Geology and Petrology of the Felsic Intrusions in the Metasedimentary

FACULTEIT WETENSCHAPPEN Opleiding Master of Science in de geologie

Geology and petrology of the felsic intrusions in the metasedimentary and metavolcanic rocks of the early Neoproterozoic Zadinian Group around Matadi (Lower-Congo Region, D.R. Congo)

Stéphanie Eeckhout

Academiejaar 2013–2014

Scriptie voorgelegd tot het behalen van de graad Van Master of Science in de geologie

Promotor: Prof. Dr. J. De Grave 1 Co-promotor: Prof. Dr. L. Tack Leescommissie: Prof. Dr. P. Van den haute, Prof. Dr. M. Fernandez-Alonso

“Rocks are records of events that took place at the time they formed. They are books. They have a different vocabulary, a different alphabet, but you learn how to read them.”

-John McPhee-

PREFACE

The realization of this thesis would not have been possible without the help of many persons. Therefore I want to thank each and everyone who helped me during this intense and interesting year.

First of all I would like to thank Professor Van den haute and Professor De Grave for offering me the opportunity to do my thesis research at the research unit of Mineralogy and Petrology. Their knowledge of petrography and geochronology, respectively, has been of great value.

Professor Tack’s knowledge was of great importance to help me understand the study area. I am very grateful for the uncountable hours of work and discussions we have had together. I want to express my appreciation for the time you have sacrificed to correct my drafts and to help improve them with your constructive criticism.

I would also like to thank Professor Baudet of the RMCA. His knowledge of the study area and his expertise of GIS and mapping projects has been very helpful.

Furthermore I also want to thank Professor Fernandez-Alonso, head of the geology department at the RMCA, for the collaboration with Ghent University. Together with Luc André, I want to thank him for the approval and financial contribution of the geochemical analyses. These analyses were carried out by Laurence Monin and Jacques Navez who also deserve my gratitude for their work and explanatory notes. Feedback “from the outside” on the results of the geochemical analyses was obtained from Sophie Decrée and Ingrid Smet. Their experience helped me in understanding and interpreting the data.

I also appreciate the help of Ann-Eline Debeer, who prepared the samples for the geochemical purposes, Elien De Pelsmaeker who helped me embed the zircons and Jan Jurçeka for the preparation of thin sections.

Without the help of Stijn Glorie, at the University of Adelaide, it would not have been possible to receive the U-Pb dating results.

Last but not least, I would like to thank my boyfriend Jurgen, my sisters and especially my parents for their support throughout my five years of study.

1. Introduction ...... 9 2. Geological setting ...... 11 2.1. Geological setting of the West Congo belt ...... 11 2.1.1. Geographic setting of the Lower-Congo region ...... 11 2.1.2. The West Congo belt in the broader context of the Araçuaí-West Congo orogen ...... 12 2.1.2.1. Araçuaí-West Congo orogen...... 15 2.2. The West Congo belt ...... 20 2.3. The West Congo Supergroup...... 22 2.3.1. Zadinian Group ...... 23 2.3.2. Mayumbian Group ...... 25 2.3.3. West Congolian Group ...... 25 2.4. Geological setting of the Matadi region ...... 26 2.5. The aim of this study ...... 30 3. Methods ...... 32 3.1. Field observations and macroscopic descriptions ...... 32 3.2. Microscopic descriptions ...... 32 3.3. Geochemistry ...... 33 3.3.1. Sample preparation ...... 33 3.3.1.1. Loss on ignition ...... 33 3.3.1.2. Alkaline fusion ...... 33 3.3.2. Major elements ...... 34 3.3.2.1. ICP-AES...... 34 3.3.3. Trace elements ...... 35 3.3.3.1. ICP-MS ...... 35 3.4. Geochronology ...... 37 3.4.1. Sample preparation ...... 37 3.4.2. LA-ICP-MS ...... 38 4. Field observations and macroscopic descriptions ...... 39 4.1. Field observations ...... 39 4.2. Macroscopic descriptions ...... 44 4.2.1. Felsic magmatic protolith ...... 44 4.2.1.1. Massive ...... 44 4.2.1.2. Foliated ...... 44 4.2.1.3. Strongly foliated ...... 45

4.2.2. Sedimentary protolith ...... 45 4.2.2.1. Slightly foliated ...... 45 4.2.2.2. Moderately foliated ...... 46 4.2.2.3. Strongly foliated ...... 46 4.2.3. Mafic magmatic protolith ...... 47 4.2.3.1. Amphibolite ...... 47 4.2.3.2. Metadolerite ...... 47 4.2.3.3. Green phyllite ...... 48 5. Microscopic descriptions ...... 49 5.1. Felsic magmatic protolith ...... 49 5.1.1. Slightly deformed ...... 49 5.1.2. Moderately deformed ...... 54 5.1.3. Strongly deformed ...... 60 5.1.4. Summary of observations ...... 62 5.2. Sedimentary protolith ...... 63 5.2.1. Slightly deformed ...... 63 5.2.2. Moderately deformed ...... 65 5.2.3. Strongly deformed ...... 66 5.2.4. Summary of observations ...... 69 5.3. Mafic magmatic protolith ...... 69 5.3.1. Summary of observations...... 75 5.4. Point counting analysis ...... 75 6. Discussion field observations, macroscopic and microscopic descriptions ...... 77 6.1. Mineral assemblage ...... 77 6.2. Relict textures ...... 79 6.2.1. Blastoporphyritic – porphyritic – texture ...... 79 6.2.2. Symplectites ...... 81 6.3. Textures induced by deformation ...... 81 6.3.1. Brittle deformation ...... 82 6.3.2. Ductile deformation ...... 82 6.3.2.1. Deformation twinning ...... 82 6.3.2.2. Kinking ...... 83 6.3.3. Recovery and recrystallisation ...... 83 6.3.3.1. Recovery ...... 83

6.3.4. Grain boundary area reduction and static deformation ...... 85 6.3.5. Core-and-mantle texture vs. porphyroclast systems ...... 86 6.4. Discussion ...... 86 6.4.1. Hypabyssal rocks ...... 86 6.4.2. Metaquartzites ...... 86 6.4.3. Rocks with a mafic magmatic protolith...... 86 6.5. Discussion regarding previous observations ...... 87 6.6. Contact metamorphism ...... 87 6.6.1. Quartzite assimilation...... 88 6.6.2. Garnet blastesis ...... 88 6.7. Metasomatism...... 88 6.8. Aegirine and riebeckite ...... 89 6.9. Lithological maps ...... 89 7. Geochemistry ...... 92 7.1. Previous research ...... 92 7.2. Major elements ...... 92 7.2.1. Classification ...... 94 7.2.2. Discrimination diagrams ...... 97 7.2.3. Harker diagrams ...... 98 7.2.4. R1 – R2 multicationic diagram ...... 100 7.3. Trace elements ...... 101 7.3.1. Variation diagrams ...... 101 7.3.2. Discrimination diagrams ...... 106 7.3.2.1. Tectonic setting ...... 106 7.3.2.2. Alphabetical classification ...... 108 7.3.3. Masuda Coryell diagrams ...... 110 7.3.4. Spider diagrams ...... 113 8. Discussion geochemistry ...... 116 8.1. Noqui granite + hypabyssal rocks versus Mpozo syenite ...... 116 8.2. Assimilation and Metasomatism ...... 116 8.3. Tectonic setting ...... 117 8.4. Fractional crystallization ...... 118 8.5. Petrogenesis ...... 118 8.6. Discussion regarding previous research ...... 119

9. Geochronology ...... 120 9.1. Zircon morphology ...... 120 9.1.1. Noqui granite ...... 120 9.1.2. White Mpozo syenite ...... 120 9.1.3. Pink Mpozo syenite ...... 120 9.1.4. Hypabyssal rock ...... 120 9.2. Dating results ...... 123 9.2.1. Noqui granite ...... 123 9.2.2. White Mpozo syenite ...... 124 9.2.3. Pink Mpozo syenite ...... 124 9.2.4. Hypabyssal rock ...... 125 10. Discussion geochronology ...... 126 10.1. Noqui granite and hypabyssal rocks ...... 126 10.2. Mpozo syenite versus Noqui granite and hypabyssal rocks ...... 126 10.3. New ages compared to earlier ages of the Lower-Congo region ...... 126 10.4. Pb-loss ...... 127 11. Discussion of the geology of the Matadi region ...... 128 12. Summary and Conclusion ...... 132 13. Nederlandse samenvatting ...... 137 14. References ...... 142 Annexes ...... 147

1. INTRODUCTION

At the beginning of the 20th century, Alfred Wegener already recognized the good geometric fit between the Atlantic margins of Africa and South America. Prior to the opening of the South Atlantic Ocean, the São Francisco craton of Brazil was united with the Congo craton of Africa. This connection was already established by the end of the Eburnian-Transamazonian (2,1 Ga) orogeny (Alkmim et al., 2006). From the Palaeoproterozoic until the Cretaceous, the São Francisco-Congo craton was incorporated in different supercontinents and remained as one entity during several cycles of continental break up and amalgamation.

One of the (super)continents formed was Gondwana of which the western part, comprising Africa and South America, was largely assembled by 600 Ma. Due to compressional events, the Araçuaí- West Congo Orogen (AWCO) was formed. Prior to its development at least six extensional events (E1 – E6) of rifting and/or anorogenic magmatism occurred (Pedrosa-Soares & Alkmim, 2011) within the area occupied by the AWCO. After the extensional events, the AWCO started to form around 630 Ma due to compressional events. These events gave rise to a series of granitoid suites (G1 – G5), formed between ca. 630 and 480 Ma (Gradim et al., 2014). In Brasil the compressional events are referred to as being the result of the Brasiliano orogeny, while in Africa this is called the Pan African orogeny. Due to the opening of the Atlantic Ocean in the Cretaceous the two parts became separated. In this work we will only focus on the African side of the AWCO, which comprises the West Congo belt.

The West Congo belt is the 1400 km long African remnant of the AWCO. This complex structural unit comprises an ENE-verging fold-and-thrust belt which was created during the Pan African orogeny. In the Lower-Congo region the peak stage of this orogeny is dated at 566 Ma (Frimmel et al., 2006). Deformation and metamorphism resulted in a tectono-metamorphic overprint.

In the Matadi region the Palaeoproteroic basement comprises the 2,1 Ga old Kimeza Supergroup, which is covered by the West Congo Supergroup. The latter can be further subdivided, from old to young, in the Zadinian, Mayumbian and West Congolian Group. Within the area, there are also two plutonic bodies exposed, being the Noqui granite and the Mpozo syenite. The Noqui granite comprises a peralkaline A-type granite which was formed during the E4 extensional event. Recent dating of the pluton resulted in an emplacement age of at 999 ± 7 Ma, evidencing a pre-orogenic emplacement (Tack et al., 2001). Compared to the Noqui granite, the Mpozo syenite is not as well documented. Delhal and Ledent (1978) have tried to date the Mpozo syenite, resulting in a poor emplacement age of 1960 ± 594 Ma (U-Pb dating on bulk zircons).

Several mapping attempts of the Matadi region have resulted in numerous, sometimes sketchy, but often contradictory geological maps. In this study we use the geological map of 2008, which is based on the observations of Tack (1975a), as a starting point. Two recent (2004 and 2011) field missions resulted in observations, which contradict to some extent the geological map of 2008. One of these contradictory aspects comprises the angular unconformity between the Kimezian basement and the overlying “Palabala Formation” (part of the Zadinian Group), described by Tack (1975a) and Tack et al. (2001). This unconformity has no longer been confirmed in 2011 (RMCA, archives). It is suggested that there is no real “Palabala Formation” overlying the Kimezian basement. New observations suggest that this mylonitic package includes various different protoliths such as the Kimezian basement, the quartzites of the Matadi Formation and the Mpozo syenite. This indicates that the

“Palabala Formation” should thus be regarded as a tectono-structural unit rather than a lithostratigraphic unit. This problem was already introduced in Behiels (2013).

During his study, Behiels (2013) focused on the geology and petrology of the Noqui granite and the Mpozo syenite. He concluded that the geological map needs adjustment of the outlines of both massifs, as they are connected to each other by a tectonic contact. Furthermore he reported an issue concerning the distribution of “felsic magmatic bodies” of limited extent, which are intrusive in the Matadi Formation, the Palaeoproterozoic basement and the Mpozo syenite body. In his work, Behiels (2013) raises the question whether these rocks were emplaced as extrusive or intrusive rocks and whether they are related to the Noqui granite.

In our study an effort is made to better constrain the geology of the Matadi region, with the main focus on the “felsic magmatic bodies” reported by Behiels (2013). During the first part of this work we focus on field observations and hand specimens sampled by three different field geologists (Hugé, Masser and Steenstra) in the same detailed region but at different times. Based on macroscopic descriptions, we identify the rocks in an appropriate and systematic way. As some of them are strongly deformed, their macroscopic observation might be insufficient to identify the protolith. To obtain a correct lithological determination and geological mapping, it is thus crucial to make microscopic observations. Therefore the second part of this work comprises a microscopic study. Combined results of the field observations, hand specimens and thin sections allow us to plot the different lithologies on a map to improve earlier incomplete sketchy maps.

The following part of this work focuses on the geochemistry of the “felsic magmatic bodies”. In this section we will compare the geochemical signature of the “felsic magmatic bodies” with those of the Noqui granite and Mpozo syenite. By doing this we will try to figure out whether these “felsic magmatic bodies” are related to the other, larger, magmatic massifs in the area, i.e. the Noqui granite or the Mpozo syenite. To answer this research question, we also incorporated a geochronologic section, to obtain the emplacement age of the “felsic magmatic bodies”, the Noqui granite and the Mpozo syenite. The results of the geochemisty and geochronology will determine whether the magmas have both resided in the same magma chamber or if they are related by a same petrogenetic process, and hence if they are respectively comagmatic or cogenetic.

The further extent of this thesis is given below. As we use different aspects of geology to approach our problems, we use a modular approach. This was also necessary because some of the geochronological data were not obtained until the end of May. To be consequent we also used a modular approach to number figures and tables, in which the first number refers to the chapter.

After this introduction, chapter two discusses the geological setting of the West Congo belt and the Matadi region. In chapter three we focus on the used methodology used. Chapter four represents the field observations and macroscopic descriptions and is followed by chapter five in which the results of the microscopic study are given. The results given in these two chapters are discussed in chapter six. Chapter seven and eight give the results and the discussion of the geochemical data and are followed by chapter nine and ten in which we describe and discuss the geochronology. Finally, in chapter eleven, we will integrate all new data in a synthetic geological scheme and map to explain the geological evolution of the Matadi region. A summary and conclusion of this work are given in chapter twelve, followed by a dutch summary in chapter thirteen. References are given in the last chapter. The annexes and a digital copy of this thesis are attached on a CD-ROM.

2. GEOLOGICAL SETTING

2.1. GEOLOGICAL SETTING OF THE WEST CONGO BELT 2.1.1. Geographic setting of the Lower-Congo region The Lower-Congo region makes up one of the eleven (2008) provinces of the Democratic Republic of Congo (DRC). The area borders in the northeast with the provinces of Kinshasa, where the capital city is located, and in the east with Bandundu. Furthermore the area is framed by the Republic of Angola in the south, by the Atlantic Ocean in the west and by Cabinda and Congo-Brazzaville in the north (Fig. 2-1). As the area is the only province of the DRC with a coastline, it offers an important economic value to the country. Its chief seaport is located in the city of Matadi, which is the capital city of the Lower-Congo region. The immediate region around the city of Matadi also covers the study area of this thesis.

Figure 2-1: The Lower-Congo region (indicated as Bas-Congo) within the Democratic Republic of Congo. The city of Matadi and the capital city Kinshasa are indicated. The neighbouring countries Angola in the south, Congo-Brazzaville (indicated as Congo) and Cabinda (indicated as Angola) in the north. After the Directorate of Human Rights and Public Relations, Civic United Front, 2008.

The name of the Lower-Congo province refers to the lowermost segment of the Congo River, which runs across the region before flowing into the Atlantic ocean. The coastline forms a rather short segment of 35 km, which is located between the former Portuguese colonies of Cabinda and Angola. From its mouth, the Congo River forms a circa 150 km long, broad and navigable estuary up to the harbor of Matadi. From there onwards transport by boat is impossible because of the rapids occurring further upstream. Therefore a 350 km long railway, a tarmac road and a petrol pipe-line link the harbor of Matadi with the capital Kinshasa.

As the Congo River crosses the Lower-Congo region it induces a lot of erosion and it has an impact on the relief. The topography is mainly hilly and some peneplanated lateritic plateaus are left. Altitudes in the area range from sea level to circa 900 m for the highest ridges and hills. Along the Congo River and in some of its tributaries, outcrop conditions can vary from excellent, to poor with scattered small boulders.

2.1.2. The West Congo belt in the broader context of the Araçuaí-West Congo orogen At the beginning of the 20th century, Alfred Wegener already recognized the good geometric fit between the Atlantic margins of Africa and South America (Fig. 2-2). Prior to the opening of the South Atlantic Ocean, the São Francisco craton of Brazil was united with the Congo craton of Africa. This connection was established by the Bahia-Gabon continental bridge, also called the São Francisco-Congo cratonic bridge. In order to explain the history of this former unit we need to go back in time to the Palaeoproterozoic.

Figure 2-2: South America-Africa fit, with indication of the São Francisco - Congo cratonic bridge. After Alkmim et al., 2006.

During the Palaeoproterozoic the assembly of the supercontinent Columbia (Nuna) arose (Goodenough et al., 2013). It is assumed that the aggregation of this supercontinent occurred during the end of the Eburnian-Transamazonian (2,1 Ga) orogeny. As a result, the proto-Congo Craton (Fig. 2-3) was formed (Fernandez-Alonso et al., 2012). Within this proto craton, the São Francisco and Gabon blocks were connected along the Bahia-Gabon bridge (Alkmim et al., 2006).

After its formation, the proto-Congo Craton stabilized at ca. 1,8 Ga. Throughout the entire Proterozoic, and thus also throughout the break-up of Columbia (1,6 – 1,2 Ga) (Fan et al., 2013), the proto-Congo craton remained united as one single entity (= “palaeoplate”). Compressional events between 1,3 and 0,9 Ga (Li et al., 2008) however caused the formation of a new supercontinent, Rodinia. During this assembly the combined Congo-São Francisco Craton (= “proto-Congo Craton”) remained an integral part of Rodinia (De Waele et al., 2008), which is illustrated in Figure 2-4.

Figure 2-3: Possible extent of the proto-Congo Craton. From Fernandez-Alonso et al., 2012.

Figure 2-4: Cartoon displaying the configuration of supercontinent Rodinia at 900 Ma. The São Francisco-Congo craton is indicated as one unit. After Li et al., 2008.

After its assembly, Rodinia lasted for about 150 Ma, before break-up began. Superplume events are believed to have caused these diachronous break-up events and are thought to have occurred around Laurentia. This caused the continental pieces to move away from Laurentia and to collide on the other side of the Earth, resulting in the formation of Gondwana (Li et al., 2008). Based on palaeomagnetic apparent polar wander paths from the world’s cratons, it is possible to reconstruct palaeogeographic possibilities of the continents after Rodinia break-up. These different reconstructions, in their Early Jurassic configuration, are given in Figure 2-5. From this figure one can conclude that different scenarios are possible, but in every configuration the Congo-São Francisco craton supposedly remained as one entity.

Figure 2-5: Reconstructions of the world’s cratons in their Early Jurassic configuration. References are: a) Dalziel (1997); b) Pisarevsky et al. (2003); c) Li et al. (2008). From Evans (2009).

Figure 2-5 continued.

West Gondwana, comprising Africa and South America, was largely assembled by ca. 600 Ma. Final Gondwana amalgamation occurred between ca. 540 – 530 Ma (Li et al., 2008) and was later incorporated in the Pangea supercontinent. During that period the São Francisco-Congo craton reached its final stages, as due to the opening of the South Atlantic Ocean in the Cretaceous, the link between the São Francisco Craton and the Congo Craton was destroyed (Alkmim et al., 2006).

2.1.2.1. Araçuaí-West Congo orogen At present times the Araçuaí orogen covers the region between the São Francisco craton and the Atlantic continental margin. It forms an orogenic edifice of approximately 1000 km long and 500 km wide. Its counterpart, the West Congo belt, is located in Africa, south west of the Congo craton (Fig. 2-6). Prior to the opening of the South Atlantic these two different but complementary counterparts made up one larger orogenic edifice; the Araçuaí-West Congo Orogen (AWCO)

The development of the AWCO started around 630 Ma, during the amalgamation of the Gondwana supercontinent. The AWCO was only a small portion of a giant orogenic network generated along the margins of the various plates that collided to form West Gondwana. At that time the Macaúbas basin was considered as a terminal branch of the Adamastor ocean. This gulf-like basin was enclosed by the São Francisco peninsula and the Congo continent. These converged and collided during the Brasiliano and Pan African orogenies and gave rise to the AWCO. As the orogen was enclosed to the west, north and east by crustal blocks, creating an upside-down “U”, it is considered a “confined” orogen (Pedrosa-Soares & Alkmim, 2011).

Prior to the development of the AWCO at least six extensional events (E1 – E6) of rifting and/or anorogenic magmatism occurred within the area occupied by the AWCO. The first three extensional events (E1 – E3) are respectively the Statherian (ca. 1,7 Ga), Calymmian (ca. 1,5 Ga) and Early Stenian (ca. 1,18 Ga) events and are only exposed in the Araçuaí belt. Evidence for the three youngest events (E4 – E6) can be found within the African West Congo belt and include respectively the Stenian- Tonian (ca. 1Ga), Tonian (930 – 850 Ma) and Cryogenian (750 – 670 Ma) event.

Figure 2-6: Indication of the São Francisco craton (SFC) and the Congo craton, which enclose the Araçuaí orogen and the West Congo belt. From Pedrosa-Soares (2013).

Within the West Congo belt, the Stenian – Tonian event (E4) is recorded by the A-type Noqui granite, which was dated at 999 ± 7 Ma (Tack et al., 2001). This anorogenic magmatic suite is probably related to the opening of the Sangha aulacogen (Pedrosa-Soares & Alkmim, 2011). The following Tonian event (E5, 930 – 910 Ma) is evidenced by a thick package of Zadinian and Mayumbian groups with related intrusions. These groups are the results of a bimodal magmatic suite, which formed during continental rifting. The Zadinian and Mayumbian groups have no correlatives on the Brazilian side, suggesting an asymmetrical rift (Fig. 2-7) with the thermal-magmatic axis in the West Congo belt (Pedrosa-Soares et al., 2008). The youngest extensional event (E6), the Cryogenian event, is on the African side best exposed in the northern portion of the West Congo belt, in the Lower-Congo region. In Gabon, rhyolites and rhyolitic tuffs appear within the La Louila Formation.

On the Brazilian side, ophiolites occur within the Ribeirao da Folha Formation (Delpomdor et al., in press; Gradim et al., 2014). These ophiolites evidence that oceanic spreading occurred in the central- southern Macaúbas basin. To the north this oceanic spreading died out, keeping the Bahia-Gabon bridge intact. Eyles and Januszczak (2004) describe this diachronous rifting process as the Zipper rift model (Fig. 2-9a). As the ophiolite is dated – i.e. crystallization age of the oceanic crust – at 660 ± 29 Ma (Queiroga et al., 2007; Delpomdor et al., in press), it can be stated that the rift-drift transition had come to an end by ca. 660 Ma (Delpomdor et al., in press).

Figure 2-7: Cartoon illustrating the asymmetric continental rift during the Tonian event. From Pedrosa-Soares et al. (2008).

After the series of extensional events a period of compression occurred. These compressive events, resulting in the amalgamation of Gondwana, are named the Brasiliano orogeny (in Brasil) and the Pan African orogeny (in Africa), which gave rise to the AWCO. As a result of the orogeny, a series of granites were produced between ca. 630 and 480 Ma. These plutonic rocks have been grouped into five supersuites (G1 – G5) and are given in Table 1. Figure 2-8 represents a cartoon displaying the generation events of these granites.

Table 2-1: plutonic supersuites of the Araçuaí orogen. From Gradim et al., (2014).

Figure 2-8: Cartoon illustrating the evolution of the AWCO during its pre-collision to syn-collision transition (a), during its collisional stage (b) and its post-collisional stage. These three stages correspond to Figures 9b, 9c and 9d respectively. The cartoon also displays the formation of the five plutonic supersuites (G1 – G5). From Gradim et al., 2014.

For the compressional history of the AWCO, Alkmim et al. (2006) suggest the nutcracker tectonics model. With this term they refer to the fact that the orogen formed when the western arm of the São Francisco-Congo craton rotated counterclockwise towards the eastern arm. By doing so the Macaúbas basin, which lies in between the two cratons, was squashed like a nut by a nutcracker. The nutcracker model explains the evolution of the AWCO in several phases. At first the extensional

18 events E4, E5 and E6, which correspond to the Zipper rift model (Eyles and Janusczac, 2004), caused the opening of the Macaúbas basin (Fig. 2-9A), which broadened progressively southward between 1000 – 700 Ma. This resulted in the formation of a narrow ocean in the southern half of the Macaúbas basin. During a second phase, at about 635 Ma, after oceanic crust production around 660 Ma, arc-related granitic suites began to form (Fig. 2-9B). These indicate that the Macaúbas basin began to close and with this closing the oceanic portion of the basin began to subduct. During closure the southern arm of the São Francisco craton rotated counterclockwise relative to the Congo craton. When the oceanic basin was closed (Fig. 2-9C), foreland-verging external fold-and-thrust belts were formed. As the nutcracker tectonics continued, a space problem developed in the southern portion of the orogen, which therefore underwent lateral escape to the southeast. The northern half which had extremely thickened underwent extensional collapse during the final phase of closure (Fig. 2-9D).

A B

C D

Figure 2-9: A) Opening of the Macaúbas basin according to the Zipper Rift Model; tectonics: B) Closure of the Macaúbas basin with formation of arc-related magmatism; C) During full development of the orogen the oceanic basin was closed; D) The southern portion of the orogen escapes to the south and the northern half undergoes extensional collapse. From Alkmim et al., 2006.

To understand the complete history of the AWCO, it is necessary to study both its Brazilian and its African counterpart. This is why in the previous section, both parts were discussed. Further on we will only focus on the African side of the AWCO, which comprises the West Congo belt.

2.2. THE WEST CONGO BELT The West Congo belt is situated subparallel to the Atlantic coast between 1° and 12° south of the equator. It is the 1400 km long and 150 to 300 km wide African remnant of the AWCO. In Figure 2-10 it can be seen that the West Congo belt exhibits a prominent flexure in its central segment (Fig. 2- 11), which overlaps with the Lower-Congo region and adjacent northern Angola. This complex structural unit, which comprises an ENE-verging fold-and-thrust belt (Alkmim et al., 2006), was created during the Pan African orogeny. This orogeny, which is locally called the “West Congo orogeny”, took place in the Lower-Congo region at 566 Ma (Frimmel et al., 2006).

Figure 2-10: Geological map of the West Congo belt. In the central segment of the fold-and-thrust belt a prominent flexure can be observed. From Tack et al., 2001.

Figure 2-11: Geological map of the central segment of the West Congo belt. The thick red line indicates the thrust front between the fold-and-thrust belt in the west and the aulacogene foreland in the east (Tack, 2014a). After Tack et al., 2001.

Field and structural data, gravimetric data, published cross-sections and lithostratigraphic charts (Tack et al., 2001 and references therein) were incorporated to draw a schematic east-west cross- section of the West Congo tectono-metamorphic domains (Fig. 2-12). This cross-section has been adapted recently and can be subdivided into an aulacogene foreland and a fold-and-thrust belt domain (Delpomdor et al., in press). The foreland domain, located in the Sangha aulacogen, comprises weakly to unmetamorphosed sedimentary rocks. To the west, the fold-and-thrust belt reveals greenschist facies metamorphism, which grades even more to the west into amphibolite facies, where it reaches its maximum metamorphism. As a result it is concluded that metamorphism, during the Pan African orogeny, decreased towards the east.

Figure 2-12: E-W cross-section of the West Congo belt, based on former concepts. Based on more recent insights (Delpomdor et al., in press) the area can be divided into a fold-and-thrust belt and a foreland domain, separated from each other by a thrust front (indicated in red). After Tack et al., 2001.

Contrary to former concepts, the fold-and-thrust belt does not grade gradually into the foreland domain, but they are separated from each other by a thrust front (Fig. 2-11 and Fig. 2-12) (Delpomdor et al., in press). The fold-and-thrust belt comprises at the base the Palaeoproterozoic basement. It contains the rocks with an age of 2,1 Ga, which belong to the Kimezian Supergroup. During the Pan African orogeny, this basement was thrust onto the Zadinian Group, which itself was thrusted onto the Mayumbian Group, which borders the younger West Congolian Group. The most western section of the West Congo belt endured the most excessive deformation, resulting in these imbricated thrust slices. Despite this it is assumed that displacement of slices along thrust faults in the Lower-Congo region was rather limited. On the other hand it can be said that Pan African deformation has locally given rise to mylonitic corridors and L-S fabrics (Tack et al., 2001).

2.3. THE WEST CONGO SUPERGROUP Studies of the West Congo belt were mainly performed during colonial times. The contribution of different countries resulted in complex and diverse terminologies. Since the 1960-ies, due to political constraints, field access may have been difficult for al long time which resulted in a nomenclature that can vary strongly from place to place. The nomenclature adopted here is proposed by Tack et al. (2001) and is supported by the rules according to the IUGS.

In the eastern West Congo belt post-Karoo and Karoo cover deposits are exposed. On the western side outcrops of the Palaeoproterozoic basement can be found. These units border the West Congo

Supergroup at its upper and lower limit. From old to young the West Congo Supergroup comprises the Zadinian, Mayumbian and West Congolian Group. A lithostratigraphic reconstruction is given in Figure 2-13.

Figure 2-13: Lithostratigraphic reconstruction of the West Congo Supergroup, which is made up of three groups. From old to young: the Zadinian Group, the Mayumbian Group and the West Congolian Group. The West Congo Supergroup covers the 2,1 Ga Kimezian basement. Used symbols: ρ = rhyolite; β = basalt; δ = dolerite; M = Mativa; BK = Bata Kimenga. From Tack et al., 2001.

The base of the West Congo Supergroup, which comprises the Zadinian Group, is separated from the basement by a major angular unconformity (Tack et al., 2001). The basement comprises Palaeoproterozoic migmatitic paragneisses and amphibolites dated at 2,1 Ga (= Kimeza Supergroup).

2.3.1. Zadinian Group Palabala Formation In the area of Matadi, the base of the Zadinian Group is “traditionally” formed by the 500 m thick Palabala Formation (Lepersonne, 1969). This formation mainly contains micaceaous quartzites and biotite schists. According to Franssen and André (1988) metarhyolites are intercalated in the upper part of the formation. At the base, as well as in the basement, sills of microgranites are present. The presence of these metarhyolites were already observed by Cahen et al. (1976) who tried to date these magmatic rocks.

Matadi Formation In the Matadi region the Palabala Formation is covered by the Matadi Formation. Outside that region the Matadi Formation makes up the base of the Zadinian Group. The formation comprises continental, siliciclastic metasediments and does not surpass a thickness of 1500 m. The sequence was deposited in a continental rift environment and contains strong lateral and vertical facies variations.

