Madagascar), a continental rift?
Master thesis
By Paul Janssen
5
thof September 2006
Index
1. Abstract ... 2 2. Introduction ... 4 3. Gondwana ... 5 4. Geology of Madagascar ... 8 4.1 Introduction ... 8 4.2 Antananarivo Block ... 8 4.3 Tsaratanana Sheet ... 104.4 Neoproterozoic Sedimentary Belts ... 11
4.5 Antongil Block ... 11
4.6 Bemarivo Belt ... 11
4.7 Deformation History ... 13
5. Fieldwork ... 15
5.1 Introduction ... 15
5.2 Geology and Morphology ... 16
5.3 Lithology ... 17
5.4 Petrology ... 18
6. Geothermobarometry ... 20
6.1 Aluminum-in-hornblende barometer ... 21
6.2 Amphibole – Plagioclase thermometer ... 23
6.3 Mineral composition ... 27 6.4 Thermobarometric results ... 30 7. Discussion ... 33 7.1 Petrology ... 35 7.2 Thermobarometry ... 38 8. Conclusions ... 43 9. Acknowledgments ... 45 10. References ... 46 11. Appendices ... 50 11.1 Sample overviews ... 50
1.
Abstract
The Bemarivo suture zone, separating the Neoproterozoic Bemarivo belt with the Archean rocks of the Antananarivo- and the Antongil block, has thus far been identified as a top-to-the south thrust contact (Collins & Windley, 2002), which according to cooling ages occurred in between 520 and 510 Ma (Buchwaldt et al., 2003). Only scarce research has been done in this suture zone. Only in the extreme east and west field research has been done, but not in the central part of the suture zone, where this research was done. In order to justify this model not only the structural geological field research is needed, but also geothermobarometric evidence is crucial.
The orthogneisses found in the Mangindrano area belong to the western part of the Sambirano-Sahantana Group, in contrast to the geological map of Besarie (1971) and the later modified map of Buchwaldt et al. (2003). Moreover, the orthogneisses experienced (partial) recrystallization during peak metamorphism, which post-dated the major deformation events. Geothermobarometry of the orthogneisses indicates a peak metamorphic grade of 4-5.5 kbar and 695-776 ºC, which indicates high amphibolite facies.
The geotherms of the Mangindrano region (44 ºC/km), Sambirano region (40.5 ºC/km) and the Lokoho region (31 ºC/km) during peak metamorphism indicate a continental rift setting along the whole Bemarivo suture zone. This does not agree with the current theory of the Bemarivo suture zone as a thrust system. The stretching lineations found in the Mangindrano area (Mulder 2006), which indicate E-W transcurrent movement along the suture zone, also do not fit with the N-S thrust movement assumed by Buchwaldt et al. (2003) and Collins et al. (2001).
The peak metamorphism in the Mangindrano area most likely post-dated the intrusion of the granitoid veins and dykes, which post-dated the major deformation events. This agrees with most of the absolute ages of peak metamorphism and leucogranites in the Sambirano and Lokoho region (Buchwaldt et al., 2003)(Buchwaldt and Tucker, 2001).It indicates that the formation of stretching lineations and the peak metamorphic conditions were not coeval. So, they are probably not formed in the same tectonic setting. In this case transcurrent movements along the Bemarivo suture zone would have caused the amalgamation of the Bemarivo belt and the older part of Madagascar, followed by a period of continental rifting.
2.
Introduction
Lying off the coast of Mozambique, Madagascar is the fourth largest island in the world and has a very rich geological history. During the Neoproterozoic to Cambrian times a supercontinent
Gondwana was formed (Collins & Windley, 2002; Kröner & Cordani, 2003; Powell et al. 1993; Shackleton 1996), which was the result of the final collision between Africa, India, Antarctica, and Madagascar (Fig. 1). During this process the Mozambique ocean closed and cratons, arcs and ocean floor material were all amalgamated into one mountain belt, namely the East African Orogeny
(EAO)(Stern, 1994), which reached from Arabian-Nubian Shield in the north to Madagascar and even Antarctica in the south (Kusky et al., 2003). Being trapped between India and Africa and in the middle of the EAO (Fig. 1), Madagascar is consequently one of the most important locations to find evidence for the processes playing a part in the formation and the break-up of Gondwana (Collins & Windley, 2002; Collins et al., 2003; de Wit et al., 2001).
The western coast line and the minority of the geology of Madagascar consists of sedimentary basins capturing sedimentation from the Carboniferous to present (de Wit et al., 2001). However, a major part of the island is made up of tectonic units consisting of Precambrian rocks which all have experienced the EAO event (de Wit et al., 2001).The major tectonic units are the Antongil block, the Antananarivo block, the Bemarivo belt and Tsaratanana sheet with the south consisting of
Neoproterozoic supracrustal belts and several shear zones (de Wit, 2003; Collins et al., 2000; Collins & Windley, 2002).
of the metamorphic event relative to the deformation and intrusion events of the tectonic contact zone dividing the Bemarivo belt, with the Archean Antananarivo- and Antongil blocks.
3.
Gondwana
Due to continental drift, several supercontinents have formed during Earth’s history. Gondwana is one of these supercontinents that was formed after the break-up of a former supercontinent Rodinia. One of the first indications for such an event was described by Kennedy (1964) using the term Pan-African, meaning all tectonothermal events with a similar age of 500 ±100 Ma in rocks of some
African countries and other Gondwana players. Further research by, among others, Kröner (1984) who included orogenic events between 950 – 450 Ma to the term Pan-African, lead to the concept of a supercontinent that formed during this period, which was named Gondwana. With the help of similar geochronology, paleomagnetics and structural- and metamorphic data found on the different continents where Cambrian and Neoproterozoic rocks were still accessible, the reconstructions were consequently made for the evolution of Gondwana (e.g. Stern, 1994). Despite the extensive research concerning this
subject, a lot of uncertainties are still present related to the exact sequence of events and the precise role and the arrangement of the different cratons and continents involved.
The formation of Gondwana is roughly viewed as the collision between West and East Gondwana during the Cambrian - Neoproterozoic. However, recent research has lead to believe that this is an oversimplification and that the final coming together was accomplished by numerous smaller steps. According to Meert (2003) the formation of Gondwana occurred in two main orogenic events. The first collision was the convergence of the arc terrains in the Arabian – Nubian region between 900 and 650 Ma which continued up to the closure of the Mozambique ocean and the consequent collision of West Gondwana (Kenya – Tanzania and further locations to the north) and a part of East Gondwana (Madagascar, Sri Lanka, Seychelles and India) at 650 – 630 Ma (Meert, 2003). This gradual event lead to the formation of the East African Orogeny (EAO)(Stern, 1994) which is believed to be a
Himalayan-type orogeny going from the Arabian – Nubian region, through Mozambique and
Madagascar to Antarctica (Kusky et al., 2003). The second step in the formation of Gondwana was the collision of East Antarctica and Australia with the still poorly defined East Gondwana. This lead to the Kuunga orogeny dated at 570 – 530 Ma (Meert, 2003). The major events are relatively well
established, which can not be said about the smaller events in between the major orogenies. According to the two-step formation of Gondwana (Meert, 2003), East Gondwana could hardly be defined by one name, considering the fact that Australia and Antarctica were not yet attached. This is supported by earlier research claiming that in between the formation of the East African and the Kuunga Orogeny, East Gondwana was made up of different parts (Meert et al., 1995; Meert & Van der Voo, 1997).
Bemarivo Belt
Tsaratanana Sheet
Antananarivo Block
Antongil Block
Bemarivo Belt
Tsaratanana Sheet
Antananarivo Block
Antongil Block
The poorly established arrangement of the continental blocks, participating in the formation of Gondwana, makes research in the area very interesting. Especially when dealing with a craton
belonging to Madagascar, which was positioned in the middle of the EAO and, therefore, contains a lot of evidence in unravelling this geological problem.
4.
Geology of Madagascar
4.1 Introduction
Besides the Phanerozoic sedimentary basins along the west coast of Madagascar, which is not of interest to this research, the majority of the geology on the island consists of rocks of Pre-Cambrian to Cambrian age belonging to several tectonic units which are divided by thrust contacts, shear- and mylonite zones. These units include: the Antananarivo block, the Tsaratanana sheet, the Antongil block, the Neoproterozoic metasedimentary belts in the southern part, and the most northern Bemarivo belt (Fig. 2; Collins, 2000; Collins & Windley, 2002; De Wit, 2001).
