Tracing the evolution of crustal-scale, transient permeability in a … · 2017. 5. 5. · Mousny, an attempt is made to frame massive quartz occurrences in the context of the Variscan

Deformation Mechanisms,

Rheology and

Tectonics

Field Trip Guidebook

Tracing the evolution of crustal-scale, transient permeability in a tectonically active, mid-crustal, low-permeability environment by means of quartz veins

Deformation Mechanisms,

Rheology and

Tectonics

Field Trip Guidebook

Tracing the evolution of crustal-scale, transient permeability in a tectonically active, mid-crustal, low-permeability

environment by means of quartz veins

Field trip organised in the framework of the 19th International Conference on Deformation mechanisms, Rheology and Tectonics, held at the KU Leuven (Belgium) from 16 to 18 September 2013

Manuel Sintubin Geodynamics & Geofluids Research Group

KU Leuven Belgium

19- 20 September 2013

This volume was edited by Manuel Sintubin Cover design by Koen Torremans Cover photo: Photo collage of early-orogenic regularly spaced layer-perpendicular veins in the Rursee area, mullions at Dedenborn, ‘boudins’ at Bastogne and late-orogenic discordant veins at Herbeumont – Photo credentials: Manuel Sintubin, Koen Van Noten & Hervé Van Baelen

© 2013 Manuel Sintubin, Geodynamics and Geofluids Research Group, KU Leuven, Celestijnenlaan 200E, B-3001 Leuven – [email protected]

All rights reserved. Except in those cases expressly determined by law, no part of this publication may be multiplied, saved in an automated data file or made public in any way whatsoever without the express prior written consent of the publishers.

Field trip guidebook Deformation Mechanisms, Rheology and Tectonics International Conference, Leuven 2013

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Welcome to the High-Ardenne slate belt …

For more than a decade our research group invested in the study of quartz veins in the High-Ardenne slate belt in the Ardenne-Eifel region. From the beginning we applied an integrated approach, not only focusing on structural, geometrical aspects of the vein occurrences, but also on the petrological, mineralogical, microstructural and geochemical aspects of the vein quartz. In particular, the study of fluid inclusions allowed us to reveal the true nature of the fluids and the pressure-temperature conditions during which the veins formed. At the end, we finally start to put the puzzle of all different vein types in the High-Ardenne slate belt together and reconstruct the story of crustal-scale, transient permeability in this low-permeability environment during the Variscan history, from the first signs of Variscan contraction affecting the Ardenne-Eifel rift basin to the waning stages of the slate belt.

During this field trip we want to illustrate this story by visiting some key outcrops, from the Rursee area in the northeast to the Herbeumont area in the southwest. The research, presented during the field trip, is primarily a compilation of the research that has been performed within the Geodynamics and Geofluids Research Group at the KU Leuven (Belgium), in particular in the framework of the PhD projects of Ilse Kenis (2004), Hervé Van Baelen (2010) and Koen Van Noten (2011), as well as a whole series of master projects (Simon Depoorter, Dominique Jacques, Yvonne Schavemaker, Hervé Van Baelen, Griet Verhaert, Helga Ferket, Ilse Kenis). A part of this research has already been published (see further); other parts are in the process of being published. This research has been performed in close collaboration with primarily the RWTH Aachen (Germany) and the Utrecht University (The Netherlands), which we gratefully acknowledge for years of constructive collaboration.

We welcome you to ‘our’ High-Ardenne slate belt, not only for its great geology and geological heritage, but also for its marvelous landscapes and intriguing Ardennais.

Enjoy the field trip,

The DRT2013 team September 2013

Deformation Mechanisms, Rheology and Tectonics Field trip guidebook International Conference, Leuven 2013

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Your guides for this field trip:

Tom Haerinck, Manuel Sintubin, Koen Torremans, Hervé Van Baelen, Koen Van Noten


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Safety information for field trip participants

The attendees of the field trip must take responsibility for their own and other participants’ safety. In order to ensure the safety of all participants the organisers reserve the right to limit or refuse attendance at the field trip.

Should you be unsure either of the risks involved or your suitability for the field trip, you must seed advice from the organisers. You must declare any disabilities or medical conditions that may affect your ability to safely attend the field trip.

The field trip leaders will give a safety briefing at the start of the field trip.

Inform the field trip leaders if you leave the field trip early.

The field trip leaders are not expected to provide First Aid. Ensure you have adequate supplies for your own needs.

Wear appropriate clothing and footwear for the locality, time of year and as recommended. Anticipate potential changes in weather conditions. The field trip organisers will advise on any personal protective equipment that participants need (e.g. hard hat, high visibility vest, goggles). These MUST be worn if asked for. Failure to do so may lead to your exclusion from the field trip. The organisers do not automatically provide personal protective equipment.

Children and animals will not be allowed on the field trip.

The leader’s decision is final on any matter relating to the field trip.

No list can cover every safety issue. Always be alert for yourself and others around you.

We visit a quarry (stop D) in exploitation. Hard hats compulsory! Please approach exploitation front with caution! Be cautious for falling rocks!

By attending the field trip, you have agreed to these terms without prejudice.


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5

Introduction

The Variscan High-Ardenne slate belt, exposed in Belgium, Germany, Luxemburg and France, can be considered as a natural laboratory to study the coupled hydrotectonic processes in low-grade metamorphic, mid-crustal, siliciclastic low-permeability environments (figure O1). Quartz veins, ubiquitously present in the slate belt, serve as a proxy to unravel the coupled fluid-pressure and stress-state evolution in these mid-crustal environments, at the base of the seismogenic crust. Both with respect to ore genesis and seismogenesis lessons can be learned.

Along the Rursee (Germany), located in the peripheral parts of the High-Ardenne slate belt, different types of quartz veins are exposed that allow to reconstruct the coupled fluid-pressure and stress-state evolution during the compressional tectonic inversion in upper-crustal conditions at the onset of orogeny. At Dedenborn (Germany) the textbook example of mullions will be visited. In the La Roche-Houffalize-Bastogne area (Belgium) attention will be paid to the mixed brittle-plastic deformation and associated quartz veining in the central parts of the High-Ardenne slate belt. At Mousny, an attempt is made to frame massive quartz occurrences in the context of the Variscan orogenic system. At Bastogne the type locality of the term boudin and boudinage will be visited. Finally, in the Herbeumont area we will focus on a late-orogenic, brittle-plastic detachment zone, materialized in particular discordant quartz veins, related to the extensional tectonic inversion period during the late-orogenic extensional destabilization of the slate belt.

Figure O1 – Schematic geological map of the Ardenne-Eifel area (Belgium, Germany, Luxemburg, France) with indication of the main tectonstratigraphical domains (Sintubin 2008). The different stops, visited during the field trip, are indicated. The classical Meuse river cross-section (A-B) is presented in figure O2.


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Programme

Wednesday 18 September 2013 • Transfer from Leuven to Schwammenauel, North Eifel, Germany (2 hour drive) • Stay at Hotel Der Seehof (www.derseehof.com)

Thursday 19 September 2013

• Stop A – walk along Rursee from Schwammenauel to Eschauler, visiting a series of outcrops along the Hubertus Höhe – Schwammenauel section

• Stop B – Dedenborn outcrop • Picnic during the transfer from North Eifel to La Roche-Houffalize-Bastogne area (2 hour

drive) • Stop C – Mousny massive quartz occurrence • Stay at Hotel Melba, Bastogne (best-western-hotel-melba.h-rez.com)

Friday 20 September 2013

• Stop D – Mardasson quarry at Bastogne • Picnic in Bertrix-Herbeumont area (1 hour drive) • Stop E – La Fortelle quarry at Herbeumont • Transfer to Leuven railway station, for connection to Brussels international airport

(Zaventem) and the Brussels South airport shuttle (Charleroi)

Figure O2 – The classical Meuse river cross-section, representing the major tectonostratigraphical domains in the Ardenne-Eifel region (see figure O1 for location). The High-Ardenne slate belt along this cross-section primarily consist of the Lower Palaeozoic Rocroi basement inlier along the backbone of the Ardenne culmination. South of the culmination, Lower Devonian rocks of the Eifel depression are exposed, which itself is limited to the south by the Herbeumont thrust, forming the sole of the Givonnne trust sheet (Sintubin 2008).

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More to read …

Stop A – Rursee, North Eifel

Van Noten, K., Van Baelen, H. & Sintubin, M. 2012. The complexity of 3D stress-state changes during compressional tectonic inversion at the onset of orogeny. In: Faulting, Fracturing, and Igneous Intrusion in the Earth’s crust (edited by Healy, D., Butler, R. W. H., Shipton, Z. K. & Sibson, R. H.). Special Publications 367. Geological Society, London, 51-69. (http://dx.doi.org/10.1144/SP367.5)

Van Noten, K., Muchez, P. & Sintubin, M. 2011. Stress-state evolution of the brittle upper crust during compressional tectonic inversion as defined by successive quartz vein-types (High-Ardenne slate belt, Germany). Journal of the Geological Society, London 168(2), 407-422. (http://dx.doi.org/10.1144/0016-76492010-112)

Van Noten, K. 2011. Stress-State Evolution of the Brittle Upper Crust during the Early Variscan Tectonic Inversion as Defined by Successive Quartz Vein-Types in the High-Ardenne Slate Belt, Germany. Aardkundige Mededelingen 28, 1-241. (https://lirias.kuleuven.be/handle/123456789/306163)

Van Noten, K. & Sintubin, M. 2010. Linear to non-linear relationship between vein spacing and layer thickness in centimetre- to decimetre-scale siliciclastic multilayers from the High-Ardenne slate belt (Belgium, Germany). Journal of Structural Geology 32(3), 377-391. (http://dx.doi.org/10.1016/j.jsg.2010.01.011)

Van Noten, K., Berwouts, I., Muchez, P. & Sintubin, M. 2009. Evidence of pressure fluctuations recorded in crack-seal in low-grade metamorphic siliciclastic metasediments, Late Palaeozoic Rhenohercynian fold-and-thrust belt (Germany). Journal of Geochemical Exploration 101, 106. (http://dx.doi.org/10.1016/j.gexplo.2008.11.040)

Van Noten, K., Hilgers, C., Urai, J. L. & Sintubin, M. 2008. Late burial to early tectonic quartz veins in the periphery of the High-Ardenne slate belt (Rursee, North Eifel, Germany). Geologica Belgica 11(3-4), 179-198. (http://popups.ulg.ac.be/Geol/document.php?id=2485)

Stop B – Dedenborn, North Eifel

Stop C – Les Blancs Cailloux, Mousny, Ardennes

Depoorter, S., Jacques, D., Derez, T., Muchez, P. & Sintubin, M. (in preparation) The Mousny Massive Quartz Occurrence – a late-orogenic dilational jog in the High-Ardenne slate belt (Belgium)?. Geologica Belgica.

Stop D – Mardasson quarry, Bastogne, Ardennes

Sintubin, M. 2008. Photograph of the month: Boudin centennial. Journal of Structural Geology 30(11), 1315-1316. (http://dx.doi.org/10.1016/j.jsg.2008.06.003)

Kenis, I. & Sintubin, M. 2007. About boudins and mullions in the Ardenne-Eifel area (Belgium, Germany). Geologica Belgica 10(1-2), 79-91. (http://popups.ulg.ac.be/Geol/document.php?id=1387)

Kenis, I., Urai, J. L. & Sintubin, M. 2006. The development of bone-shaped structures in initially segmented layers during layer-parallel extension: numerical modelling and parameter sensitivity analysis. Journal of Structural Geology 28(7), 1183-1192. (http://dx.doi.org/10.1016/j.jsg.2006.03.021)

http://dx.doi.org/10.1144/SP367.5�

http://dx.doi.org/10.1144/0016-76492010-112�

https://lirias.kuleuven.be/handle/123456789/306163�

http://dx.doi.org/10.1016/j.jsg.2010.01.011�

http://dx.doi.org/10.1016/j.gexplo.2008.11.040�

http://popups.ulg.ac.be/Geol/document.php?id=2485�





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Kenis, I., Muchez, P., Verhaert, G., Boyce, A. J. & Sintubin, M. 2005. Fluid evolution during burial and Variscan deformation in the Lower Devonian rocks of the High-Ardenne slate belt (Belgium): sources and causes of high-salinity and C-O-H-N fluids. Contribution to Mineralogy and Petrology 150, 102-118. (http://dx.doi.org/10.1007/s00410-005-0008-9)

Kenis, I., Urai, J. L., van der Zee, W., Hilgers, C. & Sintubin, M. 2005. Rheology of fine-grained siliciclastic rocks in the middle crust - evidence from structural and numerical analysis. Earth and Planetary Science Letters 233, 351-360. (http://dx.doi.org/10.1016/j.epsl.2005.02.007)

Kenis, I. 2004. Brittle-Ductile Deformation Behaviour in the Middle Crust as Exemplified by Mullions (Former "Boudins") in the High-Ardenne Slate Belt, Belgium. Aardkundige Mededelingen 14, 1-127.

Kenis, I., Urai, J. L., van der Zee, W. & Sintubin, M. 2004. Mullions in the High-Ardenne Slate Belt (Belgium). Numerical model and Parameter Sensitivity Analysis. Journal of Structural Geology 26(9), 1677-1692. (http://dx.doi.org/10.1016/j.jsg.2004.02.001)

Kenis, I., Sintubin, M., Muchez, P. & Burke, E. A. J. 2002. The "boudinage" question in the High-Ardenne slate belt (Belgium): a combined structural and fluid inclusions approach. Tectonophysics 348, 93-110. (http://dx.doi.org/10.1016/S0040-1951(01)00251-7)

Sintubin, M., Kenis, I., Schroyen, K., Muchez, P. & Burke, E. A. J. 2000. "Boudinage" in the High-Ardenne slate belt (Belgium), reconsidered from the perspective of the "interboudin" veins. Journal of Geochemical Exploration 69-70, 511-516. (http://dx.doi.org/10.1016/S0375-6742(00)00034-0)

Stop E – Carrière de la Fortelle, Ardennes

Schavemaker, Y. A., de Bresser, J. H. P., Van Baelen, H. & Sintubin, M. 2012. Geometry and kinematics of the low-grade metamorphic 'Herbeumont shear zone' in the High-Ardenne slate belt (Belgium). Geologica Belgica 15(3), 126-136. (http://popups.ulg.ac.be/Geol/document.php?id=3600)

Van Baelen, H. 2010. Dynamics of a progressive vein development during the late-orogenic mixed brittle-ductile destabilisation of a slate belt. Examples of the High-Ardenne slate belt (Herbeumont, Belgium). Aardkundige Mededelingen 24, 1-221. (https://lirias.kuleuven.be/handle/123456789/272978)

Van Baelen, H., Berwouts, I., Muchez, P. & Sintubin, M. 2009. Reequilibrated fluid inclusions as proxy for PTX-conditions of deformation and recrystallisation in vein quartz. Journal of Geochemical Exploration 101, 105. (http://dx.doi.org/10.1016/j.gexplo.2008.12.001)

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9

The geology of the Ardennes in a nutshell

The Ardennes is one of the Palaeozoic massifs in Northwest Europe. It is part of the Rhenish massif. These massifs expose rocks belonging to the Rhenohercynian zone of the Variscan orogen (figure O3). The Rhenohercynian zone represents the northern external parts of the Central European Variscides. It forms of a thin-skinned foreland fold-and-thrust belt, telescoping the Devono-Carboniferous rift and platform sediments of the southern passive continental margin of the ‘Old Red Continent’ (Oncken et al. 1999) (figure O4).

The northern part of the Rhenohercynian foreland fold-and-thrust belt, as exposed in the Ardennes (figures O1 & O2), comprises the Ardenne allochthon, which is thrusted over the Brabant parautochthonous foreland during the latest Carboniferous ‘Asturian’ stage (ca. 300 Ma) of the Variscan orogeny (e.g. Mansy & Lacquement 2003, Meilliez et al. 1991). The Brabant parautochthon (Namur synclinorium) represents the deformed Carboniferous foreland basin, covering Middle to Upper Devonian platform sediments and the Lower Palaeozoic Brabant basement. The latter is characterised by an early Palaeozoic orogenic event, the ‘Brabantian’ orogeny (see Sintubin et al. 2009 and references therein). This tectonometamorphic event took place from the late Llandovery to the Eifelian (Debacker 2001, Debacker et al. 2005), coinciding with the onset of the development of the rift basin in the Ardenne-Eifel realm during the early Devonian (Oncken et al. 2000).

Figure O3 – Terrane map of Palaeozoic Europe (Winchester & Team 2002). The Ardennes (AD), part of the Rhenish massif (RM), is situated in the Rhenohercynian zone at the northern front of the Pan-European Variscan orogen. The Ardennes is situated north of the Rheïc suture, thus belonging to the Avalonia terrane.


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The main Variscan thrust, separating the Brabant parautochthon from the Ardenne allochthon, has different names along strike: the Midi fault in the west, the Condroz thrust in the central part, and the Aachen fault in the east. In the frontal area the total displacement is estimated at 20 to 40 km (see e.g. Adams & Vandenberghe 1999).

The Ardenne allochthon consists of the Dinant fold-and-thrust belt (Dinant synclinorium) and the High-Ardenne slate belt (Ardenne anticlinorium & Eifel synclinorium). The Variscan deformation in the Dinant fold-and-thrust belt is controlled by the presence of relatively competent rock formations in the Devono-Carboniferous sedimentary sequence (Lacquement 2001, Meilliez & Mansy 1990). Contrary, the deformation in the High-Ardenne slate belt, predominantly composed of pelitic rocks of early Devonian age, is controlled by this thick, relatively homogeneous pile of incompetent rocks, which eventually led under low-grade metamorphic conditions to the development of a slate belt.

