Master Thesis, Department of Geosciences
Solitary mantle peridotite bodies in Stølsheimen, Central South
Norway
Anders S. Enger
Solitary mantle peridotite bodies in Stølsheimen, Central South Norway
Anders S. Enger
Master Thesis in Geosciences
Discipline: Structural geology and tectonics Department of Geosciences
Faculty of Mathematics and Natural Sciences
University of Oslo
20.01.2016
© "[Anders S. Enger]", 2016
Supervisor: Prof. Torgeir B. Andersen and Prof. Fernando Corfu This work is published digitally through DUO – Digitale Utgivelser ved UiO http://www.duo.uio.no
It is also catalogued in BIBSYS (http://www.bibsys.no/english)
All rights reserved. No part of this publication may be reproduced or transmitted, in any form or by any means, without permission.
Acknowledgment
I would like to thank my supervisor Torgeir B. Andersen for valuable discussions and feedback. I would also like to thank my co-supervisor Fernando Corfu and Johannes Jakob for their help during the 2014 field season. Much gratitude is due to my field partner, Øystein Kjelberg for discussions during the field work and during the writing of this thesis.
I would also like to thank the staff at the Department for Geoscience and specially; Senior Engineer Muriel Marie Laure Erambert for teaching me how to do analyses at the EMP.
Senior Engineer Maarten Aertes for teaching me how to do sample preparations for XRF analyses and also for running my analyses. Senior Engineer Salahalldin Akhavan for making my thin sections. All their help have been greatly appreciated.
The time spent during writing this master thesis would not have been the same without my fellow master students at CEED. Our discussions and breaks have been truly appreciated.
Lastly I would thank my family and friends for their support during the writing of thesis
Abstract
In the last few years the tectonostratigraphic model of the Caledonides has been revised, to account for among others the melange unit discussed in this thesis. This unit, which in the past have been overlooked, consist of deep marine schists, Alpine-type meta-peridotites, soapstone, serpentinites, dertrial serpentinites, gneisses and conglomerates. The classical way of looking at the meta-peridotites are as the lowermost portion of ophiolites, but the rest of the ophiolite stratigraphy, the sheeted dike complex and the pillow basalt, are not seen in this case. A revised model for the meta-peridotites was presented in 2012 by Andersen et.al.
They proposed that the meta-peridotites were remnants of exhumed mantle at the pre Caledonian magma poor passive margin of Baltica formed during hyperextension.
In this study two such Alpine mantle peridotites are studied, The Rauberget and Vetle Rauberget situated in Stølsheimen in Central South Norway sitting structurally below the Jotun Nappe complex and sandwiched in-between the Upper and Lower Bergsdalen Nappe complex. The peridotites are severely serpentinized, which is confirmed by petrography, no primary mineralogy is found. At the borders of the meta-peridotites blackwall alteration zones and talcified rims can be seen. The whole-rock geochemistry shows a high content of Mg and Si, and a low content of Ca and Al. The Mg and Si-content can be explained by the serpentinization process. While the Al-content points in the direction of depleted mantle origin. The trace element content of the peridotites is really low, commonly below the detection limit of the ICP-MS, which implies a dunitic origin of the peridotite. Based on this information any classification of the meta-peridotites original tectonic environment is difficult.
In order to evaluate the hyperextension hypothesis the rocks surrounding the meta- peridotites were compared with rocks seen at the present day passive margin Iberia- Newfoundland and remnants of hyperextension in the Alps and the Pyrenees.
ACKNOWLEDGMENT ... 1
ABSTRACT... 1
1. INTRODUCTION ... 1
1.1 PURPOSE OF THE STUDY ... 1
1.2 PREVIOUS WORK IN THE STUDY AREA ... 1
1.3 ENVIRONMENTS CONTAINING PERIDOTITES... 2
1.3.1 Peridotites at slow/ultra-slow spreading ridges ... 3
1.3.2 Peridotites in fore-arc environments ... 5
1.3.3 Peridotites at transform faults ... 5
1.3.4 Peridotites in the Red Sea ... 6
1.3.5 Alpine mantle peridotite massifs ... 8
High pressure (HP)/ Ultra High Pressure (UHP) massifs ... 8
Intermediate pressure (IP) massifs ... 9
Low Pressure (LP) massifs ... 9
1.3.6 Exposures of solitary Alpine mantle peridotites/serpentinite ... 9
The Pyrenees ... 9
The Alps ... 10
The Caledonides ... 10
1.4 MODELS FOR SOLITARY MANTLE PERIDOTITES IN THE CALEDONIDES ... 11
1.4.1 Ophiolites ... 12
1.4.2 The magma poor passive margin ... 14
The process of hyperextension ... 16
Exhumed mantle rocks ... 21
1.5 METAMORPHISM OF ULTRAMAFIC ROCKS ... 22
1.5.1 Serpentinization ... 22
1.5.2 Ophicarbonate rocks ... 24
1.5.3 Metasomatism ... 24
1.6 REGIONAL GEOLOGY ... 25
1.6.1 Tectonic history ... 25
1.6.2 Geologic setting and tectono-stratigraphy ... 27
1.6.3 The Melange unit ... 30
2. METHODS ... 33
2.1 FIELD METHODS ... 33
2.2 OPTICAL MICROSCOPY ... 34
2.3 SAMPLE PREPARATION FOR XRF ... 34
2.3.1 Crushing ... 34
2.3.2 Glass beads ... 34
2.3.3 Powder pellets ... 35
2.4 BASIC PRINCIPLES OF XRF AND EMP ... 35
2.5 X-RAY FLUORESCENCE (XRF) ANALYSES ... 36
2.6 INDUCTIVELY COUPLED PLASMA MASS SPECTROMETRY (ICP-MS) ... 38
2.7 ELECTRON MICROPROBE (EMP) ... 39
3. RESULTS ... 42
3.1 RAUDBERGET ... 48
3.1.1 Dunite with talc ... 53
3.1.2 Brecciated dunite ... 54
3.1.3 Foliated dunite ... 57
3.1.4 Meta-dunite with veining ... 59
3.1.5 Meta-dunite ... 62
3.2 THE VETLE RAUDBERGET BODY ... 64
3.2.1 Serpentinite conglomerate ... 65
3.2.2 Dunite with talc ... 68
3.2.3 Talc tremolite schist ... 70
3.2.4 Deformed dunite ... 72
3.2.5 Meta-dunite ... 73
3.3 WHOLE-ROCK GEOCHEMISTRY ... 74
3.3.1 Major element geochemistry ... 74
3.3.2 Trace element geochemistry ... 77
3.4 MINERALOGY AND MINERAL CHEMISTRY... 81
3.4.1 Olivine ... 81
3.4.2 Diopside ... 82
3.4.3 Magnetite/Ferrian chromite ... 84
3.4.4 Carbonates ... 85
3.4.5 Chlorite ... 85
3.4.6 Serpentine ... 85
3.4.7 Talc ... 86
3.4.8 Amphiboles ... 86
3.4.9 Brucite ... 87
4. DISCUSSION ... 88
4.1 MINERAL GEOCHEMISTRY ... 88
4.1.1 Clinopyroxene ... 88
4.1.2 Spinels ... 89
4.1.3 Olivine ... 89
4.1.4 Alteration of mineral chemistry due to serpentinisation and hydration ... 91
4.1.5 Rodingitization ... 92
4.1.6 Metamorphism of magnetite and ferian chromite ... 93
4.1.7 Classification based on the minerals from Raudberget and Vetle Raudberget ... 94
4.2 MAJOR AND MINOR ELEMENTS ... 95
4.2.1 How to use major and minor elements in classification ... 95
4.2.2 Mobility of major and minor elements during hydration ... 96