4.2 Antananarivo Block
The majority of the geology of Central Madagascar forms a part of the Antananarivo Block, which is named after the capital of the island lying roughly in the middle of this block. The basement of this tectonic block consists of Precambrian (2550-2500 Ma) granitoids, which are interlayered with younger (824-719 Ma) granites, syenites and gabbros that have a subduction related chemistry (Collins, 2006). Granulite-grade metamorphism was subsequently active between 700 and 532 Ma, overlapped by a period of granitoid intrusion between 630 and 561 Ma characterised by 100m- to km-scale bodies, a characteristic feature of the Antananarivo Block (Collins & Windley, 2002; Nédélec et al., 2000; Paquette & Nédélec, 1998). The granitoid formation was coeval with a period of extension which is associated with the Betsileo shear zone, now marking the eastern boundary the Itremo Group with the Antananarivo Block (Fig. 3). The transcurrent deformation associated with the so called ‘Zone
de Virgation’ dated at ~560 Ma by the remagnetization of the ‘stratoid’ granites and the formation of
the Angavo belt in the east (Paquette & Nédélec, 1998; Kröner et al., 2000; Meert et al., 2003). The age of this event is constrained by the youngest granitoid gneiss dated (~551 Ma) and the age of the undeformed Carion granite (~538 Ma) (Kröner et al., 2000).
The Itremo group, which lies geographically south of the Antananarivo block (Fig. 3), can be seen as an independent tectonic unit due to its different deformation history, but due to their similar age, it is considered as a part of the Antananarivo block. The Itremo group consists of a basement of
Fig. 4 Fig. 4
followed by metasediments and finally overlain by metavolcanics (Collins, 2006; Collins & Windley, 2002). The metasediments consist of quartzites, pelites and dolomitic carbonates, dated at 1855 ± 11 Ma and were therefore probably deposited during Proterozoic times before the formation of
Gondwana. Due to the similarities between sequences on the African mainland, the depositional environment could be a passive margin. However, this cannot be established due to the lack of basin boundaries within the Itremo group (Cox et al., 1998). The metasediments where deformed into large-scale recumbent folds, separated by mylonitic shear zones. The metamorphic grade increases from sub-greenschist rocks in the east to kyanite-and sillimanite-bearing rocks in the centre and the west of Itremo (Cox et al., 1998). The metasediments where intruded by syenites and gabbros at 804-779 Ma, which show a subduction characteristic chemistry similar to the Antananarivo block but are less deformed (Collins, 2000; Collins, 2006; Collins & Windley, 2002). A second deformation occurred after 789 Ma, leaving open, upright folds, divergent reverse faults and strike-slip faults, which were finally intruded by granitoid bodies (Collins, 2002; Collins, 2006). It is remarkable that no extensional deformation has been active in the Itremo Group as opposed to most of the rest of Madagascar. An explanation for this fact could be that extension was compensated by sliding along the Betsileo shear zone, i.e. the east boundary of the Itremo Group (Collins, 2000).
4.3 Tsaratanana Sheet
The Tsaratanana Sheet lies unconform on the rocks of the Antananarivo Block and is made up of three main belts, namely the Maevatanana, Andriamena and the Bforona belts (Collins & Windley, 2002).The rocks that make up these belts are mafic gneisses, tonalities, podiform chromite-bearing ultramafic rocks, and metapelites which have undergone very high temperature metamorphism
(1000°C and 10.5kb) around 2.5 Ga (Collins & Windley, 2002; Goncalves et al., 2004). Intrusions did already occur before this event (2.75-2.49 Ga). A second phase of intrusion was marked by gabbros at 800-770 Ma with coeval granulite metamorphism in part of the Tsaratanana sheet (Collins et al., 2001). Granitoid intrusions, like in the most of the rest of Madagascar, occurred at 637-627 Ma followed directly by a top-to-the-east thrusting making it the last deformation event that the
4.4 Neoproterozoic Sedimentary Belts
The southern part of Madagascar consists of Neoproterozoic sedimentary belts divided by several shear zones, namely the Ampanihy, the Vorokafotra and the Ranotsara shear zone (fig. 3). On the east the Betsileo shear zone marks the boundary of the Antananarivo block (Collins, 2006). The
supracrustals that form belts consist of the Vohibory (~850 Ma), the Androyen (before 630 Ma) and the Molo units (620-560 Ma) (fig. 3). The sediments reflect foreland basin, oceanic/arc and volcanic rift environments in the Neoproterozoic (Collins, 2006; de Wit, 2001). This is consistent with the Itremo group to the north, which was deposited earlier in passive or Atlantic-type continental shelf environment (de Wit, 2001). Between 650 Ma and 610 Ma these units experienced a compressional event in combination with granulite grade metamorphism, which was followed by orogenic collapse after 590 Ma. The shear zones that are present in this group are probably related to a transpressive regime (associated with e.g. flower structures) dated at 590-500 Ma (Martelat et al., 1999) which can probably be linked with the ‘Zone de Virgation’.
4.5 Antongil Block
Lying structurally below the Bemarivo and the Antananarivo Block (Collins, 2000) the Antongil Blocks marks most of the East coast of Madagascar. It is a block that consists of a gneiss and granitic core with metasediments along its western and northern boundary. The ortho- and paragneisses in the core were dated at 3127 Ma and were consequently intruded by granites of ~2522-2495 Ma (Tucker et al., 1999). The rocks are generally low metamorphic, namely of greenschist to lower amphibole grade (Collins, 2000; Tucker et al., 1999) but seem to have undergone no Proterozoic metamorphism like in the rest of Madagascar (Collins, 2000). A remarkable feature which lies in between the Antongil and the Antananarivo Block is a zone of graphitic pelites with podiform harzburgites, chromitites, and emerald deposits This zone has been named the ‘Betsimisaraka suture zone’, because it is believed to be the remnants of the Mozambique ocean, collected during its closure between 630 Ma and 530 Ma (Collins, 2000).
4.6 Bemarivo Belt
2000), where in the west (Sambirano region) it contains migmatized and deformed calc-alkaline gneisses (882Ma) invaded by leucogranites (~560Ma)(Buchwaldt et al., 2002). The metapelites in the Lokoho region are greenschist to granulite grade gneiss (6.5-8.5 kbar and 800-900 ºC) dated at 520-510Ma and are associated with the accretion of an island-arc terrain during the collision between Bemarivo and Malagasy mainland (Buchwaldt & Tucker, 2001; Buchwaldt et al., 2003). The
concerning gneisses in the Sambirano region are older and have indicated a subduction related origin according to their whole-rock composition (Buchwaldt et al., 2002). The rocks in the Sambirano region of the Bemarivo suture zone experienced a greenschist – amphibolite grade metamorphic grade (7 kbar and 850 ºC). To the north and structurally overlying the Sambirano-Sahantana Group are granodioritic gneisses which crystallized at ~750 Ma (Buchwaldt et al., 2003). These orthogneisses are referred to as the Mananbato-Base Group and have been subjected to amphibolite-grade
metamorphism dated at 511 ± 5 Ma by Buchwaldt et al. (2003). Meta-volcanosedimentary rocks,
Figure 4 – Simplified geological map of North Madagascar (Buchwaldt et al., 2003).
emplaced at ~715 Ma, comprise the rest of the Bemarivo and are collectively called the Daraina Group (Collins, 2000).
The emplacement of the pertinent lithologies began with a period of granite magmatism at 754-717 Ma, followed by rhyolite extrusion forming the Daraina Group at 715 Ma and a coeval sandstone and conglomerate deposition in the south. Later deformation was associated with upright isoclinal folds found in the rhyolites (Collins, 2006) and various isoclinal folds in the south-east (Buchwaldt et al., 2003). The tectonic significance of these structures is uncertain, however. Amphibolite to granulite-grade metamorphism and amphibolite-granulite-grade metamorphism at 520 -510 Ma observed, respectively, in the Sambirano-Sahantana metapelites and the Mananbato-Base Group were subsequently linked to a presumably top-to-the south thrusting placing the Bemarivo belt structurally on the Betsimisaraka suture zone, Antananarivo-, and the Antongil Block (Buchwaldt et al., 2003; Collins et al., 2001). This is evident due to the much younger ages found in the Bemarivo belt in comparison with the
Antananarivo- and the Antongil block and due to the rapid cooling rates found, which could only be accomplished by fast tectonic exhumation (Buchwaldt et al., 2003). According to Collins (2006) this granulite-grade metamorphic event is related to the final closure of the Mozambique Ocean.