The High-Ardenne slate belt is, furthermore, composed of the Ardenne culmination (Ardenne anticlinorium), which involves predominantly siliciclastic metasediments of Lochkovian age, and the Eifel depression (Eifel synclinorium), which consists of Lower Devonian siliciclastic metasediments. In the west, these sediments are of Pridolian to Pragian age, while towards the northeast the metasediments progressively young up to the Eifelian. The backbone of the Ardenne culmination is formed by a number of Lower Palaeozoic basement inliers, i.e. the Rocroi-Serpont and Givonne inliers in the west and the Stavelot-Venn inlier in the northeast (figure O1). The predominantly siliciclastic metasediments in these basement inliers are of Cambrian to middle Ordovician age, and possibly reflect an early Palaeozoic rift basin development (Verniers et al. 2002). These basement inliers are seen as footwall short-cuts of crustal ramps (Oncken et al. 2000, Plesch & Oncken 1999) (figure 05). Whether or not the Lower Palaeozoic rocks, exposed in the basement inliers, were affected by a middle Ordovician orogenic event, the so-called ‘Ardennian’ orogeny (Michot 1980), remains to date a matter of debate (see e.g. Delvaux de Fenffe & Laduron 1991, Hugon 1983, Hugon & Le Corre 1979, Lacquement 2001, Le Gall 1992, Van Baelen & Sintubin 2008, Sintubin et al. 2009).

The thick Lower Devonian sequence in the High-Ardenne slate belt reflects the rapid syn-rift basin fill, particularly active during the Pragian (Oncken et al. 2000, Oncken et al. 1999), in the northern part of the developing Rhenohercynian ocean (figure O4). In the central part of this ocean, a rift-drift transition occurred during the Emsian (ca. 400 Ma), marking the onset of the post-rift sedimentation in the Ardenne-Eifel realm. This post-rift sedimentation, characterised by the development of a carbonate platform in a passive margin setting lasted until the end of the Viséan (ca. 325 Ma). More specifically, there were two phases of carbonate platform development, the first in the Eifelian-Frasnian and the second in the Tournaisian-Viséan, interrupted by the return of terrigeneous sedimentation during the Famennian.

Figure O4 – Reconstruction of the southern passive margin of the Old Red Continent (Oncken et al. 2000) during the Middle Devonian. The Rhenohercynian ocean is situated at the eastern extremity of the section (green: oceanic crust). The early Devonian rift basin fill (brown) is covered by the Middle Devonian carbonate platform (blue). Note the total length of the section, 349 km.


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During the early Viséan (ca. 335 Ma) the Rhenohercynian ocean was closed (southern Rhenish massif). At the end of the Viséan, basin closure and the northerly prograding Variscan deformation front reached the southern parts of the Ardenne-Eifel realm. During the Carboniferous, foreland basins developed ahead of the Variscan orogen. Deformation in the High-Ardenne slate belt took place during the late Viséan – early Namurian ‘Sudetic’ stage (ca. 325 Ma) of the Variscan orogeny (Fielitz & Mansy 1999, Piqué et al. 1984). Variscan deformation eventually ended in the Stephanian (Kasimovian) ‘Asturian’ stage (ca. 300 Ma) when the Dinant fold-and-thrust belt and the Brabant parautochthon were incorporated in the Variscan orogen (figures O3 & O5).

Figure O5 – Balanced cross-section of the Rhenohercynian foreland fold-and-thrust belt (Oncken et al. 2000), bounded to the southeast by the upper plate (Saxothuringian zone) and the Rhenohercynian suture (Phyllite zone). The High-Ardenne slate belt comprises the Stavelot-Venn unit, the Eifel nappe and the Eifel depression. Northwest of the Stavelot-Venn unit the northeastern extension of the Dinant fold-and-thrust belt is exposed. Note the total length of the section, 167 km.


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STOP A – North Eifel (Germany) – Rursee: “A unique insight into the early Variscan tectonic inversion”

by Koen Van Noten & Manuel Sintubin

INTRODUCTION

The outcrops along the Rursee offers an opportunity to reconstruct the stress-state evolution within a brittle upper-crustal environment during the early Variscan positive tectonic inversion (i.e. from an extensional to a compressional Andersonian stress state). This tectonic inversion is expressed by bedding-normal and bedding-parallel quartz veins.

Logistics – Terrain

Three outcrops will be visited during a walk along the Rursee from Schwammenauel to Eschaeler: • Schwammenauel South – 50°38’02”N and 6°26’31”E • Schwammenauel North – 50°38’46”N and 6°26’13”E • Hubertus Höhe – 50°38’60”N and 6°25’31”E


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NORTH EIFEL GEOLOGY

The Rursee is located in the German part of the High-Ardenne slate belt (HASB). The Lower Devonian metasedimentary rocks rest unconformably on the Lower Palaeozoic rocks of the Stavelot-Venn basement inlier (figure A1). The thick Lower Devonian sequences reflect the rapid synrift Ardenne-Eifel basin fill on the northern passive continental margin of the Rhenohercynian ocean (figure O4), particularly active during the Pragian (Oncken et al. 1999). Sedimentation in the Ardenne-Eifel basin is diachronous, younging from the central part of the HASB towards the periphery, evolving from a Lochkovian-Pragian to a Pragian-Emsian age respectively.

The sequences in the Rursee area belong to the Upper Rurberg and Heimbach beds (Ribbert, 1992). The Upper Rurberg beds, late Pragian in age, consist of an alternation of pre-dominantly mudstones with silt- and fine-grained sanstones intercalated with thick coarse-grained sand-stones. Towards the

east, the Upper Rurberg beds gradually change into the Heimbach beds, late Pragian to early Emsian in age. These beds are characterised by grey-blue mud-stones intercalated with thin sandstone beds. In the field it is imposible to identify the litho-stratigraphic border between both units. The alternation of predominantly clayey-silty sequences and decimeter- to meter-thick, fine- to

coarse-grained, irregular sandstone beds, reflect an energetic depositional

environment. Deposition occurred in the shallow marine, deltaic to tidal environment of rapidly subsiding Ardenne-Eifel rift basin.

Figure A1 – Lithostratigraphic map of the North Eifel, SE of the Lower Palaeozoic Stavelot-Venn basement inlier (modified after Ribbert, 1992). MSZ: Monshau Shear Zone. Area of “mullion occurrence” is indicated (see Stop B – Dedenborn).


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Variscan basin inversion, eventually resulting in the HASB, is materialized in a pop-up structure in the core of the HASB, located between the frontal thrust nappes, associated with the Lower Palaeozoic Stavelot-Venn basement inlier (e.g. Eifel thrust, Venn thrust, Eupen thrust), and the diffuse

Malsbenden-Trois Vierges backthrust (figure A2). The overall

structural grain in the North Eifel area is NE-SW trending. Folds show a NW-verging asymmetry, are highly cylindrical and show close interlimb angles and rounded to angular fold hinges (Van Noten et al. 2008). An axial-planar cleavage is expressed in more incompetent material (clay- and siltstones) as a pervasive slaty cleavage, while in the more competent sandstones only a weak, spaced cleavage is apparent. Bed-on-bed movements, related to flexural-slip folding, is common. Minor out-of-syncline thrusts are often associated with centimetre-thick quartz veins, marked with slickenlines, presumably tracking out-of-syncline movements.

Two successive sets of quartz veins are typical for the North Eifel case: bedding-normal veins (BNVs), mostly confined to competent beds and regionally consistent in orientation, succeeded and crosscut by bedding-parallel veins (BPVs) at the interface of beds with contrasting lithology.

The North Eifel case is representative of the peripheral HASB (pHASB), defined in the field as the structural level in which no mullion development occurred (figure A1). Metamorphic conditions in the pHASB corresponds to the anchizone, with temperatures not exceeding 300°C (Fielitz & Mansy 1999). This very low-grade metamorphism is considered to be of burial origin, pre- to synkinematic with respect to the Variscan deformation. Published vitrinite reflectance for the Upper Rurberg unit, ranging from 4.5 to 5.5% Rmax, indicate a maximum burial paleotemperature of ca. 235 to 250°C (Fielitz & Mansy 1999; Oncken et al. 1999). Maximum depth of burial of the metasediments studied is ca. 7 km, of which the lower 5 km consists of the Lower Devonian metasediments. At the base of this 5 km thick accumulation the studied upper Pragian to lower Emsian metasediments can be found, still covering ca. 3 km of Lochkovian to Pragian metasediments. A geothermal gradient of ca. 35°C/km can thus be inferred for the rocks studied.

Figure A2 – NW-SE cross section of the North Eifel region (after Oncken et al. 1999), locating the Rursee area in a pop-up structure between the Variscan frontal thrust, cutting through the Lower Palaeozoic Stavelot-Venn basement inlier (e.g. Eifel thrust, Venn thrust), and the Malsbenden-Trois Vierges backthrust.


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HUBERTUS HÖHE – SCHWAMMENAUEL SECTION

The outcrops along the NE-SW trending section comprise a series of NW-verging, overturned, hectometer-scale, second-order folds with an associated SE-dipping, axial-planar cleavage with a mean attitude of 129/53. Along the section, a succession can be observed of steeply SE-dipping, overturned limbs and weakly SE-dipping, normal limbs, primarily based on the cleavage-refraction patterns. In the central part of the section (zone 21; see figures A3 & A4) outcrops suddenly comprise steep, almost upright, buckled, N-dipping layers, younging to the north. This bedding attitude is found nowhere else in the Rursee area. The NW-verging fold train, as observed at Hubertus Höhe (zones 17-19; see figures A3 & A4) and near the Schwammenauel dam (zones 23-25; see figures A3 & A4), are interrupted by both NW-dipping and N-dipping, upright layering. The latter observations evidence a box fold with a flat, upright hinge zone (figure A4). Surprisingly, cleavage in these N-dipping layers dips to the SE (118/51), remaining axial planar with respect to the regional fold, suggesting that the box fold development predates the regional fold-and-cleavage development (Van Noten et al. 2008). This conclusion is corroborated by a number of small-scale structures: e.g. the ‘oblique’ cleavage refraction (figure A6), cuspate-lobate folding of the interface of a siltstone bed, buckled bedding-parallel veins (figure A5). The latter observation is interesting because it suggests that box fold development postdates both veining events, but still occurred prior to the main cleavage development and regional folding.

Figure A3 – Structural map of the Rursee area. I-J section, comprising zones 17 to 25, represents the Hubertus Höhe – Schwammenauel section (see figure A4) (after Van Noten et al. 2008).


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Outcrop 1 – Schwammenauel South

The outcrop consists of an alternation of sandstones and siltstones, showing graded bedding and small-scale sedimentary structures. These rocks belong to the upper Pragian to lower Emsian Heimbach beds. Beds are steeply SE-dipping. They form part of an overturned limb (zone 13; figure A3), as corroborated by the cleavage-refraction pattern. Multiple sets of fibrous to blocky bedding-normal veins are present in sandstone beds. They show crosscutting relationships, as well as flexural-slip related deformation between different beds. Also rather continuous bedding-parallel veins are systematically present along the interfaces between the different sandstone and siltstone beds. Both thin, single-opening veins and complex, composite veins can be observed.

Outcrop 2 – Schwammenauel North

The outcrop consists of an alternation of sandstones and siltstones, belonging to the Upper Rurberg beds, Pragian in age. This outcrop is located in zone 21 (figures A3 & A4), exposing the core of the unique box fold. Beds are vertical to steeply N-dipping. Cleavage refraction is ‘oblique’, indicating that the tilting of the beds occurred prior to cleavage development (figure A6). In zone 21 layering has a attitude of 000/83. This particular attitude continues over ca. 100 m to the SE, before bending into gently NE-dipping layers (064/32; zone 22; see figures A3 & A4). The latter represents the normal limb of the fold train, consisting of two hectometre-scale overturned folds, that can be followed towards the Schwammenauel dam (zones 23-25; see figures A3 & A4). Between zones 21 and 22 the fold hinge line, with an attitude of 085/39, can be measured; between zones 21 and 20, a fold axis, with an attitude of 342/30, can be inferred based on the bedding-cleavage intersection lineation. Cleavage, with an attitude of 118/51, remains axial planar with respect to the regional fold train, indicating that the box fold formation did not alter the regional cleavage attitude.

Small-scale observations suggest that tilting of the layers predates buckling of the competent beds and occurred prior or early during the pervasive fold-and-cleavage development (figure A5). Firstly, cleavage does not refract as one would expect in the normal or overturned limbs,

Figure A4 – Schematic representation of the fold geometry of the Hubertus Höhe – Schwammenauel section (see figure A3 for location). Note that the section is not to scale. Total length of section is approximately 2 km (after Van Noten et al. 2008).

Figure A5 – Evolutionary sketch of the N-dipping layers and their particular deformation structures, exposed in zone 21 (see figures A3 & A4 for location). Figure not to scale. (a) initial stage after sedimentation. (b) early tilting of the layers towards the North, with incipient development of the box fold. (c) progressive SE-NW directed Variscan shortening, as represented by a cleavage development, causing buckled layers in the box fold and ‘oblique’ cleavage refraction across the competent beds. Cleavage remains axial-planar with respect to regional fold train. Both veining events, resulting in BNVs and BPVs, predate stage (b) of the box fold development (Van Noten et al. 2008).


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but rather ‘obliquely’ across the competent beds (figure A6). Beds must already have had their particular attitude at the time of incipient cleavage development in the competent sandstone. Secondly, the competent layers in the N-dipping limb show a cuspate-lobate morphology. Such layer buckling in a single limb, is observed nowhere else in the Rursee area, again indicating that this N-dipping attitude is rather exceptional. This buckling can be related to the cleavage development,

expressing the overall tectonic shortening. Cleavage is indeed oriented at high angle to the buckled layers. Hence, at the time of cleavage development, layers must already have been tilted, otherwise they would not start to buckle. Thirdly, a divergent cleavage fanning is present in the proximity of these small-scale buckle folds. Cleavage seems to be squeezed in the cusps, whereas in the lobate parts cleavage diverges towards the cusps, indicating that cleavage development and buckling of layers is closely related. So, buckling and cleavage development are rather contemporaneous and postdate the tilting of the beds.

The beds comprise both bedding-normal and bedding-parallel veins. Two sets of thin bedding-normal veins can be observed. The BNVs are oriented at high angle to the bedding, irrespective of the buckle folds. A consistent vein orientation is obtained after unfolding. Contrary to all other outcrops, in which the BPVs are usually planar, the BPVs, interbedded in pelitic-silty sequences, as well as below and above sandstone beds, show a buckled morphology. Veins are present in both cusps and lobes of the buckled sandstone layer. Cleavage abuts against the vein and seems to be squeezed into the hinges of small-scale folds of vein offshoots. Slickenlines, marking in particular the composite BPVs, oriented mostly at high angle to the bedding-cleavage intersection lineation, are exceptionally oriented parallel to the intersection lineation in the box fold.

All these observations provide a solid time frame for both veining events. It is obvious that the development of both bedding-normal and bedding-parallel veins occur very early in the deformation history. The predate the tilting of the beds in the box fold, which itself predates the cleavage development, associated with the development of the regional fold train.

Outcrop 3 – Hubertus Höhe

The outcrop consists of an alternation of coarse-grained sandstones and siltstones, belonging to the Upper Rurberg beds. Along the section, the regional fold train can be followed (zones 17-19; see figure A3 & A4). Two sets of bedding-normal veins can be observed, one set parallel to the intersection lineation and one set perpendicular to the intersection lineation. BNVs vary from lensoid, thin-hairline to thick-composite. BNVs can be confined to single layers, but may very well crosscut several beds. En-echelon BNVs internally show curved quartz fibres, both indicating some oblique opening. Such veins, showing a shear component, are though observed only locally. Thin, interbedded and thin, laminated, intrabedded BPVs can be observed.

Figure A6 – Sketches of different cleavage-refraction patterns along the Hubertus Höhe – Schwammenauel section. From left to right: ‘oblique’ refraction in box fold (zone 21; see figures A3 & A4); refraction in normal limb (zones 17, 19, 22 & 25; see figures A3 & A4); refraction in overturned limb (zones 18, 23, 25 & 13; see figures A3 & A4) (after Van Noten et al. 2008).


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BEDDING-NORMAL VEINS

Quartz veins, oriented subperpendicular to perpendicular to bedding, are mostly restricted to competent, fine- to coarse-grained sandstone beds. They remain at high angle to bedding across fold hinges, suggesting they predate folding. Unfolding, moreover, reveals that these originally planar veins were deformed during subsequent flexural-slip folding.

Three generations of bedding-normal veins can be recognized. The oldest generation (VA) consist of very thin, hairline veins, reflecting a single opening. They only occur occasionally in some sandstone beds. The second generation (VB) is widespread and is very consistent in orientation (after unfolding) (figure A7). These veins are organized in regularly spaced vein arrays. They are mostly composite veins, resulting from successive fracturing and sealing events. The last generation (VC) is less consistent in orientation, crosscut or abuts against the second generation veins. Vein thickness ranges from millimeter-thin, hairline, single-opening veins to centimeter-thick, composite veins with pelitic host-rock inclusion bands parallel to the vein wall. Bedding-normal veins are not affected by grain-scale heterogeneity of the host rock, as evidenced by transgranular microveins with low tortuosity, which indicates that veining occurred in an already low-porosity rock at maximum burial conditions.