4.2.3 Classification based on major and minor elements from Raudberget and Vetle Raudberget 96 4.3 TRACE ELEMENTS ... 97
4.3.1 How to use trace elements in classification ... 97
4.3.2 Mobility or immobility of trace elements during hydration ... 99
4.3.3 Classification based on trace elements from Stølsheimen ... 100
4.4 CONCLUSIONS FROM THE GEOCHEMISTRY ... 101
4.5 PASSIVE MARGIN ORIGIN OR OPHIOLITIC ORIGIN ... 101
4.5.1 The melange unit ... 102
4.5.2 Passive margins and preserved relics in mountain belts... 102
Iberia Newfoundland margin ...102
The Pyrenees ...103
The Alps ...104
The Appalachians ...105
4.5.3 Comparison ... 105
4.5.4 Conclusion on tectonic envirnoment ... 106
5. CONCLUSION AND WAY FORWARD ... 107
6. REFERENCES ... 108
APPENDIX A: SAMPLES ... 121
APPENDIX B: MAGNETITE AND FERRIAN-CHROMITE ... 122
APPENDIX C: SERPENTINE ... 124
APPENDIX D: CARBONATE ... 129
APPENDIX E: TALC ... 130
APPENDIX F: DIOPSIDE ... 132
APPENDIX G: CHLORITE ... 134
APPENDIX H: OLIVINE ... 136
APPENDIX I: AMPHIBOLE ... 141
APPENDIX J: BRUCITE ... 145
1
1. Introduction
1.1 Purpose of the study
The main goal with this thesis is to study solitary mantle peridotite bodies situated in Stølsheimen between Voss and Sognefjorden. Recent studies suggest that such solitary
“Alpine-type” peridotite may be exhumed, occasionally to the surface during extreme extension in passive continental margins, but the emplacement of such mantle rocks into a dominantly sedimentary rock complex is not well explained. In the present study two such bodies will be mapped, the Raudberget and Vetle Raudberget peridotites. These exotic rocks are associated with deep-basin schists as well as coarser grained siliciclastics including conglomerates and meta-sandstones.
During the present study a detailed map accompanied by sampling and petrography as well as geochemical investigations (X-ray fluorescence, Electron microprobe and Inductively
Coupled Plasma Mass Spectrometry) will be carried out. The data from the studied peridotites will be used to classify, which tectonic environment the peridotites originated by comparing them to data from peridotites at different tectonic environments. The rocks situated next to the peridotites in Stølsheimen will also be discussed with respect to the tectonic environment.
1.2 Previous work in the study area
During the 1980s NGU (Norwegian Geological Survey) had project together with Norwegian Talc A/S in the Raudberget area. The purpose was to see if the talc, sulphide and carbonate deposits were worth mining. The project resulted in several reports by NGU (Bakke, 1985, 1986, Trønnes, 1988, Karlsen, 1990c, b, a, Olerud, 1990) and two master thesis (Torstensen, 1981, Aarflot, 1984). The NGU reports contain information on the meta-peridotites together with drill-core description.
Kvale (1946) and (Fossen, 1993b, a) studied the metamorphism and structural geology of the Bergsdals nappes, which are in close proximity of the Melange unit. Andersen et al., (2012) looked at the meta-peridotites and the country rocks in Stølsheimen when looking for
2 evidence supporting hyperextension at the pre Caledonian margin of Baltica. Fauconnier et al., (2014) studied the peak metamorphic condition of the study area during the Caledonian orogeny.
1.3 Environments containing peridotites
Peridotites consist mainly of olivine, clinopyroxene (cpx) and orthopyroxene (opx) in variable amounts. Based on their composition the peridotite can be divided into harzburgite, dunite, wehrlite and lherzolite (Figure 1).
Figure 1. Scheme for classification off peridotites and pyroxenites, the upper part is for peridotite classification, modified after (Streckeisen, 1974).
Peridotites can be found in different environments (
Table 1), at slow spreading ridges, fore- arc systems, mountain belts, passive margins and transform faults.
3 Table 1. Showing the different environments containing peridotites with examples of localities and the peridotite classification at each environment. The different environments will be further described in the text.
Tectonic environment Some localities/examples Peridotite classification Slow spreading ridges Gakkelridge, SWIR Harzburgite and lherzolite
Fore-arc Mariana Trench, Hahajima
seamount
Harzburgite and dunite
Transform faults Vema and Garret transform fault
Harzburgite and lherzolite
Rifting Red sea Lherzolite
Alpine mantle peridotite Caledonides, the Alps, The Pyrenees
Commonly
lherzolite(Pyrenees, Alps), but also dunitic/harzburgitc (Caledonides)
Ophiolites Oman, Semail Harzburgite and lherzolite
with transformation zone with dunite
In order to evaluate the origin of the meta-peridotites studied in Stølsheimen, an overview of the tectonic environments where peridotites are likely to be found is presented below.
1.3.1 Peridotites at slow/ultra-slow spreading ridges
An ultra-slow spreading ridge is characterized by a low rate of magma production along the ridge and lack of transform faults. Mantle is placed on the seafloor continuously over large areas. The spreading rates are observed to be less than 12mm/year, but the mentioned characteristics are seen at ridges with spreading rates up to 20mm/year (Dick et al.,, 2003).
Serpentinized peridotites have been found at slow/ultra-slow to intermediate spreading ridges (Figure 2) several places and show similar features to Alpine peridotites and peridotites found in the lowermost portion of the ophiolite sequences (Dick et al.,, 1984). There are very few
4 peridotite occurrences at the sea floor in proximity of the fast spreading ridges, the reason for this is the high volcanic activity and the more steady state magmatic upwelling at fast
spreading ridges than at slow (Hekinian, 2014). Peridotites found at or near spreading ridges are referred to as abyssal peridotites. The abyssal peridotites are commonly coarse grained tectonites with porphyroblastic texture. They frequently show strong alteration due to hydration, and serpentinization alters commonly between 20-100% of the peridotites olivine and pyroxene content (Dick, 1989).
Figure 2. Showing the localities of peridotite along mid ocean ridges together with hotspots and their location on land. Notice that the peridotite located on land is named ophiolites in some cases this is thought to not be true see chapter on hyperextension (Hekinian, 2014)
The Gakkel ridge is an ultra-slow spreading ridge with spreading rates between 6-12mm/year.
The ridge extends for 1800 km from the North coast of Greenland to the North coast of Siberia (e.g. Hellebrand et al.,, 2002). The ridge show very deep rift valleys and transform faults are absent. Abyssal peridotites from this ridge have been altered due to serpentinization.