4.7 Deformation History
The deformational history relevant to this project, is linked to the Gondwana formation. Bringing together tectonic research from different localities in Madagascar, a couple of deformation phases can be recognised to have influenced most, if not all, of Madagascar. An overview of the different events in the different blocks is shown in figure 5. The first significant tectonic event active in a large part of Madagascar occurred around 800-700 Ma and concerned extensive magmatic activity (Meert, 2003). These gabbroic and granitoid intrusions can be found in the Tsaratanana-, Itremo sheet and the Antananarivo Block (Collins, 2000; Collins et al., 2001). The origin of these intrusions is still
questionable. The Itremo sheet and the Antananarivo block show subduction related chemistry which could be linked to arc magmatism due to the subduction of the Mozambique ocean (Collins et al., 2001; Kröner et al., 2000). However, magmatic underplating following the generation of plumes and magmatism related to the break-up of Rodinia have also been suggested as origins (Kröner et al., 2000).
ec et al., 2000; Paquette & Nédélec, 1998). This extensional regime is however not observed to the south and west of the Betsileo shear zone, along which the post-collisional extensional movement was probably localised (Collins, 2000). In the NeoProterozoic supracrustals compressional events, also correlated to the closure of the Mozambique Ocean, took place in a period of 647-609 Ma (de Wit et al., 2001). However, a period of ‘static annealing’ until 530 Ma was suggested to be linked to the orogenic collapse (de Wit et al., 2001). Coeval to the orogenic collapse (between 630 and ~530 Ma), the east of the Antananarivo block experienced contraction due to the collision with the Antongil block, concentrated along the Betsimisaraka suture zone (Collins, 2006).
The last major event concerning the formation of Gondwana is the Kuunga orogeny at 570-530 Ma (Meert, 2003) due to the collision of East Antarctica and Australia with the group of continental blocks named East Gondwana. This event was coeval with high- grade metamorphism in close association with transcurrent movement along N-S shear zones and the formation of the ’Zone de Virgation’ in central and south Madagascar (Kröner et al., 1999; Lardeaux et al., 1999; Martelat et al., 1999; Meert, 2003; Paquette & Nédélec, 1998), and the top-to-the south thrusting of the Bemarivo belt over the Antananarivo and the Antongil block (Collins, 2000; Tucker et al., 1999). The transpressional regime starting at ~530 Ma is observed by the rotated structures in the east and south of Madagascar (Meert, 2003; Paquette & Nédélec, 1998; Kröner et al., 2000) and by positive-flower structures and consequent exhumation prevailing until ~500 Ma in the south (Martelat et al., 1999; Lardeaux et al., 1999). Post orogenic magmatism and cooling from 550 Ma till 490 Ma marks the last Pan-African related activities (Kröner et al., 2000; Meert, 2003; Tucker et al., 1999).
5.
Fieldwork
5.1 Introduction
southwest direction. The fieldwork was carried out during September and October 2005, which coincides with the dry season in North Madagascar. The wet season namely makes this region inaccessible, especially by car. The collection of data in the field was done with the help of a global positioning system (GPS: Garmin), the 1: 100 000 maps, a miniature version of the geological 1: 500 000 map (Besarie, 1970), and additional conventional field equipment.
5.2 Geology and Morphology
Mangindrano lies at 1100m above sea level in the ‘Cuvette de Mangindrano’, which implies that it is situated in a sedimentary basin filled by the supply from the mountains of the ‘Massif de
Tsaratanana’. This immediately describes the main morphology of this area, namely hills and
mountains surrounded by cultivated sedimentary basins. The hills in the area are relatively bare, due to the deforestation for domestic purposes by the local population, and therefore undergo extensive chemical and physical weathering driven by the humid tropical climate prevailing in Madagascar. The
: Sambirano-Sahantana group – Gneiss mica schists : Sambirano-Sahantana group – Quartzites
: Granite stratoids
: Rhyolites and Trachytes (Tertiary) : Basalts (Tertiary)
: Alluvial sediments (Quarternary)
Figure 6 – Detailed geological map of the field research area (Besarie, 1971) with S1 orientations.
weathering, which presumably affects at least the first few meters of surface lithology, has given the hills a slanting appearance, depriving them of any fresh outcrops on its surface. Going further in the ’Massif de Tsaratanana’, the hills become mountains and their appearance becomes more rigid with the vegetation density increasing significantly (cover > 90%). Despite the decrease in physical
weathering, the chemical weathering and the density of the cover inhibits any data to be collected here. The only locations with reasonably fresh outcrops are along the bedding of the numerous rivers and creeks that have incised themselves in the ‘Massif de Tsaratanana’ and drain the basins to its south. The collected data are therefore spatially biased due to the fact that they only capture the geology along the rivers and streams and not within the hills and mountains. The sub-area on the other side of the first mountain ridge to the northeast, did yield some spatial variation but still consisted of data along a river.
5.3 Lithology
The geology of the study area is dominated by granite, gneiss and mylonitic rocks with a numerous amount of leucosomes and deformed and non-deformed granitoid veins and dikes. Despite the fact that different lithologies have been identified, no spatial relationship could be attached to the different types. The gneisses are fine crystalline (± 1 mm) and have a dominantly black and white appearance which is reflected by its mineral composition, namely quartz, feldspar, hornblende and biotite. Dependent on the variation of the ratio of light coloured quartz and feldspar to the dark coloured biotite and hornblende, the gneisses have various colours. The mylonites have been observed at several locations (fig. 7a) and are easily recognised by their ductile appearance and strong foliation in
combination with a very fine grained structure. The mylonites do not seem to have a different mineral composition and are therefore only structurally different.
The granitoid veins and dikes (fig. 7b) can have a thickness of a couple of cm to couple of meters and have a strong granular appearance (grain size ≥ 5mm). Its mineral assemblage consist of
The granites in the area, which are found at the mountain scarps, have the same mineral
composition as the gneisses, namely quartz, feldspar, biotite and hornblende, but hardly contain any planar fabrics or other deformational structures. These granites are referred to as rhyolites/trachytes of Tertiary age (fig. 6) by Besarie (1971).
5.4 Petrology
By microscopic analysis of the samples gathered from the field, the orthogneisses have been
divided into three sub-groups on the basis of the mineral assemblages. The first group mainly contains:
a b
G S1
S1
Figure 7 – Different structures found in the research area:(a) mylonitic outcrop describing S1; (b) granitoids (G) cutting through gneiss rock and its S1; (c) folded leucosome (Leuc); (d) granitoids (G) clearly cutting through foliation (S1).
c d
Leuc
G
Hb + Plag + Qtz + Bt (1)
Hornblende (30 – 75%), plagioclase (25 – 40%), and quartz (10 – 30%) are dominantly present where biotite is less abundant (< 10%). Further accessory minerals are apatite, titanite and very little
microcline (< 5%). The second group has a mineral assemblage as follows:
Plag + Qtz + Bt + Hb (2)
This group is especially characterised by the absence or the very little presence of hornblende (<20%) compared to the other samples. The mineral assemblage further contains: plagioclase (30 – 50%), quartz (15 – 35%) and biotite (10 – 35%), in combination with minor presence of microcline, titanite and apatite (< 5%). The last group has the mineral composition of:
Hb + Plag + Qtz + Bt + Ep + Chl (3)
This assemblage is specially marked by the presence of epidote (< 30%) and chlorite (< 5%) in small to moderate amounts in combination with hornblende (20 – 75%), plagioclase (10 – 40%), quartz (10 – 40%) and in lesser amounts biotite (< 15%). Further minerals that were observed in small amounts are microcline (< 5%), titanite (< 5%) and apatite (< 15%).
Figure 8 - Typical microscopic textures found in the samples; a) alignment of the hornblende (Hb) and biotite crystals (Bt) (19.6); also note the clear angles between the crosscutting biotite crystals; b) ‘foam texture’ of relatively small plagioclase and quartz crystals (12.9).
Qtz Bt
Plag
Hb
Concerning the spatial distribution of these three mineral assemblage, no relationship could be found especially with regard to the epidote and chlorite containing rocks. The hornblende-bearing and the hornblende-free rocks seem to show a clustering along the main river and its distributaries lying to the north-east of Mangindrano. However, the dataset is not large enough to ascertain this distribution.
Apart from the mineral assemblages in the orthogneisses it was clear that the hornblende and biotite crystals have a preferred orientation (figure 8a), but do not define a clear foliation. The biotite crystals have an angle with respect to each other (figure 8a), which could indicate a recrystallised biotite fold. The crystals in the plagioclase-quartz-rich zones of the orthogneisses often have a smaller grain size and the plagioclase and quartz crystals often form a ‘foam’ texture (figure 8b). In both the orthogneisses as the granitoid veins and dykes quartz crystals show undulose extinction (figure 9a) and some plagioclase crystals have fanning twins (figure 9b).