Microscopically, both syntaxial and stretched crystal growth morphologies have been observed. Crystals grew epitaxially on cracked grains from the host rock. Vein-filling crystals are invariably quartz, commonly occurring as stretched crystals evolving to elongate-blocky crystals, indicating growth competition In some veins even a blocky crystal morphology is found in the center of the vein, suggesting that opening rate exceeds growth rate. In the veins, host-rock fragments, reflecting episodic opening by crack-sealing, are organized both parallel and perpendicular to the vein wall. The former solid inclusion bands crosscut crystal boundaries, while the latter solid inclusion trails are confined to a single stretched crystal, tracking the vein opening. Some crystal boundaries have stylolitic appearance with concentration of insoluble particles, suggesting that dissolution occurred along crystal boundaries. Crystals are furthermore affected by low-temperature crystal plastic deformation, as exemplified by undulose extinction, deformation lamellae and subgrain boundaries.

Figure A7 – Lower-hemisphere equal-area projection showing the orientation distribution of the bedding-normal veins in North Eifel (after Van Noten et al. 2011). On the left, the actual orientation distribution, measured in the field. On the right, the original prefolding orientation distribution, after unfolding. With indication of the three generations BNVs.


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The original growth fabric of vein-filling is, however, largely unaffected. Besides quartz, chlorite occasionally occurs along vein walls. Carbonates are very rarely observed.

Primary fluid inclusions, occurring in growth zones, are absent. Pseudosecondary inclusion trails, trapped in crystals during growth, are oriented parallel to the vein wall (figure A8). They are remnants of healed microcracks, evidencing growth by incremental crack and seal. As these fluid inclusion planes result from mode I fracturing, the bedding-normal veins represent extension veins, opening in incremental crack-seal steps. Secondary inclusion trails, oriented parallel to the vein wall, corresponds to post-veining transcrystal microcracks (figure A8). Their particular orientation suggests that the stress state affecting the vein quartz has not changed with respect to the stress state in which the fractures opened and the crystals grew.

The presence of dissolution features along crystal boundaries – i.e. parallel to bedding – suggests that the maximum principal stress σ1 is still vertical during bedding-normal veining (σ1 = σV). The extension veins opened in the σ1-σ2 plane perpendicular to σ3, so that the veins are aligned parallel to the intermediate principal stress direction σ2. In their original, prefolding, attitude, a very consistent NNE-SSW to NE-SW orientation becomes apparent for the bedding-normal veins, reflecting the trend of the intermediate principal stress direction (see figure D6 – stop D). The fairly-well organized bedding-normal vein array thus infers an extensional Andersonian stress state, characterized by a vertical maximum principal stress σ1, corresponding to the overburden pressure σV, and two rather well-defined horizontal principal stresses, σH and σh, reflecting a triaxial stress state during bedding-normal veining. This stress state can be associated with the context of the Ardenne-Eifel rift basin. Veining occurred in the latest stages of the burial history of the basin. The bedding-normal vein array is, though, less well organized than the one in the cHASB, suggesting a more anisotropic stress state in the more central, deeper part of the basin.

At room temperature, both pseudosecondary and secondary inclusions consist of two phases, i.e. an aqueous liquid and an aqueous vapor phase. The measured pseudosecondary inclusions have a consistent lensoid shape, with a long axis < 10 µm. The gas volume of the majority of the inclusions is ca. 15%, with the exception of some large inclusions with a gas volume up to 35%. These large bubble sizes are most probably due to leakage. The secondary trails consist of very small spherical inclusions, often less than 5 µm, with a gas volume up to 20%.

Microthermometry shows that both pseudosecondary and secondary fluid inclusions are composed of an aqueous H2O-NaCl fluid with low salinities of 3.5 to 8 eq.wt% NaCl (Van Noten et al. 2011). A Raman analysis indicates the presence of a limited amount of gas in the aqueous inclusions. In a sample of a BNV the Raman analysis indicates 20% CH4 in addition to CO2. In literature, low- to moderate-salinity, aqueous H2O-NaCl fluids are frequently described as Variscan fluids (Muchez et al. 2000; Schroyen & Muchez 2000) (cf. type A1 fluid inclusions of Kenis et al. 2005a; see figure D4 – stop D). The initial low-salinity fluids, which most probably originate from marine seawater,

Figure A8 – Bedding-normal veins open in the σ1-σ2 plane during extension induced by the vertical load of the overburden (σ1 = σV) and an incipient tectonic stress (σT), which controls the regional vein alignment. Quartz fibres track the vein opening perpendicular to σ3 and are (sub-)parallel to the vein wall. They contain intracrystal pseudosecondary and transcrystal secondary fluid inclusion trails (Van Noten et al. 2011).


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increased in salinity to 4-8 eq.wt% NaCl by fluid-rock interaction with the early Devonian sediments during progressive burial and increasing metamorphism. These fluids change into more mature and saline aqueous-gaseous fluids towards the central, epizonal part of the slate belt in conformity with the increasing metamorphic grade (Kenis et al. 2005a) (see figure D4 – stop D).

Homogenization of all measured inclusions occurs to the liquid state. The frequency histogram shows a large variation in homogenization temperature Th between 80 and 350°C (figure A9), without significant differences in Th between pseudosecondary and secondary inclusion trails. The distribution can be classified as an asymmetrical, long-tailed distribution with skewness to higher temperatures. The wide variation in Th can to a large extent be attributed to post-entrapment inclusion re-equilibration during the deformation of the veins. The long-tailed distribution towards higher temperatures (> 180°C) can thus be interpreted as cryptic re-equilibration such as stretching of the inclusion walls. The lower Th values (< 140°C) can be attributed to cryptic re-equilibration such as necking down of an originally large inclusion into smaller inclusions. Because of re-equilibration, a

large range in Th has been measured. The distribution of homogenization temperatures that could reflect the P-T conditions has a frequency peak Th between 140 and 180°C (Van Noten et al. 2011).

BEDDING-PARALLEL VEINS

Bedding-parallel veins crosscut, truncate and offset the BNVs, clearly indicating that the BPVs postdate the BNVs. The veins extent to several tens of meters and their thickness ranges from a few centimeter to 10 cm, so that they have very high aspect ratios. The veins are continuous around the fold hinges without changing thickness. They are refracted by cleavage in the fold limbs. The veins are mostly present at the interface of beds of contrasting lithology. They show an irregular contact with the siltstone but a sharp contact with the sandstone, suggesting detachment from the sandstone during veining.

Macroscopically, the thick veins show a composite internal fabric, consisting of several distinct generations of quartz laminae intercalated with pelitic host-rock seams that vary from thin, millimeter-size slices parallel to the vein wall to brecciated pelitic host-rock fragments. The vein walls are characterized by slickenlines that are uniform in trend on a single lamina, but slightly vary in orientation from lamina to lamina, indicating bedding-parallel slip during quartz precipitation or between times of emplacement of the successive laminae. All microstructures suggest that fracturing and sealing are successive and recurrent features and that the BPVs are associated with bedding-parallel shear prior to folding and cleavage development (Van Noten et al. 2008).

The dominant mineral is quartz, but chlorite and other phyllosilicates occur near host-rock inclusions and fragments. Four types of microstructures are observed in single quartz laminae and are indicative of different types of vein growth and deformation. Crack-seal microstructures are common in this type of quartz veins. Repetitive crack and seal is expressed by fibrous to elongate-blocky quartz crystals with a continuous crystallographic orientation and with host-rock inclusion bands and host-rock inclusion trails. The inclusion bands are oriented parallel to the vein wall. Solid inclusion

Figure A9 – Histogram of homogenization temperatures of pseudosecondary and secondary inclusions for the bedding-normal veins (Van Noten et al. 2011).


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trails, which represent the offset of a host-rock grain within a single quartz crystal, are parallel to the oblique fiber margins and might track the opening of the vein. Curved fibres, on the other hand, are free of solid inclusion trails. This seems to suggest that the opening rate was higher than the rate of crystallization or that continuous opening and crystallization occurred rather discrete crack-seal events. Open-space filling microstructures are most common in the BPVs. Elongate-blocky to blocky quartz crystals are randomly distributed within a single quartz lamina and are contained between thin host-rock seams. Open-space filling suggest that opening rate exceeded growth rate. Laminae of fine-grained quartz is indicative of cataclasis during high strain-rate shear events. In thicker laminae quartz crystals show evidence of crystal-plastic deformation (e.g. undulose extinction, deformation lamellae, bulging), indicative of localized bedding-parallel shearing. Finally, anastomozing, bedding-parallel stylolites are common in BPVs, inferring pressure-dissolution along discrete planes during or after vein emplacement (Van Noten et al. 2011).

Pseudosecondary fluid inclusion trails are common in the fibrous to elongate-blocky crystals, resulting from crack-sealing (figure A10). These trails reflect fluid entrapment during vein precipitation and can thus be used to decipher the P-T conditions during quartz fiber growth. Secondary trails correspond to post-veining transcrystal microcracks. A clear relationship between orientation of secondary trails and vein walls could not be established.

At room temperature, both pseudosecondary and secondary inclusions consist of two phases, i.e. an aqueous liquid and an aqueous vapor phase. The measured pseudosecondary inclusions vary in shape from lenticular to rounded or to triangular, with sizes up to 15 µm and gas volumes of ca. 15%. Similar to the secondary inclusions in the BNVs, the secondary inclusions in the BPVs consist of very small spherical inclusions, often less than 5 µm, with a consistent gas volume of 10%.

Microthermometry shows that both pseudosecondary and secondary fluid inclusions are composed of an aqueous H2O-NaCl fluid with

low salinities of 4 to 7 eq.wt% NaCl (Van Noten et al. 2011). A Raman analysis indicates the presence of a limited amount of gas in the aqueous inclusions. In a sample of a BPV the Raman analysis indicates 27% N2 in addition to CO2. Again, these fluids can be considered as the typical Variscan fluids (cf. type A1 fluid inclusions of Kenis et al. 2005a; see figure D4 – stop D), very similar than the fluids observed in the BNVs.

Homogenization of all measured inclusions occurs to the liquid state. The frequency histogram again shows a large variation in homogenization temperature Th between 80 and 310°C (figure A11), without significant differences in Th between pseudosecondary and secondary inclusion trails. The

Figure A10 – Bedding-parallel veins open in the σ1-σ2 plane during extinction driven by tectonic stress (σT) in a stress field in which the vertical load of the overburden equals the minimum principal stress (σ3 = σV) (Van Noten 2011).

Figure A11 – Histogram of homogenization temperatures of pseudosecondary and secondary inclusions for the bedding-parallel veins (Van Noten et al. 2011).


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distribution can be classified as an asymmetrical distribution with skewness to higher temperatures. The wide variation in Th can again to a large extent be attributed to cryptic re-equilibration. Therefore, the distribution of homogenization temperatures that could reflect the P-T conditions has a frequency peak Th between 110 and 160°C (Van Noten et al. 2011).

VEIN TRAPPING CONTIONS

The obtained homogenization temperatures represent only the minimum temperature of fluid entrapment. To obtain the true trapping temperature and eventually also the trapping pressure during vein emplacement, a pressure correction should be made for the homogenization temperature. Isochores, representing a linear relationship between temperature and internal pressure of an inclusion, were therefore constructed for both type of veins. Two sets of isochores are calculated: the first set with a density range between 0.92 and 0.98 g/cm3, reflect the Th range of the BNVs; the second set with a density range between 0.94 and 0.98 g/cm3, reflect the Th range of the BPVs (Van Noten et al. 2011) (figure A12).

As independent geothermometer, published vitrinite reflectance for the Upper Rurberg unit, ranging from 4.5 to 5.5% Rmax, indicate a maximum burial paleotemperature of ca. 235 to 250°C (Fielitz & Mansy 1999; Oncken et al. 1999), is used. Maximum depth of burial of the metasediments studied is ca. 7 km, inferring a geothermal gradient of ca. 35°C/km.

Veining conditions can now be determined by considering the intersection between the calculated isochores and the temperature range from the independent geothermometer (figure A12). The results show that the maximum trapping pressure of the BNVs is ca. 190 MPa, i.e. a near-lithostatic fluid pressure at the time of fracturing and veining; the maximum trapping pressure of the BPVs is ca. 205 MPa, exceeding the lithostatic fluid pressure at maximum burial depth.

COMPRESSIONAL TECTONIC INVERSION

The BNVs developed as extension veins, as shown by the fibrous vein fill and presence of crack-seal microstructures. At greater depth, extension veins are initiated at high fluid pressures under low-differential stresses (Cosgrove 1997). In the BPVs, crack-seal microstructures and repetitive quartz laminae are also supportive for extension veins corresponding to bedding-normal uplift at supralithostatic fluid pressures. The laminated character of the veins represents multiple phases of

Figure A12 – P-T diagram showing the trapping conditions of the fluids in BNVs and BPVs. Isochores of BNVs are indicated in dashed lines (blue); isochores of BPVs in full lines (orange). Maximum metamorphic temperature of 250°C is indicated by full red line. A geothermal, lithostatic gradient of 34°C/km is indicated in green (Van Noten 2011).


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opening and closing of the fluid-filled void at the time of veining that can possibly be linked to the earthquake-related stress cycling and fluid-pressure cycling. Quartz precipitation from overpressured fluid could occur coseismically, while during post- and interseismic periods the void may collapse in response to a decrease in fluid pressure below lithostatic. This collapse is exemplified by stylolites in the BPVs. The BPVs furthermore show robust evidence of bedding-parallel shearing. These veins can thus be classified as extensional-shear, ‘hybrid’, veins.

The BNVs are thus the result of successive vertical mode I fractures induced by near-lithostatic fluid pressures (ca. 190 MPa) under low differential stresses controlling a regional consistent

alignment of the arrays of BNVs (figure D6 – see stop D). The low differential stresses demonstrate that veining occurred in the latest stages of the basin-related extensional stress regime (figure A13). The basin was at that time affected by a tectonic compressive stress induced by the Variscan orogenic convergence. Precipitation of the BNVs occurred under low-grade metamorphic conditions in low-porosity, competent rocks in a brittle, upper-crustal environment. The BPVs reflect, on the other hand, a bedding-normal uplift and bedding-parallel shear prior to the formation of folds and a cleavage. Extensional and extensional-shear fracturing occurred under supralithostatic fluid pressures (ca. 205 MPa). These fluid pressures can only be maintained at low to intermediate differential stresses during the earliest stages of compressional stress regime after the early Variscan compressional tectonic inversion (figure A13).

An initial increasing tectonic stress component, heralding the Variscan orogeny has been invoked as driving mechanism to decrease the differential stress during the latest stages of burial and to generate lithostatic overpressures (Kenis et al. 2002). A theoretical analysis of different stress-state scenarios (Van Noten et al. 2012) shows that the BNVs and BPVs can only develop in an extensional Ardenne-Eifel basin that has a predefined structural NE-SW orientation, conform to the geodynamic context of the basin on the passive continental margin of the Rhenohercynian ocean (figure O4). This particular basin configuration is corroborated by the identification of major basin-bounding NE-SW trending normal fault systems by stratigraphical mapping (Mansy et al. 1999). During extension of the basin, in which the load of the overburden corresponds to σ1 = σV, there is already a σT in a NW-SE direction parallel to σ3 that causes an increase in the magnitude of σ3 and consequently a decrease in differential stress.

Figure A13 – Brittle failure mode plot (Sibson 2000) of differential stress versus pore fluid factor, specifically adapted to the veining conditions in the Ardenne-Eifel basin at 7 km depth. BNVs developed as extension veins (ext) at differential stress condition < 4T and near-lithostatic pore fluid pressures in an extensional regime, while BPVs developed as extension veins (ext) and extensional-shear veins (es) at differential stress conditions of < 4T (ext) and < 5,66 T (es) respectively and supralithostatic pore fluid pressures in a compressional regime (Van Noten et al. 2012).


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A stress-state evolution can be reconstructed using the P-T conditions derived for the BNVs and BPVs (figure A14) (Van Noten et al. 2012). The lithostatic load at ca. 7 km depth of ca. 185 MPa is kept constant in the Mohr circle reconstruction. This means that the magnitude of the burial-related σV is kept constant. Using a Poisson’s ratio of ca. 0.25 for sandstone, the vertical overburden stress is translated in a horizontal overburden stress of ca. 62 MPa (σh). In such a configuration the differential stress is ca. 123 MPa. This value is too large to allow mode I fracturing, for which the differential stress cannot be larger than 4T (T = tensile strength of rock). Taking T = 10 MPa for sandstone, this means that differential stress must be smaller than ca. 40 MPa for mode I fracturing. Therefore, the tectonic stress σT acting on the minimum principal stress σ3, which are both oriented NE-SW, must at least increase up to ca. 83 MPa in order to sufficiently decrease the differential stress (figure A14(a)). At this stage, a fluid pressure of ca. 190 MPa is able to overcome T, allowing mode I fracturing in an extensional stress regime, resulting in the BNVs (figure 14(b)). It should be noted that in reality both processes, illustrated in steps (a) and (b) on figure A14 occur simultaneously.