Yelala Formation Between the siliciclastic metasediments of the Matadi Formation and the overlying mafic rocks of the Gangila Metabasalts, a conglomeratic unit can be found. This unit comprises the “Yelala” Formation and is described by Thonnart (1956). The conglomeratic clasts are made of the same material present in the Matadi Formation. The Formation only occurs locally and is very discontinuous.

Gangila Metabasalts On top of the metasediments, belonging to the Matadi Formation, mafic rocks can be found. These mafic rocks are considered to be continental flood basalts and have been labeled with various local names. In the type area they are described as the “Gangila” amygdaloidal metabasalts (Tack, 1975b). They have a tholeiitic composition and within the Lower-Congo region they form a 1600 to 2400 m thick sequence. To the north and to the south of the Lower-Congo region basaltic activity must have been strongly reduced, which resulted in thinner sequences. Together with the felsic magmatic rocks of the Mayumbian, this sequence is the result of a continental rift climax, described as the Tonian event in section 2.1 (Tack et al., 2001; Pedrosa and Alkmim, 2011).

Noqui granite In the central segment of the West Congo Belt a granitic body is exposed. This granitic body, which lies south of Matadi, comprises the peralkine Noqui granite. Only a minor part crops out in the DRC. The largest portion is exposed, across the border, in Angola. The Noqui granite is intrusive in the Zadinian Group in the vicinity of the Kimezian basement. For a long time there has been an ongoing debate on the field setting of the granite; whether it has a pre- or post-orogenic emplacement. New information based on sensitive high-resolution ion microprobe (SHRIMP) analysis of zircons revealed a crystallization age of 999 ± 7 Ma (Tack et al., 2001). As the peak stage of the West Congo orogeny, in the Lower-Congo region, is dated at 566 Ma (40Ar – 39Ar dating; Frimmel et al., 2006), the “early” pre-orogenic emplacement of the Noqui granite is evidenced.

Tack et al. (2004) and Behiels (2013) describe the modal mineralogy of the peralkaline Noqui granite. Na-K perthites, quartz and aegyrine, often in association with riebeckite, are abundant. Lepidomelane, an Fe-rich biotite, and/or magnetite are present in lesser amounts. Accessory minerals such as zircon, fluorine and calcite are possible. Geochemically the Noqui granite is Sr-poor which results in exceptionally high Rb/Sr-ratios.

Mpozo syenite The Mpozo River is a tributary of the Congo River and flows along the eastern side of the city of Matadi. In the vicinity of this river a small syenitic massif, the Mpozo syenite, is exposed. In this massif two varieties occur: a pink syenite, which was described by Delhal and Ledent (1978), and a white variety extensively described by Behiels (2013). Delhal and Ledent (1978) have tried to date the Mpozo syenite, but they were not able to constrain a precise age. An age of 1960 ± 594 Ma (U-Pb dating on bulk zircons) was achieved. Altough the Mpozo syenite body has never been mapped as such, it has been recognized to outcrop in the immediate vicinity of the gneisses of the Kimezian basement (Delhal and Ledent, 1973). These gneisses were also dated (U-Pb dating on bulk zircons) by Delhal and Ledent (1976) at 2088 Ma. Because of their ages Delhal and Ledent (1978) have considered that the Mpozo syenite corresponds to a late event related to the Kimezian basement.

2.3.2. Mayumbian Group The Mayumbian Group, which lies stratigraphically above the Zadinian Group, consists of a felsic volcanic-plutonic sequence with some intercalations of volcano-sedimentary and sedimentary rocks. The internal lithostratigraphy of the sequence can vary strongly from place to place. In the Lower- Congo region the Mayumbian Group consists of a 3000 – 4000 m thick felsic volcanic sequence, locally described as “Inga” metarhyolites, which is intruded by granitic bodies, referred to as the “Lufu” massif.

Two metarhyolite samples, one from the base and one from the top of the sequence, have been dated by Tack et al. (2001) using single zircon SHRIMP dating. Samples from the base and the top indicate a crystallization age of respectively 920 ± 8 Ma and 912 ± 7 Ma. The emplacement age of the granitic bodies – cross-cutting the metarhyolites – has been constrained at 924 ± 25 Ma and 917 ± 14 Ma.

2.3.3. West Congolian Group The West Congolian Group is separated from the underlying Mayumbian Group by a nonconformity. This means that the West Congolian Group was only deposited after unroofing of the Mayumbian granites. The emplacement age of the Lufu Massif (920 – 910 Ma) thus provides an age that must be older than the onset of deposition of the West Congolian Group.

At the base the Sansikwa Subgroup commences with a conglomerate which is followed by a succession of argillite, quartz arenite and arkose. U-Pb zircon dating of detrital zircons of the Sansikwa Subgroups constrains the maximum sedimentation age at 923 ± 43 Ma (Frimmel et al., 2006). A diamictite separates the Sansikwa Subgroup from the Haut Shiloango Group. The diamictite, referred to as the Lower Mixtite Formation has been considered to correspond to the Sturtian glaciation (750 – 720 Ma) (Frimmel et al., 2002). The Haut Shiloango Subgroup consists of a varied succession of conglomerate, argillite, calcpelite, quartz arenite, calcarenite and finally limestones and is overlain by a second diamictite, the Upper Mixtite Formation. Based on 87Sr/86Sr-ratios this diamictite is correlated with the Marinoan glaciations (636 Ma) (Frimmel et al., 2006). The Upper Mixtite Formation is covered by the Schisto-Calcaire Subgroup which consists of a cap carbonate at the base that develops into a carbonate ramp and platform with abundant stromatolite bioherms. The Mpioka Subgroup is a siliciclastic subgroup made up of conglomerates, quartz arenite, arkose and argillite. This succession has been interpreted as a late-orogenic fluvial molasse deposit that has experienced orogenic deformation at approximately 566 Ma.

The overlying Inkisi Subgroup consists of a predominantly coarse-grained siliciclastic sedimentary succession. Its stratigraphic and tectonic position has been problematic. The subgroup was formerly thought to be part of the West Congolian Group but Tack et al. (2001) proposed that the Inkisi Subgroup was not related to the Pan African orogeny but postdated it and thus should be considered as an individual lithostratigraphic unit.

In section 2.1 the formation of oceanic crust was mentioned. This event occurred approximately at 660 Ma. As the Upper Mixtite Formation was deposited around ca. 636 Ma, the Sansikwa Subgroup, the Lower Mixtite Formation and the Haut Shiloango Group thus belong - together with the Mayumbian and the Zadinian Groups - to a long period of break-up and rifting events (E4 – E6). The Schisto-Calcaire Subgroup evolved in a passive margin setting (Delpomdor et al., in press).

Since 2009 until 2013, the RMCA and the Geological Survey Department of the DRC in Kinshasa GRGM (Centre de Recherches Géologiques et Minières) have been working together on a mapping project to update the geological map of the Lower-Congo region. Several new lithostratigraphic names have been (re)defined. The Zadinian Group has been renamed Matadi Group, including only the Matadi Formation (metaquartzites). The Mayumbian Group has been renamed Seke Banza Group including the Gangila Formation (amygdaloidal metabasalts) underlying the Inga Formation (metarhyolites), both formations corresponding to the bimodal magmatic event (E5).

2.4. GEOLOGICAL SETTING OF THE MATADI REGION In this section, we focus on the specific geology of the Matadi region. Over the years, several mapping attempts (Behiels, 2013; Annex 1) have tried to achieve a geological representation of the region. However, these attempts were based on limited and/or scattered observations and data, without a systematic approach to integrate all available data originating from various sources (both published or unpublished). Tack (1975a) compiled an “integrated” 1:200.000 geological map of the whole Lower-Congo region to the west of the 14th meridian, thus completing the geological coverage (1:200.000) which had previously been achieved to the east of the 14th meridian. Since 1975, no systematic update of this map has been performed, although episodic geologic research within some limited areas of Tack’s map (1975a) has been conducted by various authors (results in several documents, both published or unpublished; see bibliographic references in Behiels, 2013 and in this thesis). It is thus clear that the 1975 geological map is now outdated.

In 2008, a bilateral geological mapping project between the CRGM (Centre de Recherches Géologiques et Minières) of Kinshasa (DRC) and the RMCA, Tervuren (Belgium) was launched. By lack of an alternative more recent document, the 1975 map was digitized as a starting point for modern updating purposes. An excerpt of this 2008 map for the Matadi region is used in our study (Fig. 2-14).

Two field missions (2004 and 2011) resulted in new crucial information concerning the geology of the Matadi region. These field missions include the: − UNESCO-related International Geological Correlation Programme (IGCP), project 470 (see Tack et al., 2004) − Bilateral aid project (2009-2013) between the Geological Survey Department of the DRC in Kinshasa CRGM and the RMCA, Tervuren (Baudet, Tack, Fernandez-Alonso) on the update of the geological map of the Lower-Congo region, field mission 2011, see RMCA archives.

These two missions resulted in new insights concerning the geology of the region. During these field studies, amongst others, the several hundred meters thick package of the “Palabala Formation” were revisited. Observations indicated that this package comprises mylonitic rocks, which endured metamorphism under greenschist facies conditions. These mylonites, often strongly reduced in grain size, suggest to originate from various protoliths: Kimezian migmatitic paragneisses and amphibolites, metaquartzites of the Matadi Formation and Mpozo syenite. These observations are in contradiction with the observations of Tack (1975a) and Tack et al. (2001), who described a major angular unconformity, in the Congo-da-Lamba region, between the Kimezian basement and the “Palabala Formation”. This angular unconformity has no longer been confirmed in 2011 (RMCA, archives). This suggests that there is no “Palabala Formation” at the base of the Zadinian Group unconformably overlying the Kimezian basement. The “Palabala Formation” should thus be regarded as a tectono-structural unit rather than a lithostratigraphic unit.

Figure 2-14: Geological map of the Matadi region, after Tack (1975a). Digitized (provisional and unpublished) by Rensonnet and Laghmouch (2008). The red rectangle indicates the Matadi region of interest. Indicated by the small blue quadrangle: area of studied field observations and samples. The yellow dotted line indicates the location of the panoramic view given in section 4.1 (Fig. 4-6).

Table 2-2: Temporal variations in the definition of the “Palabala Formation”, indicated in yellow. From Behiels (2013)

The discussion regarding the meaning of the “Palabala Formation” has been ongoing for several decades. Behiels (2013; Table 2) sketches the evolution of ideas on the definition of the “Palabala Formation”. Figure 2-14 still comprises the “Palabala Formation” as a lithostratigraphic unit. This illustrates that the geological map of the Matadi region needs considerable improvement, the more that in the meantime it has become clear that the geological setting of this region is much more complex than originally envisaged.

As Bertossa and Thonnart (1957) discarded the existence of the “Palabala Formation”, their 1957 map (Behiels, 2013; Annex 1) might form a good guideline for the revision of the extent of the Matadi Formation. According to Behiels (2013) further adjustments of the map of the Matadi region include the outline of the Noqui granite and the Mpozo syenite, the two bodies actually being in contact by a fault.

Behiels (2013) also raises the question concerning the distribution of “felsic magmatic bodies” of limited extent, which are mainly intercalated in the Matadi Formation (and the former “Palabala Formation”). These felsic rocks also occur within the Palaeoproterozoic basement and the Mpozo syenite body. He also wonders whether these rocks were emplaced as extrusive or intrusive rocks and whether they are related to the Noqui granite.

Because the geological map of the Matadi region is outdated, a provisional and tentative timetable of “main geological events” (MGEs), that occurred in the (broader) Matadi region (Fig. 2-14) and are essential to be taken into account as working hypotheses in any new study of the region, is summarized here below. It integrates scattered data and knowledge from literature of the last 70 years (often discussing only partial aspects of the regional geology; main references include: Tack and Baudet, RMCA archives; Behiels, 2013; Tack et al., 2001; André and Franssen, 1988; Lepersonne, 1983; Delhal et Ledent, 1976; 1978; Tack, 1975, a; b; Korpershoek, 1964; Bertossa et Thonnart, 1957; Thonnart, 1955; Mortelmans, 1948; Cahen, 1948; … and references therein), in combination with preliminary modern remote sensing observations. The timetable is based on a “convergence of evidence of geological constraints”, which will be evaluated in our study as critically and objectively as possible. It suggests a succession of (at least) four MGEs, that control the geological evolution of the Matadi region. They are respectively from younger (MGE 1) to older (MGE 4):

1) MGE 1 Discrete late general N-S (to NNW-SSE and/or NNE-SSW) trending shear zones and (broader) corridors formed under brittle conditions (cataclasis, grain size reduction of protolith, …) with (often ?) relatively steep dips (to the west ?) and – at least as observed in the Kinzao quarry in the Noqui granite body – with a left-lateral component of limited displacement. These shear zones/corridors (to be observed on remote sensing images as “lineaments”) affect all the geological rock units of the (broader) Matadi region. Thus, the region can be subdivided in a mosaic of smaller “uniform” blocks showing a continuity of the various mapped units within each of the envisaged block. Each of these “first order” blocks can tentatively be further subdivided into a series of subblocks (out of the scope of our study).

2) MGE 2 Penetrative regional tectono-metamorphic overprint under greenschist facies conditions of all protolith rocks, however with often variable intensity of deformation because of 1) rapid

variations of competence of some of the rock sequences (see lithostratigraphic successions in the Matadi region) and 2) subregionally more strongly expressed stress conditions, eventually leading to hectometric thick mylonitic packages (with gentle dips and related to a fold-and- thrust belt geometry ?).

3) MGE 3 Intrusion of the peralkaline Noqui granite body (and subsidiary microgranitic veins) into the (pre-existing) Matadi Formation (metaquartzitic host rocks). The Palabala Formation may no longer be considered a lithostratigraphic unit but corresponds to a tectono-structural unit (see former discussion above). It does not underly the Matadi Formation, whose base is not observed (and thus also a major angular unconformity with the nearby Kimezian basement is lacking).

As a result of the (forcefull ?) intrusion of the Noqui body, a broad and gently dipping dome-like structure developed in the Matadi Formation which is currently still well-expressed to the north and west of the Noqui body (as presently exposed). This setting suggests a subsurface prolongation of the Noqui granite at limited depth beneath the town of Matadi and across the Congo River along its northern (right) bank. Thermal overprint and development of a contact metamorphic aureole in the host Matadi Formation is thus to be expected and may well be noticeable in the region of the Matadi town on both sides of the Congo River. Strong alkaline (late) metasomatism (because of the peralkaline affinity of the Noqui granite) seems obvious both in the Noqui body and in the contact metamorphic aureole.

4) MGE 4 Intrusion of the Mpozo syenite body. The precise outline of the body is poorly documented and the host rock unobserved. “Microgranitic” and/or “felsic veins” of the “Noqui type” clearly cross-cut the syenite body. At least “locally” (along a ravine of a few km of length ?) the Mpozo body is in tectonic contact with the Noqui body. Obsolete very unprecise radiometric data (U-Pb on bulk zircon) point to a late Palaeoproterozoic emplacement age (?).

2.5. THE AIM OF THIS STUDY In this study an effort is made to better constrain the geology of the Matadi region in continuity with the work of Behiels (2013). However the main focus is put on the question raised by Behiels (2013) concerning the “felsic magmatic bodies” of limited extent which preferentially are exposed in the vicinity of the town of Matadi on both sides of the Congo River banks and along the Mpozo tributary river (Fig. 2-14, small quadrangle). We will try to uncover the origin of these magmatic rocks and their emplacement conditions (extrusive or intrusive). Furthermore we will investigate whether these felsic bodies are related to the other larger magmatic bodies in the region, i.e. the Mpozo syenite and the Noqui granite. For all these magmatic rocks we will try to determine their source and whether they have resided in the same magma chamber or if they are related by similar petrogenetic processes.

The choice of our specific study region (Fig. 2-14) has been made purposely by reference to MGE 1 to MGE 4:

− As far as MGE 1 is concerned, the study region (Fig. 14, small quadrangle) is roughly delimitated by the shear zone of the “Chaudron de l’Enfer” (to the W) and by the shear corridor of the

Mpozo tributary/river (to the E). It is thus located within one block to limit (and/or eliminate) as much as possible the effects of MGE 1 on the various rock units under study.

− As far as MGE 2 is concerned, our specific study region falls out of the thick mylonitic package (see point 2 of MGE 2, here above) but displays typically the conditions of variable intensity of deformation of the protolith rocks as described in point 1. As a result, careful and thorough microscopic observation of the various protolith rocks is essential to decipher the complex geological evolution of the region before any attempt to “translate” this in a reliable geological map.

− As far as MGE 3 is concerned, our specific study region comprises purposedly a large portion of the Matadi Formation, exposed to the north of the Noqui granite body in and around the town of Matadi (on both sides of the Congo River) with the Mpozo tributary/river as general boundary to the east.

− As far as MGE 4 is concerned, our specific study region includes the northernmost part of the Mpozo syenite body.

To achieve the ambitious goals of our study, we will investigate fields observations of earlier field geologists and macroscopic samples, followed by an extensive petrographic study of thin sections. A next section will focus on the geochemistry of the “felsic magmatic bodies”, which will be compared to the geochemical signature of the Noqui and Mpozo bodies. Then we will focus on the aspect of modern geochronology of magmatic rocks of the Matadi region.

In chapter eleven, devoted to a discussion of the general geology of the Matadi region, we will test our results – including the new hard data and constraints - in the light of the working hypotheses as proposed in the preliminary and tentative four-stage timetable (MGE 1 to MGE 4), summarized here above. The results will also be integrated in a (very) crude new geological sketch map and compared to the 2008 map as well as to the map of Bertossa and Thonnart (1957).

3. METHODS

3.1. FIELD OBSERVATIONS AND MACROSCOPIC DESCRIPTIONS

To improve the geological map of the Matadi region, new observations are desired. In continuity with the update work on the Noqui granite and the Mpozo syenite bodies (Behiels, 2013), our work focuses on the “felsic magmatic bodies”, which are mainly intercalated in the metaquartzites of the Matadi Formation. Therefore a region, indicated by a blue quadrangle (Fig. 2-14), has been selected. It forms a “corridor” subperpendicular to the general structures and lithologic formations in the Matadi region and is located along the Congo River and its smaller Mpozo tributary.

During the second half of the 20th century, three field geologists performed surveys in the area and made precise localizations of the studied outcrops on maps at the scale of 1:25000. In 1950 Hugé examined outcrops south of the Congo River. His observations were written down in “La géologie des environs de Matadi” (Hugé, 1950), and his samples were stored at the Royal Museum for Central Africa (RMCA) (Tervuren, Belgium). Two other collections were sampled respectively by Massar (1965) and Steenstra (1970), who both went north and south of the Congo River. As the study area of Hugé, Massar and Steenstra comprised a much larger territory than our selected study area, only parts of their collections were considered.

At the RMCA the hand specimens were subjected to a macroscopic description. To maintain the large amount of data, a database was established in Microsoft Excel. This database comprises the interpretation of the field geologists and petrographic information such as colour, grain size, textures, identifiable minerals and grains and deformation. Based on these petrographic characteristics an effort was made to describe the samples thoroughly. Besides hand specimens, Hugé, Massar and Steenstra also described their field observations, which are kept in the archives of the RMCA. These descriptions, their localizations and hand specimens resulted in a total of 308 observations.

Over the years a large amount of field photos became available as a result of various field missions, often of short duration, which include missions of 2004, 2011 and 2013 (Tack and Baudet, 2014). As field access was not possible to us, these photos are of great importance to give a clear view on both the topography and the geology of the area. Therefore they will be incorporated in this study.

3.2. MICROSCOPIC DESCRIPTIONS

As the hand specimens are affected by a tectono-metamorphic overprint, which varies in intensity of local deformation and competence-incompetence of the sedimentary or magmatic protolith. Their macroscopic observation is often insufficient to identify the protolith. To obtain a correct lithological determination and geological mapping, it is thus crucial to make microscopic observations.

When the collections first arrived at the RMCA, thin sections with a thickness of 30 µm were made of most samples. These thin sections were studied in Ghent with the Olympus BH2 polarization microscope and photographs were taken by the ColorView I camera, with the aid of the Analysis imaging-process software. A microscopic study allows us to identify the minerals and describe the microscopic textures and therefore the rocks, and their protoliths, can be properly identified.

Combined results of the field observations, hand specimens and thin sections allowed the observations to be colour coded, based on their identification. This way the different lithologies, represented by different colours, can be plotted on a map. Separate “cross-sections”, both lithological and structural, were constructed for Hugé, Massar and Steenstra and are accompanied by one composite “cross-section”. The results are then compared to some earlier maps of the region (Behiels, 2013; Annex 1) and to preliminary sketch maps based on remote sensing data (Tack and Baudet, 2014).

To determine the modal composition of the rocks a point counting analysis was carried out. During this quantitative observation the thin section is moved stepwise so that a virtual grid is created. Next, each mineral under the cross hairs of the ocular is identified and counted. A minimum total of 300 counts per thin section was achieved.

3.3. GEOCHEMISTRY

20 samples were selected for a geochemical analysis, performed at the RMCA geochemical laboratory by Navez and Monin. Major elements analysis of these whole rock samples was carried out with inductively coupled plasma atomic emission spectrometry (ICP-AES) while the trace elements were determined with inductively coupled plasma mass spectrometry (ICP-MS). In the next section the general procedure (Monin, 2014) is described. For a more detailed description, and information on problems related to analytical procedures we refer to Totland et al. (1992) and García de Madinabeitia et al. (2008).

3.3.1. Sample preparation

The first steps of the sample preparation took place at Ghent University. As most samples were quite large, they were cut into smaller blocks on the diamond saw, which was also used to remove weathered material from the edges of the samples. The remaining parts were then thoroughly washed and dried, before introducing them to the jaw crusher. The jaw crusher reduces the rock fragments to a finer grain size so that it is suitable for the disc mill which further reduces the grains to a fine powder. Further sample preparation was carried out at the RMCA.

3.3.1.1. Loss on ignition To determine the loss on ignition all samples were first dried overnight at 105°C. Next 1 g of the powdered sample, placed in a platinum crucible, was weighed. This was placed in a muffle furnace and heated to 1000°C for 1 hour. At this temperature the sample loses all of its volatile components and ferrous iron (Fe2+) is oxidized to ferric iron (Fe3+). To determine the loss on ignition the sample is weighed before and after heating. The change in mass, expressed as a weight percentage of the dry mass, is presented as LOI. Negative LOI values are possible due to the oxidation of ferrous to ferric iron which results in a slight gain in weight after heating.

3.3.1.2. Alkaline fusion To analyze the major and trace elements, the geological samples are digested by alkaline fusion. For this process, subsamples of 0,200 g were mixed with 1 g of lithium metaborate (LiBO2 Aldrich) in a platinum crucible. This mixture was then fused for 1 hour in a muffle furnace at a temperature of 1000°C. The fusion of rock samples with lithium metaborate results in the formation of glasses. These glasses can be dissolved when submersed in nitric acid. Therefore the crucible containing the hot

33 melt was then immersed in 120 ml of water which was acidified with 12,5 ml of HNO3. With the help of a magnetic stirrer, the solution was stirred for one night until complete dissolution. This solution was then transferred quantitatively to a 250 ml volumetric flask. In this volumetric flask the solution was diluted to volume so that the final product contains a 5% nitric acid solution.

The blanks and the standards were taken through the exact same process as the unknown samples. For the major element analysis seven certified reference materials were used: BHVO-1 (basalt), GA (granite), SGR-1 (shale), JB-3 (basalt), AC-E (granite), JGb-1 (gabbro) and JG-1a (granodiorite). A multi- element standard solution is used for the trace element analysis.

3.3.2. Major elements

3.3.2.1. ICP-AES Inductively coupled plasma atomic emission spectroscopy (ICP-AES) is an emission spectroscopic analytical technique in which atoms and ions are excited by inductively coupled plasma. As these excited atoms and ions return to their ground state, they emit element-specific wavelengths. The intensity of the emission is a result of the concentration of that element within the sample. The measurements were carried out by the Thermo Jarrel Ash IRIS Advantage. This type of measurement set-up (Fig. 3-1) comprises three important units: a sample introduction system, an inductively coupled plasma torch and a detector system.

A peristaltic pump is used to import the sample into the pneumatic nebulizer, where the sample, together with argon, is converted to an aerosol. Pneumatic nebulizers produce aerosols with a broad distribution of droplet diameter, up to 100 µm. Of the aerosol, only 1 % of droplets are small enough (< 10 µm) to be efficiently ionized (Linge and Jarvis, 2009). Therefore a spray chamber screens the aerosol and is used to remove large droplets through gravity and inertia. The spray chamber also removes solvent from the aerosol, which improves ionization efficiency. After conversion of the sample into an aerosol, the aerosol is injected into the ICP for ionization.

Figure 3-1: Major components of a typical ICP-AES instrument. From Boss & Fredeen, 1997.

A plasma is a highly ionized gas that is made up of ions, electrons and neutral particles, usually at high temperature. An inductively coupled plasma (ICP) is a plasma in which the transfer of energy, to create and maintain the ionized gas, is carried out via electromagnetic induction. Time-varying magnetic fields keep this process ongoing. Different gases can be used to produce plasma, but argon is the most common one used, as it is relatively inert and does not form stable compounds.

The ICP is created by a quartz torch that consists of three concentric tubes. This torch is placed inside a water cooled copper coil. The copper coil supplies the magnetic field that transfers energy to plasma. A radio frequency generator creates an alternating current within the coil and induces an intense electromagnetic field around the tip of the torch. Argon gas flows through the torch. Due to a high-voltage, spark free electrons are produced which are accelerated in the oscillating magnetic field, causing collisions and ionization of the argon gas, with plasma formation at the open end of the quartz torch as a result. As the sample is introduced into the plasma via the injector tube, each droplet is vaporized to a gas. Compounds become atomized and individual atoms are ionized.

Due to the high temperature of the plasma, approximately 7000 K (Boss & Fredeen, 1997), the sample exists now as excited atoms and ions. As the ionized elements return to their ground state, electromagnetic radiation is emitted. The energy of the emitted radiation is proportional to its frequency according to E = h*c/λ (Harris, 2005). In this equation h is defined as Planck’s constant, c equals the speed of light and λ represents the wavelength.

When the emitted light passes from the plasma through the optical spectrometer it is separated in constituent wavelengths and then focused onto the detector. The Thermo Jarrel Ash IRIS Advantage is equipped with a Charge Injection Device (CID) detector (Coelho, 2013), which is composed of doped silicon wafers, containing a two dimensional array of light. As the photons reach the detector, they liberate electrons which are then trapped in the pixel sites. The signal is digitized and displayed via the user interface.

3.3.3. Trace elements

3.3.3.1. ICP-MS Present day inductively coupled plasma-mass spectrometry (ICP-MS) is considered as the most powerful multi-element analytical technique available. For the analysis the Thermo Fisher Scientific X-Series 2 was used. Figure 3-2 shows the different components of an ICP-MS instrument. To analyze the sample it must pass four important steps. The first two steps are identical to the initial steps of the ICP-AES, after which the ions are transferred to the quadrupole mass spectrometer and eventually reach the detector system.

After ionization by the ICP, the ions are sampled through a two-stage interface. The ions first pass through a sampler cone, which is a metal disk with a small orifice of about 1 mm diameter, which is in direct contact with the plasma. After passing through the sampler cone, the plasma gas arrives into a low pressure region of between 1 x 10-2 and 1 x 10-1 kPa where it expands as a supersonic jet (Linge & Jarvis, 2009). The central section of the jet flows through a second skimmer cone which is located directly behind the sampler cone. The skimmer cone has a smaller orifice than the sampler cone (0,4 – 0,7 mm). The purpose of these cones is to sample the center portion of the ion beam coming from the ICP torch. As the sampler and skimmer cones have small diameter orifices the amount of dissolved solids in the samples should be low. It is recommended that the samples have

35 no more than 0,2 % total dissolved solids (TDS) (Linge & Jarvis, 2009). If the amount of TDS is too high the orifices in the cones can become blocked, causing decreased sensitivity and detection capability. After passing through the cones, lenses focus and transport ions to the mass analyzer. These lenses comprise a series of metal plates or rings, each with a specific voltage. Ions that have a different mass will respond different to changes in lens voltage. It is impossible to optimize the voltage on the lenses such that all ions are transported with the same efficiency. The lens voltages are optimized that way, so that a maximum sensitivity is reached for the isotopes in the middle of the mass range. This way transmission of heavier and lighter elements is sacrificed.

Figure 3-2: Schematic representation of the main components in an ICP-MS. From Linge and Jarvis, 2009.

The next step is to separate the ions from the ion beam so that each element can be quantified. A mass spectrometer can differentiate between ions based on the mass-to-charge-ratio. A quadrupole mass filter, which consists of four metal rods that are suspended in parallel to the ion beam, and which are also equidistance from the ion beam, is used. Each rod is electrically connected to the opposite rod and voltages are applied to both rod pairs. Ions which enter the quadrupole travel down the central axis. Due to the applied voltages to the rods, the ions oscillate. The magnitude of the oscillations depends on both the mass and the charge of the ion. Extreme oscillations cause the ion to strike the rods or the inside of the quadrupole housing. The rod voltages are optimized to ensure that only ions of a single m/z have a stable path and exit the quadrupole. The mass filter must be switched to sequentially filter for each m/z of interest. This switching process is very fast. Data can be collected for a range of 0-300 amu (atomic mass unit) in about 100 ms (Linge & Jarvis, 2009).

At last, the ions are counted by pulse counting. Each detected ion is converted into a discrete electrical pulse. The number of pulses depends on the number of analyte ions present in the sample and can be converted into an absolute concentration by comparing the signal from a sample with that from a calibration reference sample. An electron multiplier increases the signal to be measured.

3.4. GEOCHRONOLOGY

The magmatic rocks of the Matadi region have been subjected to former dating efforts. Cahen et al. (1976) tried to date some of the felsic rocks and in section 2.3 the results are given of the bulk zircon U-Pb analysis of the Mpozo syenite (Delhal & Ledent, 1978). More recent geochronological data (Tack et al., 2001; Behiels, 2013) include the SHRIMP single zircon U-Pb geochronology results of magmatic rocks within the Zadinian (Noqui granite) and Mayumbian Groups. These results, together with microscopic observations, allowed to select four samples for U-Pb dating, including one Noqui granite sample, one “felsic magmatic body” sample and two Mpozo syenite samples (pink and white facies).

3.4.1. Sample preparation

The selected samples were crushed and milled by a jaw crusher and disc mill. The obtained fine powder was then sieved, using 63 and 250 µm meshes. The fraction between 63 and 250 µm was selected and subjected to wet sieving, in order to remove remaining clay minerals. After drying the samples, a Frantz magnetic separator was used to remove magnetic minerals. The apparatus was used at progressively higher magnetic currents of 0,1; 0,5; 0,8 and 1,2 ampere and the non-magnetic fraction was retained. Using heavy liquids (± 2,8 kg/l), heavy minerals were separated. Following this procedure, the remaining non-magnetic heavy mineral separates were handpicked under a binocular microscope. For each sample, around 80 zircons were selected and mounted on a sticky tape. During the next step, the zircons were embedded in 25 mm epoxy mount and polished for LA-ICP-MS. These polished zircons were photographed under the binocular microscope with the aid of the ColorView I camera. These photos allow to give a good description of size, shape, colour, transparency and inclusions. Cathodoluminescence (CL) images were made using the Jeol JSM-6400 scanning electron microscope (SEM) revealing inherited cores and oscillatory magmatic zoning.

Figure 3-3: Schematic representation of a laser ablation system, using ICP-MS detection. From Russo et al., 2002.

