

Faradays 206Pb, 208Pb, 232Th, 238U MIC's 202Hg, 204Hg + 204Pb, 207Pb

**MC-ICP-MS Neptune**

Carrier flow 0.75 l/min (Ar) + 0.45 l/min (He)

**Laser operating conditos**

 Laser output power 6 J/cm2 Shot repetition rate 10 Hz Laser spot 25 and 40 µm

f206= [206Pb/204Pb]c/[206 Pb /204Pb]s f207= [207Pb/204Pb]c/[207/204]s

isotope analyses on total-rock.

**3.2 Geochronological data for different units**

exactly as has been commented by Minioli & Kawashite (1971).

age.

**Laser type New Wave UP213**

The first geochronological investigation in Maricá Formation was performed by Borba et al. (2008). These authors made analysis of zircons from pyroclastic cobbles, which yielded an age of 630.2±3.4 Ma (2σ), interpreted as the age of syn-sedimentary volcanism, and thus of the deposition itself. This result indicates, according to Borba et al. (2008), that Maricá Formation was deposited during the main collisional phase (640–620 Ma) of the Brasiliano II Orogenic system from Silva classification (2005). Paim et al (2000) suggested a depositional age between 620 and 600 Ma in a foreland basin.

Table 2. U/Pb zircon data of the sandstone sample (CK-239D) for Maricá Formation obtained by in situ LAM-ICPMS-MC.

One sandstone sample (CK-239-D) used in the U-Pb analysis was collected at North of the Estância 3 Estradas (UTM 771863- 6576404) in Coxilha do Tabuleiro Sheet (1:50.000). Twenty zircon crystals from sample CK-239-D were dated. Based on the dating results, we detected the presence of follow age set: 2606 Ma (4 zircons), 2473 Ma (3 zircons), 2050 Ma (3 zircons) and 1790 Ma (one zircon - Table 2, Fig. 4). The 601 ±13 Ma age (based on eight zircons – Table 2, Fig. 4 - Early Ediacaran) was considered as the age older from Maricá Formation (Fig. 6). All the zircons that showed older ages were considered xenocrystal zircons. Thus, the magma would have assimilated Neoarchean (NA), Paleoproterozoic (Siderian, Rhyacian and Statherian) material during its ascension. The selected zircons that allowed the Maricá age setting (Fig. 5), in general are colorless and dirty, all grains are between 175 µm and 344 µm in size and show diffuse zonation and subrounded contours. The zircons do not have impurities and inclusions, but show few fractures. In the A-I-02 zircon (Fig. 5) is observed an incomplete concentric fracturing at the edge of the crystal.

Fig. 4. U-Pb concordia data diagram with tracing considering only the younger zircon ages from Maricá Formation sandstone sample. a: U-Pb age histogram considering all the obtained ages (table 2). We detected the presence of five ages: 2606 Ma, 2473 Ma, 2050 Ma, 1790 Ma and 601 Ma. The 601 Ma age is considered as the older from Maricá Formation.

Fig. 5. Zircons microphotographies that defined the 601±13 Ma age from CK239D sample for Maricá Formation sandstone. A-I-02 = sample number, o = spot showing the place where the analyses were obtained in each zircon. Data showed in table 2.

#### **3.2.2 Hilário Formation**

84 Petrology – New Perspectives and Applications

One sandstone sample (CK-239-D) used in the U-Pb analysis was collected at North of the Estância 3 Estradas (UTM 771863- 6576404) in Coxilha do Tabuleiro Sheet (1:50.000). Twenty zircon crystals from sample CK-239-D were dated. Based on the dating results, we detected the presence of follow age set: 2606 Ma (4 zircons), 2473 Ma (3 zircons), 2050 Ma (3 zircons) and 1790 Ma (one zircon - Table 2, Fig. 4). The 601 ±13 Ma age (based on eight zircons – Table 2, Fig. 4 - Early Ediacaran) was considered as the age older from Maricá Formation (Fig. 6). All the zircons that showed older ages were considered xenocrystal zircons. Thus, the magma would have assimilated Neoarchean (NA), Paleoproterozoic (Siderian, Rhyacian and Statherian) material during its ascension. The selected zircons that allowed the Maricá age setting (Fig. 5), in general are colorless and dirty, all grains are between 175 µm and 344 µm in size and show diffuse zonation and subrounded contours. The zircons do not have impurities and inclusions, but show few fractures. In the A-I-02 zircon (Fig. 5) is observed an