The classical models for tectonic switches, as represented by the brittle failure mode plots (figure A13), only concern the maximum principal stress σ1 and the minimum principal stress σ3, ignoring the intermediate principal stress σ2. The tectonic switch thus occurs at zero differential stress, i.e. a state of isotropic pressure. This seems an oversimplification of the true 3D stress-state evolution in the Earth’s crust, which most probably always remains triaxial, even during tectonic inversions at low differential stresses. So, the progressive increase of the tectonic stress σT acting on the minimum principal stress σh = σ3, will cause a first switch between the two horizontal stresses (figures A14(c) and A15). The tectonic stress σT is now oriented parallel to the intermediate principal stress σH = σ2, oriented NE-SW (figure A15). This is followed by the first major tectonic switch between the vertical stress σV, being σ1, and the horizontal stress σH, being σ2. The result is a transitional wrench stress regime with the intermediate principal stress σ2 becoming the vertical stress σV (figures A14(d) and A15). After this stage, differential stress increases until the second major tectonic switch towards a compressional stress regime. It is remarkable that to date this intermediate regime has been ignored in an tectonic inversion

Figure A14 – Stress-state Mohr circle reconstruction for the compressional tectonic inversion taking place in the pHASB (Van Noten et al. 2012). (a) & (b) represent the stress-state evolution prior to the tectonic inversion, resulting in the BNVs. (e) & (f) represent the stress-state evolution after the inversion, resulting in the BPVs, both extensional (e) and extensional-shear (f).


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model and that no structures have ever been associated with this intermediate wrench stress regime (Van Noten et al. 2012).

After the reorientation of the principal stresses in the compressional regime, with the maximum principal stress σ1 oriented parallel to the tectonic stress σT, i.e. NW-SE oriented(σH), and the minimum principal stress σ3 oriented vertical (σV), fracturing and

veining takes place at a differential stress of < 30 MPa for extensional failure (figure A14(e)) and < 43 MPa for extensional-shear failure (figure A14(f)). Extensional failure therefore takes place at lower differential stresses in the compressional regime compared to the extensional regime, because of the lower strength of the rock due to bedding anisotropy during compression (figure A13). Tectonic stress σT rises up to ca. 155 MPa during extensional failure and ca. 167 MPa during extensional-shear failure. At this stage, supralithostatic fluid pressures of ca. 205 MPa are needed to overcome T (figure A14).

To conclude, the NE-SW bedding-normal extension veins, regionally present across the HASB (figure D6 – see stop D) can only be formed during an extensional stress state in which a NW-SE directed tectonic stress σT, oriented parallel to the opening direction (σ3) of the extensional Ardenne-Eifel rift basin, reduces the differential stress substantially in order to allow fluid-pressure driven mode I fracturing and extensional veining. Bedding-parallel veins with the observed NW-SE internal fabric can subsequently only be formed under NW-SE directed compression (σT). Also the subsequent fold and cleavage development will still be governed by the latter tectonic stress, as exemplified by the near parallelism between the cleavage-bedding intersection and the ‘mullion axes’, as we will observe in stop B at Dedenborn.

Figure A15 – Scenario illustrating the 3D stress-state changes of an extensional basin with a predefined σ3 oriented NW-SE that is shortened by a consistently NW-SE oriented tectonic stressσ3 at the onset of Variscan orogeny. BNVs are formed in the extensional regime; BPVs in the compressional regime. No structures are to date observed that could be linked to the transitional wrench regime (Van Noten et al. 2012).


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STOP B – North Eifel (Germany) – Dedenborn: “Mullions, the textbook example!”

by Koen Van Noten & Manuel Sintubin

INTRODUCTION

How can a little village in the North Eifel, with only 400 inhabitants, become world famous, at least amongst structural geologists? Simple, just look at the sign at the entrance of the village, proudly presenting their ‘Naturdenkmal’, their geological heritage site celebrating ‘mullions’.

The Dedenborn roadside outcrop is, indeed, world famous because of the textbook type example of mullions, for the first time described by Pilgers & Schmidt in 1957.


Dedenborn is situated some 9 km SW of the Hubertus Höhe – Schwammenauel section (stop A), as the crow flies. The village is located in the Rur valley. The geographical coordinates of the outcrop correspond to 50°34’53” and 6°21’08”.


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GEOLOGICAL SETTING

Dedenborn is still located in the North Eifel (figure A1). With respect to the outcrops of the Hubertus Höhe – Schwammenauel section, studied in stop A, the Dendenborn outcrop is situated towards the SW along strike, i.e. towards the deeper parts of the HASB. It represents the first clear appearance of ‘mullion’ structures, which we consider as the diagnostic field criterion to delineate the central HASB (cHASB) from the peripheral HASB (pHASB).

The rocks, exposed at the Dedenborn outcrop, belong to the Middle Rurberg beds, Pragian in age, primarily composed of fine- to coarse-grained sandstones and siltstones. Stratigraphically, they are situated underneath the Upper Rurberg beds, exposed along the Rursee (see figure A1 – stop A), representing a deeper structural level in the HASB (see figure A2 – stop A).

THE DEDENBORN OUTCROP

The Dedenborn outcrop exposes the steeply SE-dipping, overturned limb of NW-verging fold, very similar to the fold train, exposed along the Hubertus Höhe – Schwammenauel section (stop A). The bottom interface of a psammite layer composes the outcrop. The sandstone layers shows a graded bedding, inferring a younging direction towards the NW (figure B1). The adjacent pelite material, in front of the psammite interface, has disappeared.

The psammite surface shows a very regular, highly cylindrical cuspate-lobate morphology, with the cusps pointing towards the psammite (inwards in the outcrop). Due to the graded bedding, the sandstone grades into a pelite, showing a very-well developed slaty cleavage (at the backside of the outcrop) (figure B1). On the psammite surface, the bedding-cleavage intersection is apparent. This lineation makes an angle of ca. 26° CCW with respect to the long axis of the cuspate-lobate morphology. The cuspate-lobate morphology is slightly asymmetric.

Within the sandstone, a regularly and narrowly spaced, parallel array of lensoid quartz veins can be observed, at high angle to the bedding. The quartz veins systematically terminate within the cusps, suggesting a kinematic link. The quartz veins have an attitude of 128/44. After unfolding, it becomes clear that these ‘intermullion’ veins show the same

regionally consistent, prefolding orientation of all bedding-normal veins in the HASB (see figure A7 – stop A; figure D6 – stop D). No more bedding-parallel veins, so typical for the Rursee outcrops, are present at this outcrop. The absence of BPVs, indeed, typifies the cHASB.

A MIXED BRITTLE-PLASTIC DEFORMATION HISTORY

The development of the cuspate-lobate morphology, or mullions, is always associated with the presence of bedding-normal quartz veins in the psammitic layers. In layers without veins, no mullions develop. It is therefore fair to assume that the presence of the BNVs is of paramount importance for the ultimate development of the mullions.

The regional bedding-normal quartz vein arrays represent a fracturing event during the latest stages of burial in the overpressured Ardenne-Eifel basin (Urai et al. 2001; Kenis & Sintubin 2007 and

Figure B1 – An example of mullion development in a psammite layer with graded bedding. Cuspate-lobate folding only occurs at the interface with highest competence contrast (Urai et al. 2001; after Brühl 1969).


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references therein). After tectonic inversion, incipient, layer-parallel shortening eventually caused the cuspate-lobate folding in between the pre-existing quartz vein tips (see figure D11 – stop D). Quartz veins acted as rigid plates controlling the extrusion of the less competent psammite (Urai et al. 2001). During subsequent fold-and-cleavage development mullions became asymmetric due to flexural slip within the fold limbs. The angle between the mullion axes and the bedding-cleavage intersection corroborates that both structural features developed in a different setting, although closely related in space and time.

The Dedenborn outcrop already shows that the onset of Variscan shortening is expressed differently in the central and peripheral High-Ardenne slate belt, the former represented at Dedenborn (stop B) and Bastogne (stop D), the latter represented at the Rursee (stop A). In the deeper parts of the Ardenne-Eifel basin, below the brittle-ductile transition, psammites deformed in a plastic manner into mullions, pinned in between pre-existing, strong bedding-normal quartz veins. Towards the periphery of the slate belt mullions are absent and incipient shortening is expressed by bedding-parallel veins (Van Noten et al. 2008) (see figure D11 – stop D).


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BOUDINAGE VS. CUSPATE-LOBATE FOLDING

Boudinage refers to the disruption of layers, bodies or foliation planes within a rock mass in response to bulk extension along the enveloping surface (Goscombe et al. 2004). This layer-parallel extension means that the shortening direction (Z) is oriented at high angles to the enveloping surface, which infers that a contemporaneously developing tectonic foliation (XY), such as a cleavage, should be parallel to the boudinaged layers (figure B2). With respect to competence contrast, the more competent material will be affected by boudinage, resulting in alignment of boudin segments separated by boudin necks, in which veining often occurs. Boudin segments are commonly characterized by a characteristic aspect ratio of ½ (height/width) (figure B2).

Cuspate-lobate folding of an interface of two lithologies with contrasting rheology, on the other hand, is the result of layer-parallel shortening (Z) (van der Pluijm & Marshak 2004). This infers that a contemporaneous, kinematically linked cleavage (XY) is at high angles with respect to the buckled interface (figure B2). With respect to competence contrast, the cusps within the interface point towards the more competent material (e.g. psammite); the lobes towards the less competent material (e.g. pelite). Because of the analogy with architectural elements typical for the European Gothic style, the term mullion has been introduced to describe the characteristic morphology due to cuspate-lobate folding.

Figure B2 – Boudin versus mullion. Boudins – also named pinch-and-swell structures – are the result of layer-parallel extension of a competent layer in a less competent matrix. An analogue for a sequence of boudins is a string of small sausages. Boudins show a characteristic aspect ratio of ½ (height/width). Mullions – or cuspate-lobate folds – are the result of layer-parallel shortening of an interface between two lithologies with a contrasting competence. The best analogue for the ‘double-sided mullions’, typical for the cHASB (stop D), is a series of sausages – “boudins” in French – lying side by side. ‘Double-sided mullions’ show an aspect ratio, commonly larger than 1, i.e. height is equal or larger than the width of the segment between adjacent bedding-normal veins, pinning the mullions on both upper and lower interface of the layer.


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STOP C – Ardenne (Belgium) – Mousny – “Les Blancs Cailloux”: “A late-Variscan dilational jog?”

by Manuel Sintubin, Simon Depoorter & Dominique Jacques

INTRODUCTION

On the plateau of the High-Ardenne, large boulders of quartz are commonly found as decoration in gardens, as building stones for chapels, etc., suggesting the occurrence of massive quartz bodies in the poorly exposed High-Ardenne slate belt other than the veins that are observed in nearly every outcrop.

West of the small village of Mousny, a peculiar massive quartz occurrence (MQO) can be found in the locality Les Blancs Cailloux. It has long been regarded as a site of menhirs or stone circles, known as cromlechs, similar to the ones found in France and the U.K. (de Ruette 1972). Even a legend is associated with the site … The reality is, however, less romantic.


The site Les Blancs Cailloux at Mousny, is located approximately 10 km south of La Roche-en-Ardenne, 15 km west of Houffalize, and 15 km northwest of Bastogne, all as the crow flies. The geographical coordinates of the outcrop correspond to 50°06’19” and 05°36’04”.


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GEOLOGICAL SETTING

Mousny is located in the cHASB (figure O1), rather on the northern side of the Ardenne culmination, just south of the Lower Palaeozoic Stavelot-Venn basement inlier (figure E1 – stop E). Mousny is indeed still located in the zone where mullions are present (e.g. at Houffalize, Mabompré). The area is, furthermore, located at the southern extremity of the so-called ‘Ourthe Zone’ (Hance et al. 1999), a diffuse, ca. 10 km wide, NNE-SSW trending zone, bordering the Stavelot-Venn inlier to the west. This zone is interpreted as the expression of a buried buttressing lateral ramp causing transpressional

deformation within the overriding Ardenne allochthon. This transpressional deformation seems materialized by a series of inferred, dextral, NNE-SSW trending strike-slip faults (Dejonghe 2008) (figure C1).

As can be seen on the map (figure C1), the Mousny massive quartz occurrence (MMQO) is considered by Dejonghe (2008) as exemplary for the mineral deposits linked with inferred longitudinal faults, in this case the Mousny fault.

In the wider Mousny area, rocks of the Lower Devonian Villé, La Roche and Pèrnelle formations (Bultynck & Dejonghe 2001; Dejonghe & Hance 2001) are present in the subsurface. These formation are late Pragian to early Emsian in age. In this respect they can stratigraphically be correlated with the Upper Rurberg and Heimbach beds, visited along the shores of the Rursee (stop A). The Villé formation consist of dark shales intercalated with rusty brown laminated sandstones, blue sandstones and fossiliferous bluish carbonate-rich sandstones. The La Roche formation is predominantly composed of blues hales interlayered with light

Legend of « Les Cailloux de Mousny » A long time ago, there lived a shepherd in Mousny who was feared by everyone in the village because of his witchcraft. On a hot summer day, the shepherd, sitting down in the fields overlooking his herd of sheep, saw a pilgrim, on his way to the Hermitage of Saint-Thibaut (Marcourt), approaching him from the road. The pilgrim’s head was shaven and he leaned heavily on his cane, exhausted from the heat. The pilgrim sat down beside the shepherd and implored him to give him some water as the heat was unbearable. The shepherd, who had some bread and a large jar of water, despised pilgrims and urged the stranger to leave him alone. The pilgrim looked at the shepherd with tears in his eyes. Again, the shepherd insisted that the pilgrim would leave and continue his journey. The pilgrim got up and said: “Not only you refuse me some water, but you also won’t let me sit next to you for a few minutes. I’ll pray that Saint-Thibaut will bring a little charity upon your heart”. He then stood up, walked a few meters and sat down again. The shepherd went to him and ordered him to leave, if not he would strike him with his cane. The poor pilgrim got up and left once again. Enraged, the shepherd picked up a stone and threw it at the pilgrim. The pilgrim cried out in pain, took the rock and threw it back to the shepherd. As the rock hit the terrified shepherd, he, his dog and his herd of sheep turned into stone. The pilgrim turned out to be Jesus of Nazareth in disguise …

Figure C1 – Geological map of the Ortho-Mousny area (Dejonghe & Hance 2001). AMO: alluvial deposits; VIL: Villé formation; LAR: La Roche formation; PER: Pernèlle formation. Les Blancs Cailloux coincide with the trace of the longitudinal Mousny fault.


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blue sandstones and quartzitic sandstones, red sandstones and blue-green fossiliferous siltstones. The Pèrnelle formation is made up of massive, green-blue sandstones and quartzitic sandstones, interlayered with thin dark shales and siltstones (Bultynck & Dejonghe 2001; Dejonghe & Hance 2001).

Mousny is still located in the epizonal metamorphic core of the cHASB (Fielitz & Mansy 1999). Peak metamorphic temperatures of 300°C or more can thus be expected during the latest stages of burial and subsequent Variscan deformation.

MOUSNY MASSIVE QUARTZ OCCURRENCE

The Mousny massive quartz occurrence (MMQO) consists of 30 m3-size bodies of massive milky quartz (figure C2). Inside the milky quartz mass (Q2 generation), geodes of transparent quartz (Q4 generation) can be recognized. While the quartz in the milky matrix does not show crystal habit, the quartz inside the geodes shows a typical prismatic habit. Besides quartz, also chlorite, muscovite and carbonates can be found in small amounts. Chlorite commonly occur at the contact between vein quartz and the host rock.

Two, roughly orthogonal, sets of subvertical joints can be identified on the different quartz bodies: one set with an overall N40W strike, one set with an overall N40E strike. Joints consistently dip between 85 to 90°. One or both joint sets are systematically present on each of the individual bodies, supporting the basic assumption that all bodies are effectively in situ (figure C2).

Within the individual quartz bodies, also host-rock fragments can be identified. These host-rock fragments are rusty brown sandstone and grey-greenish siltstone fragments. They can most probably be attributed to the Villé formation, late Pragian in age (Bultynck & Dejonghe 2001) (figure C1) (cf. Upper Rurberg beds – stop A). Some host-rock fragments contain quartz veins (Q1 generation). Remarkable is that the host-rock inclusions are rather elongated and show a very consistent attitude, which is very similar to the average regional cleavage attitude (145/55) as measured in neighboring

Figure C2 – Les Blancs Cailloux: detailed map of the Mousny massive quartz occurrence with rose diagrams of the joint orientations for each set of boulders. Grid width is approximately 5 m. Width of the area is 65 m (Depoorter et al. in prep.).


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outcrops (figure C3). Also the consistent orientation of the host-rock inclusions corroborate the in situ assumption of the quartz bodies.

It should be noted that in the neighboring outcrops, cleavage-parallel veins (Q3 generation) are systematically present (or cutting cleavage at a very small angle (< 10°). These cleavage-parallel veins sometimes show some pinch-and-swell characteristics, indicative of some vein-parallel stretching. The specific temporal relationship of these veins with respect to the Variscan cleavage allows to identify these cleavage-parallel veins as late-orogenic veins. The particular orientation of the host-rock inclusions – parallel to the regional cleavage attitude (figure C3) – suggests that there might be a kinematic link between the cleavage-parallel quartz veins (Q3) and the MMQO (Q2 & Q4).

WHAT DOES THE QUARTZ TELLS US?