During the fieldwork in the area of Mangindrano not only lithological and petrological data was gathered, but also a lot of structural data. These data will not be presented and discussed in this report, but in the report of Mulder (2006).
Plag
Qtz
Figure 9 – Detailed microscopic pictures of structures within minerals; a) undulose extinction within quartz crystals (sample 12.9); b) fanning twins in plagioclase (sample 19.6).
A B
Plag Bt
6.
Geothermobarometry
Geothermobarometry is the name given to a set of techniques in which the pressure and temperature dependence of the equilibrium constant is used to infer metamorphic pressures and temperatures.
Due to the fact that no garnets are present in the samples, limited barometers and thermometers can be used to reconstruct the P-T conditions. In this research the aluminium-in-hornblende barometer of Johnson and Rutherford (1989) and Schmidt (1992) is used in combination with the amphibole-plagioclase thermometer of Holland and Blundy (1994). Mineral compositions were determined by analyzing six carbon-coated, polished thin sections on the JEOL JXA-8600 microprobe at Utrecht University with a maximum lateral spreading in the field research area. Operating conditions were 15 kV accelerating potential and 10 nA beam current. A focused beam diameter of <1μm was used for all minerals. Standard 10-element analyses were performed on all minerals. SiO2, Al2O3, FeO, CaO, MgO, and K2O were analyzed on the energy dispersive X-ray spectrometer and TiO2, Na2O, Cr2O3, and MnO were analyzed on the wavelength dispersive X-ray spectrometer. Core and rim analyses were made by doing multiple track analyses across amphibole and plagioclase crystals per sample.
6.1 Aluminum-in-hornblende barometer
Based on empirical evidence, Hammarstrom and Zen (1986) proposed that in calcalkaline granitic plutons, tonalites and granodiorites, the total aluminum content of hornblende (A1tot = Aliv+ AIvi) can be used as a geobarometer.
P=-3.46 (±0.24) + 4.23 (±0.13) (Altot) (1)
, where Altot is the total Al content of atoms per formula unit (apfu). The total Al content of the hornblendes is independent of the ferric/ferrous ratio within the hornblende crystals. The maximum error is ±0.6 kbar at 8 kbar and ±0.5 kbar at lower pressures. The validity of the geobarometer critically depends upon hornblende being in equilibrium with quartz. Hornblendes not in equilibrium with quartz will yield anomalously high pressures (Johnson and Rutherford, 1989).
Johnson and Rutherford (1989) concluded from experimental hornblendes that TiO2, CaO, Na2O,
K2O and MnO do not vary systematically with increasing pressure, whereas both Aliv and Alvi increase
linearly. The increasing Aliv and Alvi content implies that a Tschermak-type substitution (Mg + Si ↔
Aliv + Alvi) is controlling the pressure sensitive component of the hornblende.
However, Schmidt (1992) stated that the Tschermak-type substitution is accompanied by minor plagioclase substitution (Ca + Aliv ↔ NaM(4) + Si) and edenite-exchange (□A + Si ↔ NaA + Aliv).
Edenite exchange increases at elevated temperatures, which causesthe calibration of the barometer of
Johnson and Rutherford (1989), which was done at higher temperatures, to yield higher Altot content in hornblendes at corresponding pressures than the calibration of the barometer of Schmidt (1992), who
Fig 10 - Altot in hornblende as function of pressure. Lines of all calibrations, in which
hornblende is in equilibrium with biotite + plagioclase + orthoclase + quartz + sphene + Fe-Ti-oxide + melt + vapor, are drawn for the pressure range of each calibration. Solid lines are the field calibrations by Hammarstrom and Zen (1986) and Hollister et al. (1987), error bars are _+ 3 kbar and + 1 kbar respectively. Dotted lines are the experimental calibrations
established the following relation between pressure and total Al content in hornblendes, barometer B:
P(±0.6 kbar) = -3.01 + 4.76Altot r2 = 0.99 (2)
, where Altot is the total Al content of atoms per formula unit (apfu). This relation was calibrated at temperatures approximately 90°C lower than the calibration of Johnson and Rutherford. (~650-690°C). Figure 10 indicates the different calibrations of the aluminium-in-hornblende barometer. The
calibration of Johnson and Rutherford (1989) yields higher Altot at corresponding pressures (ΔAltot2kbar
= +0.24 apfu, ΔAltot8kbar = +0.40 apfu), or lower pressures at corresponding Altot (ΔPAl(tot)= 1.3 = 1.1kbar,
ΔPAl(tot)= 2.7 = 2.1 kbar), respectively. Therefore, the temperature during the crystallisation of the
hornblende crystals has to be known to reduce the error of the pressure determination.
However, Hammarstrom and Zen (1986) and subsequent workers calibrated this geobarometer for igneous rocks, to determine their crystallization depth, and not for metamorphic rocks, to determine the peak metamorphic pressure. However, the microscope observations in the thin sections (figure..) and their migmatic appearance in the field indicate a (semi-) complete recrystallization of the orthogneisses during peak metamorphism. According to Schmidt (1992) hornblende equilibration continues until and ceases at the wet solidus. Moreover, experimental studies (Apted and Liou 1983; Spear 1981) showed that amphibole in fluid-saturated systems is quite reactive also in the subsolidus range. Therefore, these geobarometers can also be applied to the orthogneisses addressed in this study to determine the peak pressure during metamorphism, especially when track analyses across a
hornblende crystal can indicate the presence of Altot zonation in the crystal.However, the pressure calculated from hornblende composition is indirectly also a function of fluid composition as has been shown by comparison with the experiments of Johnson and Rutherford (1989b), which should be considered when this barometer is used. It is, however, beyond the scope of this research to take this effect into account.
6.2 Amphibole – Plagioclase thermometer
Blundy and Holland (1990) developed an amphibole-plagioclase thermometer based on the following reaction:
NaCa2Mg5Si4(AlSi4)O22(OH)2 + 4 SiO2 Ca2Mg5Si8O22(OH)2 + NaAlSi3O8 (A)
This thermometer was calibrated by using silica-saturated and silica-rich materials. Blundy and Holland (1990) assumed a straightforward ideal mixing-on-sites model for amphibole and use
Darken’s quadratic formalism to describe non-ideality in plagioclase. This work was criticised by Poli and Schmidt (1992) for its over-simplicity. Therefore, Holland and Blundy (1994) reviewed their ‘old’ geothermometer by using a compositionally wider data set and found that 1) non-ideal mixing takes place in amphiboles and that 2) non-zero parameters for Na-K-□ (where □ represents a site vacancy) on the A-site of amphiboles can be expected.
Because of the complexity of multi-site solid solutions in amphibole, Holland and Blundy (1994) adopted the simplest (symmetrical) form of non-ideal interactions, which restricts the discussion to a
relatively simple chemical system involving the distribution of the species □-K-Na-Ca-Mg-Fe2+-Fe3+
-Al-Si over the A, M4, M3, M2 and T1 amphibole chrystallographic sites (fig. 11 and table 1). Mn and Ti are ignored because they appear to be insignificant as far as their effects on the thermometer are concerned. Further the substitution of Mg and Fe into the M-site is ignored, which also show no significant effect on the hornblende-plagioclase thermometer. Moreover, with current limitations on estimation of Fe3+ from microprobe data, the M4 content of Fe and Mg are extremely imprecise. However, Holland and Blundy (1994) devised a recalculation scheme for ferric iron recalculation within amphiboles, which was used in this research to estimate the Fe3+- content in amphiboles. Finally, the substitution of halogen ions (F- and Cl-) and O2- for OH- on the O(3)-site is ignored,
because not only are these components of minor importance in most natural amphiboles, there are insufficient data with which to assess the importance of O(3)-site substitutions on amphibole stability (Holland and Blundy, 1994).
The simplest set is that of non-ideal interaction assumptions of symmetrical pairwise interactions between atoms on sites (regular solution on each site together with Bragg-Williams cross-site terms) coupled with ideal configuration entropy of mixing. This concept was, therefore, used by Holland and Blundy (1994). This model is not a complete thermodynamic description of amphibole non-ideality and is merely a simplification useful for the present thermometric purposes.