Fig. 4. U-Pb concordia data diagram with tracing considering only the younger zircon ages from Maricá Formation sandstone sample. a: U-Pb age histogram considering all the obtained ages (table 2). We detected the presence of five ages: 2606 Ma, 2473 Ma, 2050 Ma, 1790 Ma and 601 Ma. The 601 Ma age is considered as the older from Maricá

Formation.

incomplete concentric fracturing at the edge of the crystal.

Ribeiro & Teixeira (1970) mentioned values between 510-535 Ma for Hilário Andesite. These values were obtained through the K-Ar method. Cordani et al. *(*1974) obtained ages of 510 and 535 Ma using the Rb/Sr method. 206Pb206/238U data (sensitive high mass resolution ion microprobe) were obtained for Hilário Andesite by Remus et al. (1999) and gave an age of 580 Ma. Janikian et al (2008) mentioned ages from 590 to 585 Ma, which are interpreted for us as representative of Hilário Formation extrusion, being U-Pb age of 590 ± 5.7 Ma obtained in the Hilário Andesite as the most representative. The ages obtained by Remus et al. (2000) using the U-Pb (sensitive high mass-resolution ion microprobe**)** for Lavras Granite (592±5 Ma, 597±5 Ma, 580±7 Ma) and Caçapava do Sul (589±5 Ma) are considered as comagmatic to the Bom Jardim volcanism. Chemale Jr. et al. (2000) reported an age of 592 Ma for Andesite Hilário.

The LV- 243 sample (alkaline basalt) dated using the U-Pb method by us was collected 15 km away from the Lavras do Sul, following by RS 357 road toward Caçapava do Sul (Fig.1). Lima & Nardi (1998) said that the out crop is a dome with size around 0.5 km2 and a circular shape. It shows a marked columnar disjunction with hexagons of variable diameter, but no exceeding 15 cm diameter. Nine zircon crystals (sample LV- 243) were dated (Table 3). The obtained age of 591.8 ±3.0 Ma was calculated based on individual zircon ages from Table 3. This age defines the lamprophyre age (Fig. 6), which belong to the Hilário volcanism manifestation. The selected zircons that allowed the lamprophyres ages setting, in general, are colorless and clean, all grains are between 150 µm and 333 µm in size and the most show very clear zoning. The grains also show prismatic and pyramidal shapes, incomplete and partially serrulated contours and few fractures (Fig. 7).

Table 3. U/Pb zircon data of the lamprophyre sample (LV-243) for Hilário Formation obtained by in situ LAM-ICPMS-MC.

Fig. 6. U-Pb concordia data diagram with tracing considering the zircon ages for Hilário Formation lamprophyre (LV-243 sample). Gray ellipse cores represent the single analysis from zircons of the table 3. Black ellipse represents the concordia age for LV-243 sample, calculated from single zircon ages of the table 3.

072-B-II-1 0,80351 3,22 0,09732 1,80 0,56 599 11 599 19 599 16 0,34 0 0,0010 072-B-II-2a 0,80971 3,22 0,09779 1,47 0,46 601 9 602 19 605 17 0,53 1 0,0005 072-B-II-2b 0,81989 3,00 0,09891 1,46 0,49 608 9 608 18 608 16 0,33 0 0,0009 072-B-II-3 0,80923 3,56 0,09799 1,58 0,44 603 10 602 21 600 19 0,34 0 0,0012 072-B-II-4 0,78691 3,08 0,09534 1,37 0,45 587 8 589 18 599 17 0,34 2 0,0004 072-B-II-7 0,78269 3,54 0,10586 1,80 0,51 649 12 638 23 602 18 0,41 -8 0,0004 072-B-II-9 0,76140 3,95 0,09280 1,59 0,40 572 9 575 23 586 21 0,41 2 0,0003 072-B-II-16a 0,74234 4,22 0,08980 3,22 0,76 554 18 564 24 602 16 0,45 8 0,0002 072-B-II-16b 0,79361 3,39 0,09618 1,57 0,46 592 9 593 20 598 18 0,34 1 0,0007

