Download Geology and petrology of the felsic intrusions in the

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
Co-promotor: Prof. Dr. L. Tack
Leescommissie: Prof. Dr. P. Van den haute, Prof. Dr. M. Fernandez-Alonso
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“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-
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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.
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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.
3.
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
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.
3.3.3.
3.4.
ICP-AES................................................................................................................... 34
Trace elements .............................................................................................................. 35
3.3.3.1.
4.
Araçuaí-West Congo orogen.................................................................................. 15
ICP-MS ................................................................................................................... 35
Geochronology ...................................................................................................................... 37
3.4.1.
Sample preparation ....................................................................................................... 37
3.4.2.
LA-ICP-MS ...................................................................................................................... 38
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
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4.2.2.
4.2.2.1.
Slightly foliated ...................................................................................................... 45
4.2.2.2.
Moderately foliated ............................................................................................... 46
4.2.2.3.
Strongly foliated .................................................................................................... 46
4.2.3.
5.
Mafic magmatic protolith .............................................................................................. 47
4.2.3.1.
Amphibolite ........................................................................................................... 47
4.2.3.2.
Metadolerite .......................................................................................................... 47
4.2.3.3.
Green phyllite ........................................................................................................ 48
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.
5.4.
6.
Sedimentary protolith ................................................................................................... 45
Summary of observations.............................................................................................. 75
Point counting analysis .......................................................................................................... 75
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
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6.3.4.
Grain boundary area reduction and static deformation ............................................... 85
6.3.5.
Core-and-mantle texture vs. porphyroclast systems .................................................... 86
6.4.
7.
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
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.
8.
Discussion .............................................................................................................................. 86
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
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
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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.
10.
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
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
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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 AraçuaíWest 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
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“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.
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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.
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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.
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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.
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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).
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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 StenianTonian (ca. 1Ga), Tonian (930 – 850 Ma) and Cryogenian (750 – 670 Ma) event.
15
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 centralsouthern 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).
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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).
17
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.
19
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. 211), 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.
20
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.
21
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 crosssection 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
22
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.
23
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.
24
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 LowerCongo 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).
25
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.
26
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).
27
Table 2-2: Temporal variations in the definition of the “Palabala Formation”, indicated in yellow. From Behiels (2013)
28
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
29
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-andthrust 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
30
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).
31
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.
32
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 (LiBO 2 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 HNO 3. 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 multielement 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.
34
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.
36
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.
37
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 (Neodymiumdoped 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.
38
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).
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 crossbedding (Fig. 4-2C). In many cases cross-bedding is enhanced by oblique laminae of heavy minerals.
39
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.
40
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.
PC
PC
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.
41
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.
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.
42
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.
43
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.
44
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).
45
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. 411B) 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. 412B) 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.
46
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.
47
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. 415B) 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.
48
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. 51D). 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.
50
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 Kfeldspar 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 Kfeldspar porphyroclast.
52
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.
53
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.
54
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.
55
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.
57
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
58
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.
59
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.
60
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. 59E). 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. 510D), 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.
61
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.
63
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.
64
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.
65
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.
66
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.
68
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. 515C), 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).
69
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.
70
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).
71
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.
72
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 518B), dominated by laths of plagioclase, alternates with a more fine-grained part (Figs. 5-18C and 518D), 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).
73
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.
74
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).
Sample
Quartz
Alkali feldspar
Plagioclase
Muscovite/Sericite
Biotite
Aegirine
Riebeckite
Epidote
Sphene
Allanite
Chlorite
Calcite
Fluorite
Opaque minerals
Total number of counts
Noqui
19223
32.60
61.30
0.50
0.00
0.80
4.20
0.30
0.00
0.00
0.00
0.00
0.00
0.00
0.30
377
Noqui
89993
30.10
61.10
1.70
0.00
3.70
1.70
1.40
0.00
0.00
0.00
0.00
0.00
0.30
0.00
352
Mpozo
19611
12.25
46.68
26.16
0.00
8.94
0.00
0.00
0.99
1.99
0.00
0.66
0.00
0.00
0.33
302
FMP
89540
30.49
32.79
14.43
4.93
10.82
0.00
0.00
2.62
2.30
0.00
0.00
0.00
0.00
1.64
305
FMP
89590
30.10
50.81
2.59
10.36
4.53
0.00
0.00
0.32
0.00
0.65
0.00
0.00
0.00
0.65
309
FMP
89876
46.36
31.13
2.65
14.57
0.00
0.00
0.00
0.33
0.99
1.32
0.00
1.66
0.00
0.99
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.
76
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.
77
Table 6-1: Metamorphic facies and mineral assemblages. From Bucher and Grapes (2011).
78
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.