Although there are 24 independent cross-site interactions energies and 12 independent regular solution parameters, leading to a maximum of 36 possible terms to be considered. In a chemically balanced reaction many terms cancel while others occur only in certain linear combinations. Of the 36 mixing parameters all but 16 are lost through cancellation, and these
Table 1 – Simple scheme of amphibole site allocations adopted in the geothermometer of Holland and Blundy (1994)
Site Multiplicity Elements
A 1 Na, K, □ M4 2 Na, Ca M1,3* 3 Fe(2+), Mg M2 2 Fe(2+), Mg, Al, Fe (3+) T1 4 Al, Si T2 4 Si
remaining 16 appear in only 8 linearly independent combinations (Holland and Blundy, 1994). Therefore, within the framework of the simple model, only the possible effects of these 8 non-ideal interaction terms on the equilibrium constant (KA) for reaction A have to be determined. This leads to the following ‘new’ expression of the edenite-tremolite thermometer (with quartz) by Holland and Blundy (1994):
T is the temperature in degrees Kelvin, P is the pressure in kbar and the XΦi denote the molar fraction
of species (or component) i in phase (or crystallographic site) Φ. The temperature uncertainty for thermometer B is stated as ± 35°C, which may enlarge for iron-rich amphiboles with very different oxidation states from those of the calibrant set. The use of this thermometer is restricted to the
following conditions: T in the range 400-900°C, amphiboles which have NaA > 0.02pfu, Aliv < 1.8 pfu
and Si in the range 6.0-7.7 pfu, and plagioclase with Xan < 0.90.
Thermometer A is based on the edenite exchange vector (□A + Si ↔ NaA + Aliv). The exchange of
the plagioclase vector component (NaA + Si ↔ CaA + Aliv) between hornblende and plagioclase has also been suggested as the basis of a thermometer by Spear (1980) who expressed this equilibrium as:
Ca2Mg3Al2Si4(Al2Si2)O22(OH)2 + 2 NaAlSi3O8 = Na2Mg3Al2Si8O22(OH)2 + 2 CaAl2Si2O8 (B)
Tschermakite albite glaucophane anorthite
This reaction can also be written in terms of the alternative endmembers:
NaCa2Mg5(AlSi3)Si4O22(OH)2 + NaAlSi3O8 = Na(CaNa)Mg5Si8O22(OH)2 + CaAl2Si2O8 (C)
Edenite albite richterite anorthite
Holland and Blundy (1994) preferred this version, because it involves only one molecule of plagioclase exchange per molecule of amphibole while Spear’s formulation involves two. A thermometer based on this reaction is likely to be useful at high temperatures where anorthite-rich feldspars predominate and/or under conditions of low silica activity (no quartz present in this reaction) , but will be of little value in low temperature metamorphic rocks containing albitic plagioclase. The edenite-richterite thermometer by Holland and Blundy (1994) may be written formally as:
, where the values for Yab term is given by: for Xab > 0.5 then Yab = 0
Otherwise Yab = 12.0 (1 - Xab)2 – 3.0 kJ
T is the temperature in degrees Kelvin, P is the pressure in kbar and the XΦi denote the molar fraction
of species (or component) i in phase (or crystallographic site) Φ. The temperature uncertainty for thermometer B is stated as ± 39°C, but this value may be larger for iron-rich amphiboles with oxidation states very different from those of the calibrant set.
According to Holland and Blundy (1994), the slightly poorer performance of thermometer B relative to A most probably reflects the greater sensitivity of B to the site fraction allocations, which in turn are sensitive to the ferric iron recalculation procedure. Thermometer B may be useful in
conjunction with the thermometer A in hornblende-plagioclase parageneses. Thermometer B should only be used in the range 500-900°C and under the following compositional restrictions: plagioclase feldspars must lie in the range Xan > 0.1 and <0.9, amphiboles must have XNaM4> 0.03, Aliv < 1.8 pfu,
and Si in the range 6.0-7.7 pfu.
Moreover, Holland and Blundy (1994) emphasize that any temptation to use the intersection of the two thermometers as a barometric indicator should be avoided as the coefficients were not optimised on the basis of minimising pressure residuals.
, where the Yab-an term is given by: for Xab > 0.5 then Yab-an = 3.0 kJ
Otherwise Yab-an = 12.0 (2Xab – 1) + 3.0 kJ
(4)
6.3 Mineral composition
The mineral compositions of the six samples used for thermobarometry are those of group (1) and group (3) of the petrology subdivision made earlier. Sample 3.3 and 7.5-2 belong to group (1) and contain no epidote nor chlorite.
Sample 2.2E, 9.2b, 14.2 and 15.2a belong to group (3) and these samples do contain epidote and/or chlorite. The locations of the samples are indicated in figure 6. The hornblende and plagioclase crystals in all of the samples show hardly any zonation, as illustrated for sample 2.2E in figure 12. However, figure 12 only shows the trends of the elements, which are used for the baro- and thermometers. Other elements, especially K and Ti, often show zonation in hornblende (fig. 13a), which can sometimes indicate two sub-zonations within a hornblende crystal (fig. 13b).
By looking at the phase assemblage of the different samples (table 2), only three samples (2.2E, 7.5-2 and 14.2) agree with the phase assemblage needed for the application of the aluminium-in-hornblende barometer. So, only these three samples can be used for geobarometry. Table 2 also indicates that all samples have an anorthite percentage which corresponds to Oligoclase/Andesine,
except for sample 15.2a which has an anomalous high anorthite percentage corresponding with Labradorite/Bytownite (table 3-7).
Moreover, two different types of hornblende were found in sample 9.2b and 15.2a. One of those types is equivalent to the hornblende crystals in the other samples and will be called Hba. The other type of
hornblende is relatively poor in Al and Fe, and relatively rich in Si, Ca and Mg (table 5 and 7). This type will be called Hbb. Especially the
high amount of Ca present in the Hbb is very
distinct. This is such a high amount that not only the M4 site, but also M1-3 and A-site are partially occupied by Ca atoms, because the
0.00 0.50 1.00 1.50 2.00 2.50 3.00 0 78 140 202 264 326 388 450 512 μm XAl XFe XCa XMg Plagioclase
Figure 12 – Trends for XAl, XFe, XCa and XMg witin a hornblende and plagioclase
crystal in sample 2.2E. Hornblende 0.000 0.002 0.004 0.006 0.008 0.010 0.012 0.014 0.016 0.018 0.020 0 31 61 92 122 153 183 214 μm XK 0.00 0.02 0.04 0.06 0.08 0.10 0.12 0.14 0 31 61 92 122 153 183 214 244 μm XTi (a) ) (b)
Figure 13 – (a) Trend for XK in a plagioclase crystal in sample 2.2E; (b) Trend
Table 2 - Phase assemblage of samples analysed for thermobarometry
Sample Hb Plag Anorthite-% Qtz Bt Sphene Fe-Ti oxide
2.2-E 60 20 An-30% 5 <1 5 <1 3.2 45 35 An-26% 5 0 1-5 0 7.5-2 20 50 An-30% 10 10 <1 5 9.2b 20 50 An-30% 10 0 0 10 14.2 30 40 An-38% 20 5 1 10 15.2a 75 10 An-84% 10 0 1 1 Table 3
Representative results of electron microprobe analyses of hornblende and plagioclase in sample 2.2E