error 207Pb / 235U

error 207Pb / 206Pb error 232Th / 238U

% Disc f206

Concordia 1 Age (Ma)

Rho 1 206Pb 238U

Table 3. U/Pb zircon data of the lamprophyre sample (LV-243) for Hilário Formation

Fig. 6. U-Pb concordia data diagram with tracing considering the zircon ages for Hilário Formation lamprophyre (LV-243 sample). Gray ellipse cores represent the single analysis from zircons of the table 3. Black ellipse represents the concordia age for LV-243 sample,

calculated from single zircon ages of the table 3.

 erro r

 Spot number 207Pb/ 235U

 error 206Pb/ 238U

obtained by in situ LAM-ICPMS-MC.

Fig. 7. Zircons microphotographies that defined the 553 ± 5.4 Ma age for lamprophyric rock from Hilário Formation (LV-243 sample). A-I-02 = sample number, o = spot showing the place where the analyses were realized in each zircon. Data showed in table 3.

#### **3.2.3 Acampamento Velho Formation**

The first geochronological investigation of the rhyolites was performed by Cordani et al. (1974) followed by Sartori (1978) and Teixeira (1982). Soliani Jr. (2000) compiled their data and obtained an age of 529 ± 4 Ma and R0 = 0,706 (Rb-Sr whole rock). Another Rb-Sr dating was performed by Almeida et al. (2002), who studied the rhyolitic flows from Cerro do Bugio area and the dykes intruding the Maricá Formation. These authors obtained two whole rock isochrons: 545.1 ± 12.7 Ma (R0 = 0.709) and 546 ± 12.9 Ma (R0 = 0.714). However, Sommer et al. (2005) presented U-Pb data (sensitive high mass-resolution ion microprobe) using eleven zircon crystals from rhyolites of Vila Nova do Sul area (Ramada Plateau), obtaining an age of 549.3 ± 5 Ma, which is more reliable. Therefore, all the obtained ages indicate that the rocks from Acampamento Velho Formation upper felsic association belong to the Early Neoproterozoic III (NP3).

The MH-13 sample (andesitic basalt) was selected to obtain the absolute age from lower mafic association of Acampamento Velho Formation base. Nineteen zircons crystals were dated (Table 4). Based on the dating results, we detected the presence of four age group: 2799 ± 21 Ma (one zircon); 2182 ± 55 Ma (one zircon), 612.6 ± 15 Ma (seven zircons) apart from anchored age, that corresponds to intersection at 2442 ± 54 Ma (three zircons) (Table 4, Fig. 8a). The 553 ± 5.4 Ma age (based on six zircons – Table 4, Fig. 8b) was considered as the lower mafic association age. All the zircons that showed older ages were considered xenocrystal zircons. Thus, the magma would have assimilated Neoarchean (NA), Paleoproterozoic (Siderian and Rhyacian) and older Neoproterozoic material during its ascension.

The selected zircons that allowed the Acampamento Velho Formation base age setting, in general are colorless and dirty, all grains are between 87 µm and 158 µm in size and show diffuse zonation, prismatic and pyramidal shapes, incomplete and partially serrulated contours, few fractures and inclusions (Fig. 9). The fain and broad zoning suggest that it may have been formed by very slow and complex crystallization of a magma body with prolonged residence time in the lower crust. The crystallization velocity appears to be the major controlling factor of the elongation ratio to the zircon. Skeletal zircon crystals are the most extreme form of rapid growth. It is relevant to point out that the Acampamento Velho Formation mafic flow dated zircons show morphologic characteristics of a slow generation in the magmatic chamber.