79
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 nondeformed 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).
80
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.
82
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.
83
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 strainfree 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.
84
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.
85
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.
86
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 straininduced 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 68C 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.
89
Figure 6-8: Lihthological maps based on the observations of: A) Hugé; B) Massar. Rocks with a felsic magmatic protolith =
hypabyssal rocks.
90
Figure 6-8 continued. C) Steenstra; D) Combined observations of Hugé, Massar and Steenstra. Rocks with a felsic
magmatic protolith = hypabyssal rocks.
91
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.
Sample
SiO2
Al2O3
Fe2O3*
MnO
MgO
CaO
Na2O
K2O
TiO2
P2O5
LOI
TOTAL
Noqui
13122
74.15
10.87
5.75
0.10
0.05
0.31
4.48
4.10
0.32
0.01
0.15
100.29
Noqui
71299
72.62
11.06
5.76
0.10
0.05
0.65
3.92
4.59
0.27
<0.01
0.04
99.06
Noqui
89974
74.39
8.99
6.55
0.07
0.05
0.08
3.92
4.19
0.29
<0.01
0.24
98.76
Noqui
89978
74.00
6.44
8.92
0.05
0.06
0.24
4.50
3.60
0.21
0.02
0.34
98.40
Noqui
89991
72.93
9.54
5.84
0.17
0.31
1.77
3.04
4.50
0.25
<0.01
0.23
98.57
Noqui
89992
73.14
9.89
5.60
0.13
0.07
1.35
3.44
4.14
0.29
0.02
0.33
98.40
Mpozo
19504
65.00
17.04
3.88
0.05
0.54
0.86
5.06
6.01
0.55
0.09
0.18
99.26
Mpozo
19611
63.72
16.90
4.43
0.06
0.48
1.37
4.63
6.43
0.56
0.09
0.27
98.96
Mpozo
89453
64.37
17.68
4.52
0.06
0.63
1.10
5.16
6.94
0.57
0.10
0.10
101.24
Mpozo
90067
67.41
15.75
3.03
0.04
0.54
1.32
4.36
5.35
0.46
<0.01
0.32
98.58
Sample
SiO2
Al2O3
Fe2O3*
MnO
MgO
CaO
Na2O
K2O
TiO2
P2O5
LOI
TOTAL
HR
89532
73.26
12.86
3.13
0.06
0.40
0.84
3.44
4.30
0.59
0.11
0.43
99.41
HR
89534
74.96
12.51
2.53
0.06
0.11
0.13
3.16
5.15
0.26
0.02
0.34
99.23
HR
89540
71.86
12.98
3.52
0.08
0.50
0.84
3.60
4.44
0.68
0.15
0.36
99.00
HR
89544
76.31
11.93
1.56
0.01
0.04
0.27
3.71
4.27
0.32
0.01
0.09
98.52
HR
89546
73.30
12.65
5.17
0.10
0.06
0.10
3.09
5.06
0.47
0.02
0.44
100.46
HR
89554
78.56
9.74
4.16
0.05
0.14
0.86
4.51
0.38
0.24
<0.01
0.26
98.90
HR
89863
80.00
8.98
4.43
0.02
0.07
0.18
4.58
0.39
0.22
<0.01
0.04
98.91
HR
89876
77.45
11.20
1.66
0.02
0.11
0.62
2.36
4.82
0.14
<0.01
0.85
99.24
HR
89979
78.21
8.45
5.74
0.03
0.06
0.27
1.36
5.10
0.22
0.01
0.03
99.47
HR
90142
77.17
11.17
2.36
0.03
0.09
0.08
2.69
4.63
0.15
0.00
0.24
98.62
92
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.
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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.
94
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.