2.2E Hb-2.1 Hb-3.1 Hb-3.2 Hb-4.1 Hb-4.2 Pl- 2 Pl -3 Pl -4 SiO2 42.46 41.08 42.49 43.56 43.82 59.84 61.11 61.15 TiO2 1.21 1.26 0.98 1.13 1.10 - - 0.01 Al2O3 9.91 10.07 9.31 9.32 9.12 24.46 23.95 23.86 FeO 18.67 17.44 18.12 17.49 17.51 0.09 0.19 0.17 MnO 0.34 0.26 0.33 0.29 0.28 0.04 - 0.05 MgO 9.85 9.92 10.53 10.74 10.91 0.02 - - CaO 11.61 11.63 11.41 11.59 11.75 6.45 5.87 5.61 Na2O 1.38 1.13 1.14 1.32 1.27 7.29 8.14 7.96 K2O 1.14 1.15 0.95 1.01 0.93 0.31 0.28 0.32 Cr2O3 0.02 0.10 0.04 0.06 0.03 - - - Total 96.59 94.05 95.28 96.50 96.73 98.51 99.56 99.12 Si 6.538 6.477 6.603 6.653 6.674 2.701 2.730 2.739 Ti 0.140 0.150 0.114 0.130 0.126 - - - Al 1.798 1.871 1.705 1.677 1.636 1.301 1.261 1.259 Fe2+ 2.404 2.300 2.354 2.235 2.230 - 0.007 0.006 Mn 0.044 0.035 0.043 0.038 0.037 - - - Mg 2.261 2.331 2.438 2.446 2.478 - - - Ca 1.916 1.964 1.900 1.897 1.917 0.312 0.281 0.269 Na 0.413 0.345 0.344 0.390 0.375 0.638 0.705 0.691 K 0.224 0.231 0.188 0.197 0.181 0.018 0.016 0.019 Cr - 0.013 - 0.007 - - - - Total 15.740 15.719 15.694 15.669 15.658 4.976 5.000 4.986 XFe 0.515 0.497 0.491 0.477 0.474 Xan 0.322 0.281 0.275
Table 4
Representative results of electron microprobe analyses of hornblende and plagioclase in sample 3.2
3.2 Hb-1.1 Hb-1.2 Hb-2.1 Hb-2.2 Pl-1.2 Pl-1.3 Pl-2.1 Pl-2.2 SiO2 43.84 43.87 44.07 44.85 61.52 61.12 61.27 62.13 TiO2 0.59 0.89 0.95 0.81 0.01 - 0.01 0.01 Al2O3 9.77 9.56 9.65 9.10 24.30 24.16 23.78 24.27 FeO 15.19 15.12 15.07 14.66 0.18 0.22 0.18 0.10 MnO 0.41 0.40 0.37 0.40 - 0.01 0.02 0.02 MgO 11.99 12.11 12.05 12.52 - - 0.01 - CaO 11.96 11.81 11.70 11.88 5.33 5.31 5.23 5.37 Na2O 1.30 1.45 1.60 1.48 8.31 7.96 8.33 8.17 K2O 1.11 1.08 1.12 0.99 0.21 0.14 0.20 0.27 Cr2O3 0.10 0.12 0.11 0.03 0.01 - - - Total 96.28 96.40 96.70 96.73 99.87 98.93 99.03 100.34 Si 6.648 6.643 6.649 6.737 2.733 2.737 2.745 2.744 Ti 0.068 0.102 0.108 0.092 - - - - Al 1.746 1.706 1.716 1.611 1.272 1.275 1.255 1.263 Fe2+ 1.926 1.914 1.901 1.842 0.007 0.008 0.007 0.004 Mn 0.053 0.051 0.048 0.050 - 0.001 0.001 0.001 Mg 2.710 2.732 2.710 2.805 - - - - Ca 1.943 1.915 1.892 1.912 0.254 0.255 0.251 0.254 Na 0.383 0.425 0.469 0.432 0.716 0.691 0.724 0.700 K 0.215 0.209 0.215 0.190 0.012 0.008 0.011 0.015 Cr 0.012 0.015 0.013 0.004 - - - - Total 15.705 15.712 15.721 15.675 4.994 4.975 4.995 4.982 XFe 0.415 0.412 0.396 0.396 Xan 0.258 0.267 0.255 0.262
For location see fig 6. Oxygenbasis of the analysis: Hb 23; Pl 8. XFe = Fe/(Fe + Mg); XAn = Ca/(Na + K + Ca)
amount of Ca atoms in the unit-formula of these hornblende crystals exceeds the known maximum of 2
atoms. Figure 14 shows that the Hbb crystals in sample 9.2b are much more broken and unstable than
the Hba crystals. The same figure also shows that near the hornblende crystals in sample 9.2b, and
especially near the Hba crystals, large amounts of FeO are present. In sample 15.2a, the Hbb crystals
even have a higher Si content (wt%) than the plagioclase crystals and the Hba crystals have Si contents
Table 4
Representative results of electron microprobe analyses of hornblende and plagioclase in sample 3.2
3.2 Hb-1.1 Hb-1.2 Hb-2.1 Hb-2.2 Pl-1.2 Pl-1.3 Pl-2.1 Pl-2.2 SiO2 43.84 43.87 44.07 44.85 61.52 61.12 61.27 62.13 TiO2 0.59 0.89 0.95 0.81 0.01 - 0.01 0.01 Al2O3 9.77 9.56 9.65 9.10 24.30 24.16 23.78 24.27 FeO 15.19 15.12 15.07 14.66 0.18 0.22 0.18 0.10 MnO 0.41 0.40 0.37 0.40 - 0.01 0.02 0.02 MgO 11.99 12.11 12.05 12.52 - - 0.01 - CaO 11.96 11.81 11.70 11.88 5.33 5.31 5.23 5.37 Na2O 1.30 1.45 1.60 1.48 8.31 7.96 8.33 8.17 K2O 1.11 1.08 1.12 0.99 0.21 0.14 0.20 0.27 Cr2O3 0.10 0.12 0.11 0.03 0.01 - - - Total 96.28 96.40 96.70 96.73 99.87 98.93 99.03 100.34 Si 6.648 6.643 6.649 6.737 2.733 2.737 2.745 2.744 Ti 0.068 0.102 0.108 0.092 - - - - Al 1.746 1.706 1.716 1.611 1.272 1.275 1.255 1.263 Fe2+ 1.926 1.914 1.901 1.842 0.007 0.008 0.007 0.004 Mn 0.053 0.051 0.048 0.050 - 0.001 0.001 0.001 Mg 2.710 2.732 2.710 2.805 - - - - Ca 1.943 1.915 1.892 1.912 0.254 0.255 0.251 0.254 Na 0.383 0.425 0.469 0.432 0.716 0.691 0.724 0.700 K 0.215 0.209 0.215 0.190 0.012 0.008 0.011 0.015 Cr 0.012 0.015 0.013 0.004 - - - - Total 15.705 15.712 15.721 15.675 4.994 4.975 4.995 4.982 XFe 0.415 0.412 0.396 0.396 Xan 0.258 0.267 0.255 0.262
For location see fig 6. Oxygenbasis of the analysis: Hb 23; Pl 8. XFe = Fe/(Fe + Mg); XAn = Ca/(Na + K + Ca)
6.4 Thermobarometric results
Table 6
Representative results of electron microprobe analyses of hornblende and plagioclase in sample 14.2
14.2 Hb-1.2 Hb-2.2 Hb-3.2 Hb-3.3 Pl-1 Pl-2.1 Pl-3.1 Pl-3.4 SiO2 42.36 41.72 42.56 42.54 58.71 58.61 58.82 58.77 TiO2 0.96 1.06 0.94 0.89 - - - - Al2O3 11.05 11.46 11.12 11.12 25.76 26.06 26.41 26.27 FeO 18.98 19.05 18.59 18.43 0.23 0.10 0.15 0.18 MnO 0.28 0.38 0.36 0.34 0.02 0.02 - 0.05 MgO 9.32 9.24 9.50 9.74 - 0.01 - - CaO 11.58 11.49 11.44 11.56 7.70 7.75 7.68 7.73 Na2O 1.26 1.32 1.30 1.19 6.60 6.91 7.01 6.98 K2O 0.70 0.76 0.68 0.61 0.13 0.17 0.14 0.05 Cr2O3 - - 0.01 - - - - 0.01 Total 96.49 96.49 96.49 96.41 99.15 99.62 100.21 100.05 Si 6.505 6.423 6.519 6.514 2.640 2.627 2.621 2.623 Ti 0.111 0.123 0.108 0.103 - - - - Al 2.000 2.079 2.008 2.006 1.365 1.377 1.387 1.382 Fe2+ 2.437 2.453 2.382 2.360 0.008 0.004 0.006 0.007 Mn 0.037 0.050 0.046 0.045 0.001 0.001 - 0.002 Mg 2.133 2.119 2.169 2.222 - 0.001 - - Ca 1.906 1.895 1.877 1.896 0.371 0.372 0.367 0.370 Na 0.376 0.394 0.385 0.353 0.576 0.600 0.605 0.604 K 0.138 0.149 0.132 0.119 0.008 0.010 0.008 0.003 Cr - - 0.002 - - - - - Total 15.641 15.686 15.627 15.616 4.969 4.990 4.993 4.990 XFe 0.533 0.536 0.515 0.515 Xan 0.389 0.379 0.374 0.378
For location see fig 6. Oxygenbasis of the analysis: Hb 23; Pl 8. XFe = Fe/(Fe + Mg); XAn = Ca/(Na + K + Ca)
uncertainty of the barometers. Therefore, the average pressure of sample 7.5-2 can also be used for thermometry. Sample 14.2 yields an average pressure of 5.13 ± 0.12 kbar with the barometer of Johnson and Rutherford (1989) and an average pressure of 6.65 ± 0.14 kbar with the barometer of Schmidt (1992) (table 8). This sample also hardly shows any zonation, which justifies the use of the average pressure for the thermometry. (fig. 15c). The errors of the pressure measurements are the mean deviations of the reported results, in which the uncertainty of the used barometer is not incorporated. When the uncertainty of the barometer is taken into account, the errors would be 0.5-0.6 kbar higher
Table 7
Representative results of electron microprobe analyses of hornblende and plagioclase in sample 15.2a
15.2a Hba-1. Hbb-1 Hba-2.2 Hbb-2 Pl-1.1 Pl-2 SiO2 46.94 53.02 47.26 53.22 47.42 46.60 TiO2 0.44 0.07 0.46 0.04 0.01 0.01 Al2O3 9.79 1.28 8.21 0.98 33.70 34.15 FeO 13.67 7.18 13.59 7.62 0.08 0.18 MnO 0.26 0.28 0.25 0.29 0.04 - MgO 13.33 13.44 13.31 13.58 - - CaO 11.73 24.06 12.18 23.98 16.93 17.23 Na2O 0.89 0.34 1.00 0.34 1.88 1.57 K2O 0.25 - 0.22 - 0.01 0.01 Cr2O3 0.01 0.04 0.03 - - - Total 97.31 99.72 96.51 100.06 100.05 99.74 Si 6.878 7.582 6.997 7.597 2.174 2.146 Ti 0.049 0.007 0.052 0.004 - - Al 1.690 0.216 1.433 0.165 1.821 1.854 Fe2+ 1.675 0.859 1.683 0.909 0.003 0.007 Mn 0.033 0.034 0.031 0.036 0.002 - Mg 2.911 2.865 2.937 2.890 - - Ca 1.842 3.686 1.932 3.668 0.832 0.850 Na 0.252 0.095 0.288 0.095 0.167 0.140 K 0.046 - 0.042 - - - Cr 0.001 0.005 0.004 - - - Total 15.377 15.348 15.398 15.364 4.999 4.997 XFe 0.365 0.231 0.364 0.239 Xan 0.833 0.859
For location see fig 6. Oxygenbasis of the analysis: Hb 23; Pl 8. XFe = Fe/(Fe + Mg); XAn = Ca/(Na + K + Ca)
Moreover, the pressure estimates of the different barometers (Johnson & Rutherford - 1989 vs. Schmidt - 1992) also vary significantly within each sample, beyond the uncertainty interval, indicating that thermometry is needed to determine which barometer should be used for each sample. However, the thermometers constructed by Holland and Blundy (1994) also need a pressure to determine the temperature (see equation 3 and 4).