Table 4. U/Pb zircon data of the andesitic basalt sample (MH-13) for Acampamento Velho Formation base obtained by in situ LAM-ICPMS-MC.

from anchored age, that corresponds to intersection at 2442 ± 54 Ma (three zircons) (Table 4, Fig. 8a). The 553 ± 5.4 Ma age (based on six zircons – Table 4, Fig. 8b) was considered as the lower mafic association age. All the zircons that showed older ages were considered xenocrystal zircons. Thus, the magma would have assimilated Neoarchean (NA), Paleoproterozoic (Siderian and Rhyacian) and older Neoproterozoic material during its

The selected zircons that allowed the Acampamento Velho Formation base age setting, in general are colorless and dirty, all grains are between 87 µm and 158 µm in size and show diffuse zonation, prismatic and pyramidal shapes, incomplete and partially serrulated contours, few fractures and inclusions (Fig. 9). The fain and broad zoning suggest that it may have been formed by very slow and complex crystallization of a magma body with prolonged residence time in the lower crust. The crystallization velocity appears to be the major controlling factor of the elongation ratio to the zircon. Skeletal zircon crystals are the most extreme form of rapid growth. It is relevant to point out that the Acampamento Velho Formation mafic flow dated zircons show morphologic characteristics of a slow generation

> **Rho 1 206Pb/ 238U**

119-A-I-02 0,73099 4,86 0,09051 2,05 0,42 559 11 557 27 551 24 0,62 -1 0,0007 119-A-I-12 0,71742 5,44 0,08906 2,96 0,54 550 16 549 30 546 25 0,51 -1 0,0008 119-A-I-24 0,71164 4,02 0,09119 3,43 0,85 563 19 546 22 476 10 0,69 -18 0,0000 119-A-I-28 0,71175 2,90 0,08921 1,89 0,65 551 10 546 16 525 12 0,99 -5 0,0000 119-A-I-32 0,72048 4,18 0,08986 2,12 0,51 555 12 551 23 535 19 0,66 -4 0,0000 119-A-I-26 0,68233 4,88 0,08835 4,35 0,89 546 24 528 26 453 10 1,06 -21 0,0000

119-A-I-05 0,77464 3,54 0,09524 1,96 0,55 586 11 582 21 567 17 0,52 -3 0,0003 119-A-I-07 0,83567 4,54 0,10066 2,19 0,48 618 14 617 28 611 24 1,32 -1 0,0005 119-A-I-16 0,79403 4,41 0,09638 1,88 0,43 593 11 593 26 595 24 0,96 0 0,0009 119-A-I-19 0,81080 3,32 0,09729 2,50 0,75 599 15 603 20 619 14 0,50 3 0,0000 119-A-I-22 0,84916 5,93 0,10603 2,60 0,44 650 17 624 37 533 28 0,49 -22 0,0001 119-A-I-23 0,83998 3,39 0,10432 1,63 0,48 640 10 619 21 545 16 0,65 -17 0,0000 119-A-I-29 0,77999 4,20 0,09460 3,70 0,88 583 22 585 25 596 12 0,66 2 0,0002

119-A-I-20 14,70031 2,19 0,54875 1,28 0,58 2820 36 2796 61 2779 49 0,72 -1 0,0001

119-A-I-01 10,30577 2,14 0,46719 1,94 0,90 2471 48 2463 53 2456 22 0,23 -1 0,0011 119-A-I-10 10,20802 2,51 0,46490 1,98 0,79 2461 49 2454 62 2448 38 0,34 -1 0,0005 119-A-I-18 9,37569 1,89 0,43936 0,99 0,52 2348 23 2375 45 2399 39 1,07 2 0,0018 119-A-I-17 4,33522 2,14 0,28940 1,58 0,74 1639 26 1700 36 1777 26 0,73 8 0,0010

119-A-I-01 7,59979 6,12 0,40841 4,51 0,74 2820 36 2185 134 2163 89 0,41 -2 0,0017

Table 4. U/Pb zircon data of the andesitic basalt sample (MH-13) for Acampamento Velho

**Concordia 1 Age (Ma) %**

**207Pb/ 235U**

**207Pb/ 206Pb**

**232Th/ 238U**

Dis. **F206**

ascension.