Sample
Noqui
13122
Noqui
71299
Noqui
89974
Noqui
89978
Noqui
89991
Noqui
89992
Mpozo
19504
Mpozo
19611
Mpozo
89453
Mpozo
90067
Q
Or
Ab
An
Hy
Di
Ap
Il
Ac
C
Ks
NaS
Mt
total
30.09
24.23
33.10
0.00
6.40
1.31
0.02
0.61
3.85
0.00
0.00
0.10
0.00
99.71
29.00
27.12
31.33
0.00
5.05
2.81
0.02
0.51
1.62
0.00
0.00
0.00
1.12
98.58
35.95
24.76
22.91
0.00
7.94
0.29
0.02
0.55
4.37
0.00
0.00
1.23
0.00
98.02
40.68
21.27
13.08
0.00
10.73
0.94
0.05
0.40
5.96
0.00
0.00
4.24
0.00
97.35
33.05
26.59
24.02
0.00
3.41
7.64
0.02
0.47
1.50
0.00
0.00
0.00
1.20
97.90
33.26
24.47
27.82
0.00
3.23
5.83
0.05
0.55
1.13
0.00
0.00
0.00
1.30
97.64
8.63
35.52
42.82
3.68
4.73
0.00
0.21
1.04
0.00
0.86
0.00
0.00
1.29
98.78
6.97
38.00
39.18
6.21
5.18
0.00
0.21
1.06
0.00
0.05
0.00
0.00
1.48
98.34
2.98
41.01
43.66
4.58
5.53
0.19
0.23
1.08
0.00
0.00
0.00
0.00
1.51
100.77
16.79
31.62
36.89
6.48
3.94
0.00
0.02
0.87
0.00
0.41
0.00
0.00
1.01
98.03
Sample
Q
Or
Ab
An
Hy
Di
Ap
Il
Ac
C
Ks
NaS
Mt
total
HR
89532
33.55
25.41
29.11
3.45
3.52
0.00
0.25
1.12
0.00
1.28
0.00
0.00
1.04
98.73
HR
89534
35.38
30.43
26.74
0.51
2.69
0.00
0.05
0.49
0.00
1.55
0.00
0.00
0.84
98.68
HR
89540
30.52
26.24
30.46
3.19
4.06
0.00
0.35
1.29
0.00
1.08
0.00
0.00
1.17
98.36
HR
89544
37.24
25.23
31.39
1.27
1.28
0.00
0.02
0.61
0.00
0.74
0.00
0.00
0.52
98.3
HR
89546
35.71
29.9
26.15
0.37
0.15
0.00
0.05
0.89
0.00
1.96
0.00
0.00
2.80
97.98
HR
89554
46.94
2.25
38.16
4.20
4.54
0.00
0.02
0.46
0.00
0.37
0.00
0.00
1.39
98.33
HR
89863
49.37
2.3
38.75
0.83
4.64
0.00
0.02
0.42
0.00
0.72
0.00
0.00
1.48
98.53
HR
89876
43.08
28.48
19.97
3.01
1.88
0.00
0.02
0.27
0.00
1.00
0.00
0.00
0.55
98.26
HR
89979
47.46
30.14
11.51
1.27
6.04
0.00
0.02
0.42
0.00
0.23
0.00
0.00
1.91
99.00
HR
90142
42.43
27.36
22.76
0.40
2.58
0.00
0.00
0.28
0.00
1.59
0.00
0.00
0.78
98.18
95
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
3a
3b
4
5
6
6*
7
7*
8
8*
Alkali feldspar granite
Syenogranite
Monzogranite
Granodiorite
Tonalite
Alkali feldspar syenite
Alkali feldspar quartz syenite
Syenite
Quartz syenite
Monzonite
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.
96
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.
Sample
ASI
PI
Noqui
13122
0.88
1.09
Noqui
71299
0.88
1.03
Noqui
89974
0.81
1.22
Noqui
89978
0.55
1.75
Noqui
89991
0.73
1.03
Noqui
89992
0.78
1.03
Mpozo
19504
1.04
0.87
Mpozo
19611
0.99
0.86
Mpozo
89453
0.98
0.91
Mpozo
90067
1.03
0.82
Sample
ASI
PI
HR
89532
1.09
0.80
HR
89534
1.13
0.86
HR
89540
1.06
0.83
HR
89544
1.06
0.90
HR
89546
1.18
0.83
HR
89554
1.04
0.80
HR
89863
1.08
0.89
HR
89876
1.09
0.81
HR
89979
1.02
0.92
HR
90142
1.17
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 yaxis. 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.
97
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 SiO 2 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
98
Figure 7-5: Harker diagrams.
99
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.
100
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.
101
Table 7-4: Trace elements in ppm.
Noqui
Noqui
Noqui
Noqui
Noqui
Noqui
Mpozo
Mpozo
Mpozo
Mpozo
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
sample
102
Table 7-4 continued.
HR
HR
HR
HR
HR
HR
HR
HR
HR
HR
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
sample
103
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.
106
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.
107
7.3.2.2.
Alphabetical classification
In 1974, Chappell and White recognized two distinct types of granitoids which they described as Iand 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.
4
Figure 7-9: Data plotted on a Zr vs. 10 Ga/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 Atype 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.
109
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.
112
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.
113
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.
115
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 domelike 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.
117
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 Vshape 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. 92C) 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
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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
1034 ± 56 Ma
RG76_24
1907 ± 33 Ma
RG76_23
1907 ± 32 Ma
RG23109_59
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 LowerCongo 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 14 th
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 seriateinterlobate 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
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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
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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 ZuidAmerika 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 “foldand-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 Atype 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
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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.
140
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).
141
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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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