Due to the compositional restrictions for the thermometers of Holland & Blundy (1994),
can be used to determine, which barometer can be used for each sample. Sample 2.2E and 15.2a are situated in the temperature interval (740-780 ºC) of the barometer of Johnson and Rutherford (1989). Sample 3.2 and 7.5-2 are situated in between the temperature intervals of both barometers. However, the temperature for sample 3.2 is closer to the temperature interval (650-690ºC) of the barometer of Schmidt (1992) and sample 7.5-2 is closer to the
temperature interval of the barometer of Johnson and Rutherford (1989) . Table 9 shows which PT estimates for each sample are used in this research. Figure 16 shows the PT conditions of the different samples in a PT-diagram, with their error bars.
Table 8
Pressure results for the orthogneisses with the barometers of Johson and Rutherford (1989) and Schmidt (1992).
Sample Pressure (kbar) Pressure (kbar)
Johnson & Rutherford (1989) Schmidt (1992)
2.2E (n=77) 4.19 ± 0.24 5.33 ± 0.45
7.5-2 (n= 13) 4.28 ± 0.18 5.69 ± 0.21
14.2 (n=56) 5.13 ± 0.17 6.65 ± 0.19
Errors are the mean deviations of reported results; n is the number of spot analyses in hornblende crystals Table 9
Temperature results for the orthogneisses with the Holland and Blundy (1994) thermometer A
Sample Pressure (kbar) Temperature (°C)
2.2E (n=9) A 4.19 743 ± 20 B 5.33 734 ± 20 3.2 (n=12) A 4.25 702 ± 9 B 5.51 695 ± 9 7.5-2 (n= 6) A 4.28 725 ± 10 B 5.69 716 ± 9 15.2a (n=11) A 5.13 776 ± 15 B 6.65 765 ± 16
Errors are the mean deviations of reported results; n is the number of 'combined' spot analyses in hornblende and plagioclase; A= barometer of Johnson & Rutherford (1989); B = barometer of Schmidt (1992); the pressure used for sample 3.2 is an average of sample 2.2E and 7.5-2; dark rows are the PT results that will be used in this research
Hba Hbb Hba Hbb Plag FeO
a) 0.00 1.00 2.00 3.00 4.00 5.00 6.00 7.00 0 13 26 40 53 66 79 92 106 119 Distance (μm) P re ss u re k b a r )
Johnson and Rutherford (1989) Schmidt (1992) b) 0.00 1.00 2.00 3.00 4.00 5.00 6.00 7.00 0 13 26 39 52 65 78 91 104 118 131 144 157 Distance (μm) P re ss u re i n k b a r )
Johnson & Rutherford (1989) Schmidt (1992) c) 0.00 1.00 2.00 3.00 4.00 5.00 6.00 7.00 8.00 0 20 40 60 81 101 Distance (μm) P re ss u re i n k b a r )
Johnson and Rutherford (1989) Schmidt (1992)
7.
Discussion
7.1 Petrology
The mineral assemblage of the orthogneisses in the Mangindrano area, which contains hornblende and plagioclase, clearly indicates amphibolite facies metamorphism. However, several samples also contain epidote and chlorite crystals, pointing to greenschist facies metamorphism. Concerning the spatial distribution of the three mineral assemblages, especially with regard to the epidote and chlorite containing rocks, no relationship could be found. The hornblende containing- and the hornblende-free rocks seem to show a clustering along the main river and its distributaries north-east of Mangindrano. However, the dataset is not large enough to use these findings for further discussion.
Furthermore, the absence of a clear foliation, the presence of recrystallised biotite folds, the ‘foam’ texture of the minerals and the smaller plagioclase and quartz in certain zones of the orthogneisses in combination with the migmatic character of the orthogneisses and the leucosomes indicate that these rocks experienced a time of (partial) recrystallization. The recrystallized minerals hardly show any strain, which is supported by the fact that hardly any foliation (on microscopic scale) is observed. Some recrystallized minerals in the orthogneisses do show small strains (figure 9a and b). Even some quartz and plagioclase crystals in the granitoid veins and dykes show a little amount of strain.
However, a lot of (outcrop-scale) deformation features were observed in the field (Mulder, 2006) and recrystallised biotite-folds are seen on a microscopic scale, indicating that the orthogneisses
experienced a significant amount of deformation before (partial) recrystallization took place. This indicates that the (partial) recrystallization of the orthogneisses and the intrusion of the granitoid veins and dykes post-dated the major deformation events described by Mulder (2006). The conclusion that the granitoids postdate the major deformation events is supported by field observations (figure 7 b and d).
According to the geological maps, originally by Besarie (1971) and later modified and further specified by Buchwaldt et al. (2003), the study area is situated in the central part of the Sambirano – Sahantana Group, which is assumed to be the suture zone between the Bemarivo belt and the
km2), which makes it most likely for these gneisses to be orthogneisses. A more probable link can be made with the western part of the Sambirano-Sahantana Group, namely the tonalitic to granodioritic gneisses, which yield two different ages 882Ma and 580Ma, and dikes/sheets of pink leucrogranites (~560Ma) cutting the host lithology (Buchwaldt et al., 2002). This is especially remarkable because the pertinent pink leucogranites seem very similar to the pink granitoid veins and dykes found in the Mangindrano area. It is not clear from the literature where the transition between the east and the west part of the suture zone is located, however, the lithologies examined by Buchwaldt et al. (2002) were found along the Sambirano and the Mahavavy rivers which lie approximately 10-20 km to the
northwest of the Mangindrano research area. However, Buchwaldt (2002) states that the whole-rock composition of the gneisses in the west indicate a subduction related origin. This does not agree with the absence, in the gneisses near Mangindrano, of needle-shaped amphiboles, which is a typical feature in subduction related igneous rocks. Due to recrystallization the signature of subduction could have been erased from the geologic record. However, no whole rock composition analyses were done to verify the igneous origin. Another possibility, which is less likely but cannot be excluded at this stage, is that the orthogneisses could belong to the rocks of the Mananbato-Base Group. Despite the fact that no leucrogranites or granitoids are found in the Mananbato-Base Group, this group does contain similar granodioritic gneisses of amphibolite facies metamorphic grade extruded at ~750Ma and subsequently metamorphosed at 511± 5 Ma (Buchwaldt et al., 2003). However, according to the accepted geological map of Besarie (1971) and Buchwaldt et al. (2003) these granodiorites lie at least 20 km to the north of the research area around Mangindrano. Dating of the orthogneisses may provide a strong link with one of the formations, however, Mulder (2006) was not able to date the monazites in
0 1 2 3 4 5 6 7 8 9 400 500 600 700 800 900 Temperature (°C) P re ss u re i n k b a r ) 2.2E 3.2 7.5-2 15.2a
Figure 16 – This pressure-temperature (PT) diagram shows the PT-conditions of the different samples of this research with their error bars. The kyanite, andalusite and sillimanite stability fields are indicated in the diagram for reference.