**Spot number**

in the magmatic chamber.

**207Pb/ 235U**

**206Pb/ 238U**

Formation base obtained by in situ LAM-ICPMS-MC.

Fig. 8. Concordia diagram from Acampamento Velho Formation lower mafic association: (a) U-Pb data diagram with tracing considering only the older zircon ages of the 553 Ma (table 4). We detected the presence of four ages: 2799 ± 21 Ma, 612.6 ± 15 Ma, 612.6 ± 15 Ma and 2442 ± 54 Ma. (b) U-Pb data diagram with tracing considering only the younger zircon ages from Acampamento Velho Formation lower mafic association (MH-13 sample). Gray ellipse cores represent the single zircon analysis (table 4). Black ellipse represents the concordia age of the MH-13 sample calculated from single zircon ages (table 4).

Fig. 9. Zircons microphotographies that defined the 553 ± 5.4 Ma age for Acampamento Velho Formation base. A-I-02 sample number, = spot showing the place where the analysis were realized in each zircon. Data showed in table 4.

#### **3.2.4 Rodeio Velho Member**

Hartmann et al. (1998) determined the U-Pb age of 470 ± 19 Ma (sensitive high massresolution ion microprobe method) for this event (Middle-Ordovician), which suggest that Rodeio Velho Member is related to the beginning oh the Paraná Basin formation. On the other hand, Chemale Jr. (2000) obtained Sm-Nd model ages (TDM) from 1.6 to 1.9 Ga, suggesting a modified mantle origin for these rocks. Rodeio Velho Member was probably the source of heat and hydrothermal solutions for Cu (Ag, Au), Pb and Zn (Cu, Ag) mineralizations in the Camaquã Mines (Lima et al., 2001). Hydrothermal illites collected in those mines also yielded K-Ar ages around 465 Ma (Bonhomme & Ribeiro, 1983).

We used the sample RLP-20 (alkaline basalt) to date Rodeio Velho event using U-Pb method. The sample was collected between Rodeio Velho and Pedra de Arara area (Fig. 1). At this area, the volcanic rocks show two lava flows, being the lower flow more vesiculated (flow top) than the upper, which is massive. This is related to the sandstones with great crossstratification, suggesting the eolian character of this rock and it would correspond to the Pedra Pintada Formation from Guaritas Group. We dated sixteen zircon crystals from sample RLP-20. Based on these dating results, we detected the presence of 4 age groups: 2190± 18 Ma (one zircon); 1079 ± 12 Ma (six zircons); 658 ± 8.3 Ma (two zircons) and 547 ± 6.3 Ma (based on five zircons), (Table 5, figs. 10a and 10b). The younger obtained age, that was calculated based on individual zircon ages from Table 5, defined the Rodeio Velho Member age (Fig. 10a). All the zircons that showed older ages were considered xenocrystal inherited zircons. Thus, the magma would have assimilated Paleoproterozoic (PP2), Mesoproterozoic (MP3) and Neoproterozoic (NP2) rocks during its ascension. The selected zircons used to obtain the Rodeio Velho Member age are in general (Fig. 11) colorless and transparent, the

Fig. 10. U-Pb concordia data diagram for Rodeio Velho Member. (a) - U-Pb concordia data diagram with tracing considering only the younger zircon ages. Gray ellipse cores represent the single zircon analysis, with exception of A-I-02 zircon (black ellipse). Dark gray ellipse represents the concordia age calculated from single zircon ages (Table 5). The A-I-02 zircon sample shows high percentage of discordance and was not included in the calculating data (black ellipse). (b) - U-Pb concordia data diagram with tracing considering only the older zircon ages of the 547 Ma (table 5).