3.4.2. LA-ICP-MS

U-Pb analyses were executed with the Laser Ablation-Inductively Coupled Plasma-Mass Spectrometry (LA-ICP-MS) at Adelaide Microscopy (University of Adelaide). This set-up (Fig. 3-3) comprises the technique (ICP-MS) explained in 3.2.3.1 but makes use of a different sample introduction system, which allows solid samples to be analyzed. The sample, which is placed in the ablation chamber, is ablated by a finely focused laser. For this purpose the New Wave 213 nm Nd-YAG (Neodymium- doped Yttrium Aluminum garnet) laser was used. This laser has a spot size of 30 µm and a typical pit depth of 30 – 50 µm. Next the ablation chamber is flushed with an inert gas to transport the ablated sample, the analyte, towards the Agilent 7500cs ICP-MS for analysis.

4. FIELD OBSERVATIONS AND MACROSCOPIC DESCRIPTIONS

4.1. FIELD OBSERVATIONS Field access was not possible to us, which makes the observations of earlier field geologists of great importance. In this section we will present some of these observations and, if possible, illustrate them with digitized sketches or pictures (Tack and Baudet, RMCA archives). These illustrations help to obtain a clear view on the topography and geology of the Matadi region. As a large amount of data is available (Tack and Baudet, RMCA archives), a selection was made of the most important aspects.

The Matadi region is characterized by a hilly topography. As the Congo River crosses the area, it has a large impact on the relief. Along the Congo River banks elevation is approximate to sea level. As we move away from the river, altitudes quickly rise, reaching up to several hundreds of meters high. The highest points in the landscape are caused by the Noqui granite (Fig. 4-1), with the 502 m high “Pic Cambier” (PC).

Figure 4-1: Topography of the Matadi region. Low elevations, near the Congo River, quickly rise towards the highest points, caused by granitic rocks, creating a hilly topography. Outcrops near the harbour are in Gangila Formation; PC = Pic Cambier; note the strong demographic pressure in the town of Matadi.

In the Matadi region, the basement comprises Palaeoproterozoic migmatitic paragneisses and amphibolites of the Kimeza Supergroup (Fig. 4-2A). They are overlain by the metaquartzites of the Matadi Formation which may be slightly to strongly deformed because of variations in intensity of tectono-metamorphic overprint. Within the slightly deformed rocks it is possible to observe features evidencing their sedimentary protolith. These features include ripple marks (Fig. 4-2B) and cross- bedding (Fig. 4-2C). In many cases cross-bedding is enhanced by oblique laminae of heavy minerals.

Figure 4-2: A) Migmatitic paragneiss and amphibolite of the Kimezian basement. B) Ripple marks in the metaquartzites of the Matadi Formation (Steenstra, 1970); C) Cross-bedding in the metaquartzites of the Matadi Formation; D) Mafic intrusion (dyke), bordered by the red dotted lines, cross-cutting the metaquartzites of the Matadi Formation.

Field geologists have observed felsic and mafic intrusions (Fig. 4-2D) within the metaquartzites of the Matadi Formation. Massar describes concordant intrusions (sills) of various thickness. Some intrusions are only 40 cm thick while others reach a width of several meters. In Figure 4-3 we present a digitized version of one of his sketches, in which the intrusive nature of the felsic and mafic magmatic rocks can be observed. This sketch indicates that felsic and mafic intrusions can occur close to each other suggesting reactivation of earlier weakness zones in the Matadi Formation during episodic intrusive events. Several other observation points also indicate that these mafic intrusions (both sills and dykes) often occur close or together with felsic intrusions. Furthermore the mafic intrusions are often intrusive in the felsic ones.

Figure 4-3: Digitized sketch of Massar’s observation point 186.

Behiels (2013) suggested that the felsic intrusions might be related to the Noqui granite. This granitic massif forms the highest topographic points in the region (Fig. 4-1). Figure 4-4A represents a view on the 502 m high “Pic Cambier” in the Noqui granite. Due to supergene weathering these rocks display a typical boulder morphology. Figure 4-4B represents relatively fresh blocks of the Mpozo syenite, indicating a white and pink facies. Unlike the Noqui granite, the Mpozo syenite is only poorly exposed along the Mpozo tributary. Behiels (2013) stated that the Noqui and Mpozo massifs are tectonically in contact with each other (Fig. 4-4C). Furthermore it is observed that the Noqui granite is intrusive in the rocks of the Mpozo massif (Tack, 2014b). This is an important aspect, as it has a significant contribution in the age relation between the two massifs, indicating that the Mpozo massif is relatively older than the Noqui granite.

Figure 4-4: A) “Pic Cambier” (PC) with typical boulder morphology of the weathered Noqui granite; B) White and pink facies of the Mpozo syenite; C) Tectonic contact between the Noqui granite and Mpozo syenite, indicated by the red dotted line, following a local ravine; The syenite body dips underneath the granite body (photo taken near Mpozo river bridge, view towards the west; In foreground railway and valley of Mpozo tributary.

The top of the Matadi Formation is locally covered by the Yelala conglomerate. Where this conglomerate is not present, the metaquartzites are directly covered, conformably to slightly unconformably, by the Gangila Formation (Fig. 4-5A), but the metaquartzites then display isolated (stretched) quartz pebbles and are more coarse-grained (Fig. 4-5B) These metabasalts form a thick package of different superimposed metric flows. Each of them is characterized by massive (competent) rocks at the base and more deformed rocks with amygdales at the top of one single flow (Figs. 4-5C and 4-5D)

Ga. F.

Ma. F.

Figure 4-5: A) Matadi Formation (Ma. F.) covered by Gangila Formation (Ga. F.); B) Isolated (and stretched) pebbles in the top layer of the metaquartzites; C and D) Superimposed metric metabasaltic flows with massive rocks at the base and amygdaloidal rocks at the top of each flow. Amygdules correspond to vacuoles formed by degassing of each basaltic unit during flow emplacement. The red lines border one flow and the dotted line separates the lower massive part from the upper incompetent part of the flow.

A section of the panoramic assemblage of photos (Fig. 4-6), from west to east, illustrates the right banks of the Congo River and part of the general geological setting of our selected region of study (Fig. 2-14; small quadrangle). The exact location of the panoramic view is indicated in Figure 2-14 by a yellow dotted line. The complete series of photos is given in Annex 1.

Figure 4-6: Panoramic assemblage of photos, from west to east (1 followed by 2), illustrating the northern right banks of the Congo River near Matadi (view taken from “Belvédère”). Note 1) the overall west dipping dip slope morphology (= slope of the landscape roughly determined by and approximately conforming with the direction and the angle of dip of the underlying rocks); 2) the relatively difficult access due to the steep slopes of the Congo River with strong current, expressed by local whirlpools; 3) predominant grass to small bush savannah with small “forest gallery” in ravines; 4) the absence of human settlements (in contrast with the very strong demographic pressure of the town of Matadi; 5) overcast sky typical of the dry season. GA. F. = Gangila Formation; MA. F. = Matadi Formation; Yellow dotted line indicates the contact between the Gangila Formation and the Matadi Formation; Red dotted lines indicate observed dip slope lines; white star indicates the access location of the former ferry.

4.2. MACROSCOPIC DESCRIPTIONS An extensive macroscopic study of the hand specimens resulted in a database (Annex 2), which comprises our own observation results of hand specimens and petrographic information such as colour, grain size, textures, identifiable minerals and grains and deformation as well as the interpretation of the field geologist. As all rocks are to some extent metamorphic, three groups of rocks have been distinguished, according to the type of protolith: 1) rocks with a felsic magmatic protolith, 2) rocks with a sedimentary protolith and 3) rocks with a mafic magmatic protolith. Furthermore it is possible to identify a variation in the degree of foliation in the collection. Based on this degree of foliation the samples were further divided into three groups: massive, foliated and strongly foliated. In this section we describe the rocks solely based on hand specimens. Because of the variety in protoliths and in the degree of deformation microscopic studies are essential for further characterization of the rocks, resulting eventually in more correct determinations and names.

4.2.1. Felsic magmatic protolith Rocks with a felsic magmatic protolith are characterized by large feldspar crystals in a fine-grained groundmass and thus display a porphyritic texture. The colour of the hand specimens mainly varies from brownish grey over pinkish grey to light and dark grey. As most rocks display a phyllitic luster they are described as feldspathic phyllites. In the next section we give an example of a massive, a foliated and a strongly foliated rock with a felsic magmatic protolith.

4.2.1.1. Massive Hand specimen RG 89.540 (Fig. 4-7) comprises a massive brownish grey rock. It is characterized by the presence of fine to medium grains in a fine-grained groundmass. These grains have an average size of 2 mm, a pink colour and represent feldspar. Besides pink crystals, aggregates of dark material occur throughout the rock. On its rough surface (Fig 4-7B), the rock displays a phyllitic luster and some small white mica flakes can be observed.

Figure 4-7: Pictures of RG 89.540. A) Cut surface; B) Rough surface.

4.2.1.2. Foliated Sample RG 89.876 comprises a pinkish grey foliated rock. This sample comprises pink porphyroclasts which are surrounded by a fine-grained (< 1 mm) groundmass. These porphyroclasts can best be observed on the cut surface (Fig. 4-8A). They comprise feldspar and have a maximum size of 5 mm. On the rough surface (Fig. 4-8B) small white mica flakes can be observed, which give the rock a phyllitic luster.

Figure 4-8: Pictures of RG 89.876. A) Cut surface; B) Rough surface.

4.2.1.3. Strongly foliated Hand specimen RG 89.595 is a strongly foliated rock which displays a light grey to green colour. The rock is characterized by the presence of medium (1 – 5 mm) to large ( > 5 mm) porhyroclasts in a fine-grained groundmass (Fig. 4-9A). Most crystals have an average size of 5 mm, show a white colour, and comprise feldspar. Some crystals, with an average size of 2 mm, display a greasy luster and might indicate the presence of high temperature quartz. This rock also displays a phyllitic luster on its rough surface (Fig. 4-9B), hence the greenish colour.

Figure 4-9: Pictures of RG 89.595. A) Cut surface; B) Rough surface. The greenish yellow spot was induced by a marker.

4.2.2. Sedimentary protolith During his fieldwork, Hugé mainly focused on the presence of rocks with a sedimentary protolith. Most of these rocks are fine-grained and display a light grey colour. In some samples, the presence of small (± 1mm) black porphyroblasts of magnetite can be observed. As most of these rocks display a phyllitic luster on their rough surface, they are described as phyllitic metaquartzites.

4.2.2.1. Slightly foliated Sample RG 89.592 comprises a slightly foliated metaquartzite. From its cut surface (Fig. 4-10A) it is impossible to identify separate grains as the rock is fine-grained. The rough surface of the rock (Fig. 4-10B) displays a light grey colour and exhibits a phyllitic luster on its cleavage plane (Fig. 4-10C).

Figure 4-10: Pictures of RG 89.592. A) Cut surface; B) Rough surface; C) Cleavage plane.

4.2.2.2. Moderately foliated Hand specimen RG 19.685 involves a moderately foliated phyllitic metaquartzite. The rock displays a beige colour and is fine-grained. Both on the cut surface (Fig. 4-11A) and the rough surface (Fig. 4- 11B) the presence of approximately 1 mm large black pyramidal porphyroblasts of magnetite can be observed.

Figure 4-11: Pictures of RG 19.685. A) Cut surface; B) Rough surface.

4.2.2.3. Strongly foliated Sample RG 19.639 includes a strongly foliated phyllitic metaquartzite. The specimen is fine-grained and has a medium grey colour. Both on the cut surface (Fig. 4-12A) and on the rough surface (Fig. 4- 12B) the presence of a vein can be noticed. In Figure 4-12B one can also observe that the rock displays incipient folding.

Figure 4-12: Pictures of RG 19.639. A) Cut surface; B) Rough surface.

4.2.3. Mafic magmatic protolith All field geologists identified the presence of rocks with a mafic magmatic protolith. After studying the hand specimens, it became clear that this group of rocks comprises different varieties. As it is out of the scope of this study to describe a massive, a foliated and a strongly foliated rock of each variety, we will describe only one sample of each variety. These varieties include amphibolites, dolerites and green phyllites. Further characterization of these rocks by thin sections could confirm or contradict the names given to these rocks.

4.2.3.1. Amphibolite RG 89.868 (Fig. 4-13) comprises a strongly foliated amphibolite. The hand specimen displays alternating lenses of black and greenish grey lenses. The black lenses presumably comprise amphiboles while the greenish grey lenses contain plagioclase. The yellowish green tone of the plagioclase is caused by saussuritization.

Figure 4-13: Pictures of RG 89.868. A) Cut surface; B) Rough surface.

4.2.3.2. Metadolerite Hand specimen RG 19.655 (Fig. 4-14) is a massive rock with a dark grey to green colour. On the cut surface one can observe the presence of two colours: dark grey and yellowish-greenish grey. The dark colours reflect mafic minerals and the light yellowish-greenish grey colours represent saussuritized plagioclase. Because the rock has a fine-grained and doleritic texture it is described as a metadolerite.

Figure 4-14: Pictures of RG 19.655. A) Cut surface; B) Rough surface.

4.2.3.3. Green phyllite RG 89.535 (Fig. 4-15) comprises a strongly foliated rock with a greenish dark grey colour. On the cut surface the presence of medium grains in a fine-grained groundmass is visible. Thin sections are necessary to determine which minerals make up these medium grains. On its rough surface (Fig. 4- 15B) a phyllitic luster can be observed. These phyllitic minerals probably comprise biotite and chlorite, which would explain the dark colour of the rock.

Figure 4-15. Pictures of RG 89.535. A) Cut surface; B) Rough surface.

5. MICROSCOPIC DESCRIPTIONS

In this section a comprehensive study of thin sections is presented. The described thin sections create an overview and try to represent the variety within the rocks. Similar to the previous chapter we categorize the rocks in three groups based on their protolith. More information on the remaining thin sections is given in Annex 3.

5.1. FELSIC MAGMATIC PROTOLITH A first group of rocks comprises metamorphic and deformed rocks with a felsic magmatic protolith. All three field geoligsts, i.e. Hugé, Massar and Steenstra, noticed the presence of these rocks in the field. Especially Massar focused on this type of rocks and collected the largest amount of samples. All of the specimens are characterized by large crystals in a more fine-grained groundmass. As the rocks are deformed and cataclastic, the term blastoporphyritic applies to them. Based on the intensity of deformation the rocks are subdivided into three subgroups: slightly, moderately and strongly deformed. Slightly deformed rocks are characterized by porphyroclasts that are only slightly fractured and that are therefore easily identified. Moderately and strongly deformed rocks contain more fractured, broken and recrystallised porphyroclasts. This makes it harder to recognize the individual former phenocrysts. Contrary to strongly deformed rocks, the largest fraction of porphyroclasts in moderately deformed rocks can still be recognized.

5.1.1. Slightly deformed RG 89.546 Thin section RG 89.546 is an alkali feldspar schist which comprises large crystals of alkali feldspar, surrounded by a fine-grained groundmass. As these large crystals are slightly fractured and broken, they can be described as porphyroclasts. These porphyroclasts are mainly euhedral (Fig. 5-1A) and subhedral. The average size of the alkali feldspar crystals is 1 mm, but maximum dimensions can reach 3 mm. Most porphyroclasts lie isolated within the groundmass, only a few are in contact with each other. Some crystals display simple Carlsbad twinning, but all crystals are characterized by the presence of exsolution lamellae. These textures comprise exsolution lamellae of sodic feldspar within potassic feldspar, which is described as a perthitic texture. Sometimes the amount of sodic and potassic feldspar is equal, which allows the textures to be called mesoperthitic. Within the sodic exsolution lamellae it is possible to detect polysynthetic twinning, indicating the presence of albite. Different shapes of perthites can be observed. According to the classification of Bard (1980) we can distinguish flames, patches, interpenetrant and chessboard-type (Fig. 5-1B) perthites in this thin section.

The porphyroclasts are partially or completely surrounded by polycrystalline rims which have a different structure than the groundmass (Fig. 5-1C). According to Passchier and Trouw (2005) these are called porphyroclast systems. If the material in the rim has the same composition as the porphyroclast, the structure is described as a mantled porphyroclast. If the rim has a different composition compared to the porphyroclast, the structure is known as a porphyroclast with strain shadows. Since the porphyroclast is composed of alkali feldspar and its recrystallised rim consists of quartz, muscovite and feldspar, the term strain shadow applies. Within these strain shadows, quartz occurs as polygonal quartz displaying triple junctions, creating a granoblastic polygonal texture. The crystals have both straight and irregular edges and only display slight undulatory extinction. Besides

49 quartz, elongated crystals of muscovite are present as well. Some of these muscovite crystals pin the quartz crystals. Muscovite crystals are also very common at the edges of the porphyroclasts (Fig. 5- 1D). Within these strain shadows, alkali feldspar displays tartan twinning, indicating the presence of microcline. Crystals with polysynthetic twinning, indicating plagioclase, can be found as well. One of the alkali feldspar porphyroclasts is broken, which indicates brittle fracturing. In between the fracture polygonal quartz is present (Fig. 5-1D). Feldspar porphyroclasts often contain inclusions of quartz and even more often inclusions of sericite. A B

C D

E F

Figure 5-1: RG 89.546: A) Euhedral porphyroclasts of feldspar; B) Chessboard perthite; C) Strain shadow of polygonal quartz around an alkali feldspar porphyroclast; D) broken alkali feldspar porphyroclasts with recrystallised polygonal quartz in the fracture and muscovite crystals around the feldspar crystals; E) Muscovite crystals are absent in pressure shadows; F) irregularly dispersed opaque minerals.

The fine-grained groundmass is mainly composed of quartz, alkali feldspar and plagioclase with a size of approximately 50 µm. Although not abundant, some plagioclase crystals display polysynthetic twinning. Quartz can often be recognized by its undulatory extinction. Some aggregates of polygonal quartz crystals, with mainly straight edges, are present. Throughout the matrix, small and elongated crystals with high interference colours occur. In plane polarized light (ppl), these crystals that are colourless to pale green, could be identified as muscovite. At a higher magnification it becomes clear that most of them display a slight preference orientation. In the pressure shadows of some K- feldspar porphyroclasts these muscovite crystals are absent (Fig 5-1E), which might indicate that the original matrix is still present here.

In plane polarized light the presence of opaque minerals (Fig. 5-1F) is evident. These minerals exhibit very irregular shapes and have a maximum size of 50 µm. Even though they are rather small, they are very abundant, though irregularly dispersed, throughout the thin section.

RG 89.590 RG 89.590 is a quartz-alkali feldspar-muscovite-biotite schist and is made up of large alkali feldspar and quartz crystals which are surrounded by a fine-grained groundmass.

Subhedral and anhedral porphyroclasts of alkali feldspar range in size between 1 and 5 mm and display perthitic and mesoperthitic textures. The shape of all these perthites can be indicated as interpenetrant or patchy. Within the sodium rich lamellae polysynthetic twins, indicating albite, can be observed. At the edges of the porphyroclasts occur polycrystalline rims that are dominated by polygonal quartz (Fig. 5-2A). Elongated crystals of muscovite are also part of these strains shadows and are very outspoken around the edges of the porphyroclasts. Besides quartz and muscovite there is also microcline (Fig 5-2B) and plagioclase present, with respectively tartan and albite twinning. Biotite also occurs within these strain shadows. This association found in the strain shadows also occurs between the broken fragments of crystals.

Porphyroclasts of quartz (Fig. 5-2C) are anhedral and are on average 1 mm large. They all show undulatory extinction and sometimes display deformation lamellae. The edges of these quartz porphyroclasts are all very irregular. Fractured porphyroclasts of quartz also occur. Some porphyroclasts are partially rimmed by polygonal quartz creating mantled porphyroclasts. As these mantled clasts do not display wings, they can be described as ϴ-type mantles. Aggregates of polygonal quartz occur as well. They display triple junctions with mainly straight edges but they can also be stepwise or irregular (Fig. 5-2D). It should be mentioned that some porphyroclasts of quartz contain inclusions of feldspar and vice versa.

The groundmass is composed of alkali feldspar, plagioclase and quartz. Polysynthetic twins can often be observed indicating the presence of plagioclase. These polysynthetic twins sometimes tend to taper out, which might indicate that these twins were formed during deformation. Throughout the thin section a few lenses of recrystallised material, with a larger grain size than the surrounding groundmass, occur (Fig. 5-2E). Within these lenses one can mainly find polygonal quartz, with both straight and irregular edges, but also plagioclase and K-feldspar.

Muscovite occurs throughout the thin section as elongated crystals. As they mainly have the same orientation they give rise to a foliation and a lepidoblastic texture. Within the recrystallised parts of the rocks, muscovite often pins the polygonal quartz crystals. In the pressure shadows of some

51 porphyroclasts no muscovite occurs. An important fraction of muscovite occurs as sericite and is found as inclusions within alkali feldspar porphyroclasts. Inclusions of sericite also occur in quartz but are less frequent. A B

C D

E F

G H

Figure 5-2: RG 89.590: A) Strain shadow of quartz and muscovite around an alkali feldspar porphyroclast; B) Microcline displaying tartan twinning in a strain shadow; C) Anhedral porphyroclasts of quartz; D) Aggregate of polygonal quartz, supposedly representing a former phenocryst; E) Recrystallised lenses of polygonal quartz; F) cluster of biotite with epidote; G) blasts of octaeder shaped opaque minerals; H) Inclusions of sericite, biotite, chlorite and epidote within a K- feldspar porphyroclast.

Besides muscovite and sericite there is also an important fraction of biotite present (Fig. 5-2F). Biotite occurs as elongated crystals that are oriented in the same way as muscovite. Some crystals deviate from this orientation and occur in clusters with epidote, opaque minerals and allanite. All biotite crystals are strongly pleochroic. Their colour varies from pale yellowish brown to dark greenish brown. Under crossed polarizers the typical birds-eye-extinction can be observed. Biotite often also occurs as inclusions within feldspar and quartz porphyroclasts.

Accessory minerals are epidote, opaque minerals, allanite and chlorite. All of these minerals sometimes occur in clusters with biotite. The opaque minerals display straight edges and can be rectangular or octaeder shaped (Fig. 5-2G) suggesting that they are magnetite. These crystals were probably formed by blastesis. Chlorite can be found as an alteration product of biotite and thus also occurs as inclusions (Fig. 5-2H) within porphyroclasts. Besides chlorite, epidote can also be found as inclusions within alkali feldspar porphyroclasts. Allanite has a high relief, a dirty brown appearance and can mainly be found in clusters together with biotite and/or epidote.

RG 89.876 Thin section RG 89.876 is a cataclastic quartz-alkali feldspar-muscovite schist. The rock is blastoporphyritic with porphyroclasts of alkali feldspar and quartz. The feldspar crystals range up to 7 mm in size. These porphyroclasts are surrounded by a fine-grained groundmass that is composed of alkali feldspar, plagioclase and quartz. As RG 89.876 is very similar to RG 89.590 only the differences and additional features will be discussed.

At the edges of a few alkali feldspar porphyroclasts vermicular symplectites occur (Fig. 5-3A). As they are intergrowths of K-feldspar and quartz, they can be called granophyric. Symplectites can also be observed within the matrix. As they are small, it is not clear if they are myrmekites or granophyric intergrowths. Intergrowths are not only constricted to the edges of crystals, but they also occur within porphyroclasts (Fig. 5-3B).

Calcite is a mineral which was not observed in RG 89.590 but that is clearly present in RG 89.876. It mainly occurs within strain shadows (Fig. 5-3C) or in fractures between crystals (Fig. 5-3D). It is also present as inclusions within feldspar porphyroclasts. Calcite can be recognized by its very high interference colours and variable relief. Large crystals often display multiple sets of twins.

Contrary to RG 89.590 some porphyroclasts in RG 89.876 have wings at their edges. These wings are φ(phi)-type wings. Wings at the edges of alkali feldspar porphyroclasts consist of polygonal quartz, muscovite and calcite. Wings around quartz porphyroclasts do not contain calcite.

Some larger white mica crystals occur. These muscovite crystals can be 200 µm large and are kinked (Fig. 5-3E). The other smaller crystals occur within the matrix and create a lepidoblastic texture.

Just as in RG 89.590 opaque minerals, allanite, chlorite (Fig. 5-3F) and epidote are accessory minerals, but in RG 89.876 there is also sphene present. All of these minerals, except chlorite, often occur in clusters. Chlorite mainly appears as inclusions in feldspar porphyroclasts or together with calcite in fractures of feldspar porphyroclasts. RG 89.876 does not contain large amounts of biotite. Biotite is present, but it only occurs as rather small crystals which are often partly altered to chlorite, and is thus also accessory.

A B

C D

E F

Figure 5-3: RG 89.876: A) Granophyric intergrowth at the edge of a K-feldspar crystal; B) Intergrowth inside a K-feldspar porphyroclast; C) Calcite within a strain shadow; D) Calcite together with quartz and muscovite within a fracture; E) Kinked biotite; F) Alteration of chlorite to biotite at the edge of an opaque porphyroblast.

5.1.2. Moderately deformed RG 89.541 RG 89.541 is a fine-grained, moderately deformed blastoporphyritic rock. The porphyroclasts have an average size of 0,5 mm and consist of alkali feldspar (Fig. 5-4A). They comprise sub- to anhedral crystals, are often broken and display exsolution lamellae creating a perthitic texture. A few crystals display Carlsbad twinning. Contrary to the other thin sections there is no recrystallised material present around the porphyroclasts. Most porphyroclasts lie isolated in the more fine-grained groundmass but clusters of crystals also appear (Fig. 5-4A), resembling a glomeroporphyritic texture.

A B

C D

Figure 5-4: RG 89.541: A) Cluster of K-feldspar porphyroclasts; B) Irregular and prismatic crystals of epidote; C) Epidote mainly at the edges of porphyroclasts; D) Presence of chlorite and biotite.

The groundmass consists of alkali feldspar, plagioclase and quartz. Polysynthetic twinning indicates the presence of plagioclase and quartz is characterized by its undulatory extinction.

Remarkable is the large amount of epidote. Small green crystals are very abundant throughout the thin section. They have a maximum size of 200 µm and often display an irregular or prismatic shape (Fig. 5-4B). The epidote crystals also occur as inclusions in the alkali-feldspar crystals or can be found around their edges (Fig. 5-4C).

Accessory minerals are sphene, biotite, chlorite and opaque minerals of which biotite and chlorite (Fig. 5-4D) often occur as inclusions in K-feldspar. Chlorite can be observed as an alteration product of biotite. The opaque minerals are small and have irregular shapes.

RG 89.544 The quartz-alkali feldspar schist, RG 89.544, is a blastoporphyritic rock in which the porphyroclasts are mainly made up of alkali feldspar, but large crystals of quartz are present as well. These two minerals together with plagioclase also constitute the fine-grained groundmass.

Porphyroclasts of alkali feldspar are mainly subhedral and can be up to 6 mm in size. On average they are 2 mm in size and often Carlsbad twins (Fig. 5-5A) can be observed. All of the alkali feldspar crystals display perthitic textures, with chessboard and patchy perthites being the dominant types. Remarkable is the large amount of symplectites (Fig. 5-5B) that occur throughout the thin section. These granophyric intergrowths mainly appear at the edges of the porphyroclasts and have a vermicular appearance. Porphyroclasts are often rimmed by polygonal crystals of quartz, alkali feldspar and plagioclase. In these strain shadows (Fig. 5-5C) microcline displays tartan twinning and some plagioclase crystals have albite twins which tend to taper. These polygonal crystals can also be found between broken fragments of crystals.

A B

C D

E F

Figure 5-5: RG 89.544: A) Alkali feldspar porphyroclast displaying perthites and Carlsbad twinning; B) Granophyric intergrowth at the edges of a K-feldspar crystal; C) Strain shadow consisting of quartz, microcline and plagioclase; D) Quartz porphyroclast; E) Biotite altered to chlorite and a cluster of sphene, allanite and opaque minerals; F) Large amount of small irregular opaque minerals.

Porphyroclasts of quartz (Fig. 5-5D) are less common and are maximum 1 mm in size. They all show undulatory extinction and are anhedral. They often display some fractures and contain inclusions of alkali feldspar. The amount of recrystallised material around these porphyroclasts is much smaller and consists mainly of quartz.

The groundmass is mainly dominated by quartz and alkali feldspar. The alkali feldspar crystals regularly display perthites and some crystals with tartan twinning can be observed as well. A third mineral making up the groundmass is plagioclase. It is much less abundant than quartz and alkali feldspar but can often be recognized by its polysynthetic twinning. As these twins often tend to taper

56 out, they are probably deformation twins. Within the matrix it is also possible to recognize large amounts of symplectites.

Accessory minerals are muscovite, sphene, allanite, epidote, chlorite, biotite and opaque minerals. Muscovite is present as small and elongated crystals within the matrix. Contrary to other thin sections, it is much less abundant and can be regarded as accessory. These small crystals do not seem to follow an outspoken orientation. Some larger muscovite crystals can be found within the strain shadows of alkali feldspar porphyroclasts. Besides white micas there is also a small amount of biotite present. Figure 5-5E shows a biotite crystal that has altered to chlorite. Right next to this crystal a cluster of sphene, allanite and opaque minerals is situated. These minerals, together with epidote, commonly occur as clusters. These cluster forming minerals can be observed within the matrix but they can also occur separately as inclusions within the alkali feldspar porphyroclasts. In plane polarized light the large amount of very small opaque minerals (Fig. 5-5F) becomes clear.

RG 89.540 RG 89.540 is a plagioclase-biotite-muscovite schist. It comprises a blastoporphyritic rock made up of plagioclase and K-feldspar porphyroclasts which are surrounded by a fine-grained groundmass. A large fraction of the porphyroclasts is broken. The fractured clasts are often grouped in clusters giving the impression of a glomeroporphyritic texture (Fig. 5-6A).

Porphyroclasts are sub- to anhedral and are maximum 4 mm in size. The largest fraction is made up of plagioclase while a smaller part consists of alkali feldspar. Crystals of alkali feldspar often display perthitic textures. Within these exsolution lamellae albite twins occur. Porphyroclasts of plagioclase are characterized by polysynthetic twinning (Fig. 5-6B). The polysynthetic twins mainly occur at the centre of the crystal. The crystals are often zoned, and the outer rim is mostly free of polysynthetic twinning (Fig. 5-6C). At the edges of some porphyroclasts one can find symplectic intergrowths (Fig. 5-6C). Vermicular intergrowths between plagioclase and quartz are called myrmekites. Around the porphyroclasts some, rather thin, rims of recrystallised material occur. These strain shadows mainly consist of polygonal quartz (Fig. 5-6D). Plagioclase and alkali feldspar are sometimes also part of these strain shadows. Plagioclase in strain shadows can often be recognized by its albite twins, which often tend to taper out.

The groundmass is made up of quartz, alkali feldspar and plagioclase. Alkali feldspar is perthitic. Plagioclase is characterized by the presence of albite twins. These twins often tend to taper which might indicate that they have formed during deformation. Recrystallised material within this groundmass occurs as well.

Muscovite, sericite and biotite are common throughout the thin section. They do not seem to form a foliation but they mainly occur as inclusions within the porphyroclasts. Biotite is pale yellowish brown to dark greenish brown pleochroic. It mainly occurs at the edges of porphyroclasts in strain shadows (Fig. 5-6E) or in fractures of porphyroclasts. This way it seems to form clusters.

Epidote, sphene, allanite and opaque minerals are accessory. It was observed that epidote makes up a large part of the inclusions occurring within the plagioclase crystals. It can also occur with the other accessory minerals, and with biotite, in clusters. The opaque minerals within this thin section are less well developed crystals with irregular edges.