The different quartz generations have been subject to an optical and cathodoluminescence analysis, fluid inclusion microthermometry and Raman microspectroscopy, (1) to test the suggested hypothesis that the cleavage-parallel veins and the MMQO may be kinematically linked, (2) to identify the composition of the fluids and (3) reconstruct P-T conditions of quartz precipitation.

All quartz generations, in both the MMQO and the cleavage-parallel veins, show a similar degree of crystal-plastic deformation, indicative of low- to moderate-temperature crystal-plastic deformation between 200 and 400°C (e.g. undulose extinction, deformation lamellae, subgrain boundaries, bulging recrystallization), which suggests that quartz precipitation occurred at low-grade metamorphic conditions at mid-crustal levels prior to the end of the Variscan orogeny.

The first quartz generation (Q1), consists of small quartz veinlets, primarily made up of elongate-blocky crystals. These veinlets, found in the host-rock inclusions in the MMQO, already crosscut the cleavage fabric, present in the host-rock inclusions. In turn, they are crosscut by the bulk quartz matrix of the MMQO (Q2). The Q1 vein quartz generation shows pink to bright blue luminescence.

The next quartz generation (Q2) composes the bulk of the MMQO. Q2 shows dark blue luminescence. Cathodoluminescence reveals, though, that Q2 quartz generation is still composed of at least four different phases of crystal growth, fracturing and infill. Microthermometry shows that the Q2 fluid inclusions are composed of H2O-CO2-N2-(CH4-NaCl) aqueous-gaseous fluids. Between different subtypes, a slight variation in gas content can be observed from enriched in CO2 to enriched in N2.

Comparing the composition of the mixed aqueous-gaseous fluid inclusions, it is clear that the fluid composition of the Q3 quartz generation in the cleavage-parallel veins, corresponds very well with the composition of one of the subphases of the Q2 quartz in the MMQO. In time, we believe these Q2-Q3 fluids can be correlated to the A2 and/or G1 type fluids of Kenis et al. (2005a) (figure D4 – see stop D).

Figure C3 – Lower-hemisphere equal-area projection of regional cleavage attitude (great circles marked by dashed line; poles marked by x; contoured; n=19) and the attitude of the elongated host-rock inclusions in the MMQO (great circles marked by full line; poles marked by triangle; n=6) (Depoorter et al. in prep.).


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The final Q4 quartz generation consists of the more transparent quartz, present within the milky quartz mass of Q2/Q3 generation. This quartz is characterized by bright blue to yellow luminescence. Q4 quartz consists of a lower-temperature, low-saline H2O-NaCl fluid. Salinities range from 0.5 to 10 eq.wt% NaCl; homogenization temperatures between 120 and 180°C. Q4 fluids are very similar to the late W1 type fluid inclusions of Kenis et al. (2005a) (see stop D), typical for the latest stages of the Variscan deformation in the entire High-Ardenne slate belt (see figure D4 – stop D).

P-T conditions of quartz precipitation

Calculating the isochores of the different fluid generations, three sets of isochores are obtained (figure C4): (1) a Q2 fluid with a density ranging between 0.91 and 0.92 g/cm3 (Q2’; red isochores on figure C4); (2) a Q2 fluid with a density ranging between 0.77 and 0.85 g/cm3 (Q2’ stretched; orange isochores on figure C4); (3) a Q2 fluid with a density ranging between 1.02 and 1.03 g/cm3 (Q2”; green isochores on figure C4); and (4) a Q4 fluid with a density ranging between 0.91 and 0.97 g/cm3 (blue isochores on figure C4).

To retrieve the original trapping conditions of the fluids, a number of assumptions are made. Based on the location of the MMQO, a peak metamorphic temperature of ca. 350°C is proposed (see figure E1 – stop E). A maximum burial depth of ca. 10 km, i.e. maximum overburden pressure of 265 MPa, is assumed primarily based on the stratigraphical context of the MMQO. Taking into account the late-orogenic nature of the MMQO and the cleavage-parallel veins, it is fair to assume that quartz precipitation occurred on the retrograde metamorphic path, characterized by a geothermal gradient of ca. 35°C/km.

It is obvious that the two isochore sets for Q2 fluids do not represent the same physical and chemical conditions of the fluid inclusion assemblage at the time of trapping. The isochore sets show a strong variation in density, with a range from 0.77 to 0.92 g/cm3, and record a strongly different internal fluid pressure at homogenization, with values of 135 to 10 MPa for set (1) and (2) respectively (figure C4). From the composition of the fluid inclusions in question, it becomes apparent that the isochores of set (2), with low densities from 0.77 to 0.85 g/cm3, represents stretching of the original fluid inclusion assemblage, represented by set (1). The isochores of set (2) can thus not be used to reconstruct P-T conditions of fluid entrapment. In a similar fashion, also the isochore set (3) cannot

Figure C4 – P-T diagram showing the trapping conditions of the fluids in the MMQO and the cleavage-parallel veins. Maximum metamorphic temperature of 350°C and a maximum lithostatic pressure of 265 MPa is considered. A geothermal gradient of 35°C/km for both lithostatic and hydrostatic conditions is indicated (Depoorter et al. in prep.).


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be used. Calculated isochores indicate internal pressures of 320 to 550 MPa at homogenization, which is completely out of range taking into account the peak metamorphic conditions considered. These values again indicate post-entrapment deformation of the fluid inclusions, or even possible leakage during the microthermometric analysis.

Hence, P-T conditions can only be derived based on the isochore sets (1) and (4) for the entrapment of Q2 and Q4 fluids respectively. P-T ranges are determined from the intersection of the isochores with the suggested, lithostatic and hydrostatic, geothermal gradient (figure C4). The resulting ranges indicates that Q2 quartz precipitated at 275 to 350°C and 120 to 245 MPa (red hatching on figure C4); Q4 quartz precipitated at 185 to 305°C and 45 to 215 MPa (blue hatching on figure C4). Because the HASB fluid system is considered rock-buffered (Kenis et al. 2005a), the P-T conditions measured for the Q2 and Q4 fluids also reflect the thermal history of the host rock along a retrograde metamorphic path. Both quartz generations precipitated in successive steps along this path, with the aqueous Q4 fluid entrapment at lower temperatures and pressures than the mixed aqueous-gaseous Q2 fluid. This difference in P-T conditions, as well as the temporal relationship, is also reflected in the microfabric of the Q4 quartz, showing less crystal-plastic deformation than the Q2 quartz. In a similar way, a relative timing between the different subphases of Q2 may be suggested. The low-density Q2’ inclusions occur commonly in less-strained quartz crystals than the Q2” inclusions. Q2’ quartz displays a more or less homogeneous, dark blue, luminescence, while Q2” quartz shows variable luminescence. Q2” inclusions are moreover extensively deformed, as shown by the aberrant isochore calculation. Possibly Q2” fluid entrapment occurred at slightly different P-T conditions than Q2’ entrapment.

Compared with the overall evolution of the fluids in the cHASB, as demonstrated by Kenis et al. (2005a) (figure D4 – see stop D), the Q2 fluids can be correlated with the progressively maturing metamorphic HASB fluid system (A1 to G1 type – see stop D), present since the latest stages of the burial history of the Ardenne-Eifel basin (BNVs – see stop A, B and D) throughout the slate belt development. The Q4 fluids can be correlated with the late-orogenic fluids of the HASB fluid system, characteristic for the retrograde metamorphic path (W1 type –see stop D).

A LATE-OROGENIC DILATIONAL JOG?

The host-rock inclusions contained in the MMQO (Q2), which are oriented following the regional cleavage attitude, in combination with the (near) cleavage-parallel milky quartz veins (Q3), showing similar microstructures, mineralogy and fluid inclusion content, suggest that the MMQO is kinematically closely related to the cleavage-parallel veins. The MMQO can therefore be considered as a particular type of the cleavage-parallel veins present throughout the HASB. The temporal relationship with the cleavage suggests that the veining occurred late in the Variscan orogeny, probably at the waning stages of the Variscan orogeny in the HASB. In this respect, the cleavage-parallel veins (Q3) show strong similarities with the discordant veins in the Herbeumont area (see stop E).

Considering the parallel to near-parallel orientation of the Q3 generation quartz veins with respect to the slaty cleavage anisotropy, it can be argued that the cleavage was activated as mode I fractures, in which subsequently quartz could precipitate. Mode I fracturing would imply that the minimum effective principal stress (σ3) should be near-perpendicular to the cleavage attitude (figure C5). Considering the similarity with the parental mode I cracks (see e.g. figure E11 – stop E), eventually evolving into the discordant veins in the Herbeumont context (see stop E), it is fair to assume that the maximum effective principal stress (σ1) should be oriented nearly parallel to the dip direction of the cleavage anisotropy (figure C5). This would infer an extensional Andersonian stress regime, with a subvertical maximum principal stress σ1 = σV, characterizing the late-orogenic extensional destabilization of the slate belt (Van Baelen, 2010). Moreover, Van Noten et al. (2011) demonstrate


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that periods of tectonic inversion, both at the onset of orogeny (i.e. positive tectonic inversion) (see stop A) and in the late stages of orogeny (i.e. negative tectonic inversion), are periods of enhanced structural permeability, primarily due to the presence of high – up to (supra-)lithostatic – fluid pressures and low differential stresses (see figure A13 – stop A). This transient enhanced structural permeability is reflected in the widespread occurrence of quartz veins within the HASB. It is fair to assume that the Q2/Q3 and Q4 quartz generations, found in the cleavage-parallel veins and the MMQO, can be related with the late, negative tectonic inversion.

We suggest that the MMQO is a excessively grown cleavage-parallel vein (figure C5). The best geometrical context to explain this localised massive quartz occurrence, is a dilational jog (figure C5). Such dilational jogs commonly develop along bend or step-overs, linking segments in fault zones (Childs et al., 1996; Cox et al. 2001; Micklethwaite et al. 2010). Classically, dilation in such fault-related jogs, is associated with episodic fluid-pressure drops caused by seismic, high strain-rate, fault movements (i.e. ‘suction pumping’; Sibson, 2005), resulting in implosion breccias of the host-rock and pre-existing precipitates (Sibson, 2000; Muchez & Sintubin, 1998). The host-rock inclusions, containing a cleavage fabric and crosscut by a Q1 generation of quartz veins, within the MMQO have, though, preserved their attitude, parallel to the regional cleavage attitude. This rather infers a more controlled, episodic, low strain-rate, opening of an ever growing cleavage-parallel veins, eventually generating a massive quartz occurrence (figure C5). The development of such massive quartz occurrence can therefore best be compared to the crack-seal opening process of a single vein.

The opening of this dilational jog will most probably be accommodated by the development of low-angle, extensional shear zones, bounding the jog (figure C5). But while the opening of fault-related dilational jogs is controlled by episodic fault movement (‘suction pumping’; Sibson, 2005), we suggest that an episodic, fluid-pressure controlled opening of the

cleavage-parallel dilation site controls the accommodation through connecting shear zones.

It is fair to assume that throughout the slate belt a system of dilational jogs, connected with extensional shear zones, acting as transfer zones of the jog opening, would develop during the extensional destabilisation of the slate belt. Vein quartz occurrences in the wider Mousny area, as well as the cleavage-parallel veins, may very well be relics of other up-dip or down-dip jogs in such a connected system. The widespread occurrence of large quartz boulders as decoration in gardens and as building material for chapels and houses all over the High Ardenne region, suggests that dilational jogs, as exposed at Mousny, may, indeed, very well be a rather common structural feature.

In this respect, the Mousny-type of dilational jog is very comparable to the discordant veins in the Herbeumont area (Van Baelen, 2010) (see stop E). The main difference is that the Mousny-type dilational jog reflects a purely brittle deformation history, while the Herbeumont-type discordant veins reflect a mixed brittle-plastic deformation history at a deeper structural level within the collapsing slate belt. Still, P-T conditions of the developing dilational jog at Mousny are such that low- to moderate crystal-plastic deformation, possibly related to the shear accommodation, is recorded in the quartz microfabric.

Figure C5 – Dilational jog model to explain the occurrence of the Mousny massive quartz occurrence. Sequential opening of a cleavage-parallel vein eventually results in the massive quartz occurrence (Depoorter et al. in prep.).


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STOP D – Ardenne (Belgium) – Bastogne – Mardasson quarry: “About boudins, veins and mullions”

by Manuel Sintubin & Koen Van Noten

INTRODUCTION

Together with the Collignon quarry, the Mardasson quarry, originally called the Ballastière de Bastogne (Gosselet 1888), is the locality where in August 1908 during a field trip organised by the Société Géologique de Belgique the term ‘boudin’ and ‘boudinage’ were used for the first time in order to facilitate the discussion with respect to particular structures in the Lower Devonian metasedimentary series, resembling sausages lying side by side (Lohest et al. 1908) (see figure B2 – stop B). These structures have been in the centre of a century of debate. Today, these structures turn out to be the expression of a brittle-plastic deformation at the onset of the Variscan orogeny.


The Mardasson quarry is located north of the war memorial of Mardasson, east of the city of Bastogne. The geographical coordinates of the quarry correspond to 50°00’50”N and 5°44’21”E.

This is a quarry in exploitation. Hard hats compulsory! Please approach exploitation front with caution!

Map: 1:25.000 sheet 60/7-8 Longchamps-Longvilly Contact: Enrobage Stockem, route de Bouillon 222, B-6700 Arlon, tel.: +32 (0)63 24 52 00; fax: +32 (0)63 22 77 13; e-mail: [email protected].

mailto:[email protected]�


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GEOLOGICAL SETTING

Bastogne is located in the core of the cHASB (figure O1), in a very similar context as the outcrops, visited in the North Eifel (stop A & B; figure A2), i.e. in the pop-up structure comprised between frontal thrust bordering the HASB to the north (e.g. Lacquement 2001; Lacquement et al. 2005) and the Malsbenden-Trois Vierges backthrust. Bastogne is located near the southwestern extremity of the map trace of the latter backthrust (see figure E1 – stop E).

The rocks, exposed in the Mardasson quarry, belong to the Mirwart formation (Bultynck & Dejonghe 2001), which is considered to have a Lochkovian to Pragian age (formerly Siegenian 1). In the quarry grey-blue metasiltstones, intercalated with grey-blue quartzites are exposed. The quartzitic layers are discontinuous, showing sedimentary wedes, dichotomy, etc. Stratigraphically, these are the oldest material encountered on this field trip.

In the Bastogne area the maximum metamorphic temperature conditions are estimated to range between 365 and 400°C (Beugnies 1986; Darimont et al. 1988). At peak metamorphism, pressure conditions are estimated to range between 180 and 200 MPa, based on fluid inclusion data (Darimont 1986) and metamorphic mineral assemblages (Beugnies 1986; Darimont et al. 1988) respectively. An overall prograde geothermal gradient of about 40 to 50°C/km is therefore considered to be present during burial in the subsiding Ardenne-Eifel rift basin and incipient deformation during Variscan orogeny (Darimont 1986; Darimont et al. 1988; Beugnies 1986; Kenis et al. 2000; Schroyen 2000).

MARDASSON QUARRY

The overall structural grain in the poorly exposed Bastogne area is NE-SW. Folds plunge weakly to the southwest. The Mardasson quarry is situated within the wide hinge of a hectometer-scale, slightly NW-verging, open antiform, with a hinge line orientation of 235/10. The axial-planar cleavage has an average attitude of 150/61.

A HISTORICAL RETROSPECT

The first geometrical description of the particular quartz-vein occurrence in the Bastogne area (Belgium) goes back to 1888 when Gosselet mentions in his chapter on metamorphism in the Ardenne that “tous ces filons de quartz … sont limités au grès et s’arrêtent à la cornéite” (figure D1). In his work on the origin of the metamorphic rocks in the Bastogne area, Stainier (1907) presents an extensive inventory of the occurrence and geometry of these veins, demonstrating the regional occurrence and significance of the veins and associated structures. To explain these particular structures, he invokes a polyphase brittle-ductile deformation sequence: “… c’est grâce à la préexistence des filons quartzeux et des crevasses que l’on peut expliquer l’allure en chapelet et l’allure en petites voussettes, unique en Ardenne, que présentes les bancs de grès de la région. Si ces bancs de grès n’avaient pas présenté des fissures, la poussée tangentielle du grand ridement les aurait plissées en voûtes ou bassins ordinaires; ou bien, si la poussée avait été trop forte, elle aurait

Bastonite K3(Mg,Fe,Fe,Al)5.5(Si,Al)8O22(OH)4.1.5H20 In quartz veins of the Bastogne area a dark mica can be observed. This hydrobiotite is described as extremely poor in potassium (Klement 1888) and is named bastonite after its type locality. It is one of the richest NH4

+-dark micas (Bos et al. 1987, Kenis et al. 2002).


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déterminé la production de failles inverses de refoulement. Au lieu de cela, chaque compartiment d’un banc de grès, devenu indépendant des ses voisins dont il était séparé par un joint quartzeux, s’est arqué, pour son compte, en une petite voussette ou bien s’est renflé en saucisson ou grain de chapelet”. To Stainier the sequence of events was evident: first fracturing and veining, then layer-parallel shortening, all prior to the main Variscan folding and cleavage development.

To demonstrate the work of Stainier, a field trip was organised by the Société Géologique de Belgique. On August 31st, 1908, at the Carrière Collignon near Bastogne, Lohest introduced the terms ‘boudin’ and ‘boudinage’ in order to facilitate discussions (Lohest et al. 1908):

“Lorsque l’on voit ces segments renflés sur une surface de stratification étendue, mise à nu par l’exploitation, on croirait voir une série d’énormes cylindres ou boudins alignés côté à côté … sur l’initiative de M. Lohest on a fréquemment utilisé, pour la facilité du langage, les néologismes de boudiner et de boudinage.”