There have also been found cores of spinels that are surrounded by alteration rims. The cores are believed to be remnants of primary minerals (Hellebrand et al., 2002).
5 The Southwest Indian-Ocean ridge (SWIR) is another example of an ultra-slow spreading ridge, with spreading less than 20mm/year. The minerals, found in abyssal peridotite from this ridge, are mainly lizardite, chlorite, carbonates and magnetite together with small amounts of talc, pyroxene and olivine (Zeng et al.,, 2012).
1.3.2 Peridotites in fore-arc environments
Serpentinized peridotite has been found by dredging and drilling in the fore-arc environment.
The peridotites can be situated in seamounts that have been serpentinized or on the landward edge of the trench (Okamura et al.,, 2006). In the lower Köli nappe, serpentinites occurs above and below meta-volcanic rocks (Trouw, 1973). This observation has been interpreted to indicate that the serpentinite could occur as hydrothermal intrusions in former fore-arc
environments (Grimmer and Greiling, 2012).
The Tonga trench, located in the Southwest Pacific, was the first place where peridotites were observed in the nearshore flank of the trench. The peridotite was observed to be fresh in some places, the minerals observed were olivine, opx, and small amounts of serpentine, the
peridotite was classified as a dunite (Fisher and Engel, 1969).
The Hahajima seamount, located in the fore-arc system of the Izu-Bonin arc, show sepentinized peridotite with harburgitic and dunitic lithologies in association with altered gabbros, dolerites and basalts. The degree of serpentination is extensive, and all the data from this locality show between 80-100% serpentinization. Chromite can be observed as cores with magnetite rims, cpx is seen in small amounts also opx and amphibole are observed (Okamura et al., 2006).
The peridotites, from the fore-arc system, are seen to have some things in common. They are most often seen to have a harburgite-dunitic lithology; they have a high degree of melt extraction and having high Cr# in the spinels (Arai, 1994).
1.3.3 Peridotites at transform faults
The transform fault represent the tectonic active parts of fracture zones, the fault makes a discontinuity and disrupt the linearity of spreading centres, and creates depressions separating two spreading ridge segments (e.g. Hekinian, 2014).
6 Peridotites have been found at transform faults, e.g. the Vema fracture zone in the North Atlantic (Cannat et al.,, 1991) and the Garret transform fault near the East pacific rise (Hébert et al.,, 1983).
The Vema fracture zone is characterized by a 10-20 km transform valley floor filled with sediments. The valley floor is bordered by two steep transform faults. Dredge hauls in the Northern flank of the Sema transverse ridge show the presence of basalts, amphiboles, serpentinized peridotites, dolerites and gabbros (Cannat et al., 1991).
The peridotites found at the Garret transform faults were classified as harzburgite, and showed medium grained nodules of peridotite located in a foliated serpentine matrix. The harzburgite nodules were seen to have diopside replacing enstatite located in a matrix of partly serpentinized olivine. Also Cr-rich spinels were observed (Hébert et al., 1983).
1.3.4 Peridotites in the Red Sea
As seen in the previous sections the peridotite found at the mentioned tectonic environments are fully or partly serpentinized and therefore it is not straight forward using them to get information on the mantle, which these rocks originated. At the Zabargad island in the Red Sea on the other hand some of the peridotites are especially well preserved with very little serpentinization.
The Zabargad Island is thought to be an uplifted part the Red Sea lithosphere. There are three groups of peridotites present at the island; a protogranular spinel lherzolite, amphibole
peridotite and plagioclase peridotite. Small outcrops of dunite and wehrlite have also been observed. Together the three peridotite bodies cover 1,5 km2 of the 5km2 island. Other lithologies on the island consist of metamorphic rocks probably of Precambrian age. Dolerite and basaltic intrusion and dikes in the peridotite and metamorphic units, and several
sedimentary rocks like reefs and beach deposits (See Figure 3 for geologic map) (Bonatti et al.,, 1986).
The mantle peridotites at this island are thought to be the result of the emplacement of mantle peridotite diapir within the Africa-Arabian continental crust during the early stages of rifting in the Red Sea (Nicolas et al.,, 1987). The distribution of three different peridotite phases is
7 evidence for small scale heterogeneities in the mantle together with small scale
metamorphism and metasomatic processes. The fo-number of olivine together with the Mg- number implies that the Red Sea peridotite is similar to the one found in ocean basins. Based on that information two theories on the origin of the peridotite has been proposed; either as ophiolites or as uplifted upper mantle, the first theory is unlikely since none of the classical characteristics of ophiolites can be found at the Zabargad Island. The first peridotite that got uplifted probably was the result of thermal upwelling that is known to happen in rift zones (Bonatti et al.,, 1981). At the time this complex on Zabargad was described (before 1990) the concept of hyperextension was not known and alternative models for exhumation of these rocks have not been suggested later. There has however, been suggested that this rocks have been exhumed in transform submarine fracture zones that is located North of the island. The fracture zone caused differential movement between the central Red Sea rift and the Northern Red Sea lift and caused the exhumation of the peridotite (Marshak et al.,, 1992).
Figure 3. Geologic map over Zabargad Island. The different symbols represents; 1: limestone from young reef, 2: limestone from old reef, 3: conglomerate and breccias, 4: evaporates, 5: Zabargad sedimentary formation, 6: metamorphic group, 7: peridotite 8: Basalt and dolerite intrusion and dikes, 9: nickle mineralization 10: faults. Figure from Bonatti et al. (1981)
It is interesting to note that these peridotites are intimately associated with old gneisses, reef and metamorphic rocks as well as young limestone and conglomerates and sandstones eroded from the exhumed rocks. All these rocks have been mixed together without compressional tectonics in a mountain belt.
8
1.3.5 Alpine mantle peridotite massifs
These peridotites range in size from a few meters to several km. The Ronda massif is an example of an orogenic peridotite massif in Spain that cover approximately 300km2 peridotite (Suen and Frey, 1987). Most Alpine mantle peridotites are not that extensive. Observations show that they are intermediate in size and range from 100 meter to a few km. It is also common to find several Alpine peridotites in the same area. The orogenic peridotite massifs are frequently lherzolites (Figure 1) with composition equilibrated in the plagioclase, spinel or garnet field (Bodinier and Godard, 2003).
Based on mineral assemblage, pressure and temperature (P-T) conditions orogenic peridotite massifs can be divided into three categories (Bodinier and Godard, 2003):
High pressure (HP)/ Ultra High Pressure (UHP) massifs
Peridotites from this group are found in many high pressure terrains in mountain belts e.g. the Western Gneiss Region (WGR) in the Caledonides. The peridotite can be of any of the
classification mentioned (Figure 1) and they commonly are associated with garnet, and therefore garnet peridotites. The garnet peridotites from such terrains show many similarities, they are strongly depleted and cut by veins of garnet pyroxenite and in some cases associated with eclogite. This group can be further divided (Brueckner and Medaris, 2000).
Prograde peridotite: peridotite thought to originate from the supra-subduction mantle wedge.