A B

C D

E F

Figure 5-6: RG 89.540: A) Broken plagioclase porphyroclasts; B) Polysynthetic twinning in plagioclase; C) Zoned plagioclase crystals. The outer rim is free of albite twins and at the edge a myrmekite can be observed. D) Plagioclase crystals with a rim of polygonal quartz; F) Biotite in strain shadows; G) Cluster of biotite and epidote.

RG 89.533 RG 89.533 displays all the same features as observed in RG 89.540 but additionally there are blasts of opaque minerals present (Fig. 5-7). These opaque minerals exhibit well developed edges, and can thus be called idioblastic. A B

Figure 5-7: RG 89.533: A) Idioblastic opaque crystal with XPL; B) with PPL

RG 89.863 RG 89.863 comprises a moderately deformed quartz-alkali feldspar-muscovite schist. The alkali feldspar porphyroclasts are sub- to anhedral, have an average size of 2 mm, display perthites and are surrounded by a fine-grained groundmass. This groundmass comprises quartz, alkali feldspar and plagioclase.

Remarkable in this thin section is the large amount of quartz. Quartz occurs as polygonal aggregates (Fig. 5-8B). These aggregates often have an elongated shape. In these aggregates the quartz crystals display stepwise and sometimes irregular edges. Crystals with deformation lamellae and fluid inclusions are present. Quartz is also the dominant mineral in the groundmass and also has a polygonal habitus with mainly straight and stepwise edges.

Muscovite occurs as elongated crystals throughout the thin section, displaying a preferred orientation. Smaller white mica crystals of sericite can be observed as inclusions in the alkali feldspar porphyroclasts.

Accessory minerals comprise allanite, sphene, epidote and opaque minerals. These opaque minerals (Fig. 5-8D) occur as very irregular patches throughout the thin section. A B

C D

5-8: RG 89.863: A) Alkali feldspar porphyroclast; B) polygonal quartz aggregate with stepwise and irregular edges; C) elongated polygonal quartz aggregate surrounded by a fine-grained groundmass, dominated by quartz, and muscovite with a preference orientation; D) Opaque minerals.

5.1.3. Strongly deformed RG 89.594 Alkali feldspar-muscovite-biotite schist, RG 89.594, is a blastoporphyritic rock in which the porphyroclasts are made of alkali feldspar. Porphyroclasts of alkali feldspar are maximum 4 mm large and have an average size of 2 mm. The crystals are sub- to anhedral and often display perthitic textures. In Figure 5-9A it can be seen that these exsolution lamellae are deformed. Some crystals do not display perthitic textures but show tartan twinning, indicating the low temperature variant microcline (Fig. 5-9B). All of the alkali feldspar porphyroclasts show undulatory extinction and are often fractured or broken. Within these fractures one can often find quartz, which can also occur as inclusions (Fig. 5-9C). Some crystals display distinct wings. These wings consist of polygonal quartz and muscovite. They can be identified as δ-type wings and they look like fringes (Fig. 5-9D). A B

C D

E F

Figure 5-9: RG 89.594: A) Deformed exsolution lamellae; B) Microcline characterized by tartan twinning in between two K-feldspar crystals; C) Small quartz grains within the fractures of a K-feldspar crystal; D) Fringed δ-type wings of quartz at the edge of a K-feldspar crystal; E) Small and parallel oriented crystals of sericite making up the groundmass; F) Accicular needles, probably rutile, within biotite.

The groundmass is different from the one observed in previous thin sections. Quartz, alkali feldspar and plagioclase are all recrystallised and surrounded by very small and elongated crystals of sericite. Quartz occurs as polygonal aggregates. It also occurs as strings of polygonal crystals creating ribbons. Alkali feldspar mainly shows tartan twinning and plagioclase can be identified by its polysynthetic lamellar twins. These twins sometimes taper out, indicating that they are deformation twins.

Sericite makes up a large part of the groundmass and occurs as small and elongated crystals (Fig. 5- 9E). They all display a distinct orientation creating a lepidoblastic texture. They can occur between broken crystals and are often bent. Sericite is also important as inclusions within feldspar porphyroclasts.

Accessory minerals are biotite and chlorite. Biotite can occur as inclusions in the feldspar porphyroclasts but it also occurs as somewhat larger, 200 µm large, crystals that are not oriented. They sometimes display pleochroic halos and can be kinked as well. They show pale yellow to orange brown pleochroism. Some biotite crystals display inclusions of acicular needlelike crystals (Fig. 5-9F) that are probably rutile.

RG 89.861 RG 89.861 comprises a very strongly deformed blastoporphyritic rock. Only two porphyroclasts, one quartz and one alkali feldspar crystal, can be observed. These porphyroclasts are surrounded by a very fine-grained groundmass. Due to the presence of phyllitic minerals, it is described as a phyllonite.

The alkali feldspar porphyroclast (Fig. 5-10A) has a size of approximately 1,5 mm. At its edges recrystallised polygonal quartz and muscovite form wings. Based on its shape one can define them as δ-type wings. Polygonal quartz and muscovite also occur around the 1 mm large quartz porphyroclast (Fig. 5-10B) but there they do not form wings.

The groundmass is mainly made up of quartz and smaller feldspar crystals. These crystals all seem to be recrystallised. In some places larger recrystallised material of polygonal quartz form structures that can be described as ribbons.

Small and elongated crystals of muscovite create a very strongly outspoken foliation and a lepidoblastic texture (Fig. 5-10C). Small sericite crystals also occur as inclusions within the feldspar porphyroclast.

Accessory minerals are opaque minerals, biotite and epidote. Biotite occurs as elongated crystals within one layer and displays the same orientation as the muscovite crystals. It is pale yellow to orangy brown pleochroic. In some places it deviates from this orientation and forms clusters with aggregates of epidote. Epidote occurs throughout the thin section as rather small and irregular crystals. In one place these epidote crystals occur together with opaque minerals. This group of epidote and opaque minerals seem to occur within what might have been a former crystal (Fig. 5- 10D), and probably did not form primarily. A few blasts of opaque minerals were found as well. Figure 5-10F represents an opaque blasts with recrystallised wings.

A B

C D

E F

Figure 5-10: RG 89.861: A) Alkali feldspar porphyroclast with recrystallised δ-type wing; B) Quartz porphyroclasts; C) ribbon quartz; D) Lepidoblastic texture created by the parallel orientation of sericite; E) cluster of epidote and opaque minerals in what might have been a former crystal; F) opaque blasts with recrystallised wings.

5.1.4. Summary of observations All rocks that are derived from a felsic magmatic protolith are characterized by porphyroclasts of alkali feldspar and quartz, surrounded by a fine-grained groundmass. In some samples there are also porphyroclasts of plagioclase present. The more fine-grained groundmass comprises quartz, alkali feldspar and plagioclase.

Porphyroclasts of alkali feldspar range in size between 0,5 and 7 mm and are often broken. They all display perthitic textures, mostly comprising chessboard-type or patchy perthites. Within the exsolution lamellae, polysynthetic twins occur, indicating albite. Around these porphyroclasts we often observe strain shadows, comprising polygonal quartz, alkali feldspar and plagioclase together

62 with muscovite and/or biotite. Granophyric intergrowths at the edges of the K-feldspar porphyroclasts are also common.

Porphyroclasts of quartz are usually less abundant than alkali feldspar. They are always anhedral with irregular edges and display undulatory extinction. The porphyroclasts are often partially rimmed by mantles of polygonal quartz or the polygonal aggregates completely replace the former phenocryst. Fractured porphyroclasts occur multiple times.

Porphyroclasts of plagioclase are only present in a few samples. The porphyroclasts are characterized by polysynthetic twinning. Zoning often also occurs within these crystals. Mostly the centre of the crystals comprises polysynthetic twins, while the outer rim does not. Myrmekites sometimes appear at the edges of the crystals.

All rocks contain micas. These micas always include muscovite/sericite and sometimes also, mostly dark brown, biotite. As these micas display a preferred orientation, creating a lepidoblastic texture. The rocks, except for RG 89.861 which is a mylonite, can best be described as schists. Sometimes muscovite and biotite crystals are present without orientation. These crystals are supposedly formed by blastesis in a static environment. All of these micas often occur as inclusions within the porphyroclasts. Sericite is the dominant inclusion in alkali feldspar porphyroclasts.

Accessory minerals comprise chlorite, epidote, opaque minerals, sphene and allanite. Chlorite is mainly present as an alteration product of biotite, and thus also appears as inclusions. The four other minerals often occur in clusters. Epidote, sphene and allanite often appear as inclusions in porphyroclasts, whereby epidote is the dominant inclusion in plagioclase crystals. Opaque minerals are mainly present as small and irregular crystals, but sometimes they display well developed crystal edges. In the latter case, they are considered blasts.

5.2. SEDIMENTARY PROTOLITH The rocks discussed in the next section are metamorphic and deformed rocks with a sedimentary protolith. These rocks consist almost completely of quartz and white mica. Microscopically it is difficult to determine whether the rocks are slightly, moderately or strongly deformed. Therefore we categorize the rocks here in these three classes based on their macroscopically observed deformation.

5.2.1. Slightly deformed RG 89.592 RG 89.592 is mainly made up of quartz and muscovite (Fig. 5-11A). Quartz occurs as polygonal aggregates, displaying triple junctions. The edges of these quartz crystals are mostly straight and stepwise. The rock has a seriate texture with the size of the quartz crystals ranging between 50 and 500 µm. Therefore the rock is best described as a seriate-polygonal metaquartzite. The polygonal quartz aggregates often display subgrains (Fig. 5-11B) and sometimes display undulatory extinction.

Muscovite occurs as elongated crystals. As they have a preferred orientation and there are nearly parallel to each other, they create a lepidoblastic texture. Some crystals pin the polygonal quartz crystals, while smaller crystals of sericite/muscovite occur as inclusions within these polygonal quartz aggregates. In Figure 5-11C a band dominated by muscovite can be observed.

Accessory minerals include opaque minerals and epidote. The opaque minerals have irregular shapes (Fig. 5-11D) and epidote occurs as small and individual crystals throughout the thin section. Epidote is not very abundant and can be recognized by its green colour and high relief. A B

C D

E F A B

G H

Figure 5-11: RG 89.592: A) Polygonal quartz and elongated muscovite crystals; B) Polygonal quartz aggregate with subgrains and triple junctions; C) Muscovite band; D) Irregular opaque minerals. RG 19.685: E) polygonal quartz with irregular boundaries creating a seriate-interlobate texture; F) polygonal quartz and parallel oriented muscovite; G) accessory epidote; H) octaeder shaped opaque blasts, supposedly of magnetite.

5.2.2. Moderately deformed RG 19.685 RG 19.685 is dominated by quartz grains, displaying a granoblastic polygonal texture. These polygonal quartz crystals often display irregular edges (Fig. 5-11D), but stepwise and straight edges also occur. Therefore the texture can be described as seriate-polygonal to seriate-interlobate.

Elongated crystals of muscovite are abundant. They are oriented parallel (Fig. 5-11E) to each other, making up a lepidoblastic texture. Small inclusions of sericite/muscovite occur within the quartz crystals.

Opaque minerals and epidote (Fig. 5-11F) make up the accessory fraction of the thin section. Two groups of opaque minerals can be observed. A first group comprises irregular patches of opaque material. The other group comprises crystals with well developed edges, often displaying an octaeder shape (Fig. 5-11G).

RG 19.643 Just like RG 19.685, RG 19.643 is made up of quartz, muscovite, accessory epidote and small irregular opaque minerals. The thin section is mentioned because it contains a small vein (Figs. 5-12A and 5-12B) in which the quartz aggregates have a different structure. The quartz grains within the vein are larger than the surrounding polygonal quartz crystals. They show undulatory extinction and contrary to the surrounding quartz, they do not contain inclusions of muscovite/sericite. The edges of the different crystals are very irregular and show features of bulging.

RG 19.656 RG 19.656 is somewhat different from all the other rocks discussed in this section. The rock is mainly composed of quartz, describing a seriate-polygonal to seriate-interlobate (Fig. 5-15C) texture.

Biotite is mainly present as anhedral crystals. They are pale brown and have a dirty appearance (Fig. 5-15D). Most of them do not display a distinct cleavage. A few crystals are better developed and are pale yellow to dark greenish brown pleochroic.

Remarkable is the presence of a colourless mineral with high positive relief, identified as zoisite (Figs. 5-15E and 5-15F). It occurs mainly as prismatic crystals with a well developed cleavage. Under crossed polarizers first order grey, blue and yellow interference colours are observed. The crystals display anomalous interference colours and parallel extinction. Most of these crystals are oriented parallel to each other.

Muscovite can be regarded as accessory. It only occurs as small elongated crystals with a preferred orientation. Other accessory minerals are irregularly shaped opaque minerals, sphene and epidote. Chlorite occurs as an alteration product of biotite.

A B

C D

E F

Figure 5-12: RG 19.643: A) quartz vein; B) The grains constituting the quartz vein are free of sericite inclusions and have very irregular edges. RG 19.656: C) seriate – polygonal to seriate interlobate quartz; D) pale brown to dark greenish brown biotite; E) prismatic zoisite crystals; F) anomalous interference colours of zoisite.

5.2.3. Strongly deformed RG 19.669 RG 19.669 is a phyllitic metaquartzite, comprising approximately equal amounts of quartz and muscovite (Fig. 5-13A). Quartz displays a polygonal shape with mainly stepwise, but also irregular, edges. The different grains have approximately the same size and thus describe an equigranular – polygonal to equigranular – interlobate texture.

Muscovite occurs as elongated crystals, which are oriented parallel to each other forming a lepidoblastic texture. Due to the large amount of muscovite, the quartz crystals are often separated from each other.

A B

C D

Figure 5-13: RG 19.669: Equal amounts of polygonal quartz and sericite with a lepidoblastic texture; (B) Kinked biotite; (C) Biotite partially altered to chlorite; (D) Syntectonic poikiloblasts of garnet with quartz inclusions.

Accessory minerals are biotite, garnet, chlorite, epidote and zoisite. Contrary to muscovite, biotite does not exhibit a preferred orientation. They have an average size of approximately 400 µm and display pale yellow to orange brown pleochroism. The biotite crystals are often kinked (Fig. 5-13B) and sometimes altered to chlorite (Fig. 5-13C). The random orientation of the biotite crystals might indicate that these crystals were formed by blastesis.

Garnet occurs as a pale brown mineral. As garnet is an isotropic mineral, it appears black under crossed polarizers. These crystals contain inclusions of quartz, and can therefore be described as poikiloblasts. Inside these blasts the inclusions are slightly rotated (Fig. 5-13D), which might indicate that these garnet crystals have formed as syntectonic to posttectonic blasts. At the edges of these garnet poikiloblasts there is commonly biotite present.

RG 19.639 RG 19.639 is dominated by quartz with irregular edges (Fig. 5-14A). The thin section can be divided into two different parts. The two parts are displayed in Figure 5-14B. The lower half contains very irregular quartz aggregates with muscovite crystals between the quartz grains. The upper section contains more coarse-grained quartz. Muscovite is absent in this part of the thin section.

The first part of the thin section thus comprises muscovite. These muscovite crystals are elongated and display a preferred orientation, creating a lepidoblastic texture (Fig. 5-14C). Quartz in this section is characterized by irregular edges. These sutured contacts often bulge into each other. Due to their irregular shapes the quartz aggregates can be described as interlobate to amoeboid. Accessory

67 minerals within this section are epidote, biotite, chlorite and a few opaque minerals. Aggregates of small epidote crystals (Fig. 5-14D) are common. Biotite occurs as a pale yellow to dark greenish brown pleochroic mineral with mainly the same orientation as the muscovite crystals. Alteration of biotite to chlorite sometimes occurs. The opaque minerals are all characterized by irregular shapes.

The part of the thin section without muscovite is displayed in Figures 5-14E and 5-14F. The quartz grains are larger than the ones in the other part of the thin section, but they also display very irregular edges and features of bulging. Figure 5-14F displays a close-up of these irregular grain contacts. Within this section almost no other minerals than quartz occur. Very few crystals of epidote and chlorite occur. A B

C D

E F

Figure 5-14: RG 19.639: A) Irregular quartz grains and muscovite; B) two different parts within one thin section: upper part: coarse-grained without muscovite, the lower part is more fine-grained and contains muscovite; C) lepidoblastic texture created by the parallel orientation of muscovite; D) aggregate of epidote; E) Irregular quartz grains indicating features of bulging; F) close-up of the irregular quartz grains.

5.2.4. Summary of observations The rocks with a sedimentary protolith are mainly composed of quartz and muscovite. Quartz occurs mainly as seriate-polygonal and seriate-interlobate aggregates. In between the quartz crystals, muscovite is present, often pinning the quartz grains. These muscovite crystals have an elongated shape and are oriented parallel to each other, creating a lepidoblastic texture.

Accessory minerals often comprise epidote and opaque minerals. Opaque minerals occur as small irregular crystals, but they also occur as blasts with well developed crystal faces. Less frequent accessory minerals are chlorite, biotite and zoisite. Garnet was only described in one thin section.

5.3. MAFIC MAGMATIC PROTOLITH All three field geologist, i.e. Hugé, Massar and Steentra, sampled metamorphic rocks with a mafic magmatic protolith. Hence, a lot of hand specimens of these rocks were available. Remarkably is that there were almost no thin sections made of these rocks, leaving us with only 13 thin sections. Therefore it is impossible to objectively divide the samples in groups of different degree of deformation. For this reason we try to give a representative account of the variation within this group of rocks, which moreover does not represent the main topic of our study (see sections 5.1 and 5.2).

RG 89.868 RG 89.868 is a moderately deformed rock. The dominant mineral in this thin section is characterized by strong pleochroism (Fig 5-15A and 5-15B). Colours vary from pale yellowish green to dark bluish green. Interference colours range up to second order yellow and pink. The minerals occur as subhedral prismatic crystals of which a large fraction shows a preference orientation. Besides prismatic crystals, we also observe minerals with an acicular habit. Throughout the thin section we also observe minerals with a deeper dark green colour, displaying typical amphibole cleavage (Fig. 5- 15C), which is characterized by 60° - 120° cleavage angles. This suggests the co-existence of actinolite and/or hornblende and the rock is best described as an actinolite schist and/or amphibolites.

In between the actinolite crystals, intergranular (occupying the space between the larger crystals) quartz appears. Most of these quartz crystals are anhedral, have an average size of 100 µm and display undulatory extinction. In some places polygonal quartz (Fig. 5-15D) with triple junctions can be observed. Quartz crystals within these polygonal aggregates do not display undulatory extinction. In all of the quartz crystals small fluid inclusions can be observed.

Epidote is also very abundant as an intergranular mineral. It has a pale yellowish green colour and occurs as small crystals with an average size of 50 µm. Due to its bright birefringence colours it is easily noticed under crossed polarizers.

Besides quartz and epidote, plagioclase (Fig. 5-15E) also appears as an intergranular mineral, but it is much less abundant than the other two minerals. Albite twins often characterize these minerals.

Under plane polarized light, the presence of sphene and opaque minerals becomes clear. Sphene occurs as polycrystalline masses which are often elongated. These polycrystalline masses of sphene can regularly be found around irregular patches of opaque material (Fig 5-15F).

A B

C D

E F

Figure 5-15: RG 89.868: A and B) Pleochroic actinolite; C) amphibole cleavage in actinolite; D) Polygonal quartz; E) intergranular plagioclase displaying polysynthetic twinning; F) Elongated cluster of sphene and opaque material.

Biotite is present as an accessory mineral. It occurs as small crystals which display pale yellowish brown to dark orange brown pleochroism.

RG 19.684 RG 19.684 is a moderately deformed actinolite schist that is rich in epidote. The rock thus mainly comprises epidote and actinolite (Fig 5.16A). Similar to the previous thin section, actinolite displays pale yellowish green to dark bluish green pleochroism, but here it occurs more in its acicular form. In this thin section the presence of epidote is prominent. Large masses of polycrystalline aggregates with a pale yellowish brown colour occur.

A B

C D

Figure 5-16: RG 19.684: A) Epidote bearing actinolite schist; B) Polygonal quartz; C) Accessory biotite; D) Irregular mass of opaque minerals.

Epidote and actinolite make up a large part of the rock. Besides these minerals, there is also a large amount of quartz present. The quartz crystals mainly occur as intergranular crystals between epidote and actinolite. Furthermore these quartz crystals display a polygonal habit (Fig. 5-16B), sometimes displaying undulatory extinction.

Biotite and opaque minerals are observed as accessory minerals. Biotite (Fig. 5-16C) displays pale brown to orange brown pleochroic colours. The opaque minerals (Fig. 5-16D) form irregular masses, spread throughout the thin section.

RG 89.535 Thin section RG 89.535 shows a strongly foliated biotite-chlorite-calcite schist. The phyllitic minerals form the dominant part of the rock, and comprise biotite and chlorite (Fig. 5-17A). These minerals are oriented parallel to each other, creating a lepidoblastic texture. Alterations of biotite to chlorite can be observed several times.

Throughout the thin section elongated lenses or aggregates of polygonal calcite, displaying triple junctions, occur. Within these calcite crystals it is possible to observe different sets of deformation twins (Fig. 5-17B). Within these aggregates of polygonal crystals, quartz occurs as well. This mineral is also present as inclusions within the calcite crystals (Fig. 5-17C).

Between the phyllitic minerals, intergranular quartz appears. These crystals are mostly polygonal and sometimes form ribbons (Fig. 5-17D). A cluster of polygonal quartz was observed as well (Fig. 5-17E).

A B

C D

E F

G H

Figure 5-17: RG 89.535: A) Chlorite and biotite are the dominant minerals; B) Twinning in calcite; C) Aggregate of polygonal calcite and quartz; D) Ribbon quartz; E) Aggregate of polygonal quartz; F) Corroded plagioclase with inclusions of quartz, calcite and epidote; G) randomly oriented muscovite crystal; H) Opaque mineral surrounded by epidote.

Plagioclase appears, just as quartz, as an intergranular mineral between the phyllitic minerals. Some crystals display polysynthetic twinning. One larger, but strongly corroded, crystal of plagioclase is displayed in Figure 5-17F. This crystal comprises a lot of quartz, calcite and epidote inclusions.

Muscovite, opaque minerals and epidote are accessory. Muscovite generally is subhedral and elongated (Fig. 5-17G). Contrary to the other phyllitic minerals, these muscovite crystals do not follow a preference orientation. They display one good cleavage and are characterized by their birds eye extinction. Opaque minerals occur as irregular patches throughout the thin section and are often surrounded by aggregates of epidote (Fig. 5-17H). Furthermore epidote is spread throughout the thin sections as small crystals.

RG 89.593 RG 89.593 exhibits two different and alternating textures. A coarse-grained part (Figs. 5-18A and 5- 18B), dominated by laths of plagioclase, alternates with a more fine-grained part (Figs. 5-18C and 5- 18D), which mainly comprises quartz.

The coarse-grained part includes laths of plagioclase, which display a random orientation. Polysynthetic twinning is common in these crystals and the edges are very irregular. Inclusions of epidote, quartz and sericite occur within the plagioclase crystals. In between the laths, intergranular quartz can be found. These quartz crystals often display a polygonal habitus with straight edges. Biotite and chlorite are also very dominant. The biotite crystals mainly display a random orientation and thus give rise to a decussate texture. Calcite also makes up a large fraction of the rock. Accessory minerals include epidote, sphene and opaque minerals. The latter ones are mainly irregularly shaped.

The fine-grained part of the rocks comprises the same minerals but displays a different texture. Even though this part of the rock also contains plagioclase, it is much less abundant and quartz is the dominant mineral. In these sections, quartz displays a polygonal texture. Biotite and opaque minerals are much less abundant compared to the coarse-grained part.

RG 19.655 In RG 19.655 (Fig. 5-18E) epidote is the dominant constituent. It displays a pale yellow colour. Therefore the rock might be best described as an epidosite. Besides epidote, there is also actinolite (Fig. 5-18F) present. These minerals display a pale yellowish green to dark bluish green pleochroism and mainly display an acicular habitat. In between the large epidote masses we also observed intergranular quartz. Occasionally aggregates of polygonal quartz do occur (Fig. 5-18G). These quartz grains often contain small and elongated inclusions, probably of sericite. Allanite is present as an accessory mineral, forming irregular brown patches (Fig. 5-18H).

A B

C D

E F

G H

Figure 5-18: RG 89.593: A) Coarse-grained part with XPL; B) Coarse-grained part with PPL; C) Fine-grained part with XPL; D) Fine-grained part with PPL. RG 19.655: E) RG 19.655, dominated by epidote; F) Bluish green actinolite; G) Polygonal aggregate of quartz with sericite inclusions; H) Accessory allanite.

5.3.1. Summary of observations. In the previous section we described several rocks with a mafic magmatic protolith. As these rocks display various textures, it is difficult to make a summary of our observations. Although the rocks display different textures, their mineralogy is similar. We summarize the observed mineralogy as follows.

Actinolite is present in the majority of the rocks. It is characterized by its pale yellowish green to dark bluish green pleochroism and displays an acicular or prismatic habit. As actinolite and hornblende display relatively similar optical characteristics, it is impossible to differentiate unambiguously between the two.

Epidote occurs in all samples. It is characterized by its high relief and generally has a pale yellowish green colour. Epidote can be very abundant, but in some rocks it occurs only as an accessory mineral.

In some thin sections we observe plagioclase. This mineral often displays polysynthetic twinning and is mainly lath-shaped.

Biotite sometimes makes up a large fraction of the rock. It often displays light to dark brown pleochroism and is mostly elongated. Biotite repeatedly occurs together with chlorite, which can also be very abundant. In some thin sections these minerals only occur as an accessory constituent.

In some rocks calcite is present. Calcite is easily recognized by its very high birefringence, its changing relief and its multiple sets of twins.

The amount of quartz strongly varies from rock to rock. It often occurs intergranular and is almost always polygonal.

Common accessory minerals are sphene, allanite, opaque minerals and muscovite/sericite.

5.4. POINT COUNTING ANALYSIS In the previous section we described the different types of rocks observed in the study area. From here on we will focus on the rocks with a felsic magmatic protolith. As we want to figure out whether these rocks are related to the rocks of the Noqui granite and the Mpozo syenite, we will compare them.

Using a point counting analysis, it is possible to determine the modal mineralogy of the rocks. Such point count analyses are time-consuming. Therefore we selected a few representative samples, which were subjected to a point counting analysis. For the Noqui granite, the results of Behiels (2013) were used. Furthermore one sample of the Mpozo syenite and three rocks with a felsic magmatic protolith were analyzed. During these analyses a minimum number of approximately 300 grains were counted and normalized to 100. The results of these analyses are given in Table 5-1.

Knowing the modal composition of the rocks, it is possible to plot the rocks in a QAPF diagram, also called Streckeisen diagram. To plot the samples the amount of quartz, alkali feldspar, plagioclase and feldspathoids are normalized. The results are displayed in Figure 5-19. Samples of the Noqui granite, indicated in blue, plot within field 2 and can thus be described as alkali feldspar granites. The red triangle is representative for the Mpozo syenite (RG 19.611). It comprises more plagioclase, less alkali feldspar and less quartz than the Noqui granite. It plots on the borderline of field 7* and 8* and is

75 therefore best described as a quartz syenomonzonite. In green the rocks with a felsic magmatic protolith are plotted. RG 89.590 and RG 89.876 plot, just like the Noqui granite, in the alkali feldspar granite field. RG 89.540 comprises more plagioclase, and plots therefore more to the right in the diagram in field 3a and is described as a syenogranite

Table 5-1: Modal mineralogy of two Noqui samples, one Mpozo sample and three rocks with a felsic magmatic protolith (FMP).

Noqui Noqui Mpozo FMP FMP FMP Sample 19223 89993 19611 89540 89590 89876 Quartz 32.60 30.10 12.25 30.49 30.10 46.36 Alkali feldspar 61.30 61.10 46.68 32.79 50.81 31.13 Plagioclase 0.50 1.70 26.16 14.43 2.59 2.65 Muscovite/Sericite 0.00 0.00 0.00 4.93 10.36 14.57 Biotite 0.80 3.70 8.94 10.82 4.53 0.00 Aegirine 4.20 1.70 0.00 0.00 0.00 0.00 Riebeckite 0.30 1.40 0.00 0.00 0.00 0.00 Epidote 0.00 0.00 0.99 2.62 0.32 0.33 Sphene 0.00 0.00 1.99 2.30 0.00 0.99 Allanite 0.00 0.00 0.00 0.00 0.65 1.32 Chlorite 0.00 0.00 0.66 0.00 0.00 0.00 Calcite 0.00 0.00 0.00 0.00 0.00 1.66 Fluorite 0.00 0.30 0.00 0.00 0.00 0.00 Opaque minerals 0.30 0.00 0.33 1.64 0.65 0.99 Total number of counts 377 352 302 305 309 302

RG 19.223 – Noqui granite

RG 89.993 – Noqui granite

RG 19.611 – Mpozo syenite

RG 89.540 – FMP

RG 89.590 – FMP

RG 89.876 – FMP

Figure 5-19: Modified QAPF diagram (after Streckeisen, 1974) with modal mineralogy. 2) Alkali feldspar granite; 3a) Syenogranite; 3b) Monzogranite; 4) Granodiorite; 5) Tonalite; 6*) Alkali feldspar quartz syenite; 7*) Quartz syenite; 8*) Quartz monzonite.

6. DISCUSSION FIELD OBSERVATIONS, MACROSCOPIC AND MICROSCOPIC DESCRIPTIONS

As both field observations, macroscopic and microscopic observations were necessary to fully characterize the rocks, all three aspects are integrated in this section. However main focus is put on the discussion of microscopic descriptions.

6.1. MINERAL ASSEMBLAGE Both hand specimens and thin sections evidence that tectono-metamorphic processes have affected the rocks. Therefore the metamorphic mineral assemblage might be very useful in defining the facies of regional metamorphism.

For the rocks with a felsic magmatic protolith, the dominant minerals are alkali feldspar and quartz. These two minerals generally make up the large porphyroclasts, but they are also abundant in the fine-grained groundmass. Furthermore plagioclase is present in the groundmass of almost all specimens, however in some rocks it occurs as large porphyroclasts. Besides these three main constituents we also observe a lot of muscovite and/or sericite. Biotite sometimes occurs as a very abundant mineral, whilst in some of the rocks it is accessory or even absent. As accessory minerals we generally observe epidote, chlorite, sphene, allanite and opaque minerals.

We can now compare this mineral assemblage to the ones given in Table 6-1 for the metagranitoids. According to this table the greenschist facies is characterized by albite, alkali feldspar, chlorite, quartzite and sometimes biotite, actinolite and epidote (Bucher and Grapes, 2011). We observe these minerals in the group of rocks with a felsic magmatic protolith, except for the ferromagnesian amphibole (actinolite), which we do not expect because of the felsic composition of these rocks. Therefore, we conclude that the mineral assemblage indicates regional greenschist facies metamorphism.

The effects of greenschist facies metamorphism are also visible in the rocks with a sedimentary protolith, which comprise metaquartzites of the Matadi Formation. These rocks display accessory epidote, and in the rocks with a mafic magmatic protolith which often contain abundant actinolite, epidote and chlorite. Furthermore we observe, mainly in the rocks with a felsic magmatic protolith, the effects of sericitization and saussuritization. These two processes typically occur during greenschist facies conditions and replace alkali feldspar by sericite and plagioclase by epidote, Figure 6-1: Pressure - temperature fields of metamorphic facies. From Bucher and Grapes (2011). respectively.