At this early stage, these terms were used purely descriptive to indicate the series of sausage-like structures, lying side by side (figure D2; see also figure B2 – stop B). Already during the field trip, a discussion unfolds on the relative timing of the different deformation features, setting the stage for a century long discussion on the development of these particular structures observed at Bastogne.

In the first half of the 20th century numerous models were proposed with variable chronologies with respect to the development of the different deformation features (Corin 1932, Holmquist 1931, Quirke 1923, Wegmann 1932), clearly demonstrating the uncommon nature of the structures (cf. Quirke 1923). Although these authors often mentioned that the ‘boudins’ at Bastogne were seemingly related to ‘pinch-and-swell’ structures as described by e.g. Ramsay (1866) and Harker (1889), it was generally agreed that they did not fully resemble, primarily because of the strongly different aspect ratio of the segments (Holmquist 1931, Walls 1937). Cloos (1947), however, did not make this clear distinction, and the term ‘boudinage’, applied to both types of structures, obtained its current kinematic definition referring to a process of layer-parallel extension.

Ever since confusion only got worse, in particular when very similar structures in the North Eifel region were described as ‘mullions’ (Pilgers & Schmidt 1957) (see stop B – Dedenborn). In an attempt to relate the ‘mullions’ of the North Eifel – considered as layer-parallel shortening structures (i.e. cuspate-lobate folding) – and ‘boudins’ of Bastogne – at that time considered as extension structures – terms such as L-boudins or ‘auslängungs-boudins’ and K-boudins or ‘verkürzungs-boudins’ were introduced (Brühl 1969). The ‘mullions’ of Pilgers and Schmidt (1957) became ‘half-boudins’ (Brühl 1969). Quickly, it was acknowledged that this new terminology was definitively not the solution to the confusing, taking into account that the association of the term ‘boudinage’ with bulk extension was

Figure D1 – First sketch of the brittle-ductile structures in the Mardasson quarry (‘Ballastière de Bastogne’) by Gosselet in 1888. c: cornéite; g: grès; q: quartz vein.

Figure D2 – An illustration of the original ‘boudins’ in the report of the field meeting in Bastogne of the Société Géologique de Belgique (Lohest et al. 1908).


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already too entrenched in English literature (Mukhopadhyay 1972). Mukhopadhyay (1972) and Spaeth (1986), more recently corroborated by Urai et al. (2001), considered the structures in the Ardenne-Eifel region as mullions. Apart from them the idea of ‘shortened boudins’ gained, however, ground, primarily based on the atypical aspect ratio of the segments (Jongmans & Cosgrove 1993, Lambert & Bellière 1976, Rondeel & Voermans 1975), an idea still persistent to date (Vanbrabant & Dejonghe 2006).

A century of debate primarily focused on the cuspate-lobate – sausage-like – geometry of the sandstone-pelite bedding interface, largely ignoring the quartz veins. Renewed interest in the quartz veins allowed constraining the formation conditions and kinematic significance of the veins with respect to the overall structure development (Kenis et al. 2005a, Kenis et al. 2002). A numerical modelling approach (Kenis et al. 2004) ultimately corroborated the mullion model and even allowed to use the particular cuspate-lobate morphology to constrain the rheology of the deforming middle crust (Kenis et al. 2005b). Eventually, the hypothesis initially postulated by Stainier in 1907 becomes remarkably up to date!

THE BRITTLE-PLASTIC STRUCTURES

Describing the particular structures in the Bastogne area, we make a clear distinction between two distinct structural features (figure D3): (1) the veins, and (2) the cuspate-lobate morphology of the bedding interface. Although we consider both features separately, a regional survey shows that both features are spatially closely related and are kinematically linked. If veins are not present, the cuspate-lobate morphology of the bedding interface is lacking (see also stop B – Dedenborn).

These brittle-plastic structures are limited both in stratigraphic and regional extend (Kenis 2004, Urai et al. 2001). They only occur in formations of Upper Lochkovian to Pragian age. Regionally, their occurrence is limited to the southern limb of the Ardenne culmination (see figures O1, A1 & E1), coinciding with the region of highest metamorphic grade (Fielitz & Mansy 1999) and representing the deepest parts of the Eifel-Ardenne rift basin (Oncken et al. 1999).

Veins – geometrical characteristics

The veins show the following geometrical characteristics (figure D3): (1) commonly lens-shaped; (2) oriented perpendicular to bedding, independent of orientation of bedding; (3) limited to the most competent parts of the multilayer sequence, i.e. the sandstone layers; mostly limited to one single bed; (4) systematic array of parallel veins, very continuous in the longitudinal direction; (5) regular

Figure D3 – Synthetic figure of the brittle-ductile structures in the Ardenne-Eifel area. Q: quartz vein; Bpd: bedding-parallel dissolution seam; gray lines represent the tectonic cleavage (Kenis & Sintubin 2007).


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spacing; spacing in relation to bed thickness; and (6) aspect ratio (height/width) of the segments in between two adjacent veins ranges from 1 to 7 (even up to 20).

A number of features allow constraining the relative timing of the veining: (1) prior to Variscan folding and cleavage development (always perpendicular to bedding; cleavage crosscuts veins); (2) occurred in beds in a horizontal, depositional, disposition (perpendicular to bedding); (3) prior to maximum burial conditions (bedding-parallel dissolution seams crosscutting the veins). It is therefore fair to assume that veining occurred during the last stages of the deep burial of the sediments within the Ardenne-Eifel rift basin, thus most probably during the early Viséan (ca. 340 Ma).

Veins – mineralogy and fabric

The veins predominantly consist of quartz, but also contain chlorite, muscovite, biotite (bastonite), pyrite, chalcopyrite, ilmenite, and feldspar. Native gold and silver have also been described in the veins at Bastogne (Hatert et al. 2000). In general, the vein minerals all are well-crystallized and no indications for a different timing of the formation of the different minerals could be observed, such as broken mineral fragments enclosed by other minerals, crosscutting relationships, etc.

The vein quartz shows a number of microstructures characteristic of crystal plasticity (Kenis et al. 2005a). Large cores of quartz grains abundantly occur. Grain sizes range between 500 µm and 1 cm. Quartz grains show low-temperature crystal-plastic deformation microstructures, such as undulose extinction, subgrain boundaries and deformation lamellae. Sometimes, the vein quartz shows a core-mantle structure. Grain size of smaller grains ranges between 40 and 120 µm. The vein-quartz fabric is indicative for dislocation creep, evolving to recovery and dynamic recrystallization by subgrain rotation and grain-boundary migration. These microstructures comply with regime 2 of experimental dislocation-creep regimes that corresponds to natural temperature conditions of 350 to 500°C (Stipp et al. 2002). This temperature range is conform with the metamorphic temperatures of the host rock (Fielitz & Mansy 1999).

In contrast with the vein quartz, the host rock only contains evidence for quartz dissolution and overgrowth, suggesting that solution-precipitation creep was the dominant deformation mechanism responsible for the deformation of the sandstone. The absence of notable dislocation creep deformation of quartz in the sandstone suggests that the differential stress remained too low for dislocation creep to be competitive with solution-precipitation creep. The flow stress in the quartz veins was much higher, causing an inhomogeneous stress field in the different materials. The latter difference will cause buckling of the interface at the onset of the Variscan shortening (Kenis et al. 2005a).

Veins – fluid inclusion characteristics

Fluid inclusions (ca. 510 in 13 samples – size ranging from 2 to 20 µm) occurring in veins from ten different localities enabled to reconstruct a temporal and spatial evolution of the chemical composition of the fluids during veining and subsequent Variscan deformation. Using petrography, microthermometry (using a USGS-modified fluid-inclusion stage) and Laser Raman Spectroscopy, four types of fluid inclusions were identified in the vein quartz:

Type 1 inclusions (A1) found in samples of vein quartz in both the epizone and anchizone; occur in growth zones and are therefore assigned a primary origin; also occur as isolated

inclusions; mixed aqueous-gaseous fluid inclusions (H2O-NaCl-CO2-other gases); low to high salinity; between ca. 0 and 3.5 eq.wt% NaCl in achizone; between 0.6 and 17 eq.wt%

NaCl in epizone;


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Type 2 inclusions (A2) only found in samples of vein quartz in the epizone; occur spatially associated with A1 primary inclusions and are morphological similar; also

observed in growth zones; gaseous (CO2-N2-(CH4)) fluid inclusions;

Type 3 inclusions (G1) only found in samples of vein quartz in the epizone; inter-grain, trail-bound inclusions occurring along the surface of former microcracks and are thus

described as secondary fluid inclusions; the composition of this gaseous fluid type varies from a CO2-N2-(CH4) composition to a pure N2-

(CH4) composition;

Type 4 inclusions (W1) found in samples of vein quartz in both the epizone and anchizone; inter-grain, trail-bound inclusions; W1 trails crosscut G1 trails; aqueous (H2O-NaCl) fluid inclusions; this type of fluid inclusions is frequently described in the

Devonian and Carboniferous sequences in the entire Variscan foreland fold-and-thrust belt exposed in the Ardenne-Eifel area (Darimont et al. 1988, Kenis et al. 2000, Muchez et al. 1997, Schroyen 2000).

The first (A1) and second (A2) type of fluid inclusions are considered primary inclusions that reflect the earliest phase of vein-quartz growth. The timing of these primary inclusions is correlated with the final stages of the burial history of the Lower Devonian metasediments.

The third (G1) and fourth (W1) type of inclusions are closely related with microcracks that formed after vein formation and can most probably be linked with the deformation of the veins during the Variscan orogeny. Type 3 (G1) fluid inclusions are interpreted to have formed prior to the formation of the latest type of fluid inclusions (type 4 – W1).

Figure D4 – Temporal and spatial evolution of the chemical composition of the fluid inclusions within the epizonal metamorphic cHASB and the anchizonal pHASB (see stop A) (Kenis et al. 2005a).


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Veins – chemical variation of the fluids

The primary aqueous/gaseous fluid inclusions are common in the metamorphic zone of the High-Ardenne slate belt (Darimont 1986, Schroyen & Muchez 2000). Their compositional range is compatible with metamorphic fluids produced by intense dehydration and decarbonisation under low-grade metamorphic conditions in sedimentary sequences. The presence of a large amount of non-polar species (e.g. CO2, N2, CH4) in the primary inclusions is indicative of an intense fluid-rock interaction during metamorphism. CO2 likely originated from decarbonisation of sedimentary sequences or oxidation of organic material, while CH4 is thought to be produced by the maturation of organic matter. The origin of N2 could be attributed to two main sources: (1) N2 could be released during maturation of NH4

+-bearing minerals in the host rock (e.g. bastonite) due to prograde metamorphism; (2) N2 can be bound on organic matter.

The A1 inclusions occur in both the epizonal and anchizonal domain, while the A2 inclusions only occur in the epizonal domain (figure D4). In general, the proportion of non-polar species with respect to water decreases from the epizonal domain to the anchizonal domain. In addition, the portion of non-polar species other than CO2 also decreases from the epizonal towards the anchizonal domain. The salinity of the primary fluid in the anchizonal domain is rather low (between 0 and 3.5 eq.wt% NaCl), while in the epizonal domain the salinity ranges up to 17 eq.wt% NaCl. Salinity decreases from the epizonal towards the anchizonal domain. The increase of salinity with increasing metamorphism in the closed fluid flow system of the Ardenne-Eifel basin may be attributed to the hydrolysis of Cl-bearing silicates.

The aqueous/gaseous fluid inclusions (A1) in the anchizonal domain all show a limited variation in composition, bulk molar volume and phase-transition temperature. This suggests a homogeneous trapping of a single fluid above the immiscibility surface for mixed aqueous/gaseous fluids. In contrast, the inclusions in the epizonal domain show a wide range in bulk composition, a large spread of phase-transition temperatures within the same fluid inclusion assemblage and variable phase proportions at a wide range of temperatures. Total homogenisation is both to the liquid and gaseous state. Two possible explanations are suggested: (1) the mechanical mixing of immiscible fluids heterogeneously trapped in the two-phase region; the phase separation could have been triggered by strong pressure fluctuations during veining or could be induced by the higher salinity; or (2) mixing of fluids of different composition and possibly different origin.

In the secondary gaseous fluid inclusions (G1) different compositions can be recognised, ranging from CO2-N2-CH4 to pure N2(-CH4) (figure D4). Since all three-gas components are miscible at almost all conditions, it is fair to assume different origins for the gases. CO2 likely originated from the decarbonisation under low-grade metamorphic conditions. CH4 is thought to have been produced by the maturation of organic matter. N2 can be released during the maturation of NH4

+-bearing minerals or released from organic matter. The circulation of gaseous Variscan fluids is restricted to the epizonal domain. The proportion of N2 increases towards the centre of the metamorphic area.

The latest aqueous fluid type (W1) shows a decrease in salinity from the epizonal towards the anchizonal domain (figure D4). Here, chemical processes, described as acting during retrograde metamorphism (e.g. water loss by hydration reactions) could have caused the increase in fluid salinity of the secondary fluids.

Veins – fluid transport

The presence of non-polar species in the primary inclusions indicates an intense fluid-rock interaction. To define the nature of the mass-transport process (open or closed fluid flow system), a stable oxygen isotope analysis has been performed on vein-quartz and their immediate host rock (sandstone and pelite). The similarity of δ180 values between vein and host rock suggests a rock-buffered fluid system, resulting in isotopic equilibrium at peak metamorphic conditions between vein


46

and host rock (figure D5). Moreover, the differences in isotopic composition of the vein-host rock couple at different localities demonstrates that the isotopic equilibrium existed on a very small scale.

Diffusion is considered the dominant transport mechanism during the formation of the veins. This is supported by the presence of bedding-parallel dissolution seams within the sandstones. Because some of these dissolution seams also crosscut the veins, vein formation and dissolution are considered to have occurred simultaneously. The dissolution in the sandstone is considered the main source of the vein-filling material. A

diffusive closed fluid flow system as the main driving mechanism implies that lateral and temporal chemical variation observed in the fluid inclusions can only be attributed to fluid-rock reactions occurring in the direct vicinity of the veins and can thus directly be linked with the metamorphic grade of the host rock. The latter complies with the chemical evolution of the fluids across the metamorphic area of the High-Ardenne slate belt, with more mature fluids in the epizonal central domain in comparison with the anchizonal peripheral domain of the Ardenne-Eifel rift basin.

The closed fluid system is, moreover, supported by the presence of near-lithostatic fluid pressures that are most probably responsible for the fracturing and vein development (Kenis 2004, Kenis et al. 2002).

Veins– P-T conditions at time of veining

For the determination of the P-T conditions at the time of veining, the isochores of the fluid inclusions that occur in growth zones and thus certainly having a primary origin are used. In the anchizonal domain (Bütchenbach; see figure A1 – stop A for location) the isochores of the aqueous/gaseous inclusions are used. A homogeneous trapping of a single fluid is suggested. In the epizonal domain (Bertrix; see figure E1 – stop E for location) the mixed aqueous/gaseous primary fluid inclusions often decripitate before total homogenisation. Variable total homogenisation temperatures, moreover, suggest post-entrapment deformation of these inclusions, so that they cannot provide sufficient reliability for constraining the P-T conditions. Therefore, in the epizonal domain the isochores of the gaseous (A2) inclusions are used. The original P-T conditions of entrapment can, however, not be reconstructed from isochore evidence alone. The pressure and total homogenisation temperature serve only as the minimum constraint of entrapment condition. An independent geobarometer or geothermometer is required to determine the exact condition by intersection with the isochores.

In the case of the quartz veins in the High-Ardenne slate belt chlorite geothermometry has been used on chlorites within the veins. Microscopy suggests a syngenetic origin of both quartz and chlorite. At Bertrix (epizonal domain) the analysis shows that chlorites formed at a temperature of ca. 390°C. This temperature corresponds to the maximum temperature of metamorphism, constrained from mineral assemblages (Fielitz & Mansy 1999). It also complies with the temperature obtained from chlorites in the host rock, again suggesting a close association between vein filling and metamorphism and a thermal equilibrium of the fluids with the host rock. The latter association infers that the metamorphic temperatures may be used to constrain temperature conditions in the entire metamorphic zone (Bütchenbach). In both cases a variation in position of the maximum and minimum isochores is observed. This variation can be indicative of pressure variations or can be due to post-entrapment deformation. Microthermometric evidence (Kenis 2004) allow us to fairly assume

Figure D5 – δ18O (SMOW) of the quartz veins and their host rock at different localities (Kenis et al. 2005a).


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that the different position of the isochores reflect a difference in pressure at the time of trapping. The data can thus be applied to estimate pressure variations in the mineralising fluids.

The maximum and minimum fluid pressure in the anchizonal domain (Bütchenbach) correspond to values of ca. 190 MPa and ca. 100 MPa respectively. The maximum and minimum fluid pressure in the epizonal domain (Bertrix) corresponds to ca. 255MPa and ca. 115 MPa respectively. The difference in pressure indicates a maximum pressure fluctuation of ca. 140 MPa during veining. Considering a geothermal gradient of ca. 40°C/km (Helsen 1995, Schroyen & Muchez 2000), veining occurred at a depth ranging between ca. 7 and ca. 10 km. The maximum fluid pressure thus approximates lithostatic confining pressure. The minimum fluid pressure suggests that fluid pressure during veining varied between near-lithostatic and near-hydrostatic. Both the near-lithostatic maximum fluid pressure and the pressure fluctuation between near-lithostatic and near-hydrostatic, suggests that hydraulic fracturing is the responsible mechanisms for the formation of the veins.