The peridotites are exhumed to the surface when continental rocks underthrust in the mantle due to continent-continent collision. The continental rocks are more buoyant and then starts to rise. At this point the peridotite will intrude the slab, and get garnet bearing through
subduction and prograde metamorphism (Bodinier and Godard, 2003). An example of this category of peridotites can be found in the Western Alps, the Alpe Arami (Green II et al.,, 2010). Also the peridotites found in the Seve Nappe complex in Sweden have been suggested to be of this origin (e.g. Brueckner et al.,, 2004, Brueckner and Van Roermund, 2007).
Relict peridotite: This category of peridotites consists of old subcontinental lithosphere that shows no evidence of subduction processes. The creation of this group of peridotites is still unclear. The garnet peridotites observed in the WGR have been suggested to be of this origin
9 (Brueckner, 1998). The different assignment of the WGR peridotites shows that their origin is still unclear.
Intermediate pressure (IP) massifs
Peridotites in this category are equilibrated in the spinel field. This category of peridotites commonly show preserved mantle structures and for this reason data from IP massifs are preferred to use for geochemical studies of Alpine peridotites . The mineralogy is also often well preserved since IP peridotites have not been altered by prograde recrystallization in the garnet or retrograde plagioclase peridotite field. Examples of this category of peridotites can for example be found in the Ronda massif in Spain and the lherz peridotite in the Pyrenees (Bodinier and Godard, 2003).
Low Pressure (LP) massifs
Consist of commonly fertile lherzolite that was exhumed during continental rifting and exposed as denudated mantle on the seafloor at passive continental margins (Bodinier and Godard, 2003). This group of peridotite is commonly found in the Western parts of the Alps where they form a belt between the Alpine and the North Appenine arc. Ophicarbonated and oceanic sediments are regularly found juxtaposed with the serpentinized peridotite. The peridotites frequently have strong alteration due to hydrothermal activity, and frequently peridotites are situated in cores surrounded by serpentinite (e.g. Beltrando et al.,, 2014).
1.3.6 Exposures of solitary Alpine mantle peridotites/serpentinite
The Pyrenees
The peridotites in the Pyrenees consist of approximately 40 bodies ranging from a few metres to 3 km across. The peridotites can be found in an
approximately 400 km long, but only a few km wide zone that is located parallel to the
Figure 4. Simplified map showing the locations of alpine mantle peridotites in the Pyrenees mountain belt from Lagabrielle et al. (2010)
10 North Pyrenean fault zone (Fabriès et al.,, 1991) (Fig. 2). Most of the peridotites exposed in this zone are well preserved with little serpentinization. The Lherz peridotite body, which is the type locality for lherzolite show characteristics that imply it is of subcontinental origin (Lagabrielle and Bodinier, 2008)
Based on mineralogy and geothermobareometry the peridotites in the Pyrenees can be divided into the Eastern massifs (EP) and the central Western massifs (CWP). The CWP massifs shows an abundance of coarse grained structures and the peridotites consist of spinel
lherzolites with a large amount of cpx (Fabriès et al.,, 1998). All the massifs in this part of the Pyrenees have been affected by hydrothermal alteration, this alteration result in
serpentinization. The degree of serpentinization varies from one massif to another; some are completely serpentinized while some others are only slightly affected. Cores of relict olivine have been found in mesh textures of serpentine. The opx has commonly been replaced by pseudomorps often by talc and actinolite amphibole (Fabriès et al., 1998).
The EP massifs consist of layered spinel lherzolites with harzburgite layering, the cpx content in these massifs are low. Also spinel websterite can be found. The borders of the massifs are often brecciated, but inside the lherzolite it can be found areas of rather fresh peridotite (Burnham et al.,, 1998).
The Alps
The peridotites seen in the Western parts of the Alps probably had their origin in the Tethys Ocean (Manatschal and Müntener, 2009). They all underwent a great amount of
serpentinization, and thus consist mainly of antigorite and small amounts of magnetite
(Barnes et al.,, 2014). The serpentinites are juxtaposed to different rock types including meta- pillow-basalt, meta-gabbro and meta-sediments (Beltrando et al., 2014). At the borders of the different lithologies reaction rims due to metasomatism can be witnessed, e.g. in the Piedmont unit where talc, tremolite and carbonate can be seen at the borders of the ultramafic bodies.
No relict fresh peridotite is found at this locality, the olivine and the opx has been altered to brucite and antigorite (Beltrando et al.,, 2012).
The Caledonides
The ultramafic rocks found in the Caledonides can be found as cumulates in layered intrusions, Alpine peridotites and as detrial serpentinites (e.g. Qvale and Stigh, 1985). The
11 Alpine-type peridotites are most abundant, and can be seen as; ultramafic associated with ophiolites and solitary Alpine peridotites. The solitary bodies lay in metasedimentary sequences and are lens-shaped. The contact zone is commonly highly deformed. In South Norway, the unit with abundant solitary meta-peridotites have been referred to as a melange (Andersen et al., 2012, Corfu et al.,, 2014)
Based on mineralogy and textures the solitary Alpine mantle peridotites can be divided into three groups; category one can be found in rocks of Lower Paleozoic age. They may have primary olivine, cpx, opx (rare) and chromite, but they are often fully serpentinized. Type two consists of polymetamorphic meta-peridotites. The minerals present are olivine, enstatite, carbonates, talc and amphiboles. The host rock is medium to high grade metamorphic. Type three occurs in gneisses, and the peridotites have a metamorphic mineral assemblage
consisting of olivine, opx and small amounts of chromite. Cpx, amphibole and chlorite can also be present (Moore and Qvale, 1977).
The Alpine mantle peridotites can be seen as massifs with low Al-content with relict metamorphic structures found in harzburgitic, dunitic or as layered bodies thus with a relict cumulate structure (Qvale and Stigh, 1985).
In the Caledonides most of the modern studies on meta-peridotites have been on the high- and ultra-high pressure rocks in the WGR and the Seve Nappe Complex. Less attention has been paid to the melange, and they have mostly been regarded as remnants of ophiolites and not much studied until recently (Beinlich et al.,, 2010) and this work.
1.4 Models for solitary mantle peridotites in the Caledonides
There are two possible models for the peridotites in the Caledonides. The traditional model explains the peridotites as the lowermost part of ophiolites, expect the ones found in WGR (e.g. Qvale and Stigh, 1985). The revised model explains the peridotites as originating at the magma poor passive margin of Baltica during hyperextension before the Caledonian orogeny (Andersen et al., 2012). It is important to notice that the peridotites in question are those that are situated in the melange unit and not the more famous high/ultra-high pressure garnet peridotites found in WGR and in the Seve Nappe. Below follows firstly a description of the
12 term ophiolite, then the magma poor passive margin. The term hyperextension will also be presented in order to understand the revised model.
1.4.1 Ophiolites
According to the Penrose field conference (Anonymous, 1972) ophiolites consist of (from bottom to top) upper mantle peridotites, layered cumulates of ultramafics to gabbros, layered to isotropic gabbros, sheeted dikes, extrusive rocks and a sedimentary cover. The upper mantle rocks are commonly layered with harzburgite and lherzolite and lenses of dunite with chromites.