Table 6-1: Metamorphic facies and mineral assemblages. From Bucher and Grapes (2011).

The mineral assemblage thus indicates that all of the rocks in the Matadi region were affected by regional greenschist facies metamorphism. As greenschist facies temperatures generally range between 300 and 500 °C (Fig. 6-1) and pressures remain rather low, we have a good estimation of the P-T conditions that the rocks endured. At slightly higher temperatures, already above 450°C, the transition to amphibolite facies occurs. As this transition is gradual, it can be possible to find minerals characteristic for both the greenschist facies and amphibolite facies, in one specimen. This explains the concomitant occurrence of both actinolite and hornblende in RG 89.868.

6.2. RELICT TEXTURES The tectono-metamorphic overprint did not only cause a change in mineralogy, but also in textures. Before we discuss features induced by deformation, we focus on relict textures. These are textures inherited from the original protolith. The information in this section is based on Vernon (2004), unless mentioned otherwise.

6.2.1. Blastoporphyritic – porphyritic – texture Macroscopic and microscopic studies have shown that the rocks with a felsic magmatic protolith are characterized by large crystals in a more fine-grained groundmass. Dealing with deformed rocks, this texture is described as blastoporphyritic. However, before deformation occurred, these rocks were porphyritic. The formation of this texture traditionally consists of two phases. During a first phase the large crystals (phenocrysts) form, followed by a second phase of rapid crystallization of smaller crystals resulting in the fine-grained groundmass. It is however possible to form this type of texture in a single, uninterrupted, cooling phase, but for this specific case we refer to Vernon (2004). In the next paragraphs we consider the cooling history of the rocks more in detail.

Crystallization of liquid melts occurs as temperature drops. Before crystallization can occur, nucleation is necessary, a process which requires a certain degree of undercooling. Undercooling can be explained as a drop in temperature below the equilibrium freezing temperature, without the rock becoming solid. As the process of nucleation is still poorly understood, we will not discuss it. Once a certain amount of nuclei have formed, these nuclei will grow at the expense of smaller nuclei. This process is called ageing and during this stage no more new nuclei are created.

In Figure 6-2 both the growth rate (G) and nucleation rate (N) curves are displayed. As temperature drops at first, and undercooling increases, both G and N increase. But as the growth rate is larger than the nucleation rate, only a small amount of large crystals can form. During a next phase, when temperature decreases even more, both N and G decrease. This is due to the effect that diffusion goes slower with decreasing temperatures. As the growth rate drops faster than the nucleation rate, a lot of nuclei can be formed resulting in a lot of small crystals which constitute the groundmass.

Porphyritic textures in intrusive rocks can thus be explained by a slow crystallization at depth to form phenocrysts. This is then followed by a phase of rapid cooling to form the groundmass. Field observations and sketches of our felsic magmatic rocks of the Matadi region suggest that these rocks occur as intrusions of limited extent in the surrounding rocks. The large temperature difference between the cold surrounding rocks and the hot intrusive melt can explain the second phase of rapid cooling, which resulted in the fine-grained groundmass.

As the rocks are all deformed, it is possible that mylonitization of the rocks caused the large crystals to form smaller fragments, leaving a few remaining porphyroclasts. As we find some euhedral crystals, this hypothesis is probably not valid. Furthermore we observe, in all of the rocks, micas. These micas are sometimes absent in the pressure shadows of porphyroclasts. This might indicate that the original matrix can be observed in these areas. This aspect suggest that even before deformation, there was a fine-grained groundmass, favouring the hypothesis of originally non- deformed porphyritic rocks.

Based on the discussion above we no longer have to describe the rocks with a blastoporphyritic texture as “rocks with a felsic magmatic protolith”, but we can conclude that they are “hypabyssal rocks”. According to the definition of the IUGS (Fettes and Desmons, 2007) hypabyssal rocks are described as follows: “pertaining to an igneous intrusion, or to the rock of that intrusion, whose depth is intermediate between that of abyssal (= plutonic) and the surface”.

Figure 6-2: Curves displaying the variation of nucleation rate (N) and growth rate (G) with increasing undercooling. The general grain sizes and shapes produced at each stage are given. After Vernon (2004).

6.2.2. Symplectites Symplectite is a general term which refers to fine-grained intergrowths of two or more minerals. In the group of the hypabyssal rocks we observed two types of symplectites: granophyric intergrowths and myrmekites.

Granophyric intergrowths A granophyric texture is a micrographic intergrowth of quartz and alkali feldspar. According to Vernon (2004) several studies have pointed out that granophyric intergrowths approximate the composition of the ternary minimum in the Or-Ab-Qtz system. Therefore it is suggested that the texture is formed by simultaneous and rapid crystallization of the two mineral phases (alkali feldspar and quartz), although there are other possible ways to form this texture (Vernon, 2004).

In the hypabyssal rocks these granophyric textures generally occur at the edges of large K-feldspar crystals. In one of the thin sections, RG 89.544, the groundmass consists mainly of granophyric textures. This would mean that the larger crystals formed during a first phase of relatively slow cooling. As granophyric textures are induced by quick cooling, a second phase of rapid cooling followed. According to Bard (1980) this texture generally occurs in subvolcanic or hypabyssal rocks.

Intergrowths between quartz and K-feldspar often create a graphic texture, which looks similar to old runic writing. In our thin sections we do not observe this type of intergrowths but more vermicular shaped intergrowths. Vernon (2004) indicates that intergrowths in deformed rocks can become more rounded and ellipsoidal to spherical. This process is generally induced by heating, which lower the total interfacial free energy by reducing the total grain-boundary area. This process is discussed more in detail in a section 6.3.4.

Myrmekites Myrmekites comprise vermicular intergrowths of quartz and sodic plagioclase. The formation of myrmekites has led to numerous studies, resulting in various hypotheses. Contrary to the cotectic formation of granophyric textures, myrmekites are believed to form as a subsolidus process. Myrmekites can form directly during crystallization, but they are most commonly formed during deformation. Therefore it might be incorrect to discuss this texture as a relict texture, but we include it here as it is a form of symplectite.

Myrmekites regularly appear in high-grade metamorphic rocks and igneous rocks as a breakdown product of K-feldspar during retrograde metamorphism. As K-feldspar is replaced by plagioclase an excess amount of silica is released as quartz, causing the intergrowth of plagioclase and quartz.

6.3. TEXTURES INDUCED BY DEFORMATION All of the rocks in the Matadi region were affected by deformation. This resulted in complex rocks with various microstructures. In this section each of the observed microstructures will be explained in detail based on Vernon (2004) and Passchier and Trouw (2005). This allows us to determine which deformation processes affected the rocks.

There are various ways of classifying deformation mechanisms, but a primary distinction, at the microscope scale, should be made between brittle and ductile deformation. Brittle deformation is characterized by the presence of fractures across and/or between grains, in which the resulting

81 fragments often move relative to each other. This is in contrast with ductile deformation, where grains change their shapes or move relative to each other without fracturing at the grain scale.

6.3.1. Brittle deformation Brittle fracturing occurs at low temperature or at high strain rate and causes the rocks to change by fracture formation. This type of deformation is called cataclastic flow. As fracturing causes new surfaces, they are easily recognized by their sharp and straight nature. These microfractures are common within the hypabyssal rocks. The largest fractures occur within the alkali-feldspar crystals, but broken quartz crystals are also present. Feldspars, both alkali feldspar and plagioclase, generally deform by brittle fracturing over a wide range of conditions. For example, at low metamorphic grade (< 400 °C) feldspar is thought to deform mainly by brittle fracturing and cataclastic flow. This results in grain fragments with a wide range of grain size. Often patchy undulatory extinction and subgrains with vague boundaries can be present. At slightly higher conditions, but still low-medium grade (400 – 500 °C), feldspar still deforms by microfracturing, but it is accompanied by minor dislocation glide. Therefore we can observe tapering twins, bent twins, undulose extinction, deformation bands and kink bands with sharp boundaries. As the rocks were affected by greenschist facies conditions (section 6.1), these two temperature ranges apply, which is also confirmed by the thin sections. Furthermore we observed minerals as quartz, muscovite, calcite and chlorite within some of the fractures. These minerals are secondary and therefore the fractures can be described as healed fractures.

6.3.2. Ductile deformation Crystals can also deform internally without brittle fracturing. This type of deformation, called intracrystalline deformation, is caused by the migration of lattice defects, which comprise point defects and line defects or dislocations. As a result of intracrystalline deformation, various microstructures arise.

6.3.2.1. Deformation twinning One way of intracrystalline deformation comprises deformation twinning, which is most common in plagioclase. Generally twinning occurs in the lower temperature range of deformation. In our thin sections we only observed deformation twins in plagioclase and calcite. Microscopically deformation twins can easily be distinguished from growth twins (Fig. 6-3). Deformation Figure 6-3: Twinning in plagioclase: A) stepwise growth twins; twins never display simple twinning and are B) Wedge-shaped deformation twins which taper out. From Passchier and Trouw, 2005. usually wedge-shaped. Both growth twins and deformation twins can cross whole grains but if not, growth twins terminate abruptly with planar terminations while deformation twins taper out. Growth twins usually have a uniform width or are stepped (Fig. 6-3A). As deformation twins are wedge-shaped (Fig. 6-3B), this does not apply. Growth twins are often bounded by zoning. Deformation twins are generally concentrated at high strain sites like the rim of a crystal and generally taper out towards the centre of crystal.

6.3.2.2. Kinking Kinking is somewhat similar to twinning, but twinning is restricted to specific crystallographic planes and directions, which is not the case for kinking. Kinking is most observed in minerals with strongly anisotropic crystal structures. Therefore it occurs in minerals with only one slip plane such as micas. As a consequence of deformation, slip occurs on the slip plane. As this plane is inadequate to maintain homogeneous deformation, the grain sharply bends or kink and deformation localizes into kink bands. These kink bands enable shortening of the grain to continue.

In the observed thin sections several kinked biotite, and sometimes also muscovite, crystals were observed, both in the hypabyssal rocks and the metaquartzites.

6.3.3. Recovery and recrystallisation Deformation can build up a certain amount of dislocations within the crystal lattice. A larger concentration of dislocations results in a higher “internal strain energy”, as the increase in internal energy is proportional to the increase in total length of dislocations per volume of crystalline material. As dislocations in slip planes can interfere with each other, they form tangled dislocations. These inhibit further movement of the dislocations and thus also further deformation, a process described as strain hardening.

Two processes, recovery and recrystallization, reduce the concentration of dislocations. This way they allow the material to continue to deform. Therefore ductile deformation is often considered as a competition between strain strengthening and recovery processes.

6.3.3.1. Recovery Recovery is a term which comprises all processes which attempt to return crystals to their undeformed state without forming high-angle and high-energy boundaries. This means that no new grains are formed.

In the hypabyssal rocks undulatory extinction in quartz is common. Undulatory extinction is the result of spread dislocations (Fig. 6-4). As recovery tries to return the crystal to its undeformed state, deformation bands and ultimately subgrains form. All of these stages can be observed in the hypabyssal rocks. A certain amount of the polygonal subgrains are thus formed by recovery.

Figure 6-4: Recovery: A bended quartz grain, in which dislocations give rise to undulatory extinction, is recovered to a quartz grain with deformation bands and ultimately subgrain boundaries are formed. After Passchier and Trouw, 2005.

Contrary to recovery, where no new grains are formed, recrystallization reduces the amount of dislocations by the creation and/or movement of grain boundaries. Recrystallization produces strain- free volumes by forming aggregates of new grains. During the process no new minerals are formed. In minerals with relatively uniform three-dimensional lattice structures (quartz, feldspar and calcite), recrystallization will lead to polygonal grains.

Recrystallization can be subdivided in three processes: bulging (BLG), subgrain rotation (SR) and grain boundary migration (GBM). These processes can be distinguished with increasing temperature and decreasing strain rate. As they occur during deformation they are described as dynamic recrystallization.

Figure 6-5: Three main types of dynamic recrystallization, with increasing temperature bulging, subgrain rotation and: grain boundary migration. From Passchier and Trouw, 2005.

Bulging The first type of recrystallization, bulging, occurs at low temperatures. Therefore it is often also described as low-temperature grain boundary migration. Under these conditions, the mobility of the grain boundaries is restricted. If two neighbouring grains contain a different concentration of dislocations, the grain boundary can start to bulge into the grain with the highest density. By doing so new and small crystals are formed. Because of this process, remains of old grains are often surrounded by a rim of recrystallised grains, resulting in core-and-mantle structures.

Subgrain rotation During the process of recovery, dislocations are concentrated in subgrain boundaries. As dislocations are added to the subgrain boundaries, their complexity and misorientation increases. Adding dislocations to the subgrain boundaries causes the angle between the crystal lattice on both sides of the subgrain boundary to increase until the subgrain can no longer be classified as a part of the same grain. As new grains are developed by the misorientation of subrains the process is known as subgrain rotation-recrystallisation. Contrary to bulging, this process occurs at higher temperatures, and generally forms elongated new grains.

Grain boundary migration At high temperatures grain boundary mobility is higher and allows the grain boundaries to sweep through the entire crystal. By doing so the dislocations are removed and this process is called grain boundary migration recrystallization. As a result of this process grains become variable in size and grain boundaries become irregular in shape. At very high temperature, grains have highly loboid or amoeboid boundaries. The resulting grains are almost strain free and do not display undulatory extinction and subgrains.

6.3.4. Grain boundary area reduction and static deformation As the process of bulging, subgrain rotation and grain boundary migration occur during deformation, they are grouped under the term dynamic recrystallization. After deformation slows down or stops, grain boundary migration may continue by grain boundary area reduction.

After all, grain boundaries can also be considered planar defects, and therefore they also contribute to the internal free energy of the rock. A decrease in the total surface area of grain boundaries would thus result in a reduction of the free energy. Therefore straight grain boundaries and large grains are favoured. Polycrystalline material will try to achieve large polygonal grains with straight boundaries. This process of grain growth and straightening of the grain boundaries (Fig. 6-6) is called grain boundary area reduction (GBAR).

If after deformation the polycrystalline material has not reached a state of minimum internal free energy, this process may continue and is then known as static recrystallization. This process requires much water to be present along grain boundaries or for temperatures to remain high after deformation stopped.

Figure 6-6: Illustration of grain boundary area reduction. Irregular grains, formed during deformation and dynamic recrystallization, undergo grain growth and straightening of the grain boundaries. From Passchier and Trouw, 2005.

6.3.5. Core-and-mantle texture vs. porphyroclast systems In deformed rocks, large single crystals surrounded by a more fine-grained groundmass, usually of polymineralic composition, are best described as porphyroclasts. As a result of recrystallization, these porphyroclasts often have attached polycrystalline rims that differ in structure and composition from the matrix. This texture used to be called a mortar texture but is nowadays described as a core-and-mantle texture. The texture is believed to be formed by dynamic recrystallization. Although this term is adopted in Passchier and Trouw (2005), they use different terminology when dealing with mylonites. Then the texture is described as porphyroclast systems and they make the following subdivision. If the rim has the same composition as the porphyroclast, the rim is described as a mantle and the overall structure is called a mantled porphyroclast. If the surrounding rim has a different composition than the porphyroclasts, then the, often tapering, domains around the porphyroclast are known as strain shadows. The total structure is then called a porphyroclasts with strain shadows. According to Passchier and Trouw (2005) strain shadows are often composed of carbonate, quartz, mica and opaque minerals, which applies to our rocks. They state that these minerals are often not formed by reaction with the porphyroclasts, but by precipitation from solution.

6.4. DISCUSSION 6.4.1. Hypabyssal rocks Brittle and ductile deformation in the hypabyssal rocks is evidenced e.g. by, respectively, fractured porphyroclast and kinked crystals. Recovery and recrystallization is evidenced by bulged edges, subgrains and sometimes irregular edges. The dominance of polygonal crystals suggests that deformation was accompanied or followed by high persistent temperatures.

6.4.2. Metaquartzites Within the metaquartzites we do not observe large porphyroclasts, and therefore it is hard to say whether these rocks were affected by brittle deformation. As kinked biotite crystals occur, we are convinced that ductile deformation occurred. All aspects of bulging, subgrain rotation and grain boundary migration can be observed within these rocks, evidencing recovery and recrystallization. Similar to the hypabyssal rocks there are very abundant polygonal crystals present. Therefore we suggest that the elevated temperature event was also recorded in these rocks.

6.4.3. Rocks with a mafic magmatic protolith. Mafic rocks have different textures compared to the felsic hypabyssal rocks, and only limited attention was given to these rocks in our study. Due to the orientation of acicular, prismatic and platy minerals we see aspects of foliation. This foliation reveals that ductile deformation occurred.

In some places we find aggregates of polygonal quartz. Most of them display very straight edges again pointing to late elevated temperatures. This is supported by the fact that we sometimes find randomly oriented micas. For example, RG 89.593, is a rock with alternating textures. These alternating textures were probably caused by segregation as deformation increased. Remarkable in this thin section, is the random orientation of micas and the polygonal quartz grains. Both aspects support late elevated temperatures.

6.5. DISCUSSION REGARDING PREVIOUS OBSERVATIONS As some previous geologists have studied the regional metamorphism of the rocks of the (broader) Matadi region, we can compare our own results with their observations.

Tack (1975b) focussed on the amygdaloidal metabasalts of the Gangila Formation. He concluded that the mineral assemblage of these rocks, comprising tremolite-actinolite, epidote, chlorite, leucoxene, saussuritized plagioclase, biotite, quartz and calcite, indicates the influence of regional greenschist facies metamorphism, more specifically at the limit of the chlorite- to biotite subfacies. This was also observed in the felsic Mayumbian rocks (Tack, 1979), were the greenschist subfacies varies between the chlorite and/or biotite subfacies to almandine subfacies. Also the Kimezian basement (Delhal and Ledent, 1976) and the Mpozo syenite (Delhal and Ledent, 1978) have suffered greenschist facies retrograde metamorphism.

Franssen and André (1988) focussed on the regional metamorphism from Boma to Matadi (Fig. 6-7), including the transition from amphibolite (Boma) to greenschist facies (Matadi) conditions. In Figure 6-7 we notice that in the Matadi region, greenschist facies conditions vary from one subfacies to another and are in agreement with our own observed mineral assemblage. However, in this sketchy section steep dips of the various lithostratigraphic units are (strongly) exaggerated as well as the prominent diapiric character attributed to the Noqui body (see our own results).

Figure 6-7: East-west section of metamorphic facies. From Franssen and André, 1988.

Moreover, Franssen and André (1988) used the geochemical signature of the observed amphiboles along their Boma-Matadi section to define the regional p-t metamorphic conditions. They conclude that the composition of the amphiboles in th Matadi region indicates low pressure and high temperature regimes during regional metamorphism. They also mention features indicating strain- induced recrystallization (cfr. static recrystallization).

Therefore, we suggest that the low pressure regional metamorphism was accompanied by a persistent high thermal regime with elevated temperatures still proceeding during deformation.

6.6. CONTACT METAMORPHISM Contact metamorphism, sometimes also called thermal metamorphism, occurs when an intrusive magma heats the surrounding host rocks and causes changes in its mineralogy and texture. The zone in which this contact metamorphism is expressed is called the contact aureole. The intrusion of the

87 hypabyssal rocks and – in particular – the Noqui granite body (Mortelmans, 1948; Behiels, 2013), have caused contact metamorphism. In our study region, the contact metamorphism related to peralkaline granitic magmatism (MGE 3) obviously precedes regional metamorphism and deformation (MGE 2). An in-depth study of the contact metamorphic processes falls out of the scope of this thesis.

6.6.1. Quartzite assimilation One of the hypabyssal rocks, RG 89.863, displays a very high amount of quartz. As this rock is intrusive in the metaquartzites of the Matadi Formation, we suggest that this magmatic rock has assimilated metaquartzites. Similarly, Behiels (2013) describes assimilated metaquartzites in relation with the Noqui granite.

Bulk assimilation of crustal fragments with variable size (millimeters to 1 km) has been proposed as an efficient mechanism to cause large chemical and mineralogical modifications in granitoids (Beard et al., 2005). This was evidenced by Erdmann et al. (2007) who carried out melting experiments involving metasedimentary and granitic rocks to study reactions and products of granite contamination by assimilation of metasedimentary material. The basic principle is that the fragments of the surrounding country rocks, in our case the metaquartzites, become xenoliths as they are trapped by the intruding magma. This might have occurred in our situation, causing aggregates of quartz crystals to appear in the hypabyssal rocks. The volume of assimilated materials mainly depends on the proximity to the contacts with the surrounding materials. Therefore a more precise localization of the sampling points is demanded. Samples taken at the core of the intrusion will probably not display aspects of quartzite assimilation, while samples from the edges, close to the quartzite contact, could easily have been contaminated.

6.6.2. Garnet blastesis One thin section of a relatively undeformed rock was examined during this study. Due to time constraints, no description of this thin is given in our study. However, it shows an excellent example of idioblastic garnet porphyroblasts. As this rock is only slightly affected by deformation, and because of the abundant idioblastic garnet porphyroblasts, we conclude that this rock has experienced contact metamorphism. Similar observations are discussed by Mortelmans (1948) and Behiels (2013).

However, we maintain that the described syn- to post-tectonic porphyroblast of garnet in RG 19.699 (see section 5.2.3) is attributed to tectono-metamorphic processes (see sections 6.4 and 6.5).

6.7. METASOMATISM According to the IUGS classification (Le Maitre, 2002), metasomatism is : “a metamorphic process by which the chemical composition of a rock or rock portion is altered in a pervasive manner and which involves the introduction and/or removal of chemical components as a result of the interaction of the rock with aqueous fluids (solutions)”. Metasomatic processes, often related to contact metamorphism, particularly in the case of peralkaline magmatism, are discussed from the geochemical point of view in chapter seven.

We suggest that the presence of calcite in some of the hypabyssal rocks (e.g. RG 89.876) is linked to metasomatism and formed due to infiltrating fluids rich in CO2. However, as these calcite crystals are also present in the recrystallised rims and fractures of porphyroclasts, they also must have

88 experienced deformation and/or recrystallization. Thus carefulness in their interpretation is necessary, the more that calcite can indeed be induced by greenschist facies metamorphism.

6.8. AEGIRINE AND RIEBECKITE Within the Noqui granite aegirine and riebeckite are characteristic minerals. Nevertheless, these minerals were not observed within the hypabyssal rocks.

Behiels (2013) described the agpaitic texture of the Noqui granite with formation of the dark minerals of the granite only in a late stage of crystallization. Thus, as our hypabyssal rocks endured a phase of rapid cooling, resulting in the fine-grained groundmass, minerals such as aegirine and riebeckite are not expected to have formed and have indeed not been observed.

6.9. LITHOLOGICAL MAPS As all of the rocks are to some extent deformed, a macro- and microscopic study was necessary to correctly identify the protolith of some rocks. Based on the observations in chapter four and five and the discussion here above, we can classify the rocks into three groups of rocks: 1) rocks with a felsic magmatic protolith, better described as hypabyssal rocks; 2) rocks with a sedimentary protolith, better described as metaquartzites; 3) rocks with a mafic magmatic protolith (at least in some cases also with a hypabyssal setting). Based on this classification all of the samples were colour coded with respectively a red, a yellow and a green colour.

All of the samples were plotted on a map, displaying the lithological information. Figures 6-8A to 6- 8C represent the lithological information obtained from our observations of the samples of respectively Hugé, Massar and Steenstra. A composite map, containing all information of these three field geologists is given in Figure 6-8D. For clarity reasons the metaquartzites are not plotted on these maps, but they are plotted on the schematic maps in Annex 4.

The map in Figure 6-8D gives a good idea of the abundance of felsic and mafic intrusions (both sills and dykes) within the metaquartzites of the Matadi Formation. Several times we observe felsic and mafic intrusions occurring together or in each other’s vicinity. Furthermore, the mafic sills and dykes often show intrusive features within the felsic intrusions (section 4.1), indicating that they are younger. Based on these observations, we conclude – in agreement with Massar (1965) – that the emplacement of the mafic magmatic rocks was often controlled by reactivation of earlier weakness zones, where the felsic hypabyssal rocks had already been emplaced.

The field geologists have also regularly noted the strikes and dips of the sedimentary layers and hypabyssal intrusions. These structural data are plotted on maps given in Annex 5. They show persistent nearly similar values of ca. 30° dip to the west, which are well-illustrated by the dip slope morphology along the panoramic section given in section 4.1. For detailed information we refer to Annex 2.

Figure 6-8: Lihthological maps based on the observations of: A) Hugé; B) Massar. Rocks with a felsic magmatic protolith = hypabyssal rocks.

Figure 6-8 continued. C) Steenstra; D) Combined observations of Hugé, Massar and Steenstra. Rocks with a felsic magmatic protolith = hypabyssal rocks.

7. GEOCHEMISTRY

7.1. PREVIOUS RESEARCH Earlier limited geochemical data on the various magmatic rocks of the Matadi region are given in Behiels (2013; Annex 6). They include data of Polinard (1934), Mortelmans (1948), Korpershoek (1964), Delhal and Ledent (1978), Baert (1995) and Makutu et al. (2004). Although more analytical data of major and trace elements of some 20 rocks (Franssen; Archives G400, RMCA Tervuren) are available, their use for interpretation is hampered by the lack of their precise location, unlike reported in Franssen and André (1988, p. 221).

In order to allow an updated geochemical interpretation of the various magmatic rocks of the Matadi region, twenty relevant samples have been subjected to a major and trace element analysis. They include six samples of the Noqui granite (NE part of the massif), four samples of the Mpozo syenite and ten felsic hypabyssal rocks. The exact location of the sample points is given in Figure 7-1.

7.2. MAJOR ELEMENTS The twenty rock samples were analyzed with ICP-AES, as discussed in section 3.3, to retrieve data on their major element concentrations. The results of this analysis are listed in Table 7-1.

Table 7-1: Major element content (in wt%) of the Noqui granite, Mpozo syenite and hypabyssal rocks (HR). * Total Fe as Fe2O3.

Noqui Noqui Noqui Noqui Noqui Noqui Mpozo Mpozo Mpozo Mpozo Sample 13122 71299 89974 89978 89991 89992 19504 19611 89453 90067

SiO2 74.15 72.62 74.39 74.00 72.93 73.14 65.00 63.72 64.37 67.41 Al2O3 10.87 11.06 8.99 6.44 9.54 9.89 17.04 16.90 17.68 15.75 Fe2O3* 5.75 5.76 6.55 8.92 5.84 5.60 3.88 4.43 4.52 3.03 MnO 0.10 0.10 0.07 0.05 0.17 0.13 0.05 0.06 0.06 0.04 MgO 0.05 0.05 0.05 0.06 0.31 0.07 0.54 0.48 0.63 0.54 CaO 0.31 0.65 0.08 0.24 1.77 1.35 0.86 1.37 1.10 1.32

Na2O 4.48 3.92 3.92 4.50 3.04 3.44 5.06 4.63 5.16 4.36 K2O 4.10 4.59 4.19 3.60 4.50 4.14 6.01 6.43 6.94 5.35

TiO2 0.32 0.27 0.29 0.21 0.25 0.29 0.55 0.56 0.57 0.46 P2O5 0.01 <0.01 <0.01 0.02 <0.01 0.02 0.09 0.09 0.10 <0.01 LOI 0.15 0.04 0.24 0.34 0.23 0.33 0.18 0.27 0.10 0.32 TOTAL 100.29 99.06 98.76 98.40 98.57 98.40 99.26 98.96 101.24 98.58

HR HR HR HR HR HR HR HR HR HR S ample 89532 89534 89540 89544 89546 89554 89863 89876 89979 90142

SiO2 73.26 74.96 71.86 76.31 73.30 78.56 80.00 77.45 78.21 77.17 Al2O3 12.86 12.51 12.98 11.93 12.65 9.74 8.98 11.20 8.45 11.17 Fe2O3* 3.13 2.53 3.52 1.56 5.17 4.16 4.43 1.66 5.74 2.36 MnO 0.06 0.06 0.08 0.01 0.10 0.05 0.02 0.02 0.03 0.03 MgO 0.40 0.11 0.50 0.04 0.06 0.14 0.07 0.11 0.06 0.09 CaO 0.84 0.13 0.84 0.27 0.10 0.86 0.18 0.62 0.27 0.08

Na2O 3.44 3.16 3.60 3.71 3.09 4.51 4.58 2.36 1.36 2.69 K2O 4.30 5.15 4.44 4.27 5.06 0.38 0.39 4.82 5.10 4.63 TiO2 0.59 0.26 0.68 0.32 0.47 0.24 0.22 0.14 0.22 0.15 P2O5 0.11 0.02 0.15 0.01 0.02 <0.01 <0.01 <0.01 0.01 0.00 LOI 0.43 0.34 0.36 0.09 0.44 0.26 0.04 0.85 0.03 0.24 TOTAL 99.41 99.23 99.00 98.52 100.46 98.90 98.91 99.24 99.47 98.62

Figure 7-1: Localization of the twenty samples selected for geochemical analysis. Noqui granite: 1) RG 13.122; 2) RG 71.299; 3) RG 89.974; 4) RG 89.978; 5) RG 89.991; 6) RG 89.992; Mpozo syenite: 7) RG 19.504; 8) RG 19.611; 9) RG 89.453; 10) RG 90.067; Hypabyssal rocks: 11) RG 89.532; 12) RG 89.534; 13) RG 89.540; 14) RG 89.544; 15) RG 89.546; 16) RG 89.554; 17) RG 89.863; 18) RG 89.876; 19) RG 89.979; 20) RG 90.142.

7.2.1. Classification

Data from Table 7-1 are plotted in a TAS diagram (Fig. 7-2), in which the total alkalis (Na2O + K2O) are plotted against SiO2. Originally the TAS classification diagram was developed for volcanic rocks. Dealing with plutonic and/or hypabyssal rocks we cannot apply this classification, albeit that the diagram is often adapted for intrusive rocks (Rollinson, 1993).

Within the diagram, the red triangles represent the Mpozo massif. The blue squares are indicative for the Noqui granite and the green triangles represent the hypabyssal rocks. From Figure 7-2 to Figure 7-15, the same symbols will be used unless mentioned otherwise. Data points of the Mpozo massif plot within the trachyte and trachydacite field. Adapted for plutonic rocks, this results respectively in syenite and quartz monzonite. Samples of the Noqui granite and the hypabyssal rocks both plot within the rhyolite field which corresponds to granites in plutonic rocks.

Furthermore the diagram illustrates that all samples can be considered acid as they contain more than 63 wt% SiO2. Some of the hypabyssal samples contain remarkably high SiO2 contents reaching up to 80 wt%. Samples with values higher than 78 % are indicated in bold in Table 7-1. These high values usually do not appear in magmatic rocks. Therefore carefulness in the interpretation of these data points is required.

Figure 7-2: TAS classification diagram.

To classify plutonic rocks the QAPF diagram, also called Streckeisen diagram, is recommended. This classification makes use of the mineralogical composition of the light-coloured minerals including quartz, alkali feldspar, plagioclase and feldspathoids.

In section 5.4 the modal composition of some rocks was determined based on a point counting analysis. Based on the chemical composition it is also possible to calculate the normative composition. At the beginning of the twentieth century W. Cross, J.P. Iddings, L.V. Pirsson and H.S. Washington proposed the CIPW norm calculation scheme, which allows to estimate the standard mineral assemblage of an igneous rock based on its geochemistry. To calculate the normative composition of the rocks we applied the rules proposed by Kelsey (1965) of which the results are presented in Table 7-2.