Veins– reconstruction of the stress-state evolution

Mineralogical, microthermometrical and geochemical evidence demonstrate that veining occurred at peak metamorphic conditions in a closed fluid flow system in thermal equilibrium with the host rock. Veining occurred at the latest stages of the burial history of the Palaeozoic Ardenne-Eifel rift basin. Veining occurred at near-lithostatic fluid pressures by hydraulic fracturing and generated a regionally consistent, NNE-SSW trending, fracture/vein array (figure D6).

Figure D6 – Compilation of the orientation distribution of the bedding-normal veins in the High-Ardenne slate belt, based on data of Kenis (2004) and Van Noten et al. (2012). The lower-hemisphere equal-area projections show the original orientation of the veins prior to folding and reflect a consistent extensional stress field in the Ardenne-Eifel basin at the time of veining.


48

The regionally occuring bedding-normal veins, initially vertical at the time of vein formation, range in orientation from NE-SW in the SW and cHASB to NNE-SSW towards the NE (North Eifel – stop A). This structural change is probably related to a post-veining oroclinal bending of the slate belt during the main Variscan contraction because of the presence of a rigid basement buttress in the foreland (i.e. Brabant basement) (Mansy et al. 1999). Nothwithstanding this oroclinal bending, this consisten array of extension veins, i.e. mode I fractures, can be used as a regional paleostress indicator for the stress state during the extensional regime. The orientation of the principal stress axes is derived from the distribution of the extension veisn that are assumed to have formed perpendicular to the minimum principal stress σ3 and align according to the σ1-σ2 plane. Bedding-normal veins therefore reflect an triaxial Andersonian stress state that is characterized by a vertical σ1, corresponding to the overburden stress σV, and two well-defined horizontal principal stresses, σH and σh.

A stress-state reconstruction can be performed, very similar to the one elaborated for the BNVs in the pHASB (North Eifel – stop A). In the Bastogne case, a lithostatic load at ca. 10 km depth of ca. 265 MPa (σV) has to be considered. Again, a tectonic stress σT acting on the minimum principal stress σ3, both oriented NW-SE, is needed to sufficiently decrease the differential stress (figure A14(a) – stop A). At that stage, the fluid pressure of ca. 255 MPa is able to overcome T, allowing mode I fracturing in the extensional stress context, resulting in the BNVs (figure A14(b) – stop A).

Veins– vein spacing

In the Mardasson quarry it is very obvious that vein spacing shows a relationship with layer thickness. This is exemplary for all arrays of BNVs throughout the HASB (In the Rursee area (stop A) this is in particular apparent in the Wildenhof section (figure A3)). A comparative field-based (Van Noten & Sintubin 2010) study demonstrates that vein spacing is indeed strongly shows a linear correlation with layer thickness in thin competent sandstone layers (< 40 cm thick beds), embedded in a pelitic matrix, but becomes non-linear in thicker sandstone layers (> 40 cm thick beds). Vein spacing tends to increase to a maximum value becoming more or less independent of layer thickness (figure D7). The resemblance with fracture spacing suggests that in an unfractured rock vein saturation can occur. As with fractures, veins may inhibit the formation of new adjacent veins at a certain distance during initial development. High fluid pressures are responsible for vein nucleation but the stress state around the initial veins control the spacing pattern. At the stage of fracture saturation, high fluid pressures can still overcome the compressive stresses in between the saturated initial veins causing regularly spaced veins at distances smaller than those predicted at fracture saturation. The presence of crack-seal microstructures supports the idea that vein-filling material is weaker than the host rock, so that the initial vein will rather thicken by crack-sealing, than that new veins will form

Figure D7 – Vein spacing (in cm) versus layer thickness (in cm) relationship for quartz veins in thick metasedimentary multilayer sequences (Van Noten & Sintubin 2010), compared to published fracture spacing data in greywackes and in limestone. Fracture spacing in greywacke shows a comparable non-linear relationship (I: thickness of interlayers between competent beds). Spacing related to boudinage is much wider.


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near pre-existing ones. Whether a single consistent set of veins or crosscutting generations will develop, with only the oldest generation regularly spaced, depends largely on the consistency of the regional stress field (Van Noten & Sintubin 2010).

Veins– a proxy for high-pressure compartments

Vein quartz originated from fluids that were linked to the regional low-grade burial metamorphism (ca. 400°C) in the deepest part (ca. 10 km depth) of the Palaeozoic Ardenne-Eifel sedimentary basin (Kenis 2004, Kenis et al. 2005a). Veining occurred at near-lithostatic fluid pressures by hydraulic fracturing (Kenis et al. 2002) and generated a regionally consistent, NNE-SSW trending, fracture/vein array. At this depth, in the brittle-plastic transition zone, this hydraulic fracturing event is considered to be related to a change in tectonic regime (Sibson 2001) at the onset of Variscan orogeny. Transient co-seismic stress fluctuations (Nüchter & Stöckhert 2006, 2007) may eventually have caused the hydraulic fracturing.

Quartz veins serve as a proxy for high fluid pressures, enabling to outline high-pressure compartments in the metasedimentary sequence of the Palaeozoic Ardenne-Eifel basin. In this respect the Ardenne-Eifel region exposes an analogue of a ca. 340 Ma old, fractured reservoir, composed of sequences of interbedded high-pressure compartments (Hilgers et al. 2006).

Cuspate-lobate morphology – the geometrical characteristics

The second feature of the brittle-plastic structures is the cuspate-lobate geometry of the interface between materials of different composition/competency (figure D3). The layer-parallel dissolution seams, which crosscut the veins, follow this cuspate-lobate morphology, clearly indicating that the cuspate-lobate buckling occurred after veining and should be considered a separate – plastic – event. In addition, internal cleavage fanning inside pelitic layers of multilayer structures (figure D8) demonstrates that the cuspate-lobate morphology is a consequence of layer-parallel shortening,

occurring at the onset of the Variscan orogeny.

The development of this cuspate-lobate morphology of the interface between sandstones and pelites due to layer-parallel shortening complies with the kinematic definition of mullions as nowadays acknowledged by the geological community.

The presence of the mullions is, however, entirely constrained by the presence of the veins, although both structures are the result of separate events. Without the presence of the veins, no cuspate-lobate geometry will occur at the interfaces. The formation of the mullions is caused by a strong rheological difference between the vein quartz (deforming by dislocation creep) and the sandstones (deforming by solution-precipitation creep). Without this rheological difference no mullions would have formed. Mineralogical and microstructural evidence show that mullion formation occurred at very similar conditions as the vein formation (i.e. 350-400°C, ca. 260 MPa). The formation of the mullions at the sandstone-pelite interface is pinned by the pre-existing quartz veins.

Figure D8 – Divergent cleavage fan pattern in the pelitic parts of a multilayer structure (Boeur) (Kenis 2004).


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Cuspate-lobate morphology – the numerical model

The simple geometry of the mullions provides an unusually well-constrained set of initial geometries and boundary conditions. A simplified analytical model (Urai et al. 2001) suggests that the morphology of the mullions is a strong function of the stress exponent of the power-law creep equation of the sandstone. A more representative plane-strain geomechanical model using finite element techniques has been developed (Kenis et al. 2004). A viscoplastic formulation in which the strain-rate potential can be written as a function of equivalent stress, as implemented in the ABAQUS software package, has been used.

A parameter sensitivity analysis (Kenis et al. 2004) shows that the shape of the individual mullions is controlled by four parameters: (1) total layer-parallel shortening; (2) initial aspect ratio of sandstone/psammite segments before shortening; (3) stress exponent of the sandstone; and (4) the competence difference between vein quartz and sandstone. The curvature of the interface is primarily a function of the stress exponent.

The structural observations of a mullion in the field provide the opportunity to construct models specific to each mullion, using a suitably chosen set of parameters. The sensitivity analysis suggests that there might exist a unique set of parameters, which makes the model match all the observations related to specific mullions. This provides a tool to solve the inverse problem and determine the rheology of the sandstones.

Cuspate-lobate morphology – a palaeorheology gauge

To date, we have performed seven case studies of mullions observed in various outcrops (figure D9). Using the numerical parameter estimation method, in all cases the model converges to a single set of parameters, which define the flow law of fine-grained siliciclastic rocks, such as sandstone, at low-grade metamorphic conditions (350-400°C) and geological strain rates (Kenis 2004, Kenis et al. 2005b). We consistently find an approximately ten-fold contrast in strength between the sandstone and vein quartz (wet quartz) (figure D10), together with a stress exponent of 1.

Finite Strain ε∆σ (MPa)Undeformed Shape Deformed ShapeField observationa b c d e56

26

34

19

12

4

0.3

0.00

0.26

0.17

0.50

Strain vein 10%, n=1ps

Mardasson

Figure D9 – Best-fit solution of the caste study of a mullion at Bastogne Mardasson (Kenis 2004), showing the field observation (a), undeformed (b) and deformed mesh (c) of the numerical model and contour plots of the differential stress (d) and finite strain (e) of the mullion. The result of the numeric solution fits well with observations of the mullion shape.

This indicates that sandstone at low-grade metamorphic conditions deforms in a linear or Newtonian way. A Newtonian constitutive equation for the deformation behaviour of fine-grained siliciclastic rocks in the middle crust is in agreement with microstructural observations in the sandstone, indicating that solution-precipitation creep was probably enhanced by the presence of phyllosilicates within the fine-grained rocks. The absence of notable deformation by dislocation creep of quartz in


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the sandstone indicates that the differential stress remained too low for the deformation mechanism to be competitive with solution-precipitation creep. In contrast, the vein quartz is dominantly deformed by dislocation creep indicating high flow stress in these parts and causing an inhomogeneous stress field in the different lithologies, leading to the process of mullion formation at the onset of Variscan shortening (Kenis et al. 2005b).

As a consequence, the strength of polyphase quartz-rich rocks located in the middle crust is much lower than predicted by conventional models based on flow laws from dislocation creep (figure D10). Because fine-grained siliciclastic rocks control the rheology of the middle crust in many sedimentary basins, these results provide new quantitative parameters for geodynamic modelling in which a flow law for solution-precipitation creep is essential. Moreover, mullions occur worldwide in deformed sediments, and the method is therefore applicable to quantify rock rheology in other areas and geological settings offering further perspective for the quantification of rheological flow laws.

A KINEMATIC MODEL

The brittle-plastic structures in the cHASB can be interpreted as the reflection of the compressional tectonic inversion at the onset of Variscan orogeny in the deeper parts of the Ardenne-Eifel rift basin (figure D11). In this respect the mullions in the cHASB can be seen as the kinematic equivalent of the bedding-parallel veins in the pHASB (see stop A), both resulting of layer-parallel shortening.

A two-stage history can thus be reconstructed. During the latest stages of the basin development and burial (ca. 340 Ma ago), a fracturing/veining event occurred at low-grade metamorphic mid-crustal conditions (350-400°C; ca. 260 MPa) due to hydraulic fracturing in high fluid-pressure compartments. This hydraulic fracturing event seems related to a change in tectonic regime, heralding the Variscan orogeny. Transient co-seismic loading (Nüchter & Stöckhert 2006, 2007) may eventually have triggered brittle fracturing below the regional brittle-plastic transition. At the onset of the Variscan orogeny (ca. 330 Ma), layer-parallel shortening resulted in the formation of the cuspate-lobate morphology of the pelite-sandstone interface, i.e. mullion formation. The strength contrast between the vein quartz and sandstone controlled the shape of the mullion, which allows us to use the mullions as a palaeorheological gauge. The ongoing Variscan deformation (ca. 325 Ma – Sudetic stage) resulted in the folding and cleavage development, giving rise to the High-Ardenne slate belt.

Figure D10 – Crustal strength curve with indication of the strength of the sandstone/psammite with respect to the hydraulic fracturing event and to the mullion development. In the latter case the sandstone/psammite is much weaker than the vein quartz.


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CLOSING REMARKS

Nearly a century after the terms ‘boudin’ and ‘boudinage’ have been coined in the Bastogne area, the ‘original’ structures in the Ardenne-Eifel region turn out to be very particular, expressing a polyphase deformation history in brittle-ductile deformation conditions at the onset of the Variscan orogeny. Brittle deformation resulted from high fluid pressures causing hydraulic fracturing of the overpressured sandstone layers. Once formed, the quartz veins acted as ‘rigid’ bodies, causing buckling of the sandstone-pelite interfaces between the veins due to layer-parallel shortening, creating textbook examples of ‘mullions’. In this respect, the Ardenne-Eifel region may be considered an interesting natural laboratory to study the brittle-ductile deformation behavior of siliciclastic metasediments in mid-crustal tectonometamorphic conditions. Notwithstanding the fact that these particular structures are not related to the process of boudinage, historical justice has to be done to the original ‘boudins’ in the Ardenne-Eifel region by aspiration to create a geological heritage site at Bastogne.

Figure D11 – Comparison of structural evolution during the compressional tectonic inversion at the onset of Variscan orogeny in the pHASB (North Eifel; see stop A) and the cHASB (after Kenis 2004). In the Ardenne-Eifel basin a consistent array of bedding-normal veins originated in the latest burial stages, under the influence of a tectonic stress. After the tectonic switch, bedding-parallel veins develop in the pHASB (see stop A), while in the cHASB mullions develop (see also stop B).

Figure D12 – An original drawing of the brittle-plastic structures in the High-Ardenne slate belt by Stainier (1907).


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STOP E – Ardenne (Belgium) – Herbeumont – “Carrière de la Fortelle”: “Late-Variscan discordant veins in a destabilized slate belt”

by Hervé Van Baelen & Manuel Sintubin

INTRODUCTION

The Route des Ardoisières, along the valley of the Ruisseau d’Aise, some 5 km south of Bertrix, is at the core of the prosperous roofing slate industry in the 19th and early 20th Century, both in open pit mining and underground.

Exposing the core of the Eifel depression, these man-made exposures allow the study of a wide range of structural aspect, typical for a slate belt: from multiple cleavage development and kinking, progressive shearing, to discordant veins. At the Carière de la Fortelle we will discuss the kinematics of these discordant veins.

The mine of La Fortelle was in activity from 1849 till a first closure in 1874 and a second closure in 1886, due to collapse of a pillar. After resumption of exploitation the exploitation was finally closed in 1912 (Voisin 1987).


The abandoned La Fortelle quarry is located some 2 km north of Herbeumont, along the banks of the Semois river. The geographical coordinates of the outcrop correspond to 49°47’54” and 5°13’14”.


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THE EIFEL DEPRESSION

The HASB is composed of the Ardenne culmination, comprising the Lower Palaeozoic basement inliers along its backbone, and the Eifel depression, formerly known as the Neufchâteau synclinorium (Asselberghs 1946) (figure E1). The overall structure of the Eifel depression is a matter of controversy. Commonly, it is seen as an, N-verging isoclinal fold with a pervasive axial-planar cleavage development. Due to the very poor degree of exposure, the reconstruction of the overall structure was primarily based on stratigraphy (Asselberghs 1946). In the core of the Eifel depression, in the Herbeumont area (figure E2), the repetition of Siegenian 2 and Siegenian 3 rocks is commonly seen as the result of thrusting and/or isoclinal folding. The hinge zones of these isoclinals folds have, however, never been observed properly. This is partly due to the difficulty of structural mapping in the area, primarily because of the parallelism between bedding and cleavage, making the recognition of the structural polarity nearly impossible. The structure is furthermore obscured by the superposition of kink bands and multiple crenulation cleavages.

To the south, the Eifel depression is bordered by the ‘Herbeumont shear zone’ (Schavemaker et al. 2012), a weakly S-dipping thrust bringing the Givonne culmination, bearing the Lower Palaeozoic Givonne basement inlier in its core, on top of the Eifel depression (figure O2). Originally, this structure was described as the faille de Herbeumont, the Herbeumont fault (Fourmarier 1914), primarily based on intensely disrupted and folded zones, brecciation, veining. Revisiting the type locality at Herbeumont resulted in a novel kinematic interpretation of a polyphase progressive top-to-the-north brittle-plastic shear zone, developing in the late stages of the Variscan orogeny (Schavemaker et al. 2012).

Figure E1 – Geological map of the western part of the HASB, exposed in the Ardennes (after Asselberghs 1946). The Herbeumont area (indicated by a square – see figure E2) is located just south of the region in which mullions occur (Kenis 2004). It is, though, still located in the epizonal metamorphic zone (Fielitz & Mansy 1999) (Van Baelen 2010).


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THE HERBEUMONT AREA

As already mentioned, the Herbeumont area exposes the core of the Eifel depression in the direct footwall of the Herbeumont thrust (figure E2). Stratigraphically this core zone is characterized by the repetition of Siegenian 2 and Siegenian 3 rocks. Asselberghs (1946) relates most of the repetitions to reverse faults. In this respect, the La Fortelle quarry, exposes the southern fault contact of the so-called Bande de la Fortelle (figure E2). The overall structural grain is nearly EW. In the core of the Eifel depression, the foliation has an overall attitude of 170/40-45.