Ophiolites are considered to represent a part off the upper mantle and the oceanic crust, which after an orogeny is positioned in a mountain belt (Dilek and Furnes, 2011). They are
commonly found in suture zones in mountain belts. Based on their characteristics, ophiolites can be divided into subduction related and subduction unrelated environments (Table 2). The supra-subduction ophiolites form in the extending upper plates of subduction zones. They may evolve in the fore-arc to back-arc tectonic settings. These ophiolites commonly show the Penrose architecture. The volcanic arc ophiolite differs from the previous mentioned by thicker and better evolved arc crust (Dilek and Furnes, 2011).
The subduction unrelated ophiolites can be divided into continental margin, mid-ocean ridge and hotspot ophiolites (see Table 2.). The continental margin ophiolites are characteristic of magma poor passive margins formed in the early stages of the basin evolution. (See section on exhumed mantle rocks)
13 Table 2. Showing the different tectonic environments that have ophiolites. The ophiolites have been divided into subgroups based on the classification from (Dilek and Furnes, 2011). The table also list localities with the mentioned ophiolites together with their geochemical characteristics and
crystallization order of minerals. From Dilek and Furnes (2014).
The mid-ocean ridge ophiolites can form at oceanic spreading ridges in close proximity to plumes like Iceland or at ridges positioned away from plumes. The makeup of ophiolites from this tectonic environment is dependent on the spreading rates of the ridges. At fast spreading ridges, the magmatism follows the speed of the plate separation and the result is an ophiolite that fits the Penrose classification. Ophiolites from slow spreading ridges have a thinner layer of mafic rocks since the magmatism at these ridges are smaller than the fast spreading ridges.
At slow to ultra-slow spreading ridges the ophiolites will show exhumed serpentinized mantle peridotites (Dilek and Furnes, 2014).
The plume ophiolite can form at ridges in close proximity to plumes or at ocean plateaus; they often show thick volcanic and plutonic sequences.
The ultramafic section in an ophiolite that is derived from ridges is of harzburgitic affinity or lherzolitic. The harzburgite can be found in ophiolites from ridges with spreading rates larger than 1 cm/year, while lherzolites are found at very slow spreading ridge ophiolites (Boudier and Nicolas, 1985).
Above the harzburgite or lherzolite section there is a transition zone that consists of dunite with wehrlite, pyroxenite, troctolite layering together with lenses of chromite. The lowermost portion of the transition zone shows an irregular contact with the underlying peridotite and a rapid graduation from peridotite to dunite commonly with the loss of opx. At the top of the
14 transition zone there is no straight forward discrimination between the ultramafic and the mafic section, but in the field the border is often seen as a mappable line (Nicolas and Prinzhofer, 1983).
The origin of the transition zone is still unclear and debated. One theory states that this is cummuls olivine coming from melts rich in MgO. The second theory states that the dunite is mantle peridotite that has been relocated by melts and transformed into its present day origin.
In the Oman ophiolite there has been proposed that both these theories can be correct (Abily and Ceuleneer, 2013).
The peridotites found in the Caledonides have previously been and in some cases still are interpreted to be the lowermost part of ophiolites. The peridotites found North of Otta have recently been described to be of ophiolitc affinity (Nilsson and Roberts, 2014) while the ones found South of Otta in the melange unit have been described as remnants of hyperextension (Andersen et al., 2012, Nilsson and Roberts, 2014). The ophiolite model is however
problematic since the sheeted dike complex, the pillow basalt and the rest of the ophiolite pseudostratigraphy are missing.
1.4.2 The magma poor passive margin
Magma poor passive margins have been suggested to consist of four main domains, created by the rifting processes that take place at such margins. The different domains and their position relative to each other can be seen in
Figure 5. Below follows a short description of each following the terminology of Manatschal and coworkers;
The proximal domain: Consists of the landward part of the margin and is formed in basins, where normal faults can be seen (Manatschal et al.,, 2010). The faults are dipping towards the centre of the basins, resulting in a near symmetrical graben structure in the centre and tilted blocks at the borders (Manatschal, 2004). The faults apparently penetrates down to middle crustal levels (Manatschal et al., 2010).
The necking zone: This zone is defined as the area where the thickness of the crust is reduced to less than 10 km. The zone separates the landward proximal domain and the seaward distal domain (Péron-Pinvidic and Manatschal, 2009).
15 The distal domain: Is the oceanward part of passive margins. It is characterized by
detachment faults reaching down in to the mantle. These, low angle normal faults, are believed to be an important factor in the exhumation of ultramafic mantle rocks. The
exhumation of mantle gives the lithosphere in the distal margins distinctive properties, since its characteristics neither are continental or oceanic (Manatschal et al.,, 2007). The crust in this domain is hyperextended commonly to less than 10 km thickness (Manatschal et al., 2010). In the Iberia-Newfoundland margin the detachment faults are seen to separate continental crust from serpentinized mantle (Pérez-Gussinyé, 2013).
The ocean continent transition zone: It consist of serpentinized mantle rocks together with breccia’s of ophicarbonated rocks as witnessed on the Iberia Newfoundland margin (Manatschal et al., 2010).
The classical locality for description of modern passive margin architecture is the Iberia Newfoundland margin since this margin has been dredged and drilled (Boillot et al.,, 1980) together with geophysical surveys. This gives it the status as the best studied present day passive margin in the world. This is however, not the only magma poor passive margin available. Below follows some other examples;
Brazil-Angola margin e.g. (e.g. Aslanian et al.,, 2009)
North Western Australia margin (e.g. Direen et al.,, 2007)
The Norwegian Sea margin (e.g. Osmundsen and Ebbing, 2008)
South China Sea (Zhou et al.,, 1995, Hayes and Nissen, 2005, Yan et al.,, 2006, Zhou and Yao, 2009).
16 Figure 5. 2D model of margin architecture at magma poor passive margins. Modified after (Péron- Pinvidic and Manatschal, 2009)
The recent recognition of hyperextended passive margins around the world points towards its important role in thinning of continental lithosphere and its important role in the kick-off stage in the Wilson cycle.
The process of hyperextension
A commonly used, but rather simplistic model to describe the thinning of continental crust is the pure shear model (e.g. McKenzie, 1978). This model proposes a rapid and symmetrical stretching of the crust, resulting in thinning and upwelling of hot asthenosphere followed by faulting of blocks and thermal subsidence.
The second model was based on field observations. Wernicke and Burchfield (1982) observed faults that were separating middle and lower crustal metamorphic rocks from upper crustal rocks. This kind of faults implied that extension was dominated by simple shear. These faults were named detachments faults.
Exactly how hyperextension do occur at magma poor passive margins are not well understood, and therefore several models for the process exist. As mentioned above
detachment faults have been imaged in seismic profiles at passive margins (e.g. Osmundsen and Ebbing, 2008) and they are thought to play an important part in the hyperextension process. The detachment faults are believed to be active at low angels and a polyphase rifting model (Figure 6) has been proposed (e.g. Péron‐Pinvidic et al.,, 2007). In this model the crust is first stretched by high angel brittle faults. The second part involves detachment fault
17 forming along the lithosphere, which thins the crust. This may lead to mantle exhumation and crustal rupture (Manatschal, 2004). The model presented reproduce the deformation
sequences that can be witnessed on the Iberia-Newfoundland margin and the Alpine Tethys margin as well (Péron‐Pinvidic et al., 2007).