Table 7-2: Normative mineralogy in %. Q = quartz; Or = orthoclase; Ab = albite; An = anorthite; Hy = hypersthene; Di = diopside; Ap = apatite; Il = ilmenite; Ac = acmite; C = corundum; Ks = potassium metasilicate; NaS = sodium metasilicate; Mt = magnetite.

Noqui Noqui Noqui Noqui Noqui Noqui Mpozo Mpozo Mpozo Mpozo Sample 13122 71299 89974 89978 89991 89992 19504 19611 89453 90067 Q 30.09 29.00 35.95 40.68 33.05 33.26 8.63 6.97 2.98 16.79 Or 24.23 27.12 24.76 21.27 26.59 24.47 35.52 38.00 41.01 31.62 Ab 33.10 31.33 22.91 13.08 24.02 27.82 42.82 39.18 43.66 36.89 An 0.00 0.00 0.00 0.00 0.00 0.00 3.68 6.21 4.58 6.48 Hy 6.40 5.05 7.94 10.73 3.41 3.23 4.73 5.18 5.53 3.94 Di 1.31 2.81 0.29 0.94 7.64 5.83 0.00 0.00 0.19 0.00 Ap 0.02 0.02 0.02 0.05 0.02 0.05 0.21 0.21 0.23 0.02 Il 0.61 0.51 0.55 0.40 0.47 0.55 1.04 1.06 1.08 0.87 Ac 3.85 1.62 4.37 5.96 1.50 1.13 0.00 0.00 0.00 0.00 C 0.00 0.00 0.00 0.00 0.00 0.00 0.86 0.05 0.00 0.41 Ks 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 NaS 0.10 0.00 1.23 4.24 0.00 0.00 0.00 0.00 0.00 0.00 Mt 0.00 1.12 0.00 0.00 1.20 1.30 1.29 1.48 1.51 1.01 total 99.71 98.58 98.02 97.35 97.90 97.64 98.78 98.34 100.77 98.03

HR HR HR HR HR HR HR HR HR HR Sample 89532 89534 89540 89544 89546 89554 89863 89876 89979 90142 Q 33.55 35.38 30.52 37.24 35.71 46.94 49.37 43.08 47.46 42.43 Or 25.41 30.43 26.24 25.23 29.9 2.25 2.3 28.48 30.14 27.36 Ab 29.11 26.74 30.46 31.39 26.15 38.16 38.75 19.97 11.51 22.76 An 3.45 0.51 3.19 1.27 0.37 4.20 0.83 3.01 1.27 0.40 Hy 3.52 2.69 4.06 1.28 0.15 4.54 4.64 1.88 6.04 2.58 Di 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Ap 0.25 0.05 0.35 0.02 0.05 0.02 0.02 0.02 0.02 0.00 Il 1.12 0.49 1.29 0.61 0.89 0.46 0.42 0.27 0.42 0.28 Ac 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 C 1.28 1.55 1.08 0.74 1.96 0.37 0.72 1.00 0.23 1.59 Ks 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 NaS 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Mt 1.04 0.84 1.17 0.52 2.80 1.39 1.48 0.55 1.91 0.78 total 98.73 98.68 98.36 98.3 97.98 98.33 98.53 98.26 99.00 98.18

The CIPW norm values of quartz, alkali feldspar and plagioclase are normalized and plotted in the QAPF diagram (Fig. 7-3). Rocks from the Mpozo body plot within fields 8 and 8*. Based on this classification the rocks can be described as monzonites or quartz monzonites. The blue squares, which represent the Noqui body, are all located within field 3b, and can thus be classified as monzogranites. Most of the hypabyssal rocks also plot within this field but three samples deviate. One sample, RG 89.979, contains a lower plagioclase content and is therefore plotted within the syenogranite field. The two other samples, RG 89.554 and RG 89.863, contain almost no alkali feldspar according to the CIPW norm and are therefore plotted in the tonalite field.

2 Alkali feldspar granite 3a Syenogranite 3b Monzogranite 4 Granodiorite 5 Tonalite 6 Alkali feldspar syenite 6* Alkali feldspar quartz syenite 7 Syenite 7* Quartz syenite 8 Monzonite 8* Quartz monzonite

Figure 7-3: Modified QAPF diagram (after Streckeisen, 1974) based on normative minerals.

Furthermore, the normative mineralogy reveals that samples of the Noqui granite and the hypabyssal rocks display similar amounts of quartz, ranging approximately between 30 and 40%. Some of the hypabyssal samples display values above 40%. Three samples even show values higher than 45 % and are indicated in bold in Table 7.2. The amount of normative quartz within the Noqui granite and the hypabyssal rocks, is in contrast with the low values of the Mpozo syenite, where values do not exceed 17%.

The total amount of feldspar (Or+ Ab+ An) varies around 50% for the Noqui granite, 75% for the Mpozo syenite and 55% for the hypabyssal rocks. Within the hypabyssal rocks, two samples differentiate themselves from the other samples based on their orthoclase content. According to the norm, RG 89.554 and RG 89.863 contain no more than respectively 2,25 and 2,3% orthoclase.

Contrary to the rocks of the Mpozo syenite and the hypabyssal samples, rocks of the Noqui granite are characterized by the presence of acmite and sometimes also some sodium metasilicate. These minerals typically occur in peralkaline rocks. A normative mineral typifying peraluminous rocks is corundum, as it indicates an oversaturation in Al2O3. This normative mineral does not occur within the Noqui granite but is present in all hypabyssal rocks. Except for one sample, RG 89.453, the Mpozo syenite also contains normative corundum. This samples, contrary to all other Mpozo samples, has diopside in its norm. In the next section the terms peralkaline and peraluminous will be explained in more detail.

7.2.2. Discrimination diagrams Granites and granitoids can be classified based on their aluminum saturation index (ASI) (Shand,

1943), defined as the molecular ratio [Al2O3/(CaO+Na2O+K2O)]. Rocks with an excess of aluminum over alkali have an ASI > 1,0 and are called peraluminous. Rocks which are under-saturated in aluminum with respect to alkali have an ASI < 1,0 and are said to be metaluminous.

Besides peraluminous and metaluminous, rocks can also be peralkaline. This term is restricted to rocks in which the molecular amounts of Na2O plus K2O exceed Al2O3. To decide whether this applies, the peralkaline index (PI) can be used. This index is defined as the molecular ratio [(Na2O+K2O)/Al2O3] and peralkaline rocks thus have a PI > 1. The ASI and PI were calculated for all rocks and are given in Table 7-3.

Table 7-3: Calculated values of ASI and PI.

Noqui Noqui Noqui Noqui Noqui Noqui Mpozo Mpozo Mpozo Mpozo Sample 13122 71299 89974 89978 89991 89992 19504 19611 89453 90067 ASI 0.88 0.88 0.81 0.55 0.73 0.78 1.04 0.99 0.98 1.03 PI 1.09 1.03 1.22 1.75 1.03 1.03 0.87 0.86 0.91 0.82

HR HR HR HR HR HR HR HR HR HR Sample 89532 89534 89540 89544 89546 89554 89863 89876 89979 90142 ASI 1.09 1.13 1.06 1.06 1.18 1.04 1.08 1.09 1.02 1.17 PI 0.80 0.86 0.83 0.90 0.83 0.80 0.89 0.81 0.92 0.85

In Figure 7-4 the rocks are plotted in a diagram with ASI on the x-axis and the inverse of PI on the y- axis. In this diagram rocks of the Noqui massif plot within the peralkaline field. This is in agreement with the PI-values of Table 7-3, where all Noqui granites have a PI larger than 1,0 and can thus be called peralkaline. All other rocks have a PI smaller than 1,0. The hypabyssal rocks plot within the peraluminous field. Peraluminous rocks have an ASI > 1,0 which is valid for these rocks. Samples taken from the Mpozo syenite do not plot within one field. Two of its samples, RG 19.504 and RG 90.069, plot together with the hypabyssal rocks in the peraluminous field. The two other samples are metaluminous. In Table 7-3 it can be seen that the ASI values for the Mpozo samples are close to 1,0. Two samples display a value slightly smaller than 1,0 while the two other samples are slightly bigger.

Figure 7-4: Plot of the ASI and inverse PI, indicating metaluminous, peraluminous and peralkaline rocks.

7.2.3. Harker diagrams To give a clear representation of the major element geochemistry of the rocks, Harker diagrams are plotted (Fig. 7-5). These diagrams are bivariate diagrams in which the vertical axis represents weight percents of major element oxides. On the x-axis the SiO2 content is displayed. These Harker diagrams allow to observe variations and overall trends of the major element oxides. In the next section we will discuss the Harker diagrams of the different elemental oxides.

A first observation from these Harker diagrams is that there are two groups. Samples from the Noqui granite and hypabyssal rocks plot together as one group and differ from the rocks of the Mpozo syenite. These groups are separated from each other by a chemical gap in the SiO2 content, which can also be observed in the TAS diagram (Fig. 7-2). Rocks from the Mpozo body are characterized by a lower SiO2 content. Values range between 63,72% and 67,41%. Rocks of the Noqui granite and the hypabyssal rocks contain more SiO2 and values range between 71,86% and 80,00%.

In Figure 7-5A it can be observed that rocks of the Mpozo body have a higher Al2O3 content compared to the Noqui granite and the hypabyssal rocks. Al2O3 values of the Mpozo syenites vary between 15,75% and 17,68% while values of the other rocks do not exceed 12.98%. Looking at the green triangles, it can be stated that the hypabyssal rocks display an inverse trend. As the SiO2 content increases, the Al2O3 content decreases. This also applies to the blue squares which represent the Noqui granite.

The CaO content of the rocks is displayed in Figure 7-5B and is relatively low. Rocks of the Noqui granite and hypabyssal rocks have a CaO content generally lower than 1% but two samples, RG 89.991 and RG 89.992 deviate as they display somewhat higher CaO values (respectively 1,77% and 1,35%). Except for these two samples, rocks of the Mpozo body contain slightly more CaO compared to the other rocks.

The Na2O contents (Fig. 7-5C) of the Mpozo syenites are slightly higher than for the other rocks. A maximum value of 5,16% can be observed. Values for the hypabyssal rocks and for the Noqui granites are all lower than 5%. Values for the Noqui granite range between 3,04% and 4,50%. The range of the hypabyssal rocks is bigger and values vary between 1,36 and 4,58%. Within these hypabyssal rocks an inverse trend can be observed. As the rocks get more felsic, the Na2O contents decrease. Two samples, RG 89.554 and RG 89.863 deviate from this trend and also display somewhat higher Na2O values compared to the other hypabyssal rocks.

The K2O content, displayed in Figure 7-5D, shows some similarities with Na2O. The amount of K2O within the Mpozo syenites is larger than for the more felsic rocks and ranges between 5,35% and

6,94%. Rocks of the Noqui granite and the hypabyssal rocks contain slightly less K2O. Average values range between 4% and 5%. The same two samples which deviate for their Na2O content also aberrate for K2O. These anomalies are characterized by very low amounts of K2O, respectively 0,38% and 0,39%.

The total amount of iron within the rocks, given as Fe2O3, shows a very different pattern (Fig. 7-5E) compared to the other oxides. The hypabyssal rocks are characterized by values ranging between 1,56% and 5,74%. Rocks of the Mpozo syenite plot within a similar range, varying between 3,03% and

Figure 7-5: Harker diagrams.

4,52%. Higher Fe2O3* contents can be found within the Noqui granite. These rocks display values between 5,60% and 8,92%.

The MgO content of the rocks is displayed in Figure 7-5F. Taking into account the scale of the y-axis it is clear that the rocks are poor in MgO. All rocks contain less than 0,7% MgO. The highest values can be found within the Mpozo syenite. Here values range between 0,54% and 0,63%. For most of the more felsic rocks values are lower than 0,2%. One of the Noqui granite samples, RG 89.991, shows a slightly higher value of 0,31%. This is also the case for two hypabyssal rocks, RG 89.532 and RG 89.540, which display respectively a value of 0,40% and 0,50%.

Just as Mgo, MnO (Fig. 7-5G) is present at very low concentrations. Rocks of the Mpozo syenite and hypabyssal rocks all contain less than 0,1% MnO. Rocks of the Noqui granite contain slightly more MnO. Values for these rocks range between 0,05% and 0,17%.

Concentrations of TiO2 (Fig. 7-5H) are also very low. For the Mpozo syenite values range between 0,46% and 0,57%. Values for the Noqui granite are lower and vary between 0,21 and 0,32. The range of TiO2 content within the hypabyssal rocks is wider than for the other rocks. Here values range between 0,14% and 0,68.

For P2O5 values are often below the detection limit of 0,1 wt%, and therefore this element oxide is not plotted in a Harker diagram.

7.2.4. R1 – R2 multicationic diagram The use of wt% oxide data has been criticized by Chayes (1964) and Pearce (1969). The principal criticism is that wt% oxides do not faithfully represent the cation distribution in the sample. Therefore a different approach, using cationic values, can be helpful. This concept has been used by de la Roche et al. (1980) who proposed a classification scheme for volcanic and plutonic rocks. The proposed diagram uses the plotting parameters R1 and R2. R1, plotted on the x-axis is defined as [4Si - 11(Na+K) – 2(Fe+Ti)] and R2, plotted along the y-axis is defined as (Al+2Mg+6Ca).

The R1 – R2 multicationic diagram of de la Roche et al. (1980) was used by Bachelor and Bowden (1985) to define the tectonic setting of granitoids. The diagram also allows us to suggest processes such as fractional crystallization and mixing.

In Figure 7-6 it can be seen that all of the Mpozo syenite samples plot within field 4 (late-orogenic setting) but close to the limit of field 5 (anorogenic). Most of the Noqui granite samples plot within field 5 (anorogenic setting). Some of the hypabyssal samples also plot within this field, while others suggest a rather post-orogenic setting (field 7). Evidence for magma mixing is lacking.

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Figure 7-6: R1 – R2 multicationic diagram of Batchelor and Bowden (1985).

7.3. TRACE ELEMENTS The same samples which were analyzed for major elements were also subjected to a trace element analysis. This analysis was carried out with ICP-MS, which is explained in section 3.2. The results are given in Table 6-4 and are expressed in parts per million (ppm).

7.3.1. Variation diagrams Variation diagrams of trace elements are plotted in Figure 7-7. From Figure 7-7A one can deduce that only small concentrations of Ba are present within the Noqui granite. Larger amounts of this element can be found within the Mpozo syenite and within the hypabyssal rocks. These hypabyssal rocks display a negative correlation. Two of the hypabyssal samples, RG 89.540 and RG 89.532, plot at higher Ba values of approximately 1000 ppm.

Just as Ba, Strontium (Fig. 7-7B) is not very abundant in the Noqui granite. Concentrations are slightly bigger within the hypabyssal rocks, where values range between 29 and 114 ppm. Compared to these rocks, the Mpozo body is rich in strontium. Concentrations vary between 157,5 and 262 ppm.

Rubidium (Fig. 7-7C) displays strongly different patterns compared to Barium and Strontium. Noqui granite samples comprise large concentrations of Rb compared to the other rocks. Values within these samples range between 276 and 580 ppm. Abundances within the Mpozo syenite and the hypabyssal rocks are similar to each other and lower than within the Noqui granite. Two of the hypabyssal samples, RG 89.554 and RG 89.863, display very low values of respectively 11,6 and 24 ppm. These are the same two samples which also display anomalies for Na2O and K2O.

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Table 7-4: Trace elements in ppm.

Noqui Noqui Noqui Noqui Noqui Noqui Mpozo Mpozo Mpozo Mpozo

sample 13122 71299 89974 89978 89991 89992 19504 19611 89453 90067

Sc 9.9 10.5 9.9 25.18 11.03 12.19 13.5 12.9 10.84 7.6 V 4.2 4.7 6.2 7.99 4.56 5.39 9.6 8.0 7.14 12.2 Cr 70 60 280 395.11 73.97 252.73 71 54 54.85 45 Co 1.33 0.97 3.3 0.89 1.54 0.75 2.6 3.1 2.70 2.9 Ni 5.1 9.5 4.7 5.11 6.75 3.86 4.7 5.7 4.01 4.5 Cu 14.2 14.1 10.4 17.40 180.15 48.43 25.8 27.3 25.20 174 Zn 235 275 417 384.59 284.10 359.87 33 51 36.00 22 Ga 42 39 42 41.95 42.27 44.92 16.3 18.1 17.66 15.6 Ge 3.0 3.1 3.0 4.60 3.83 3.59 1.28 1.32 1.28 1.39 Rb 276 343 443 580.00 308.62 312.62 91 112 120.03 75 Sr 2.6 15.1 3.1 13.26 30.37 33.66 179 164 157.53 262 Y 206 316 216 685.57 409.73 702.03 11.5 10.5 8.27 13.4 Zr 2124 2213 2642 10660.69 2467.10 3681.66 971 216 225.78 372 Nb 170 196 267 505.10 289.13 364.39 13.2 7.2 6.66 14.6 Cs 1.54 0.69 0.67 0.35 0.45 0.49 0.90 1.15 1.15 0.50 Ba 48 39 36 55.20 50.31 45.11 379 827 837.14 734 La 63 173 96 968.08 220.13 395.08 23 23 21.65 29 Ce 120 387 196 1623.36 464.62 709.42 43 39 36.27 57 Pr 11.9 41 23 210.18 54.65 91.04 4.7 4.6 4.31 6.8 Nd 44 149 83 715.40 198.27 325.92 16.2 16.6 15.21 22 Eu 0.89 1.65 0.97 6.20 2.28 3.66 0.92 0.98 0.93 0.70 Sm 14.1 33 18.3 131.63 44.38 69.95 2.8 2.9 2.52 3.3 Gd 17.1 36 21 122.44 50.48 79.33 2.3 2.4 2.13 2.7 Dy 30 47 34 108.43 65.39 98.86 1.94 1.90 1.57 1.98 Ho 6.3 9.8 7.3 19.95 13.56 20.75 0.37 0.36 0.28 0.40 Er 22 32 25 58.68 44.16 67.06 1.23 1.09 0.82 1.34 Yb 25 31 27 52.11 42.14 60.70 1.36 0.99 0.84 1.69 Lu 3.6 4.43 4.0 7.46 5.89 8.27 0.23 0.14 0.13 0.26 Hf 52 57 46 256.81 67.65 95.53 16.0 4.3 4.48 8.5 Ta 9.1 12.0 13.5 33.27 16.70 20.80 0.80 0.36 0.33 1.02 W 2.1 2.3 17.2 23.12 3.11 14.79 1.77 1.64 1.79 1.48 Pb 25 27 39 78.77 31.95 37.67 14.4 21 16.58 24 Th 26 20 45 34.94 42.98 84.13 10.7 4.3 9.09 12.4 U 7.6 7.2 8.5 30.08 9.88 21.98 2.9 1.28 1.73 5.0

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Table 7-4 continued.

HR HR HR HR HR HR HR HR HR HR

sample 89532 89534 89540 89544 89546 89554 89863 89876 89979 90142 Sc 12.9 8.5 14.1 8.1 8.4 8.9 12.1 8.4 12.4 7.1 V 9.0 3.5 11.7 27.7 4.3 15.3 26.9 2.6 9.1 3.9 Cr 107 190 111 126 135 262 460 158 273 42 Co 2.6 2.8 3.2 2.2 1.68 5.3 2.5 2.9 1.03 0.97 Ni 522 972 567 631 690 6.4 6.0 800 3.6 4.3 Cu 19.3 9.0 7.4 10.5 12.5 44.8 11.0 16.7 6.4 9.2 Zn 61 54 78 31 113 203 164 24 79 58 Ga 16.9 26 19.6 17.4 21 31 29 20 28 21 Ge 1.21 1.43 1.47 1.07 1.42 1.33 1.87 0.93 1.36 0.97 Rb 150 256 177 87 118 11.6 24 145 245 197 Sr 63 31 93 43 114 60 36 46 29 30 Y 43 84 50 106 81 262 391 75 349 76 Zr 478 618 471 898 603 2626 3026 336 3345 361 Nb 29 63 30 74 68 168 314 59 285 70 Cs 1.01 0.80 1.26 0.28 0.29 0.33 0.25 0.24 0.57 0.46 Ba 987 258 1099 346 549 194 40 111 72 72 La 48 97 61 104 80 188 345 88 249 98 Ce 106 201 123 217 162 408 815 170 595 186 Pr 11.1 19.6 13.8 23 17.7 43 77 18.3 58 21 Nd 38 66 49 80 63 152 269 60 205 68 Eu 1.67 0.77 2.14 0.43 0.71 1.35 3.24 0.28 2.29 0.30 Sm 7.7 12.5 9.2 15.6 12.8 32 55 11.1 41 12.8 Gd 7.4 13.1 8.9 15.1 12.8 33 56 10.6 41 12.2 Dy 7.4 13.7 8.7 16.9 13.8 41 70 12.2 55 13.1 Ho 1.43 2.7 1.67 3.4 2.7 8.1 14.1 2.5 11.5 2.5 Er 4.5 8.5 5.2 11.3 8.6 26 46 7.8 38 8.0 Yb 4.6 8.4 4.8 11.3 8.9 26 44 8.0 41 8.2 Lu 0.66 1.21 0.70 1.63 1.33 3.6 6.2 1.15 5.9 1.17 Hf 11.9 15.3 11.7 21 15.0 46 77 10.5 84 12.2 Ta 1.97 3.8 1.91 5.1 4.0 10.3 18.7 3.9 19.6 4.6 W 1.74 1.93 1.52 1.53 1.79 14.5 28 2.4 22 2.5 Pb 19.0 11.8 12.2 23 29 27 52 11.2 33 22 Th 24 32 24 38 26 55 64 40 68 43 U 4.8 6.9 5.2 7.4 3.2 10.4 15.3 9.0 14.8 5.1

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Figure 7-7: Trace element variation diagrams.

104

Figure 7-7 continued.

In Figure 7-7D Vanadium is displayed. For all rocks values are approximately similar and range between 2,6 and 15,3. Two samples of the hypabyssal rocks (RG 89.544 and RG 89.863) display slightly higher values of respectively 27,7 and 26,9.

Tantalum and Niobium both belong to the group of high field strength elements (HFSEs) and display similar characteristics. In Figures 7-7E and 7-7F it can be observed that the patterns for these two elements are analogous and that there are very low abundances of these elements within the Mpozo syenite. Values for Nb vary between 6,66 and 14,6 ppm while Ta concentrations are even lower and range between 0,33 and 1,02 ppm. The Noqui granite samples contain higher abundances of Ta and Nb with maximum values of 33,27 and 505,10 ppm for respectively Ta and Nb. For the hypabyssal rocks we observe a trend in which Ta and Nb increase as the rocks become more acidic. For seven out of ten hypabyssal rocks the values of Ta range between 1,91 an 5,0 ppm and for Nb they vary between 29 and 74 ppm. In Figures 7-7E and 7-7F we observe that three of the hypabyssal rocks separate themselves from the other hypabyssal rocks. These three samples comprise RG 89.979, RG

89.554 and RG 89.863 which are characterized by very high SiO2 contents. They display high values for Ta and Nb that are similar to the values of the Noqui granite.

Zirconium and Hafnium (Figs. 7-7G and 7-7H) exhibit similar chemical properties and thus present similar patterns. Compared to the Noqui granites, samples of the Mpozo syenite contain rather low abundances of Zr and Hf with maximum values of respectively 971 and 16 ppm. Most of the hypabyssal rocks also contain low abundances of Zr and Hf but three samples display much higher

105 values. These three samples are the same ones who display an anomaly for Ta and Nb. For most of the hypabyssal rocks values of Zr and Hf do not exceed, respectively, 898 and 21 ppm. The anomalous samples display values higher than 2626 ppm for Zr and higher than 46 ppm for Hf. These values lie in a similar range as the one observed for the Noqui granite. In these samples values of Zr range between 2124 and 10660.69 ppm and for Hf they vary between 46 and 256.81 ppm.

In Figures 7-7I and 7-7J it becomes clear that Yttrium and Holmium display similar patterns which also have some affinities with the Zr and Hf patterns. Values of Y and Ho remain low in the Mpozo syenite samples and do not exceed, respectively, 13,4 and 0,40 ppm. These low values are in contrast with the high values of the Noqui granite which displays minimum values of 206 ppm for Y and 6,3 ppm for Ho. For the hypabyssal rocks we observe again three anomalous samples. Most of the hypabyssal rocks contain relatively low abundances of Y and Ho with values not exceeding, respectively, 106 and 3,4 ppm. The three anomalous samples display higher values which fall in the range of the Noqui granite samples.

Gallium is plotted in Figure 7-7K. Values for the Mpozo syenite range between 15,6 and 18,1 ppm. As the amount of SiO2 increases, Ga decreases for these samples. For the hypabyssal samples, the abundances of Ga increase as the rocks become more acidic. Values range between 16,9 and 31 ppm and are thus slightly higher than the values of the Mpozo syenite. The highest concentrations can be observed within the Noqui granite. Values within these samples range between 39 and 45 ppm.

Germanium, plotted in Figure 7-7L, displays low and similar values for the Mpozo syenite samples and the hypabyssal rocks. These concentrations range between 1,07 and 1,87 ppm. Although values remain low, the Noqui granite is slightly enriched in Ge with variations between 3,0 and 4,6 ppm.

7.3.2. Discrimination diagrams 7.3.2.1. Tectonic setting In 1984 Pearce et al. came up with trace element discrimination diagrams to interpret the tectonic setting of granitic rocks. In his classification he defined “granites” as plutonic rocks that contain more than 5% of modal quartz. Based on the discrimination diagrams, granitic rocks can be subdivided into four main settings: ocean ridge granites (ORG), volcanic arc granites (VAG), within plate granites (WPG) and collision granites (COLG).

These discrimination diagrams are based on the following trace elements: Rb, Nb, Ta, Y and/or Yb. Rubidium belongs to the group of the large ion lithophile elements (LILEs). It is characterized by a large ionic radius and low charge. These properties cause Rb to be very soluble in aqueous fluids. Niobium and tantalum belong to the group of high field strength elements (HFSEs). These elements are characterized by very high charges and tend to be incompatible. Contrary to the LILEs they are insoluble in aqueous solutions. Yttrium and Ytterbium belong to the rare earth elements (REEs) which are, just as the HFSEs, insoluble in aqueous solutions.

Discrimination diagrams are given in Figure 7-8. The first two plots display Rb on the y-axis. Y + Nb and Yb + Ta plot respectively in Figures 7-8A and 7-8B on the x-axis. Nb and Ta belong to the same group of elements and thus display similar characteristics. This is also the case for Y and Yb. Because of their similar characteristics, Figure 7-8A and 7-8B display approximately the same patterns. This also applies for Figure 7-8C and 7-8D.

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As Rb belongs to the LILEs it is very soluble in aqueous fluids. This means that Rb is easily removed due to weathering of the rock. To avoid this effect it is better to display Nb and Ta on the x-axis, which are alteration-independent (Pearce et al., 1984).

Based on the classification of Pearce et al. (1984) samples of the Mpozo syenite plot as VAG. This is in contrast with samples of the Noqui granite and the hypabyssal rocks. These rocks both plot within the same field, being the WPG field. In Figures 7-8A and 7-8B two hypabyssal samples, RG 89.863 and RG 89.554, deviate and plot as ocean ridge granites. In Figures 7-8C and 7-8D these two samples do not deviate so vigorously. The deviation is rather small and cause the two samples to plot within the overlap zone between ocean ridge granites from anomalous ridge segments and within plate granites from attenuated continental lithosphere, which is marked by the dashed line.

Figure 7-8: Discrimination diagrams for WPG, VAG, syn-COLG and ORG of Pearce et al. (1984). A) Rb – (Y + Nb); B) Rb – (Yb + Ta); C) Nb – Y; D) Ta – Yb. The dashed line indicates an overlap zone between ocean ridge granites from anomalous ridge segments and within plate granites from attenuated continental lithosphere.

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7.3.2.2. Alphabetical classification In 1974, Chappell and White recognized two distinct types of granitoids which they described as I- and S-type granitoids. According to their classification the I-type is metaluminous to weakly peraluminous, is rather sodic and has a broad range of SiO2 content. These types of rocks are considered to have an igneous or meta-igneous source. This is in contrast with the S-type granitoids, which are strongly peraluminous, relatively potassic and which are characterized by higher silica contents. These types of rocks supposedly form by melting of metasedimentary rocks. Half a decade later Loiselle and Wones added another type of rock to the “alphabet soup”, being the A-type granitoids. The letter “A” denotes anorogenic, anhydrous and alkaline. These relatively potassic rocks typically have high FeO / (Feo + MgO) ratios. They are also characterized by high Zr, Nb, Ga, Y and Ce contents. The ratio of Gallium over Aluminum is typically also very high. The amount of CaO and Sr within these rocks is assumed to be low (Whalen et al., 1987). Based on the fact that these granitoids were almost never deformed it was thought that they had intruded after the deformation events of an orogeny or without a link with an orogenic setting and they were therefore called “anorogenic”. Today there still is considerable dispute on the definition, origin and evolution of these types of granitoids (Frost et al., 2001).

Based on the characteristics of A-type granites, which strongly differ from I- and S-type granites, Whalen et al. (1987) figured that it was possible to make good discrimination diagrams. They stated that the best diagrams are Ga/Al ratios on the x-axis, plotted against Y, Ce, Nb, or Zr on the y-axis. These diagrams are believed to be relatively insensitive to moderate degrees of alteration. They also remarked that highly fractionated, felsic I- and S-type granitoids possibly contain Ga/Al ratios and some major and trace element abundances, which cause overlap with the concentrations in typical A-type granitoids.

Figure 7-9: Data plotted on a Zr vs. 104Ga/Al diagram of Whalen et al. (1987).

108

Figure 7-9 represents one of the discrimination diagrams proposed by Whalen et al. (1987). Zirconium is plotted on the y-axis, while the ratio of 10.000 * Gallium over Aluminum is plotted on the x-axis. Samples taken from the Noqui granite and the hypabyssal rocks clearly plot within the A- type field. For the samples of the Mpozo syenite there is some ambiguity. Two samples, RG 19.504 and RG 90.067, also plot within the A-type granite field while the other two samples plot within the I&S-types field, although close to the A-type field.

According to Eby (1992) the A-type granitoids, based on certain trace element distributions, can be divided into two subgroups. These subgroups comprise granitoids which plot in the A1 or in the A2 group. The first group of granitoids were interpreted as differentiates of basalt magma derived from an ocean island basalt (OIB) like source. The A2 granitoids are thought to be derived from subcontinental lithosphere or lower crust. Eby (1992) also suggests that the A1-types are associated with anorogenic settings, while A2- types are often emplaced in post-collisional, post-orogenic settings.

The A1 and A2 discrimination diagrams should only be used for granitoids that plot both in the within plate granite field of Pearce et al. (1984) and the A-type granite field of the Ga/Al plots of Whalen et al. (1987). As the rocks of the Mpozo syenite, plot as VAGs in the diagrams of Pearce et al., they are not considered in the A1 and A2 discrimination diagrams in Figure 7-10A. As we want to keep an open perspective we also plotted the Mpozo samples in the diagram (Fig. 7-10B). RG 89.554 and RG 89.863 were also retained as they plotted in the ORG field of Pearce et al. (1984) and they have been considered numerous times as anomalous samples.