Andalusite
the orthogneisses and the granitoid veins and dykes in the Mangindrano area. If the orthogneisses in the research area would belong to the Mananbato-Base Group, this would have a major influence on the location of the Bemarivo suture zone. According to the current geological maps of North
Madagascar (Besarie, 1971; Buchwaldt et al., 2003) the sutue zone is defined by the Sambirano-Sahantana Group. If the orthogneisses found in this research are part of the Mananbato-Base Group then this would affect the boundaries of the contact zone. However, linking them to the western part of the Sambirano-Sahantana Group, which is most likely, would only alter the subdivision within the suture zone. It is suggested that the boundaries of the different formations within and around the suture zone need at least to be specified and/or changed in agreement with the findings of this research and supported by further petrological and geochronological research.
According to the microprobe analysis of hornblende and plagioclase crystals samples 9.2b and 15.2a appeared to have a significantly different mineralogy compared to the other samples. Both samples contain two types of hornblend crystals. According to Leake et al. (1997) the chemistry of the Hbb-crystals is more similar to (Ca-rich) tremolite than hornblende. The composition of the Hbb
hornblende can not be explained by a change in the amount of deformation during the formation of the different types of hornblende crystals, because deformation does not affect the amount of Ca in a hornblende (Spear, 1981). However, Spear (1981) also stated that a decreasing grade of metamorphism tends to produce changes in amphibole composition, i.e. decrease of Altot, Ti, Fe3+, Na and K, and an
increase in Si, Fe2+, Mg and Ca. In the case of Hbb crystals with respect to Hba crystals, this would
mean that the Hbb crystals are formed at a significantly lower metamorphic grade. It has to be noted
that the rock composition, fO2 and temperature can influence this trend seen in hornblende crystals.
This could indicate that part of the hornblende crystals experienced some recrystallization during retrograde metamorphism. The presence of epidote and chlorite crystals in the various samples confirms this idea. The size and abundance of these minerals indicate that transport of the rocks from
15-20 km depth to the surface was not very rapid. On the basis their unstable appearance the Hbb
crystals could also have grown during prograde metamorphism (fig. 14). However, the microprobe analyses give too little information to elucidate this problem
Moreover, sample 15.2a also contains plagioclase crystals, which are significantly different from the plagioclase crystals found in the other samples. The relatively high anorthite percentage of plagioclase crystals in sample 15.2a (Labradorite/Bytownite) could be explained by relatively high temperatures at which the plagioclase crystals were formed (table 9). However, Wenk et al (1991) state that high An-content does not necessarily mean that the plagioclase was formed at high temperatures. Another explanation could be, that the high An-content was formed by decomposition of an
1991). However, in the analysis of sample 15.2a no Ca-poor plagioclases were found. Moreover, such kind of decomposition is rarely seen in amphibolites, and occurs mostly in metamorphic carbonate rocks (Wenk et al, 1991). Therefore, it is difficult to find a good explanation for this phenomenon.
The trends for various elements across the hornblende and plagioclase crystals show that the minerals have hardly any zonation. However, zonation in the hornblende and plagioclase was
occasionally found with respect to the distribution of K and Ti (figure 13). The two-zonations in figure 13b could indicate subgrain formation, but the other elements do not show such a zonation. An
explanation for the K and Ti zonation with respect to changing P-T conditions and formation of subgrains has not been researched in the past and is beyond the scope of this research. The lack of zonation was expected for hornblende, which is a hydrous mineral with a high diffusivity, but plagioclase does not have such a high diffusivity. This indicates that the complete crystallization of plagioclase, and probably also hornblende, took place under the same metamorphic conditions and was not disturbed by later metamorphic events.
7.2 Thermobarometry
Geothermobarometry of the selected samples indicates a low-pressure (4 – 5.5 kbar) high-temperature (695 – 776 ºC) amphibolite-facies metamorphism or even low granulite-facies metamorphism (figure 16). However, with microscopic analysis of the samples no clino-,
orthopyroxenes nor other higher-grade minerals were found, which should be present if the rocks experienced granulite-facies metamorphism. In view of the fact that no clino- or orthopyroxenes nor
other relics of high grade metamorphic facies were found in the orthogneisses, the
pressure-temperature conditions probably indicate the peak metamorphic conditions experienced by the rocks. If the measured PT conditions would be retrograde then one would expect to find some remnants of higher PT conditions. Unfortunately, no P-T-t paths could be determined, due to the fact that the minerals hardly show any zonation. Garnet-related thermometers, in this respect, are better suited to construct P-T-t paths, due to the fact that garnets are often zoned. The granodioritic composition of the orthogneisses, however precluded garnet growth.The peak metamorphic conditions established here are not consistent with the peak metamorphic conditions of the suture zone found by Buchwaldt et al. (2002,2003) in the Lokoho and Sambirano region (figure 17). The peak metamorphic conditions of the metapelites in the Lokoho and Sambirano region are much higher than those of the orthogneisses in the Mangindrano. The retrograde conditions found in the Lokoho region, however, show a similarity with the peak conditions of this research. Due to the fact that the metamorphic grade of the orthogneisses are assumed to be peak metamorphic conditions, it is difficult to relate its metamorphic grade with the retrograde metamorphic grade of the metapelites in the Lokoho region. However, by comparing the geotherms during the peak metamorphism of the Lokoho region (31 ºC/km), Sambirano region (40.5 ºC/km) and Mangindrano region (44 ºC/km) (figure 18), they appear to be fairly similar. When the geotherms are compared with figure 19, it can be seen that all three geotherms roughly indicate a similar tectonic setting, namely a continental rift (figure 19). The similarity of the geotherms and the tectonic setting that they indicate, makes it very probable that the peak metamorphic grades of the
Figure 18 - This pressure-temperature (PT) diagram shows the PT-conditions from the Mangindrano region, Lokoho region (Buchwaldt et al., 2003) and Sambirano (Buchwaldt, et al., 2002) and their associated geotherms. The blue geotherm of the Lokoho region is 31°C/km, the green geotherm of the Sambirano region is 40.5 °C/km and the red geotherm of the Mangindrano region is 44 °C/km; The kyanite, andalusite and sillimanite stability fields are indicated in the diagram as reference frame; for the geotherms the following relation was used: 1 kbar = 3.5 km
different sub-regions within the Bemarivo suture zone formed coevally in the same tectonic setting of continental rifting. The relatively deep position of the metapelites in the Lokoho and Sambirano region with respect to the orthogneisses in the Mangindrano region during peak metamorphism is the most plausible explanation for the difference in metamorphic grade. Unfortunately, because of the intense deformation they experienced, it is difficult from the current data to judge if the formations are positioned on top of each other,. Due to the fact that the three regions are distributed all along the axis of the Bemarivo suture zone (in the west, in the centre and in the east), it can be concluded that the whole suture zone experienced a similar peak metamorphic event which can be associated with a continental rift. However, this can also be associated with a period of heating caused by processes like slab detachment. More field research in between the three regions is needed to link those regions.
Dating of monazites in the orthogneisses and granitoid veins and dykes in the Mangindrano area by Mulder (2006) failed due to the fact that the orthogneisses did not contain any monazites and the monazites in the granitoid veins and dykes were unstable and yielded no ‘realistic’ ages. Buchwaldt et al (2003) was able to date various lithologies in the Lokoho region, and Buchwaldt et al (2002) and Buchwaldt and Tucker (2001) did the same in the Sambirano region. However, it not clear in these papers if the ages are obtained from rocks of the Sambirano region or from the Lokoho region. It will be assumed that both regions are dated separately. The peak metamorphism both in the Sambirano region as in Lokoho region was dated at 520-510 Ma. In view of the similar geotherms, it is, therefore, reasonable to assume that the peak metamorphism in the Mangindrano region also took place in the same time frame. Buchwaldt et al (2002) also dated the leucogranites in the orthogneisses in the western part of the suture zone, which show a lot of similarities with the granitoid veins and dykes found in the area of Mangindrano, at ~560 Ma. This implies that the peak metamorphism post-dated the leucogranitic intrusion by ~40 Ma, which makes it less probable that the intrusion of the
leucogranites caused the HT/LP metamorphism in the orthogneisses or were formed in the same tectonic setting. The leucogranitic intrusion in turn post-date the major deformation events observed in the orthogneisses by Buchwaldt et al (2002) and Buchwaldt and Tucker (2001). In the area of
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