Figure 6. Simplified 2D model of how hyperextension is thought to happen at magma poor passive margin through a polyphased rifting model and the following exhumation of serpentinized peridotite.
Modified after Mohn et al., (2010)
Another model that aims to explain hyperextension was proposed by Ranero and Pérez- Gussinyé (2010) and Pérez-Gussinyé (2013) and explain hyperextension of faults not being active at low angles. The faults are active at high angles, resulting in rotation of previous faults and they then appear as being active at low angles. This way of describing
hyperextension explains it within an Andersonian framework (normal faults active at high angles 60o and revers faults being active at low angles 30o (Anderson, 1951)). The model has two stages and explains hyperextension as starting in basins;
18 The basin stage: the rifting inside basins often start with a number of unconnected small faults that are dipping both outward and inward. During faulting they may rotate a few degrees; this leading to the faults getting bigger and they connect. The strain now focuses on the faults dipping inwards toward the basin centre and the ones dipping outward become inactive. At the end of this stage, the strain focuses only on one major fault that defines the centre of the basin (
Figure 7a-c.) (Ranero and Pérez-Gussinyé, 2010).
The margin stage: Stress still causes the plates to move apart, thus the crust in the hanging wall of the last active fault in the basin stage behaves brittle. New faults will then start to form in the hanging wall, slipping and further thinning the crust. Successively more faults will be created in the same way (
Figure 7 and Figure 8) (Ranero and Pérez-Gussinyé, 2010). Exhumation of mantle rocks will follow when the crust has been thinned sufficiently (Pérez-Gussinyé, 2013).
The transformation from planar to listric to detachment faults are another characteristic imaged in seismic interpretation at passive margin. This can be explained by the model described above. When crust gets thinned, new faults form at smaller spaces; the inactive faults will start rotating and get a similar geometry of the previous inactive fault at depth. The no longer active faults will rotate to a lower angle and produce the classical listric geometry (Pérez-Gussinyé, 2013).
As mentioned earlier the detachments faults are believed to play an important part in the exhumation of mantle rocks. It has also, however, been proposed that the mantle rocks may be ascending due to their own buoyancy when they become serpentinized (Pérez-Gussinyé, 2013).
In the recent year several places containing serpentinized peridotites have been seen to be former passive margins. Below follows a few examples;
The Caledonides as already mentioned, but also the lherzolites in the Pyrenees has been interpreted as hyperextended exhumed mantle rocks. The exhumation in this case is not the result of opening of an ocean as was the case in the Caledonides, since no samples of oceanic lithosphere can be found at this location. The exhumation of subcontinental mantle happened in pull apart basins (Fig. 9) during the separation of Iberia and European plates (Lagabrielle and Bodinier, 2008, Lagabrielle et al.,, 2010)
19 Figure 7. Detailed model for hyperextension based on angles observed at magma poor passive
margins. The grey part of the model is pre-rift sediments deposited under shallow water level. The red lines mark active faults, while the pink lines mark inactive faults. The thin red lines indicate initial fault geometries before the faults get rotated. At the bottom of the model is a blue line, this line represent moho.
20 Figure 8. Continued from
Figure 7 the blue and green dots marks the position of structures and Moho respectively before and after faulting. The orange shaded areas represent the lower part of the crust.
21 Figure 9. Model of the hyperextension process in the Pyrenees. a: show how and when the pull a part basins formed and their location within the Pyrenees. b: model of the basins and how the
hyperextension happens in pull a part basins. C: profile of one of the hyperextended basins in the Pyrenees. Figure from (Lagabrielle and Bodinier, 2008).
Exhumed mantle rocks
The Iberia Newfoundland margin is the only place where a complete OCT has been drilled at a magma poor passive margin. This margin has also extensively been analysed by geophysical methods, but detailed information on OCT rock types and their relationship are not found from the drilling or the seismic. For this information we need to turn to ancient magma poor margins and look at the remnants of the OCT rocks. The ancient magma poor rifted margins can be found in mountain belts as the Caledonides and the Appalahines (Andersen et al., 2012, Chew and van Staal, 2014), the Alps (e.g. Beltrando et al., 2014) and the Pyrenees (Lagabrielle and Bodinier, 2008, Lagabrielle et al., 2010). The Alpine Tethys ophiolites have been of interest, since they contain OCT rocks of Jurassic Thethyan ophiolites (150 Ma).
It should be mentioned that the Alpine Tethys are not true 3 layered ophiolites (Manatschal and Müntener, 2009), the Alpine ophiolite do not show the characteristics of slow spreading ridges or transform environments, but show characteristics that are commonly seen at the OCT at the magma poor passive margins.
22 In some Alpine ophiolites, detachment faults and low angle normal faults can be observed (Froitzheim and Rubatto, 1998). The low angle normal faults are characterized by
cataclastites and gouges that are often ophicarbonated. The ophicarbonates rocks have been taken as evidence of detachment faults, they are often found in clast supported breccia’s consisting with exhumed serpentinized peridotites. The faults are overlain by tectono- sedimentary breccia’s that continue up in the post-rift sedimentary layers (Manatschal and Müntener, 2009). Below the detachment fault, serpentinized peridotites together with gabbros can be observed. Primary olivine, cpx and opx are rarely found inside the peridotite since the serpentinization process in Alpine peridotites commonly is almost complete. In the hanging wall of the detachment fault, veining with chlorite and serpentine can be observed. The intensity of the brittle deformation increases upwards into a fault core zone of serpentinite gouges. In some cases, there have been found clasts of dolerite inside the fault core, which imply that the detachment fault was followed by magmatic activity (Manatschal and Müntener, 2009).
At the passive continental margin in the Alpine Thethys, the exhumed mantle rocks show compositional differences with respect to their location inside the OCT. The exhumed rocks in close proximity to the continent are serpentinized spinel peridotites with abundant
pyroxene layering. The exhumed mantle rocks located further away from the continent do not show any pyroxene layers at all (Schaltegger et al.,, 2002).
1.5 Metamorphism of ultramafic rocks
Ultramafic rocks can either display the same metamorphic grade as the neighbouring rocks or a different metamorphic grade. Rocks that show the same metamorphic grade as the
surrounding rocks are referred to as isofacial, the rocks that have different metamorphism are in some papers named allofacial rocks. Isofacial ultramafics can contain some primary minerals and/or structures, the allofacial ultramafic more rarely show any primary structures (Bucher and Grapes, 2010).
1.5.1 Serpentinization
Serpentinization occurs in a diversity of tectonic environments and is a result of the hydrothermal alteration of ultramafic and mafic rocks. The serpentinization process is an
23 important process because it alters both the chemical and physical properties of rock e.g. the density and the volume. The density decreases in peridotites from approximately 3,3g/cm3 to approximately 2,65g/cm3 (O'Hanley, 1992), the volume may increases by as much as 40%
(Schroeder et al.,, 2002) due to hydration. For the rocks to become fully serpentinized an incorporation of a water content of 12-14% are needed. The serpentinization also alters the rock’s rheology and weakens them; therefore this process also is important in the creation of detachment faults (Escartin et al.,, 1997). At the ocean floor a magnetic anomaly around peridotites can be observed. This is the result serpentinization because the process produces large amount of magnetite, which increases the magnetic susibility as well as decreases the seismic velocity.