Eby (1992) suggests that diagrams in which the element ratios of Yb, Ta, Y, Nb and/or Ce are plotted, can distinguish between the two groups. The first diagram (Fig. 7-10) comprises a triangular diagram in which the Y, Nb and Ce content is plotted. Within this plot the A1 and the A2 field are separated from each other by a line which corresponds to an Y/Nb ratio of 1,2. In Figure 7-10 it can be seen that the samples plot on the borderline between the two fields.

A B

Figure 7-10: A1 and A2 discrimination diagrams (Eby, 1992). A) Plot only including the Noqui granite and the hypabyssal rocks; B) Same plot as A but with additional Mpozo syenite.

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In his paper, Eby (1992) describes that the A1 group has similar element ratios as OIBs. The other group has certain similarities to average crust and island-arc basalts (IABs). For further discrimination and discussion he proposed two other diagrams, which are given in Figures 7-11A and 7-11B. The first diagrams plots the Yb/Ta ratio vs. the Y/Nb ratio. The second diagrams comprise the same x-axis but displays the Ce/Nb ratio on the y-axis. A1 group granitoids plot within or near the OIB fields, suggesting a source similar to OIBs. The A2-types will plot within the IAB field or form a trend between the OIB and the IAB field extending from average continental crust to island-arc basalts. This would suggest that these granitoids formed by subduction or continent-continent collision. Some samples plot near the OIB field but most samples follow a trend in between the OIB and IAB field.

Figure 7-11: A) Yb/Ta vs. Y/Nb diagram; B) Ce/Nb vs. Y/Nb diagram (Eby, 1992). A1-type granitoids plot within or near the OIB fields. A2-type granitoids form a trend between the OIB and the IAB field, extending from average continental crust to IAB. C) and D) represent the same diagrams but with the Mpozo samples plotted as well.

7.3.3. Masuda Coryell diagrams Rare earth elements (REEs) are plotted in Masuda Coryell diagrams. To avoid the Oddo-Harkins effect the REE-values are chondrite-normarlized by data used from Sun and McDonough (1989). The elements are arranged from the most incompatible on the left, to the least incompatible elements on the right.

Figure 7-12A represents a Masuda Corryell diagram of six Noqui granite samples. This plot displays a prominent negative Eu-anomaly. The Eu/Eu*-ratio was calculated for every sample and resulted in average value of 0,15 which is in agreement with the negative peak. All samples show an enrichment in light rare earth elements (LREEs) and a depletion in heavy rare earth elements (HREE). Samples RG 89.874 and RG 13.122 show a very slight increasement in Yb and Lu.

110

Figure 7-12: Masuda Coryell diagrams. A) Noqui granite; B) Mpozo syenite; C) Hypabyssal rocks.

111

Samples of the Mpozo syenite are plotted in Figure 7-12B. Three samples display a slightly positive Eu-anomaly with an average Eu/Eu*-ratio of 1,12. RG 90.067 displays a slightly negative Eu-anomaly with an Eu/Eu*-ratio of 0.70. All Mpozo samples are enriched in LREEs while they contain lower abundances of HREEs. Two samples, RG 19.504 and RG 19.611, display a spoon shape and are enriched in Yb and Lu. A detailed examination of Figure 7-12 reveals that there are three samples which differentiate themselves from the other samples by their higher REE content. These three samples comprise RG 89.979, RG 89.554 and RG 89.863 which are also the anomalous samples discussed in section 7.3.1.

Hypabyssal rocks are plotted in Figure 7-12C and display a prominent negative Eu-anomaly. The average Eu/Eu*-ratio is constrained at 0,24. Two samples, RG 89.532 and RG 89.540 display a less prominent, but still negative, Eu-anomaly. High abundances of LREEs can be observed while the rocks are more depleted in HREEs.

When all samples are plotted in one diagram (Fig. 7-13), the differences and/or similarities between the Noqui granite, the Mpozo syenite and the hypabyssal rocks become clear. The Noqui granite and the hypabyssal rock display similar patterns. They are both enriched in LREEs and depleted in HREEs. Their negative Eu-anomaly is very prominent. The pattern of these two groups is different from the Mpozo samples which do not display a prominent Eu-anomaly. Comparing the amount of REEs in the rocks it becomes clear that rocks of the Noqui granite and the hypabyssal rocks contain higher abundances of REEs compared to the Mpozo syenite where values generally do not exceed 100.

Figure 7-13: Composite Masuda Coryell diagram of the Noqui granite, Mpozo syenite and the hypabyssal rocks.

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7.3.4. Spider diagrams As an extension of the chondrite-normalized REE diagrams, normalized mulit-element diagrams exist. In these diagrams, trace elements are added to the REE diagram, and normalized over mantle values or chondrites. As the term mantle-normalized multi-element does not roll of the tongue, the term spider diagram or spidergram is used.

Different types of spider diagrams exist. In Figure 7-14 we used the spidergram proposed by Pearce (1983) in which the data are normalized over Mid Ocean Ridge Basalts (MORB). The elements are ordered so that the most mobile elements (Sr, K, Rb and Ba) are placed at the left of the diagram and in order of increasing incompatibility. The immobile elements are arranged from right to left in order of increasing incompatibility.

Figure 7-14A displays the spidergram of the six Noqui granite samples. All samples display the same pattern, with distinct negative peaks for Sr, Ba, P and Ti.

Four Mpozo syenites are plotted in Figure 7-14B. The samples display roughly the same pattern but there are some small differences. All samples display a negative peak for P, but for RG 90.067 this is more outspoken.

The hypabyssal rocks (Fig. 7-14C) show negative peaks for Sr, Ba, P and Ti. For two samples, RG 89.540 and RG 89.532 these peaks are less outspoken. High values of K and Rb appear in most rocks, but RG 89.554 and RG 89.863 reveal much lower values for these elements. Similar to Figure 7-12C the three anomalous samples display the highest values and for some elements they differentiate themselves from the hypabyssal rocks.

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Figure 7-14: Spidergrams. A) Noqui granite; B) Mpozo syenite; C) Hypabyssal rocks.

114

Figure 7-15 represents a composite spidergram. In this diagram one can see that the Noqui granite and the hypabyssal rocks display the same pattern with negative peaks for Sr, Ba, P and Ti. Rocks of the Mpozo syenite differentiate themselves from the others samples as they do not display a negative peak for Ba and Ti. This composite diagram thus reveals that rocks of the Noqui granite and the hypabyssal rocks display a similar chemical signature, which is different from the Mpozo syenite.

Figure 7-15: Composite spidergram.

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8. DISCUSSION GEOCHEMISTRY

The chemical composition of igneous rocks offers an important tool to find out which genetic processes occurred at greater depths. Based on their geochemical signature, rocks can be classified into groups. Plotting rocks in diagrams and classifying them, instead of looking at tables with geochemical data, helps us to organize our observations and ideas. During the interpretation of these diagrams, and their classifications, carefulness is necessary as these classifications should not lead to rigid thinking. During a geological and geochemical study, an open view of multiple working hypotheses and/or processes is required, taking into account all relevant aspects of available data (field observations, hand specimens, microscopy, geochemistry, geochronology, remote sensing, etc.). In this section, we will discuss the results given in chapter seven.

8.1. NOQUI GRANITE + HYPABYSSAL ROCKS VERSUS MPOZO SYENITE In addition to Figures 7-2 and 7-3, the Harker diagrams (Fig. 7-5) of the major elements and the variation diagrams of the trace elements (Fig. 7-7) reveal that the rocks can be divided into two groups. A first group comprises the Noqui granite and the hypabyssal rocks, which differ from the

Mpozo syenite by a chemical gap in their SiO2 content. Based on the diagram in Figure 7-6 the two groups of rocks are not related by any mixing process.

8.2. ASSIMILATION AND METASOMATISM Major element data (section 7.2) revealed that some of the hypabyssal rocks (RG 89.979, RG 89.554 and RG 89.863) display very high amounts of SiO2 (e.g. Fig. 7-2). These rocks are characterized by a

SiO2 content of more than 77,5% (Table 7-1; see bold figures). These high values are also reflected in the normative mineralogy of the rocks, which contain more than 45% normative quartz (Table 7-2; see bold figures). Such high values generally do not occur in magmatic rocks, therefore we assume that these chemical data do not reflect the signature of the original magmatic rock.

In section 5.1 we discussed the high amount of modal quartz in thin section RG 89.863. This was explained by the fact that the rocks are intrusive in the metaquartzites of the Matadi Formation, and showed effects of assimilation of the magmatic rocks. As RG 89.979 and RG 89.554 are also intrusive within these rocks, the same process (quartzite assimilation) probably resulted in remarkably high

SiO2 contents. Two other hypabyssal rocks, RG 89.876 and RG 90.142, also display high amounts of

SiO2. As these values do not exceed 77,5%, they might reflect the original signature of the rocks but carefulness is necessary in their interpretation.

The three samples with very high SiO2 values also display anomalous trace element values compared to the other hypabyssal rocks (Table 7-4; see bold figures). These anomalous values are observed for the following elements: Y, Zr, Nb, La, Ce, Pr, Nd, Sm, Gd, Dy, Ho, Er, Yb, Lu, Hf, Ta, W, Pb, Th and U. The values of these elements are similar to those of the Noqui granite (Table 7-4; Figures 7-7, 7-12 and 7-14). Therefore we suggest that alkaline metasomatism, due to the Noqui granite, has affected these hypabyssal samples in line with the possible subsurface extension (to the north) of the dome- like Noqui body beneath the Matadi Formation (see chapter three, MGE 3).

Our hypothesis of a metasomatic stage is supported by the study of Behiels (2013). His microscopic study of the Noqui granite has revealed the influence of a late metasomatic stage, evidenced by fibrous riebeckite overgrowing aegirine. The transformation of aegirine to riebeckite requires

116 hydroxyl groups to form the amphibole after the pyroxene. This process can be attributed to percolating fluids (metasomatism).

We thus suggest that the samples RG 89.979, RG 89.554 and RG 89.863 were affected by quartzite assimilation and metasomatism and therefore do not reflect the original signature of the hypabyssal rocks. As a result we will not include these samples in any of the further interpretations.

8.3. TECTONIC SETTING Pearce was one of the pioneers who tried to fingerprint magmas from different tectonic settings based on their chemical signature. The first attempts were made on basalts, but in 1984 discrimination diagrams also became available for rocks with a granitic composition. A few years later Whalen et al. (1987) came up with another discrimination diagram to distinguish A-type granites from I- and S-type granites. In these diagrams the Noqui granite and the hypabyssal rocks are classified as A-type granites which formed in a within plate (WPG) tectonic setting. In the diagram of Whalen et al. (1987), some of the Mpozo syenite samples plot in the A-type field, while others plot just outside of the field, but in its vicinity. As the samples plot as volcanic arc granites (VAGs) in the diagrams of Pearce et al. (1984) their interpretation is less straight forward. In the beginning of this chapter we mentioned that the diagrams should not lead to rigid thinking. Therefore we do not necessarily conclude that the Mpozo syenite formed in a VAG and keep an open mind for the further interpretation of the data of the Mpozo syenite.

For the Noqui granite and the hypabyssal rocks the previous diagrams thus support the hypothesis of A-type granites. For the Mpozo syenite this is not the case, but we cannot exclude that these samples might also reflect an A-type granitoid. When dealing with A-type granites, it is possible to distinguish them further into A1- and A2-type granites. According to Eby (1992; 2011) the A1-type granites are formed in an anorogenic setting while A2-type granites originate in post-orogenic environments.

Figure 7-10B shows that almost all samples plot on the borderline of the A1- and A2-field, therefore their interpretation is biased. A similar situation is obtained in Figure 7-10A where only the Noqui granite samples and hypabyssal rocks have been plotted. Literature revealed that it is not uncommon for A-type granites to plot on the borderline (Eby, 2011: Keivy Alkaline Province and Chilwa Alkalite Province). In his paper Eby (2011) describes that these granitoids formed indeed in an extensional setting and are thus anorogenic. However, to some extent they have a post-orogenic signature and therefore the samples plot on the borderline. Such post-orogenic signature can be explained by crustal contamination of older crust.

Eby (1992) also suggests other diagrams to discriminate between petrogenetic processes of the granitoids. These Yb/Ta vs. Y/Nb and Ce/Nb vs. Y/Nb diagrams are given in Figure 6-10. None of the samples plot immediately in the OIB or the IAB field, but they plot in between, and in the vicinity of the OIB field. According to Eby (2011) samples displaying a trend between these two fields have been affected by crustal contamination. Thus it is suggested that the A-type granitoids originate from an OIB-type magma, and thus have a mantle component. Afterwards crustal contamination caused a shift in the geochemical composition.

The Kimezian basement, with a post-orogenic signature, might thus have influenced the chemical composition of the rocks and thus explains Figures 7-10A and 7-10B.

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8.4. FRACTIONAL CRYSTALLIZATION Trace elements are also very useful in testing assimilation – fractional crystallization processes (AFC). We will both look at the Rare Earth Elements and the other remaining trace elements, which are presented by Masuda-Corryell diagrams in section 7.3.3. and spidergrams given in section 7.3.4. Figure 7-13 and 7-15 reveal that the Noqui granite and the hypabyssal rocks display a similar pattern that is different from the signature of the Mpozo syenite, a feature which was already observed in several other geochemical diagrams (see point 8.1).

The Masuda-Coryell diagrams of the Noqui granite and the hypabyssal rocks are characterized by a V- shape which is induced by the strong negative Eu-anomaly. This anomaly has not been observed in the Mpozo syenite and is probably caused by fractionation of plagioclase. This hypothesis is supported by the modal mineralogy observed in thin sections. Behiels (2013) observed only very slight amounts of plagioclase in the Noqui granite and also the hypabyssal rocks do not contain a lot of plagioclase. RG 89.532 and RG 89.540 display a much smaller negative Eu-anomaly, and this is also evidenced in their thin sections, which contain more abundant plagioclase. Even larger amounts of plagioclase are observed in the Mpozo syenite samples. Therefore the melt, giving rise to the Noqui granite and the hypabyssal rocks, must have suffered fractional crystallisation of plagioclase.

The spidergrams of the Noqui granite and the hypabyssal rocks both display negative anomalies for Sr, Ba, P and Ti. These diagrams thus indicate that feldspar (negative Sr and Ba anomalies), apatite (negative P anomaly) and Fe-Ti oxides (negative Ti-anomaly) were fractionated from the magmas. In the spidergram of the Mpozo syenite we observe a negative anomaly for P and a less dominant anomaly for Sr and Ba. Just as for the Noqui granite and the hypabyssal rocks, the Mpozo syenite thus underwent fractionation of apatite and maybe some fractionation of feldspar.

8.5. PETROGENESIS The diagrams of Eby (1992) suggest that both groups of rocks (Noqui granite + hypabyssal rocks versus Mpozo syenite) originate from an OIB-type source which was affected by crustal contamination to produce the melt at the origin of groups of rocks. According to him, two models may account for the further evolution of this melt.

In a first (comagmatic) model (Fig. 8-1A) we assume that the melt gave rise to evolved liquids after precipitation of (ultra)mafic cumulates. Fractional crystallization of the evolved liquid then produced successively the Mpozo syenite followed by the Noqui granite and the hypabyssal rocks.

A second (cogenetic) model (Fig. 8-1B) suggest that both groups of rocks would following a different path. After precipitation of (ultra)mafic cumulates, evolved liquids would give rise at different moments to both groups of rocks separately.

Solely based on the available geochemical data it is impossible to favour one of both models. More information on the emplacement ages of both the Mpozo syenite and the Noqui granite might give additional support for one hypothesis. Therefore we refer to chapter nine and ten in which U-Pb isotope geology is used to constrain the emplacement ages of the rocks.

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Figure 8-1: Schematical presentation of the petrogenesis. A) model 1: comagmatic; B) model 2: cogenetic.

8.6. DISCUSSION REGARDING PREVIOUS RESEARCH In section 7.1 we have mentioned the availability of earlier limited geochemical data. Therefore we compare our own results with earlier preliminary studies.

Franssen and André (1988) analysed several of the hypabyssal rocks, which they described as metarhyolitic sills and microgranitic veins, and compared them to the Noqui granite. The Mpozo syenite was not taken into account in this study. As spidergrams of both groups of rocks displayed a similar pattern, they suggested their comagmatic character. A model assuming that the rocks were derived by fractional crystallization of the same parental magma was discarded on the base of the available radiometric age of the Noqui granite in 1988, which later proved to be incorrect (Tack et al., 2001). A process of alkaline metasomatism (enrichment of Th, Nb, Y, Zr and REEs) limited to the microgranitic veins of the Noqui granite was also envisaged. We have already discussed metasomatic effects in section 8.2.

Makutu et al. (2004) performed a geochemical study of the Noqui granite and the Mpozo syenite. Based on their results they assume that the Noqui granite and the Mpozo syenite are related to each other. They explain the petrogenesis of these rocks by fractional crystallization with element fractionation attributed to different degrees of partial melting (relatively higher for the Noqui granite compared to the Mpozo syenite) of an enriched source (possibly an enriched mantle) with only limited to negligible crustal contamination.

Behiels (2013) suggests that the Noqui granite was formed by an early crystallization of the K-Na-Si-Al system giving rise to a “crystal mush” of mesoperthites and quartz followed by late – i.e. agpaitic – crystallization of dark minerals, i.e. aegirine, biotite, opaque minerals and allanite. As a result of percolating fluids, metasomatism caused the chemical alteration of aegyrine to fibrous riebeckite.

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9. GEOCHRONOLOGY

The magmatic rocks of the Matadi region have been subjected to former dating efforts which are described by Delhal and Ledent (1978), Tack et al. (2001) and Behiels (2013; Annex 7). Here we present new zircon U-Pb ages of four magmatic rocks of our study region, selected on the level of our earlier envisaged “main geological events” (MGE 1 – MGE 4), and relate them to their emplacement. They include one Noqui granite sample (RG 23.109), two samples of the Mpozo massif, including one pink (RG 89.709) and one white rocktype (RG 76), and one hypabyssal rock (RG 89.590 = same sample location as RG 90.142) (Fig. 9-1). The analyses were carried out with LA-ICP-MS (section 3.4).

9.1. ZIRCON MORPHOLOGY In the next section, we describe the zircon morphology of the four samples, illustrated by photographs of the zircons under the binocular microscope and associated CL-images (Figs. 6-2 and 6-3). Additional images and information on the analysed spots are given in Annex 6.

9.1.1. Noqui granite Zircons within the Noqui granite (RG 23.109) are mainly subhedral to anhedral, but a few euhedral zircons can be observed as well. When they have well developed crystal faces, they display a prismatic-bipyramidal habit. The zircons range up to 400 µm in length and have length to width ratios between 2:1 and 3:1. They have a light brown (Fig. 9-2A) and sometimes orange colour. Almost all zircons contain inclusions and display cracks. CL-images (Fig. 9-3A) reveal the presence of cores and concentric zoning.

9.1.2. White Mpozo syenite The zircons of sample RG 76 are rather large, ranging up to 550 µm in size with a length to width ratio between 2:1 and 3:1. They mainly display a sub- to anhedral morphology. The typical bipyramidal shape is not well expressd and most crystals are rounded. Their colour varies from pink to purple (Fig. 9-2B). Furthermore, the crystals display cracks and contain inclusions which sometimes give them an opaque appearance. Concentric zoning was not observed under the binocular microscope but was revealed by the CL-images (Fig. 9-3B), which also indicate the presence of cores.

9.1.3. Pink Mpozo syenite RG 89.709 comprises the white variety of the Mpozo syenite. The zircons within this sample (Fig. 9- 2C) are mainly sub- to euhedral with some anhedral crystals. Even though the crystals display somewhat rounded edges, their prismatic-bipyramidal shape is often recognizable. Furthermore they are characterized by a dark brown to pink colour with a dark edge. They range in size between 150 and 250 µm. The length to width ratio varies between 2:1 and 3:1. Inclusions often give the crystals an opaque aspect. The large amount of inclusions is confirmed by the CL-images (Fig. 9-3C), which also reveal concentric zoning.

9.1.4. Hypabyssal rock The zircons of RG 89.590 are sub- to euhedral and exhibit prismatic-bipyramidal forms. Their length generally varies between 150 and 250 µm with length to with ratios between 2:1 and 4:1. They display a light to dark brown colour and are largely transparent but opaque zones, caused by

120 inclusions, occur. Under transparent light (Fig. 9-2D) it is sometimes possible to detect features of euhedral concentric zoning. This is even better expressed in the CL-images (Fig. 9-3D). The latter also indicate, in some zircons, the presence of a core.

Figure 9-1: Localization of the samples. 1) Noqui granite (RG 23.109); 2) White Mpozo syenite (RG 76); 3) Pink Mpozo syenite (RG 89.709); 4) Hypabyssal rock (RG 89.590); 5) Noqui granite (RG 71.299) analyzed by Tack et al. (2001).

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Figure 9-2: Zircons under the binocular microscope. A) Noqui granite; B) White Mpozo syenite; C) Pink Mpozo syenite; D) Hypabyssal rock.

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RG23109_58 RG76_24 1034 ± 56 Ma 1907 ± 33 Ma

RG76_23 RG23109_59 1907 ± 32 Ma 1017 ± 48 Ma

RG89590_85 1070 ± 81 Ma

RG89709_26 1925 ± 44 Ma

Figure 9-2: CL-images. A) Noqui granite; B) White Mpozo syenite; C) Pink Mpozo syenite; D) Hypabyssal rock. Red circles indicate the analysed spots.

9.2. DATING RESULTS In this section we give the results of the LA-ICP-MS zircon U-Pb dating of the four magmatic samples. Although we observe cores in some of the zircons, there are no significant age differences between the cores and the rims of the crystals. For the exact location of the analyzed spots and for the detailed results we refer respectively to Annex 6 and Annex 7.

9.2.1. Noqui granite Twenty-nine analyses were obtained from sample RG 23.109. These analyses resulted in a range of U/Pb ages between 933 - 1100 Ma. The data are displayed in the U-Pb concordia plot in Figure 9-3A. As all ages fall in a relatively small range, the 2σ ellipses cluster together. A weighted mean of all data gives an emplacement age of 1018 ± 19 Ma.

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A B

C D

Figure 9-3: LA-ICP-MS zircon U-Pb concordia diagrams. A) Noqui granite; B) White Mpozo syenite; C) Pink Mpozo syenite; D) Hypabyssal rock.

9.2.2. White Mpozo syenite Thirty-nine analyses were obtained from sample RG 76. Except for three analyses, all the data cluster together on concordia and the U/Pb ages range between 1907 - 1977 Ma. A weighted mean of 1948 ± 10 Ma gives its emplacement age.

9.2.3. Pink Mpozo syenite Twenty-one analyses from sample RG 89.709 show an age range of 1781 - 2033 Ma. A weighted mean of 1947 ± 30 Ma gives its emplacement age. The concordia plot shows that several zircons underwent Pb-loss. These points plot on a discordia line of which the lower intercept gives a poorly constrained age of 610 ± 150 Ma.

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9.2.4. Hypabyssal rock Twenty-six < 10% discordant analyses from sample RG 89.590 show a range of U/Pb ages between 968 - 1167 Ma. On the U-Pb concordia plot (Fig. 9-3B), the cluster of data reveals an emplacement age of 1043 ± 25 Ma. Three samples deviate from the concordia curve and were probably drawn down a discordia line by Pb-loss. Howevr, as only three analyses deviate, it is impossible to draw a reliable discordia line. Solely based on these three analyses a “possible” discordia line points to an intercept with concordia at approximately 500 Ma.

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10. DISCUSSION GEOCHRONOLOGY

10.1. NOQUI GRANITE AND HYPABYSSAL ROCKS In chapter eight we concluded that the rocks of the Noqui granite and the hypabyssal rock display a similar geochemical signature, which therefore points to a comagmatic character of these rocks. The results given in section 9.2 confirm this statement, as both groups of rocks present a similar emplacement age of approximately 1,0 Ga.

10.2. MPOZO SYENITE VERSUS NOQUI GRANITE AND HYPABYSSAL ROCKS Both the white and pink Mpozo syenite display a similar emplacement age of respectively 1948 ± 10 Ma and 1947 ± 30 Ma. Such a ca. 2,0 Ga age indicates that the emplacement of the Mpozo body was by no means related to the ca. 1,0 Ga emplacement of the Noqui body and accompanying hypabyssal rocks. On the contrary, the Mpozo ages show that emplacement of these rock types occurred at a late stage of the ca. 2,1 Ga migmatisation event of the Kimezian basement and thus are related to its late geological evolution.

Therefore, the discrepancies discussed in the geochemical part of this study (opposition of two groups of magmatic rocks, i.e. Noqui granite and accompanying hypabyssal rocks versus Mpozo syenite) may convincingly be explained by our new radiometric results, showing that both groups of magmatic rocks are linked to completely different geological events.

Thus, the discussion of the apparently aberrant plot in the VAG field of the Mpozo syenite (Fig. 7-8) – at variance with the WPG field for the Noqui granite and accompanying hypabyssal rocks and discussed more in detail in chapter eight - is no longer relevant. The meaning of the Mpozo rocks plotting in the VAG field has to be (re)considered in the light of the (late) geological evolution of the Kimezian basement. This falls out of the scope of our study.

Finally, the ca. 2,0 Ga age of the Mpozo syenite is a late marker of the geological evolution of the Kimezian (i.e. pre-AWCO) basement (on the African side) prior to the first extensional event (E1) starting the evolution of the AWCO (on the Brasilian side) around 1,7 Ga (Pedrosa-Soares and Alkmim; 2011).

10.3. NEW AGES COMPARED TO EARLIER AGES OF THE LOWER-CONGO REGION The Noqui granite has given a U-Pb zircon SHRIMP emplacement age of 999 ± 7 Ma (Tack et al. , 2001). Within error, our new age of 1018 ± 19 Ma overlaps with the 999 ± 7 Ma age, although they have been obtained by slightly different methods.

For the hypabyssal rocks, our new 1043 ± 25 Ma age is in line with the very poorly constrained and obsolete U-Pb bulk zircon age of ca. 1050 Ma of Delhal and Ledent (1978).

Similarly, for the Mpozo syenite our new ages (1947 ± 30 Ma and 1948 ± 10 Ma) significantly improve the extremely poorly constrained and obsolete U-Pb bulk zircon age of 1960 ± 594 Ma (Delhal and Ledent, 1978).

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10.4. Pb-LOSS In the concordia plot of the hypabyssal rocks (Fig. 9-3B) three analyses deviate from the concordia line, suggesting Pb-loss. Based on only three points the obtained “discordia line” is relatively poorly constrained, with a lower intercept of discordia at approximately 500 Ma.

For the white Mpozo syenite (RG 76) we also observe three points deviating from the concordia line (Fig. 9-3C). The scatter of these points does not allow to determine a lower intercept along the discordia curve. For the pink Mpozo syenite, several samples deviate from concordia and converge along discordia line to a lower intercept of 610 ± 150 Ma (Fig. 9-3D). This suggests a geological event that might have caused Pb-loss around that time.

Both the hypabyssal rocks and the Mpozo syenite, affected by lead loss (respectively around 500 Ma and 610 ± 150 Ma), point to have been affected by the same event. It overlaps with the Pan African orogeny, which took place around 550 Ma (Tack et al., 2001) and of which the peak in the Lower- Congo region is constrained by – only – the 566 Ma 40Ar - 39Ar age of Frimmel et al. (2006).

Abundant ages of the same order of magnitude (ca. 450 to 600 Ma) are available for various episodes of the AWCO: on the African side: see Cahen et al. (1984 and references therein; various methods since the 1950’s!), and on the Brasilian side: see Gradim et al. (2014 and references therein; very recent methods). The discussion of all these results falls out of the scope of our study.

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11. DISCUSSION OF THE GEOLOGY OF THE MATADI REGION

In this section we integrate all our the results – field observations, macroscopic descriptions, microscopic descriptions, geochemistry and geochronology – to better constrain the general geology of the Matadi region.

In section 2.4 a preliminary and tentative four-stage timetable (MGE 1 to MGE 4) was proposed as a working hypothesis. Our data allow to better understand the geological history of the region and to coplete this timetable (Fig. 11-1).

Figure 11-1: Timetable summarizing the geological history of the (broader) Matadi region. Light grey text indicates aspects out of the scope of our study.

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Integration of our own results with data of Behiels (2013; “Palabala Formation”; tectonic contact between the Noqui granite and the Mpozo syenite), allows us to make a lithostratigraphical reconstruction of the Matadi region (Fig. 11-2), which is an important update of the lower part of Figure 2-13 (Tack et al., 2001).

Figure 11-2: Lithostratigraphic reconstruction of the Matadi region. Mylonites and/or M/A correspond to the former “Palabala Formation” (pro parte).

Furthermore, our results also allow substantial improvements to a tentative geological sketch map of the Matadi region. A first update concerns the non-existence of the “Palabala Formation”, which corresponds to a tectono-structural unit and should therefore no longer be represented on the geological map as a lithostratigraphic unit. Secondly, abundant felsic and mafic hypabyssal intrusions occur within the Matadi region.

As Bertossa and Thonnart (1957) already discarded the existence of the “Palabala Formation” and strengthened the occurrence of (only) the Matadi Formation, their 1957 map forms an appropriate guideline for the revision of the northern part of the geological map of the Matadi region. Therefore, we have plotted the felsic and mafic intrusions on this map (Fig. 11-3).

As Behiels (2013) described the mineralogical distribution within the Noqui granite and Mpozo syenite and evidenced – indeed only locally – a tectonic contact between the Noqui and Mpozo bodies, the outline of these bodies on the 2008 geological map may also be substantially improved

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(i.e. the southern part of the geological map of the Matadi region). The partial and tentative outline for this region is proposed in Figure 11-4. South of the DRC-Angola border, the adopted modified outline refers largely to the earlier geological map of Korpershoek (1964). The outline of the prolongation of the Kimeza basement in Angola falls out of the scope of this thesis.

Figure 11-3: Felsic and mafic hypabyssal rocks plotted on the map of Bertossa and Thonnart (1957).

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Figure 11-4: Red dotted lines = N-S trending shear zones and corridors; only main “lineaments” are indicated after preliminary data from remote sensing and DEM (digital elevation model) as given in Behiels (2013); note: offset of Noqui granite along several of the (late) N-S trending shear zones; thick red dotted line = tectonic contact between Noqui and Mpozo bodies with dip to the west; green lines = roads; black lines (full and/or dotted) = borders of the geological units; 1 = Pic Cambier; 2 = Kinzao Quarry. Geological units: Ki = Kimeza basement; Ma. F. = Matadi Formation; Ga. F. = Gangila Formation; No = Noqui granite; Mp = Mpozo syenite-monzonite. Base map from Behiels (2013): blue = riebeckite; green = aegirine; brown = lepidomelane; black = riebeckite + aegirine + lepidomelane; localities of observed contact metamorphism and/or metasomatism are not given; similarly, at this stage of study, no outline of contact metamorphic nor metasomatic aureole can be proposed on the map.

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12. SUMMARY AND CONCLUSION

Already by the end of the Eburnian-Transamazonian orogeny (2,1 Ga) there was a connection between the São Francisco craton of Brazil and the Congo craton of Africa. This connection remained unbroken until the opening of the South Atlantic Ocean in the Cretaceous. During that period the São Francisco-Congo craton became incorporated in supercontinents and endured cycles of continental break-up and amalgamation. Around 600 Ma, compressional events related to Gondwana amalgamation, resulted in the formation of the Araçuaí-West Congo Orogen (AWCO). In Brasil these compressional events are referred to as being the result of the Brasiliano orogeny, while in Africa this is described as the Pan African orogeny.