An alteration in the main elements of peridotite is also known to happen, e.g. the Solund basin in South-Western Norway clast of peridotite occurs. The clast can be seen to be altered, serpentinized and ophicarbonated. In the ophicarbonate stage a depletion in MgO occurs (drop from 40Wt.%-2,5Wt%) and an increase in CaO ( from 1wt.% -35Wt.%) (Beinlich et al., 2010). A detailed description on how the serpentinization/hydration process affect the mineral, major, minor and trace elements will be given in the discussion page 88-101.
The following are the basic reactions for the serpentinization and hydration process. For a further discussion of serpentinite reactions see the discussion chapter, where the process will be discussed on the basis of results from the Raudberget and Vetle Raudberget Alpine peridotites.
2 Mg2SiO4 + 3 H2O → Mg3Si2O5(OH)4 + Mg(OH)2 forsterite serpentine brucite 6 MgSiO3 + 3 H2O → Mg3Si2O5(OH)4 + SiO2 Enstatite serpentine silica 3Mg2SiO4 + SiO2aq + H2O→2Mg3Si4O10(OH)2
Forsterite Talc
24 Serpentine minerals cannot incorporate Fe in their structure, therefore Fe is incorporated in magnetite (Winter, 2010).
During the formation of brucite, the Fe-content get removed from the forsterite and
incorporated in magnetite grains. This process may liberate H2 in gas and aqueous form (Sleep et al.,, 2004). Methane expiration is known to happen during low temperature serpentinization ( below 100oC) (Etiope et al.,, 2013).
1.5.2 Ophicarbonate rocks
Carbonates like dolomite, calcite, magnesite etc. are commonly formed by the metamorphic and metasomatic processes in peridotites. The carbonation process is similar to the hydration process; the important difference is that the peridotite minerals react with CO2. The
carbonates are in most cases a secondary mineral and a result of carbonation reactions (e.g.
Bucher and Grapes, 2010).
Serpentine minerals can become carbonated, and when they do they get classified as ophicarbonates. This happens when the serpentine reacts with CO2 from the crust that has been liberated through metasomatism of meta-sediments with carbonates found there.
Serpentine can easily incorporate CO2, and the peridotite becomes ophicarbonated (e.g.
Bucher and Grapes, 2010).
1.5.3 Metasomatism
Due to the strong contrast in composition between the peridotites and the surrounding rocks, metasomatic zones are very common to find along contacts within such rocks. There are commonly sharp chemical gradients in MgO, FeO and SiO2, this are a result of infiltration and diffusion metasomatism taking place. How many and the width of metasomatic zones
witnessed at a peridotite outcrop is controlled by how stable the minerals present are and how easy they reacts with fluids. The processes that control this are the P-T conditions, which rock types that are juxtaposed and the composition and availability of fluids in the area (Winter, 2010).
Already Read (1934) described what has become known as the typical mineral zonation around ultramafic bodies (Figure 9) In modern literature this is often referred to as “blackwall
25 alteration” and is commonly seen at the borders of ultramafic bodies. The blackwall zones often consist of dark and soft biotite. The thickness of the zones varies, but they commonly are some tens of centimetres. Other minerals such as amphiboles and chlorites can also be a part of the blackwall alteration, which minerals present are dependent on the metamorphic processes that take place (Bucher et al.,, 2005). Blackwall alteration zones have been recognized in many parts of the world e.g. in the Valmalenco in Northern Italy where a chlorite blackwall is present at the border of the ultramafic body (Puschnig, 2002), and at the ultramafic bodies in Barramiya Area, Egypt (Takla et al.,, 2003). In the study area blackwall alteration zones are commonly present adjacent to the meta-peridotites (see chapter 3)
Figure 9. Ideal mineral zonation around pod like peridotites, in modern literature these phenomena are named blackwall alteration (Winter, 2010)
1.6 Regional Geology
1.6.1 Tectonic history
Much of Norway’s geology is influenced by the Caledonian orogeny. Before describing the different units in the Caledonides a brief summary of the plate motions (Fig. 10) that caused the orogeny will be given with reference to the reconstructions by Cocks and Torsvik (2006).
26 Figure 10. Plate reconstructions of Baltica and Laurentias position and movement from 550-420 Ma when they collided and formed the Scandian Orogeny modified after (Cocks and Torsvik, 2006) A:
Representing the time after the separation between Baltica and Laurentia B: Representing the time when Avalonia started to rift a part from Gondwana and thus started to close the Tornquist Ocean and starting to open the Rheic Ocean (Cocks and Torsvik, 2006). C: Represent the time when Avalonia collided with Baltica, resulting in the closure of the Tornquist Sea, prior to the main collision D:
Represent the time when Baltica-Avalonia collided with Laurentia, a collision that last from 430Ma to 400Ma (Corfu et al.,, 2006). This collision has been named the Scandian Orogeny (Gee, 1975). The end of the collision was followed by an extensional collapse (Andersen, 1998).
The early part of the rifting that produced the pre Caledonian margin of Baltica is not well- constrained. There are two models for the Baltica-Laurentia configuration prior to the collision, one where Baltica was inverted with respect to Laurentia, and a second model where orthogonal opening and closure of the Iapetus took place (Hartz and Torsvik, 2002).
Reconstruction becomes more reliable from the Cambrian- to early Ordovician when
Avalonia started drifting apart from Gondwana (Fig. 10B). The collision between Baltica and Avalonia are thought to happened in the late Ordovician to early Silurian, at proximately 445Ma (Fig. 10C) (Torsvik and Rehnström, 2003).
27 The continental collision between Baltica-Avalonia and Laurentia started at 430Ma (Fig.
10D) and ended at 400Ma (Corfu et al., 2006). This was the result of the closing of the Iapetus Ocean.
1.6.2 Geologic setting and tectono-stratigraphy
The traditional way to look at the Caledonian geology in Scandinavia is as a four layer model;
the Lower, Middle, Upper and Uppermost Allochthons. The Lower and Middle Allochthons have been interpreted to originate at the margin of Baltica. The Upper Allochthon has been seen as originating in or in close proximity of the Iapetus Ocean. The Uppermost Allochthon consist of rocks from the Laurentian margin (Roberts and Gee, 1985).
This classical model for the Caledonides has recently received criticism for being too rigid and simplistic. Corfu et al. (2014) pointed out several points that were problematic and that a new model should be applied;
In southern Norway the Jotun, Lindås and Dalsfjorden nappe overlay a melange unit with peridotite lenses inside. The two nappes were previously assigned to the Middle Allochthon. The melange unit however, has been seen as a part of the Upper
Allochthon. This points out that there is no simple tectono-stratigraphic model as described in Gee and Sturt (1985) and can no longer be used in its original form.
In the Northern parts of Norway, the Kalak nappe complex has been classified as remains of Baltica, and has been classified as Middle Allochthon. This nappe complex however shows both signs to be of Baltican affinity (Middle Allochthon) and outboard affinity (Upper Allochthon). This makes it difficult to fit in the traditional scheme and questions its validity (Corfu et al.,, 2007); (Kirkland et al.,, 2007).