The Pan African orogeny gave rise to the West Congo belt, which is situated subparallel to the Atlantic coast, between 1° and 12° south of the equator. This structural unit is 1400 km long , 150 to 300 km wide and comprises an ENE-verging fold-and-thrust belt. In its central segment, the West Congo belt displays a prominent flexure which overlaps with the Lower-Congo region. In that area, the peak stage of the orogeny is constrained at 566 Ma (40Ar – 39Ar dating; Frimmel et al., 2006).

In the Matadi region, i.e. our study region, the Palaeoproteroic basement comprises the 2,1 Ga old Kimeza Supergroup, which is covered by the West Congo Supergroup. This unit can be subdivided, from old to young, in the Zadinian, Mayumbian and West Congolian Group. Within our region, there are also two plutonic bodies exposed, being the Noqui granite and the Mpozo syenite. The Noqui granite comprises a peralkaline A-type granite. Recent dating of the pluton resulted in an emplacement age of at 999 ± 7 Ma, evidencing a pre-orogenic emplacement (Tack et al., 2001). Compared to the Noqui granite, the Mpozo syenite is not as well documented. Delhal and Ledent (1978) have tried to date the Mpozo syenite, resulting in a poor emplacement age of 1960 ± 594 Ma (U-Pb dating on bulk zircons).

Over the years, several mapping attempts (Behiels, 2013; Annex 1) have tried to achieve a geological representation of the region. However, these attempts were based on limited and/or scattered observations and data, without a systematic approach to integrate all available data originating from various sources (both published or unpublished). Tack (1975a) compiled an “integrated” 1:200.000 geological map of the whole Lower-Congo region to the west of the 14th meridian, thus completing the geological coverage (1:200.000) which had previously been achieved to the east of the 14th meridian. Since 1975, no systematic update of this map has been performed. The 1975 geological map is now outdated and needs considerable improvement. By lack of an alternative, more recent document, the 1975 map was digitized in 2008 as a starting point for modern updating purposes.

Two recent field missions (2004 and 2011) resulted in new crucial information concerning the geology of the Matadi region. Contrary to the observations of Tack (1975a), no angular unconformity was observed between the basement and the overlying base of the Zadinian Group (“Palabala Formation”). The discussion regarding the meaning of the “Palabala Formation” has been ongoing for several decades. During the two field missions the “Palabala Formation” was revisited and observations indicated that this package comprises mylonitic rocks which originate from various protoliths: Kimezian migmatitic paragneisses and amphibolites, metaquartzites of the Matadi Formation and Mpozo syenite. This suggests that there is no “Palabala Formation” at the base of the

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Zadinian Group unconformably overlying the Kimezian basement. Therefore the “Palabala Formation” should thus be regarded as a tectono-structural unit rather than a lithostratigraphic unit.

In order to better constrain the geology of the Matadi region, in continuity with the work of Behiels (2013), we mainly focused on the question concerning the “felsic magmatic bodies” of limited extent which are mainly intercalated in the Matadi Formation (and the former “Palabala Formation”). Behiels (2013) wonders whether these rocks were emplaced as extrusive or intrusive rocks and whether they are related to the Noqui granite.

With the purpose of answering these questions the field observations and samples of three field geologists, i.e. Hugé (1950), Massar (1965) and Steenstra (1970), were examined. As field access was not possible to us, their field observations, together with more recent information of Tack and Baudet (2014), were of great importance. Completed by a macroscopic and microscopic study of the rocks, three groups of rocks, all of them to some extent deformed, were distinguished in the region: 1) rocks with a felsic magmatic protolith; 2) rocks with a sedimentary protolith and 3) rocks with a mafic magmatic protolith. Based on their identification the samples were colour coded an plotted to create a lithological map of the area. Microscopically the three groups of rocks comprise the following characteristics:

1) Rocks with a felsic magmatic protolith are characterized by a blastoporphyritic texture in which the porphyroclasts are made up of perthitic alkali feldspar, quartz and sometimes plagioclase. The surrounding, more fine-grained, groundmass is also made up of alkali feldspar, quartz and plagioclase. Furthermore all of the rocks contain micas. These micas always include muscovite/sericite and sometimes also biotite. Accessory minerals comprise chlorite, epidote, opaque minerals, sphene and allanite. In some of the samples we also observe accessory calcite.

2) Rocks with a sedimentary protolith are mainly composed of quartz and muscovite/sercite. In some of the thin sections quartz displays very irregular crystal edges, describing a seriate- interlobate texture, while in most samples quartz displays polygonal crystals, described as seriate-polygonal. Accessory minerals comprise epidote, opaque minerals, chlorite, zoisite, biotite and garnet. In one thin section, randomly oriented garnet porphyroblasts were observed, indicating contact metamorphism.

3) Within the group of rocks with a mafic magmatic protolith various textures occur. The mineralogy on the other hand is generally very similar. The most abundant minerals are actinolite, epidote and plagioclase. The abundance of biotite, calcite and quartz is strongly variable from sample to samples. Accessory minerals include sphene, allanite, opaque minerals and muscovite/sericite.

The mineral assemblage of these three groups reveals that the rocks in the Matadi region were affected by regional greenschist facies metamorphism, presumably related to the Pan African orogeny. This orogeny also caused deformation. Within all of the rocks we find evidence of ductile deformation by e.g. kinked micas or deformation twins. The rocks with a felsic magmatic protolith often also display fractured porphyroclasts suggesting that also brittle deformation occurred. Furthermore there is also evidence of recovery and recrystallization within all of the rocks. These effects are easily observed in the first group of rocks in which the porphyroclasts are surrounded by recrystallised rims of mainly quartz and alkali feldspar. Additionally, quartz with subgrains and very

133 irregular edges, evidencing respectively recovery and dynamic recrystallization occur in all of the rocks. Moreover a lot of grains display a polygonal aspect as a result of grain boundary area reduction and/or static deformation. We suggest that the tectono-metamorphic overprint of the three groups of rocks occurred under a low pressure regime with high persistent temperatures.

A main focus was put on the rocks with a felsic magmatic protolith. Field observations indicate the limited size of the bodies and the intrusive nature of the rocks in the metaquartzites. Their blastoporphyritic texture suggests that the rocks were porphyritic before deformation occurred. This porphyritic texture indicates a two phase cooling history of the rocks. Therefore we assume that the rocks were intrusive at intermediate depth and are therefore best described as hypabyssal rocks. As some of these samples display very high amounts of quartz we suggest that some quartzite assimilation also occurred.

To determine whether these hypabyssal rocks are related to the Noqui granite and/or the Mpozo syenite, a geochemical and geochronological study was carried out.

Geochemical data point out that the geochemical signature of the hypabyssal rocks is similar to that of the Noqui granite and is different from the Mpozo syenite. Within the group of the hypabyssal rocks we observe a few samples with very high SiO2 values which also display anomalous trace element (Y, Zr, Nb, La, Ce, Pr, Nd, Sm, Gd, Dy, Ho, Er, Yb, Lu, Hf, Ta, W, Pb, Th and U) values compared to the other hypabyssal rocks. The values of these elements are similar to those of the Noqui granite. Therefore we suggest that alkaline metasomatism, due to the Noqui granite, has affected these hypabyssal samples in line with the possible subsurface extension (to the north) of the dome-like Noqui body beneath the Matadi Formation. Alkaline metasomatism can best be observed in the Masuda Coryell diagrams and the spidergrams.

According to the diagrams of Pearce et al. (1984) and Whalen et al. (1987) these hypabyssal rocks and the Noqui granite are A-type granites, formed in a within plate tectonic setting (WPG). In the diagrams of Eby (1992) these samples plot on the borderline between the A1 and A2 fields. This suggests that there might be crustal contamination of older crust (Eby, 2011). The process of crustal contamination is confirmed by the Yb/Ta vs. Y/Nb and Ce/Nb vs. Y/Nb diagrams (Eby, 1992). According to Eby (1992; 2011) the hypabyssal rocks and the Noqui granite originate from an OIB-type source that was affected by crustal contamination resulting into evolved liquids. Based on the Masuda Coryell diagram and the spidergrams we assume that these evolved liquids then fractionated feldspars, apatite and Fe-Ti oxides to eventually result in the present day rocks.

The geochemical signature of the hypabyssal rocks and the Noqui granite is different from the one of the Mpozo syenite. In the diagram of Pearce et al. (1984) the syenite plots as a volcanic arc granite (VAG). Furthermore fractional crystallization in the Mpozo syenite only comprises fractionation of apatite and minor feldspars.

The assumptions based on the geochemical data are confirmed by the geochronological information. LA-ICP-MS zircon U-Pb dating resulted in new emplacement age data of both the hypabyssal rocks, the Noqui granite and the Mpozo syenite. For the hypabyssal rocks a new emplacement age of 1043 ± 25 Ma was obtained, which is in line with the very poorly constrained and obsolete U-Pb bulk zircon age of ca. 1050 Ma of Delhal and Ledent (1978). For the Noqui granite a new emplacement age of 1018 ± 19 Ma was obtained. Within error, our new age overlaps with the U-Pb zircon SHRIMP

134 emplacement age of 999 ± 7 Ma (Tack et al., 2001). These data thus support that the hypabyssal rocks comagmatic with the Noqui granite, as both types of rocks were formed at approximately 1,0 Ga.

Both the white and pink Mpozo syenite display a similar emplacement age of respectively 1948 ± 10 Ma and 1947 ± 30 Ma. Such a ca. 2,0 Ga age indicates that the emplacement of the Mpozo body was by no means related to the ca. 1.0 Ga emplacement of the Noqui body and accompanying hypabyssal rocks. On the contrary, the Mpozo ages show that emplacement of these rock types occurred at a late stage of the ca. 2.1 Ga migmatisation event of the Kimezian basement and thus are related to its late geological evolution.

Together with information obtained by earlier studies we are now able to sketch the evolution of the Matadi region in a chronological way, in which four “main geological events” (MGEs) may be considered, from old to young respectively.

The basement in the Matadi region comprises the Palaeoproteroic Kimeza Supergroup. After its deposition, this Supergroup became migmatised (2088 Ma; Delhal and Ledent, 1976) during the Tadilian orogeny (Eburnian-Transamazonian-aged orogeny). At 1947 – 1948 Ma the Mpozo syenite, which should - based on microscopic observations - better be described as the Mpozo syenomonzonite, was intruded within this basement (= MGE 4).

This event was followed by deposition and lithification of both the Matadi Formation and the, locally exposed, Yelala conglomerate. The ca. 1,0 Ga peralkaline Noqui granite was intruded into the host rocks of the Matadi Formation (= MGE 3). As a result of the (forceful ?) intrusion, a broad and gently dipping dome-like structure developed in the Matadi region. This setting suggests a subsurface prolongation of the Noqui granite at limited depth beneath the town of Matadi and across the Congo River along its northern (right) bank. Together with the Noqui granite, the hypabyssal rocks were emplaced as sills and dykes.

This event was followed by mafic intrusions in the Matadi Formation and are believed to have formed the feeder dykes for the overlying metabasalts of the Gangila Formation. The emplacement of these mafic intrusive rocks and sills is often controlled by reactivation of the weakness zones where the felsic hypabyssal rocks were emplaced.

All of these rocks were affected by regional greenschist facies metamorphism ( = MGE 2), which was induced by the Pan African orogeny. As a result of this orogeny, all of the rocks in the region are affected by a tectono-metamorphic overprint however, with often variable intensity of deformation.

The “last” main geological event (= MGE 1), in the Matadi region, comprises the formation of N-S (to NNW-SSE and/or NNE-SSW) trending shear zones and broader corridors formed under brittle conditions. These shear zones are characterized by generally steep dips to the west and affect all the geological rock “units” of the (broader) Matadi region.

Our study has made important contributions to the geology of the Matadi region. Based on field observations, macroscopic and microscopic descriptions, the “felsic magmatic bodies” of the region prove to be intrusive and can best be described as hypabyssal rocks. These rocks are together with the younger mafic intrusions, intrusive in the metaquartzites of the Matadi Formation. Furthermore geochemical and geochronological data have confirmed their relation with the Noqui granite.

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Contrary to some earlier suggestions, this study revealed that the Noqui granite and the Mpozo syenite are not related to each other as the emplacement of the Noqui granite and Mpozo syenite bodies are separated by ca. 1,0 Ga. Finally, our observations show that the geological map of the Matadi region need substantial improvement.

We are convinced that this thesis forms a good starting point for such a new mapping study of the Matadi region. To do so, additional information from remote sensing and/or geophysics is necessary. Additional constraints on the tectono-metamorphic overprint of the various rocks of the Matadi region would also be helpful. Finally, we suggest that isotope geochemistry (e.g. Sm-Nd, Rb-Sr) is necessary to constrain, even better, the petrogenesis of the magmatic rocks in the region.

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13. NEDERLANDSE SAMENVATTING

Reeds in het begin van de 20ste eeuw merkte Alfred Wegener op dat de kustlijnen van Afrika en Zuid- Amerika in elkaar passen. Vόόr de opening van de Atlantische Oceaan, in het Krijt, waren het São Francisco craton van Brazilië en het Congo craton van Afrika met elkaar verbonden. Deze verbinding was reeds verwezenlijkt aan het einde van de “Eburnian-Transamazonian” (2,1 Ga) orogenese (Alkmim et al., 2006). Vanaf het Paleoproterozoicum tot het Krijt bleef het São Francisco-Congo craton één eenheid, die in verschillende supercontinenten opgenomen werd en één geheel bleef doorheen verschillende cycli van opbreken en amalgamatie.

Eén van de (super)continenten, Gondwana, vormde zich rond 600 Ma. Dit ging gepaard met collisie van plaatranden, wat leidde tot de vorming van het Araçuaí-West Congo orogeen (AWCO). Deze compressionele fase werd in het gebied voorafgegaan door minstens zes extensionele fases (E1 – E6). Deze extensionele gebeurtenissen resulteerden in rifting en anorogeen magmatisme. De compressionele fase die erop volgde, resulteerde vanaf 630 Ma in de vorming van het AWCO. Aan de Braziliaanse zijde vat men deze compressionele gebeurtenissen samen onder de term “Braziliaanse orogenese”, terwijl men in Afrika verwijst naar de Pan Afrikaanse orogenese. Hier spitsen we ons enkel toe op de Afrikaanse zijde, waar de Pan Afrikaanse orogenese aanleiding gaf tot de “West Congo belt” die deel uitmaakt van het AWCO.

De “West Congo belt” bevindt zich subparallel aan de Atlantische kustlijn, tussen 1° en 12° Zuid. Het complex is 1400 km lang en 150 tot 300 km breed. Bovendien omvat het een ONO-gerichte “fold- and-thrust belt”. In het centrale gedeelte van de “West Congo belt” bevindt zich de Neder-Congo regio. In dit gebied werd de maximale intensiteit van de orogenese vastgelegd op 566 Ma (40Ar – 39Ar datering; Frimmel et al., 2006). Ten gevolge van deze orogenese zijn alle gesteenten in het gebied gekenmerkt door een tectono-metamorfe overprint.

In de Matadi regio bestaat de sokkel uit de 2,1 Ga oude Kimeza Supergroep die gekenmerkt wordt door migmatitische gneissen en amfibolieten. Deze gesteenten worden bedekt door de West Congo Supergroep waarin men van jong naar oud de volgende groepen terugvindt: de Zadiniaan Groep, de Mayumbiaan Groep en de West Congo Groep. In het gebied ontsluiten zich ook twee plutonische massieven, namelijk de Noqui graniet en de Mpozo syeniet. De Noqui graniet is een peralkalische A- type graniet. De vormingsouderdom van deze graniet werd recent vastgegelegd op 999 ± 7 Ma (Tack et al., 2001), en geeft de pre-orogene vorming van het massief weer. In vergelijking met de Noqui graniet, is de Mpozo syeniet minder gedocumenteerd. Delhal en Ledent (1978) hebben geprobeerd deze syniet te dateren (“U-Pb dating on bulk zircons”), wat resulteerde in de onprecieze ouderdom van 1960 ± 594 Ma.

In de laatste decennia werden verscheidene pogingen ondernomen om de Matadi regio in kaart te brengen (Behiels, 2013; Annex 7). Deze geologische kaarten zijn echter vaak gebaseerd op gelimiteerde datasets met weinig observatiepunten. Tack (1975a) construeerde een geologische kaart, met een 1:200.000 schaal, van het Neder-Congo gebied ten westen van de 14de meridiaan. Hiermee vervolledigde hij de geologische kaart ten oosten van de 14de meridiaan. Deze kaart is echter verouderd, aangezien er sinds 1975 geen nieuwe pogingen werden ondernomen om de geologie in kaart te brengen. Bij gebrek aan recentere documenten, werd de kaart van 1975 in 2008 gedigitaliseerd om een vertrekpunt te bieden voor nieuwe karteringen.

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Tijdens twee recente veldexpedities (2004 en 2011) werd nieuwe, cruciale informatie omtrent de geologie van de Matadi regio, verworven. In tegenstelling tot de observaties van Tack (1975a), werd er geen hoekdiscordantie aangetroffen tussen de sokkel en de bovenliggende basis van de Zadiniaan Groep (“Palabala Formatie”). Tijdens de veldexpedities van 2004 en 2011 werd deze “Palabala Formatie”, waarvan de definite al enkele decennia voor discussie zorgt, bestudeerd. Hierbij stelde men vast dat deze “formatie” een pakket van mylonieten omvat. Deze mylonieten bevatten verscheidene protolieten: migmatitische paragneissen en amphibolieten van de Kimeza Supergroep, metakwartsieten van de Matadi Formatie en Mpozo syeniet. Op basis hiervan wordt verondersteld dat er geen “Palabala Formatie” aanwezig is aan de basis van de Zadianiaanse Groep. De “Palabala Formatie” moet dus eerder beschouwd worden als een tectono-struturele eenheid in plaats van een lithostratigrafische eenheid.

Behiels (2013) merkte op dat er binnen de Matadi Formatie “felsische magmatische lichamen” met beperkte afmetingen aanwezig zijn. Om de geologie van het gebied beter te begrijpen, werd er hoofdzakelijk op deze “felsische magmatische lichamen” gefocust. Behiels (2013) vroeg zich bovendien af of deze “felsische magmatische lichamen” extrusieve of intrusieve gesteenten omvatten en of ze gerelateerd zijn aan de Noqui graniet.

In de hoop deze vragen te kunnen beantwoorden, werden in deze studie de veldnota’s en monsters van drie veldgeologen, i.e. Hugé (1950), Massar (1965) en Steenstra (1970), bestudeerd. Aangezien het studiegebied voor ons niet toegankelijk was, zijn de veldnota’s samen met meer recente veldobservaties van Tack en Baudet (2014), van groot belang. Deze observaties, samen met een macro- en microscopisch onderzoek van de gesteenten, laten ons toe drie types van gesteenten, allemaal in zekere mate vervormd, te onderscheiden: 1) gesteenten met een felsische magmatische protoliet; 2) gesteenten met een sedimentaire protoliet en 3) gesteenten met een mafische magmatische protoliet. Op basis hiervan werd een kleur toegekend aan elke type gesteente. Dit liet ons toe de gesteenten weer te geven op een lithologische kaart.

Microscopisch omvatten de drie gesteentetypes de volgende eigenschappen:

1) Gesteenten met een felsische magmatische protoliet worden gekenmerkt door een blastoporfyritische textuur. De porfyroclasten bestaan uit perthitische alkali veldspaat, kwarts en soms plagioklaas. De omgevende, meer fijnkorrelige, grondmassa is opgebouwd uit dezelfde mineralen. Bovendien bevatten alle gesteenten ook mica’s. Deze mica’s omvatten altijd muskoviet/sericiet en soms ook biotiet. Accesorische mineralen zijn chloriet, epidoot, opake mineralen, titaniet en allaniet. In enkele gevallen werd ook calciet waargenomen.

2) Gesteenten met een sedimentaire protoliet bestaan hoofdzakelijk uit kwarts en muskoviet/sericiet. In enkele slijpplaatjes vertoont kwarts heel onregelmatige kristalranden, beschreven als een seriële-interlobate textuur. In de meeste gesteenten treft men echter kwarts aan met een seriële-polygonale textuur. Accesorische mineralen zijn epidoot, opake mineralen, chloriet, zoisiet, biotiet en granaat.

3) Binnen de groep van gesteenten met een mafische magmatische protoliet, treft men verscheidene texturen aan. De mineralogische samenstelling van de gesteenten is echter zeer gelijkaardig. Actinoliet, epidoot en plagioklaas zijn het meest voorkomend. De hoeveelheid

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biotiet, calciet en kwarts varieert van monster tot monster. Accesorische mineralen zijn titaniet, allaniet, opake mineralen en muskoviet/sericiet.

De mineraalassemblages van de drie gesteentegroepen tonen aan dat de gesteenten in de regio rond Matadi beïnvloed zijn door regionaal greenschist metamorfisme. Dit metamorfisme is waarschijnlijk gerelateerd aan de Pan Afrikaanse orogenese. Deze orogenese ging gepaard met vervorming, waarvan we aanwijzingen terugvinden in de gesteenten. Alle gesteenten zijn ductiel (“ductile”) vervormd, wat onder meer aangetoond wordt door “kinked” mica’s en deformatietweelingen. De gesteenten met een felsiche magmatische protoliet vertonen vaak gebroken porphyroclasten die wijzen op broze (“brittle”) deformatie. Verder treft men in deze groep van gesteenten ook texturen aan die wijzen op herkristallisatie. Zo vindt men namelijk porphyroclasten weer die omgeven zijn door een gerekristalliseerde rand die hoofdzakelijk bestaat uit kwarts en alkali veldspaat. Gerekristalliseerd kwarts vertoont soms onregelmatige randen, wat wijst op dynamische herkristallisatie. De kwartskristallen vertonen echter meestal en polygonale textuur, was duidt op statische herkristallisatie. Om deze texturen te verklaren, nemen we aan dat de tectono-metamorfe overprint van de drie gesteentegroepen het resultaat is van vervorming onder lage druk, die gepaard ging met aanhoudend hoge temperaturen.

Tijdens deze studie werd de focus vooral op de gesteenten met een felsische magmatische protoliet gelegd. Veldobservaties hebben aangetoond dat deze gesteenten intrusief zijn in de metakwartsieten van de Matadi Formatie. Hun blastoporfyritische textuur suggereert dat deze gesteenten oorspronkelijk, en dus vόόr de vervorming, porfyritisch waren. Deze textuur toont aan dat de smelt, waaruit het gesteente zich vormde, eerst een trage afkoeling ondervond, gevolgd door een fase met snelle afkoeling. Dit wijst erop dat de gesteenten zich intrusief vormden op intermediaire diepte. Daarom kunnen ze best beschreven worden als hypabyssale gesteenten. Enkele van de gesteenten vertonen echter extreem veel kwarts, waardoor we aannemen dat assimilatie van de omgevende metakwartsieten plaatsvond.

Om na te gaan of deze hypabyssale gesteenten gerelateerd zijn aan de Noqui graniet en/of de Mpozo syeniet, werd een geochemische en geochronologische studie uitgevoerd.

De geochemische samenstelling van de hypabyssale gesteenten is gelijkaardig aan die van de Noqui graniet, maar verschilt van de samenstelling van de Mpozo syeniet. Binnen de groep van de hypabyssale gesteenten zijn er een aantal gesteenten met zeer hoge SiO2-waarden. In vergelijking met de andere hypabyssale gesteenten, vertonen deze monsters ook abnormale waarden voor een reeks sporenelementen (Y, Zr, Nb, La, Ce, Pr, Nd, Sm, Gd, Dy, Ho, Er, Yb, Lu, Hf, Ta, W, Pb, Th and U). Deze abnormale waarden zijn echter wel gelijkaardig aan de waarden geobserveerd in de Noqui graniet. Daarom veronderstellen we dat alkalisch metasomatisme, ten gevolge van de Noqui graniet, de samenstelling van de gesteenten heeft gewijzigd. Dit fenomeen kan men verklaren door de mogelijke ondergrondse verdere verbreiding (naar het noorden) van de koepelvormige Noqui graniet.

De diagrammen van Pearce et al. (1984) en Whalen et al. (1987) tonen aan dat de hypabyssale gesteenten en de Noqui graniet A-type granieten zijn die zich vormden in een intraplaat setting. De diagrammen van Eby (1992) laten toe om een verder onderscheid te maken tussen A1- en A2-type granieten. De gesteenten plotten echter op de grens tussen beide domeinen, wat verklaard kan worden door crustale contaminatie (Eby, 2011). Het proces van crustale contaminatie wordt

139 bevestigd door de Yb/Ta vs. Y/N en de Ce/Nb vs. Y/Nb diagrammen (Eby, 1992). Op basis van de modellen van Eby (1992; 2011) veronderstellen we dat de hypabyssale gesteenten en de Noqui graniet afgeleid zijn van een OIB-type bron. Deze bron werd gewijzigd ten gevolge van crustale contaminatie en vormde een geëvolueerde smelt. Op basis van de Masuda Coryell diagrammen en de spidergrams weet men dat deze smelt vervolgens veldspaten, apatiet en Fe-Ti oxides fractioneerde en zo aanleiding gaf tot de huidige gesteenten.

De geochemische samenstelling van de hypabyssale gesteenten en de Noqui graniet is verschillend van die van de Mpozo syeniet. In de diagrammen van Pearce et al. (1984) plot de syeniet in het domein van vulkanische eilandboog granieten. Bovendien tonen de Masuda Coryell en de spidergrams aan dat er enkel apatiet en een kleine hoeveelheid veldspaat gefractioneerd werd tijdens de vorming van het gesteente.

De veronderstellingen gebaseerd op de geochemische data worden bevestigd door de geochronologsiche data. “LA-ICP-MS zircon U-Pb dating” leverde nieuwe vormingsouderdommen op voor de felische hypabyssale gesteenten, de Noqui graniet en de Mpozo syeniet. De vormingsouderdom van de hypabyssale gesteenten werd vastgelegd op 1043 ± 25 Ma, wat overeenkomt met de eerder voorgestelde, weinig precieze ouderdom van ca. 1050 Ma (Delhal and Ledent, 1978). Voor de Noqui graniet werd een nieuwe vormingsouderdom bekomen van 1018 ± 19 Ma. Rekening houdend met deze foutenmarge, komt deze ouderdom overeen met de “U-Pb zircon SHRIMP” ouderdom van 999 ± 7 Ma (Tack et al., 2001). Deze gegevens bevestigen dat de felsische hypabyssale gesteenten comagmatisch zijn met de Noqui graniet, en dat ze zich ca. 1,0 Ga geleden vormden.

Zowel het wit als het roos facies van de Mpozo syeniet geven eenzelfde vormingsouderdom weer van respectievelijk 1948 ± 10 Ma en 1947 ± 30 Ma. Deze ca. 2,0 Ga ouderdom, toont aan dat de Mpozo syeniet niet gerelateerd is aan de ca. 1,0 Ga oude Noqui graniet en de bijhorende hypabyssale gesteenten. Bovendien wijst de ouderdom van de Mpozo syeniet erop dat de vorming ervan gebeurde tijdens een late fase van de migmatitisatie van de ca. 2,1 Ga oude Kimeziaanse sokkel.

Deze gegevens, samen met informatie van voorgaande studies, laten ons toe de geologische evolutie van de Matadi regio chronologisch weer te geven aan de hand van vier “main geological events” (MGEs), hieronder besproken van oud naar jong.

In de Matadi regio bestaat de sokkel uit de Paleoproterozoische Kimeza Supergroup. Deze afzetting werd gemigmatitiseerd (2088 Ma; Delhal en Ledent, 1976) tijdens de Tadiliaan orogenese (= Eburniaan-Transamazoniaan orogenese). Omstreeks 1947 – 1948 Ma werd de Mpozo syeniet – op basis van ons onderzoek beter beschreven als Mpozo syeno-monzoniet – geïntrudeerd in deze sokkel (= MGE 4).

Dit werd gevolgd door de afzetting van de Matadi Formatie en het, lokaal ontsloten, Yelala conglomeraat. De ca. 1,0 Ga oude peralkalische Noqui graniet werd geïntrudeerd in de metakwartsieten van de Matadi Formatie (= MGE 3). Ten gevolge van deze intrusie ontstond een brede, lichtjes hellende, koepelvormige structuur in de Matadi regio. Dit suggereert een ondergrondse verdere verbreiding van de Noqui graniet op beperkte diepte onder de stad Matadi en langs de noordelijke (rechter)oever van de Congostroom. Samen met de Noqui graniet werden ook de hypabyssale gesteenten geïntrudeerd als sills en dykes.

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Hierna ontstonden mafische intrusies in de Matadi Formatie, die vermoedelijk de aanvoerpijpen vormden voor de bovenliggende metabasalten van de Gangila Formatie. De plaatsen waar deze mafische intrusies zich vormden, werden vermoedelijk gecontroleerd door reactivatie van zwaktezones waarin eerder de felsische hypabyssale gesteenten werden geïntrudeerd.

Vervolgens herkristalliseerden alle gesteenten onder invloed van regionaal greenschist facies metamorfisme (= MGE 2), gerelateerd aan de Pan Afrikaanse orogenese. Ten gevolge van deze orogenese hebben alle gesteenten in het gebied een tectono-metamorfe overprint met variabele intensiteit.

Tenslotte vormden zich in de Matadi regio N-Z (tot NNW-ZZO en/of NNO-ZZW) gerichte “shear zones” (= MGE 1). Deze “shear zones” vertonen steile hellingen naar het westen en komen voor in alle geologische eenheden in het gebied.

Onze studie heeft een belangrijke bijdrage geleverd aan de kennis omtrent de geologie van de Matadi regio. Op basis van veldobservaties, macro- en microscopische beschrijvingen, weet men dat de “felsische magmatische lichamen” in het gebied intrusief zijn en best beschreven worden als hypabyssale gesteenten. Samen met de mafische intrusies, zijn deze gesteenten geïntrudeerd in de Matadi Formatie. Bovendien hebben geochemische en geochronologische data aangetoond dat deze hypabyssale gesteenten gerelateerd zijn aan de Noqui graniet. In tegenstelling tot enkele vroegere hypotheses, weet men nu dat de Noqui graniet en de Mpozo syeniet niet aan elkaar gerelateerd zijn aangezien er tussen hun vorming een tijdsspanne van ca. 1,0 Ga loopt. Bovendien heeft deze studie aangetoond dat de geologische kaart van de Matadi regio bijgewerkt moet worden.

We zijn ervan overtuigd dat deze thesis een goed vertrekpunt vormt voor de grondige kartering van de Matadi regio. Om deze kartering echter tot een goed einde te brengen is – buiten modern veldwerk – extra informatie nodig, afkomstig van “remote sensing” en mogelijk ook van geofysische prospectie. Bijkomend onderzoek omtrent de tectono-metamorfe overprint van de gesteenten in de Matadi regio zou eveneens bijdragen tot een betere kennis van de geologie. Tenslotte suggereren we dat isotopen geochemie (bv. Sm-Nd, Rb-Sr) nodig is om de petrogenese van de magmatische gesteenten in het gebied nog beter te bepalen (o.a. de bron van het magmatisme).

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ANNEXES

Annex 1: Panoramic assemblage of photos illustrating the northern right banks of the Congo River near Matadi

Annex 2: Macroscopic descriptions

Annex 3: Microscopic descriptions

Annex 4: Sketched lithological maps

Annex 5: Structural maps

Annex 6: Microscopic images and CL-images of zircons

Annex 7: Geochronological data

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