The traditional model, however, only consider Baltica, Avalonia and Laurentia.
Recent plate reconstruction shows that Baltica may have been inverted and therefore could have faced many tectonic environments before the final collision. It is possible that tectonic slivers foreign to Baltica could have been incorporated into the collision before the final emplacement of nappes.
28
The traditional models also do not take into account the complicated stacking and thrusting events that juxtapose the different units in complicated patterns. This also points out that a simple four event model does not fit.
Finally the 1985 model does not incorporate other models than traditional ophiolites for Alpine peridotites and the concept of mantle exhumation by hyperextension was not developed at this time.
To address these problems of correlation a new model for the Caledonides has been proposed.
This model describes the Caledonides in terms of structure and composition of different units that are not classified based on the traditionally scheme (Table 3) (Corfu et al., 2014).
Figure 11. Map showing the location of the melange unit with solitary meta-peridotites.It can be seen that the meta-peridotites situated close to Gulefjellet and Baldersheim are in an area with largescale anti- and synforms (Andersen unpublished, 2015)
29 Table 3. Showing the classification of the Caledonides with respect to Corfu et al. (2014)new way of describing the units to the left in the table. To the left is Gee and Sturt (1985) interpretation of the same tectonostratigraphic units based on the traditional classification. From Corfu et al. (2014)
30
1.6.3 The Melange unit
The peridotite bearing Melange unit can be traced nearly continuous from the Bergen arcs in Southern Norway and at least as far as Lom in central South Norway (Fig.12), but may continue North-Eastwards towards and beyond Røros (Andersen et al., 2012). More than 80 peridotites with a length of ˃10m have been found in the area between the Bergen Arcs and Stølsheimen (Fig. 11). The unit is placed structurally below the Lindås, Jotun and Dalsfjord nappe in the South-Western parts of the Caledonides in Norway. The peridotites are
metamorphosed, and are many places altered to serpentinite and soapstone. They range in size from a few metres to a few km. Hydration zones and ophicarbonate alteration zones are commonly seen at the borders of the bodies. The peridotites are located in meta-sediments consisting of siliciclastic meta-sediments, phyllites, garnet mica schists, graphic schists and coarser grained meta-sandstones as well as Precambrian gneisses.
The Melange unit was metamorphosed at greenschist to lower amphibolite facies during the Caledonian orogeny (Andersen et al., 2012).
A characteristic rock formed in the melange are detrial serpentinites, which occur in this unit as both conglomerate and sandstone (Qvale and Stigh, 1985). Most commonly they occur as well rounded serpentinite conglomerates. Examples of conglomerates can be found in Reiggarhaugen and at Høyvatnet (Bøe, 1993, Alsaif, 2015). In Reiggarhaugen it is a
serpentinite conglomerate consisting of alternating layers of sandstone and conglomerate. The matrix of the conglomerate consists of serpentine, chromite, magnetite, magnesite, dolomite and talc. The clasts are of red and green serpentinite together with talc and carbonate clasts.
The green serpentinite clasts are commonly bigger than the red clast. They show a zonation consisting of talc rimes with cores of carbonate. The red serpentinite clasts are smaller than the green and also show talcification, but not rims as were the case with the green serpentine clasts. The talc clasts are often well rounded, but have varying shapes; the talc is often fine grained. They are less abundant than the serpentine clasts. The carbonate clasts show varying textures and the amount of carbonates also varies. The mineralogy of the carbonate clasts show magnetite, dolomite, serpentine and talc (Alsaif, 2015). Detrial serpentinite also occur in the Bergen Arcs (Hana) and in Baldersheim (Fig. 11) (Qvale and Stigh, 1985).
31 Figure 12. Tectonostratigraphic map of South Western Norway. The map shows the position of the mélange unit, marked by red stars. The mélange unit is structurally positioned below Jotun and Bergsdals nappe. The mélange is distributed from the Bergen area and is seen up Røros. Figure modified from Andersen et al. (2012). The biggest square marks the area of Fig. 11. While the smaller square marks the study area.
32 The best preserved detrial serpentinites occur in close proximity to Otta. In this place an Ordovician fauna is located (Strand, 1951, Bruton and Harper, 1981) in a serpentinite
conglomerate. The fauna consist of fossils gastropods, brachiopods and trilobites (Bruton and Harper, 1981), which show affinities that are both of Baltican and Laurenitan origin. Such a fauna has been called “Celtic” fauna. It has been suggested that the serpentinite conglomerate was deposited on the Vågemo ophiolite after obduction onto Baltica (Sturt and Ramsay, 1999). This conglomerate represents the same rocks as the meta-peridotites presented in this section, and therefore this is probably not the case (Andersen et al., 2012). The occurrences of the fossils of both Laurentian and Baltican affinity also question this (Bruton and Harper, 1981, Andersen et al., 2012).
Any satisfactory explanation of the occurrence of the melange unit has not yet been given. It has been proposed that it is of ophiolitic origin (Qvale and Stigh, 1985), but no sheeted dike complex has been found in connection with the solitary Alpine mantle peridotites. Gabbro, gneisses and dykes previously assumed to be part of the ophiolites have now been shown to be middle Proterozoic in age (Jacob et.al in prep). These missing parts of the ophiolites question this interpretation. In recent years a new interpretation of the Alpine mantle peridotites has been given. Andersen et al. (2012) suggested that the peridotite bodies were exhumed at the pre Caledonian margin of Baltica during hyperextension. This would better explain the observed lithologies. In this study a detailed mapping, petrography, mineralogical and geochemical study of the main Alpine peridotites in Stølsheimen have been undertaken in order to see if a better interpretation of these little studied rocks can be obtained.
33
2. Methods
The following methods have been used in this thesis: detailed field mapping and sampling, X- ray fluorescence (XRF) for bulk rock major and minor elements, electron microprobe (EMP) for mineral identification and petrographic purposes and optical microscopy for detailed petrography description. All the analyses were carried out at the Department of Geoscience at the University of Oslo. Each method required different sample preparation, a description of the preparations and the principles behind each method will be given below. The sample preparation for XRF was done under the guidance of Senior Engineer Maarten Aerts. The XRF analyses were carried out by Senior Engineer Maarten Aerts. The analysis on the EMP was done under the guidance from Senior Engineer Muriel Marie Laure Erambert.
In addition 5 samples were analysed at the ICP-MS at Actlabs in Canada for trace element composition. Actlabs conducted the sample preparation and also run all the analyses.
2.1 Field methods
Several methods were used during fieldwork. All strike and dip measurements were done with a compass following the right hand rule. Compass was also used when measuring banding, foliation and orientations of veins. For navigational purposes a GPS, aerial photographs and topographical maps were used. The GPS was mainly used to record the different position of the different lithologies and to record where the different outcrops were located. The data stored on the GPS was later used when constructing the geologic maps of Raudberget and Vetle Raudberget.
Sampling was carried out to get data that later would be used for the XRF, the ICP-MS and thin sections. Samples were taken from each of the different litholgies. When collecting samples, the average rock from each lithology was chosen, but also rocks showing special characteristics, like veining and banding.