The Siegenian rocks, exposed in the Herbeumont area, corresponds in age to Lochkovian to Pragian formations in the present-day stratigraphical framework (Bultynck & Dejonghe 2001). Siegenian 1 (Sg1), also the Anlier facies (Asselberghs 1946), consist of finely laminated, blue to black pelites and light-coloured sandstones. The Anlier facies may be considered the southern equivalent of the Mirwart or Bois d’Ausse formation (Bultynck & Dejonghe 2001), Lochkovian to Pragian in age. Siegenian 2 (Sg2), also the Bouillon facies (Asselberghs 1946), consists of blue calcite-rich pelites. The resulting slates are of bad quality for roofing slates due to the high carbonate content. This facies can be considered the lateral equivalent of the Villé formation (Bultynck & Dejonghe 2001), Pragian in age (see also stop C). Siegenian 3 (Sg3), also the Neufchâteau facies (Asselberghs 1946), primarily consist of blue or black pelites, with exceptionally silty and sandy intercalations. The cleavage is nearly perfect and bedding is very hard to recognize. Therefore, this facies has been exploited for roofing slates along the route des Ardoisières (figure E2). Siegenian 3 most probably corresponds to the La Roche and Pèrnelle formation (Bultynck & Dejonghe 2001), late Pragian to early Emsian in age (see also stop A and stop C).

Figure E2 – Geological map and NS cross-section of the Herbeumont area (after Asselberghs 1946). Indication of the Carrière de la Fortelle. Sg1: Siegenian 1; Sg2: Siegenian 2; Sg3 : Siegenian 3 (Van Baelen 2010).


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LA FORTELLE QUARRY

The slates exploited in the southern part of the quarry were of bad quality due to the high carbonate content and the intensive kinking (Voisin 1987). Most probably, the underground mine never reached the underlying slates, of good quality, which are exposed in the northern part of the open pit quarry. The former slates most probably belong to the Siegenian 2, while the latter slates belong to the Siegenian 3. Asselberghs (1946) situates a reverse fault between both lithologies (figure E2).

Based on proper mapping of both lithologies, Van Baelen (2010) could deduce that the contact is EW-trending and dips 30 to 35° to the south, largely coinciding with the overal structural grain in the area. Since this lithology contrast most likely represents the original stratigraphy, foliation is considered subparallel to bedding. Neither stratigraphical, nor structural polarity could be identified in the quarry.

The rocks, exposed in the southern part of the quarry (figures E3 & E4), can be described as bluish-black or bluish-grey, coarse-grained, mica-rich slates or laminated psammites. On foliation planes small mica flakes are visible with the naked eye. The spacing of the foliation is generally larger than 1mm, up to a few centimeters. The planes are spotted, wavy and anastomose. In thin section, compositional differences can be observed between quartz-rich microlithons and mica-rich cleavage domains. In general, no lineation can be observed in the foliation plane. The rocks

underlying the contact (figures E3 & E4), are finely laminated, bluish-black slates with a very well developed penetrative cleavage. The foliation is planar and does not contain mica flakes, visible with the naked eye. Contrary to the spaced and anastomozing cleavage in the coarse-grained rocks, a continuous (spacing less than 1 mm) and planar cleavage is present in the fine-grained rocks, often marked with a NS-trending lineation.

Figure E3 – Simplified geological and topographical map of the La Fortelle quarry (Van Baelen 2010).

Figure E4 – Synthetic section across the La Fortelle quarry with projection of the different outcrops. White lines indicate the general attitude of the veins (Van Baelen 2010).


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According to Asselberghs (1946) the lowermost lithology corresponds to the Siegenian 3 and the uppermost lithology corresponds to the Siegenian 2. Both lithologies are supposed to be separated by a fault. Most probably the presence of quartz veins were interpreted as evidence for the fault. At the level of the lithological boundary traced, there is no clear evidence of a reverse fault and the décollement horizons observed seems to represent small-displacement, top-to-the-south, extensional structures. In the latter case, the stratigraphic polarity would be inverse. Because of the parallelism between bedding and cleavage nothing can be said, though, about the structural polarity.

In general an evolution from steeply-inclined foliation below the level with the veins to gently-inclined above the level of the veins can be observed (figure E4). Below the level of the veins, foliation attitude is rather constant and disturbances are absent. This may be a reason why the slates were worth to be exploited below the level of the veins. The anomalous attitude of the foliation above the level of the veins will be related to the development of the veins itself.

The La Fortelle quarry is subdivided into 5 outcrops (figures E3 & E4). During the field trip we will focus on outcrop [D], exposing the discordant veins. A very short summary is given of the main observations on the other outcrops.

La Fortelle – outcrop [A]

Outcrop [A] displays an evolution from monoclonal kink bands in the north, to symmetric or slightly asymmetric conjugate systems with mostly parallel, but exceptionally crossing kink axis in the central sector (part I & II on figure E5), to finally monoclonal kink bands in the southern sector (part IV on figure E5). In between (part III on figure E5) irregular and open chevron-like kink bands occur. Parallel to this evolution a gradual change in degree of shortening can be observed (figure E5) with a maximum towards part II. A décollement is supposed between part I and II because of the sudden change in foliation-parallel shortening. Based on the classical paleostress interpretation of kink bands (cf. Debacker et al. 2008 and references therein), Van Baelen (2010) suggests that in part I and IV the maximum principal stress σ1 is more (> 48°S), respectively less (< 42°S) inclined than the foliation and at 90° to the intermediate principal stress σ2, parallel to the kink axis. For the conjugate systems, in part I and II, the maximum principal stress σ1 dips ca. 50°S. The maximum principal stress σ1 remains more or less in the plane of foliation.

La Fortelle – outcrop [B]

An paleostress analysis of the kink bands in outcrop [B] provides a comparable result. The maximum principal stress is inclined to the south and varies from more to less inclined than the foliation.

La Fortelle – outcrop [C]

Most of the structures observed in outcrop [C] are difficult to separate in time from each other, because of unclear crosscutting or overprinting relationships. Therefore a rather progressive deformation history, including faulting, kinking and veining, is preferred (Van Baelen 2010). Contrary to the previous outcrops, the faults observed do crosscut foliation locally. Where they crosscut the foliation a fault gouge is present, consisting of reworked host-rock fragments. Two crosscutting relationships can be recognized: one with a foliation that is more inclined to the south than the fault; and the opposite one with a foliation less inclined to the south than the fault. In the former case, the foliation is always kinked; in the second case foliation is generally unaffected by kinking, but sometimes displays microfracturing interpreted as boudinage, thus inferring a NS-extension of the

Figure E5 – Summary of the types of kink bands in the different parts of outcrop [A], the associated degree of shortening and the theoretical derived principal stresses (Van Baelen 2010).


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foliation. The increase of foliation-parallel shortening, as expressed by kink bands, close to the faults and the increase in intensity and complexity of the kink bands, suggest that kink bands and faults are kinematically linked. A number of veins are present. Their relative timing with respect to the other structures is ambiguous.

A reconstruction of the original juxtaposition of the different fault blocks and an attempt to unravel the order of faulting is presented on figure E6. The restoration of the fault blocks is based on the changing crosscutting relationships along the faults. The evolutionary model fits in a top-to-the-south kinematics, suggesting an overall weakly S-dipping detachment zone. In the model, the main fault plane is considered to correspond to the plane of zero length change of the finite strain ellipsoid. Material planes that are inclined more to the south are subsequently in the shortening sector (kink bands), while material planes that are inclined less to the south are in the extension sector (boudinage). EW-trending veins are in this context optimally oriented for further opening. This conceptual shear model suggest a moderately south-dipping maximum principal stress, i.e. parallel to the EW-trending veins and optimally oriented for activation of the discrete faults. This

orientation is not at all in accordance with the orientation of the maximum principal stress for the kink bands in the different fault blocks.

La Fortelle – outcrop [E]

Outcrop [E] has not been studied in detail because the south-dipping vein array crosscutting the outcrop is inaccessible for its major part. It is not clear whether it concerns one branching vein or several smaller aligned veins. The veins are up to 30 cm thick and have numerous offshoot veins. They seem very irregular and complex in the third dimension. The host rock is entirely composed of the fine-grained, black slates. A few meters below the vein array, the foliation is steeply dipping to the south with a relatively constant dip(figure E4). Locally widely spaced kink bands are present. The slates overlying the vein array are intensely disturbed by kink bands, veins and a secondary foliation (figure E4). The foliation is generally less dipping to the south than the foliation below the vein array.

La Fortelle – outcrop [D]

The outcrop [D] consists of a ca. 20 m long exposure at the back wall of the quarry. The lithological boundary can be traced across the outcrop. In the coarse-grained rocks, veins are rare or limited to thin, generally planar, south-dipping veins. In the fine-grained slated the veins are thicker and have very irregular shapes. Two detachment horizons can be identified in the outcrop. A first detachment

Figure E6 – Tentative evolutionary model for the décollement structure in outcrop [C] (Van Baelen 2010).

Figure E7 – Line drawing of the outcrop [D]. The outcrop is subdivided in the northern, central and southern zone (Van Baelen 2010).


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is present in the southern zone and appears to crosscut the foliation at a very low angle. It is fair to assume that this detachment is the northward prolongation of the detachment zone in outcrop [C] (fault A on figure E6). A second detachment horizon separates the northern and central zones, dipping somewhat steeper (218/48) than the overall foliation (180/36), but subparallel to a secondary foliation, characterizing

the northern zone. The attitude of the foliation gradually changes from south-dipping in the northern zone, to subhorizontal in the central zone, to again south-dipping in the southern zone. A large number of veins with different geometrical characteristics is exposed in outcrop [D]. With the exception of a set of offshoot veins in vein array 1 and 2, most vein-wall segments are EW-trending, but with variable dips. The same evolution in dip as described for the foliation can be observed for the vein-wall orientation. The northern and southern zones have south-dipping veins that are mostly planar, thin and steeper than the foliation. In the main portion of the central zone, the veins are horizontal to weakly dipping. The veins are lens-shaped and the crosscutting relationship with the foliation is rather complex (figure E8). These veins have typical blocky infill, with quartz as major component.

As exemplary for the discordant veins we focus on vein array 1 (figure E8). The vein is a more than 2 m long, lens-shaped vein with a maximum thickness of 15 cm in its central part. Its southern extremity is not exposed. The vein is composed of blocky quartz, but in the southern exposed part chlorite, ankerite/Fe-dolomite and feldspar become dominant, especially at the vein walls. A few offshoot veins are present. Concerning the vein wall-foliation crosscutting relationship only one internal switching point (ISP) is observed on the upper side of the main vein. This ISP marks the transition of a north-joining vein-wall segment, north of the ISP, to a south-joining vein-wall segment, south of the ISP. The exposed part of the bottom side of the vein entirely falls within a north-joining vein-wall segment. An ISP can be inferred in the unexposed part of the vein wall. Generally the vein wall is ca. 15° more inclined to the south than the foliation, expect between the northern vein tip and the ISP, where the situation is opposite. A fanning of the foliation accommodated by a change in foliation thickness can be observed at the northern vein tip. At the direct contact with the vein, slates often display small-scale kink bands or crenulation cleavages. Conjugate systems and monoclonal sets of kink bands have been observed at the lower, northern side of the vein, while the crenulation cleavage has been observed between the northern vein tip and the IS, at the upper side of the vein.

Translated to the initial fracture, it is clear that the south-joining vein-wall segment, south of the ISP, and the north-joining vein-wall segment at the bottom side of the vein are the ‘regular’ parts of the

Figure E8 – Line drawing of central zone of outcrop [D] (see figure E7) with indication of the different vein arrays (Van Baelen 2010).

Figure E9 – line drawing of vein array 1 (figure E8) in the central zone (figure E7) of outcrop [D]. Dots on the vein wall indicate the internal switching points (ISPs) (Van Baelen 2010).


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vein. The north-joining vein-wall segment between the northern vein tip and the ISP on the upper side of the vein is the exception. The foliation also becomes more inclined above this segment of the vein wall, contrary to all other parts where the foliation is gently inclined to the south. It is clear that this vein is not simply the result of opening of a fracture, since the foliation planes truncated at the bottom side of the vein cannot directly be continued at the upper side of the vein, expect south of the ISP. The initial situation of the fracture should have a symmetrical vein wall-foliation crosscutting relationship, in which the ISPs coincide with the vein tips (figure E10(a)). It is suggested that the ISP located at the northern fracture tip shifted down parallel to the upper vein wall (figure E10(b)) to create the final geometry observed in the field (figure E10(c)). This can be achieved by foliation-parallel shearing in the slates north of the initial fracture, which will be transferred over the vein tip from the bottom vein wall to the upper vein wall. The transported vein wall segment remains a north-joining vein-wall segment. The position of the northern vein tip is believed to be temporary, analogously to the rear wheel of a caterpillar-traced vehicle moving forward, where the caterpillar itself serves as analogue for the mobile vein wall. What exactly happens at the southern extremity of the vein is unknown in the present case, since it is not exposed. In the model it is suggested that the southern fracture tip propagates southward parallel to the foliation, so that the displaced hanging wall is floored on this shear plane. The amount of displacement taken up by this overall shearing is estimated at ca. 2 m.

So, the model proposed is that the vein initiated as a mode I fracture, oriented parallel to the maximum principal stress. A plane of weakness oriented oblique to the principal stresses is optimally oriented to be activated as shear planes. Foliation-parallel shear at the northern vein tip can be considered the active mechanism. The overlying material is transported over the descending vein tip into a ‘shadow zone’ behind the vein tip (figure E10).

Top-to-the-south extensional kinematics

The structures analyzed, namely the EW-trending veins, kink bands and faults, all postdate the development of the south-dipping, pervasive cleavage, which reflects the main, contraction-dominated stage of the Variscan orogeny. It is therefore fair to assume that all structures are late- to post-orogenic with respect to the Variscan orogeny.

The first structure that develop, is a subvertical to south-dipping set of thin, planar mode I fractures – the parental cracks – generally making an angle of 10 to 30° with the pre-existing south-dipping cleavage (figure E11). These fractures are supposed to have formed at low differential stresses and high pore fluid pressures (see figure A13 – stop A). Failure was induced by hydraulic fracturing, induced by fluid pressures exceeding the subhorizontal minimum principal stress, oriented perpendicular to the vein walls.

During subsequent deformation, the foliation was activated as shear planes with a top-to-the-south kinematics, so that the vein shape modified. Vein-shape modification is characterized by a component of vein-parallel shear and a component of opening. Their rheological weak behavior suggests that they were not (entirely) filled at the moment of vein-parallel shearing. It is furthermore

Figure E10 – Kinematic interpretation of the final vein-foliation geometry of vein array 1 (figure E9), central zone, outcrop [D]. (a) initial mode I ‘parental crack’. (b) evolving foliation-parallel shear. (c) final configuration (see figure E9). Dots represent ISPs. Black arrows the orientation of the principal stresses. Grey zone is zone affected by shearing (Van Baelen 2010).


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suggested that this transition from tension failure to foliation-parallel shear is related to an increase in differential stress. Foliation-parallel shearing is favored when the maximum principal stress is oriented oblique to the foliation plane, so that there is no need for a reorientation of the stress field with respect to the initial fracturing, i.e. maximum principal stress oriented parallel to the vein wall.

For another end-member type of veins (not discussed in field guidebook), the foliation-parallel shear zone crosscuts the vein (figure E11). The vein, always located in the shortening sector of the incremental strain ellipsoid, is shortened by folding, kinking or fracturing. These veins are generally thin veins, and their rigid behavior suggests that they were filled at the moment of vein-shape modifications.

The kink bands are supposed to have developed during a late stage or posterior to the vein-shape modifications. A kinematic model, explaining the generation of the different types of kink bands in one single, progressive top-to-the-south shearing, infers a south-dipping maximum principal stress, inclined more steeply than the south-

dipping cleavage, which corresponds to the orientation of the maximum principal stress causing the parental cracks.

Most kink bands predate the brittle reactivation of the foliation as a fault plane. These south-dipping faults also have a top-to-the-south kinematics (e.g. figure E6) and infer a subvertical maximum principal stress. Locally these faults represent narrow zones of intensely shortened kink folds. The faults represent the last observed step in a progressive deformation history, which is governed by top-to-the-south kinematics and in which the presence of the cleavage as initial anisotropy played a crucial role.

WHAT DOES THE QUARTZ TELLS US?

The mineral assemblage observed in the veins (quartz, ankerite/Fe-dolomite, feldspar, chlorite, muscovite) is no different than the bulk mineral assemblage of the Lower Devonian host rocks, and even the Lower Palaeozoic basement rocks. This mineralogical homogeneity does not allow to conclusively identify a potential source for the vein infill. The first precipitate in the discordant veins (type (b) on figure E11) is ankerite/Fe-dolomite and feldspar in minor amounts, which crystallized on the vein walls. Precipitation of both minerals was restricted to the initial stages of infill. Afterwards quartz precipitation became dominant. Due to vein-parallel shearing the rim of carbonate and feldspar crystals locally detached from the vein wall and accumulated as fragments within the vein infill. Both feldspar and ankerite/Fe-dolomite show fracturing and intracrystalline deformation by twinning. Ankerite/Fe-dolomite displays a sweeping extinction related to the genesis of saddle

Figure E11 – General kinematic model for the development of the discordant veins in the well-foliated rocks in the Herbeumont area, initiated as a ‘parental crack’/’precursor vein’ that is modified due to foliation-parallel top-to-the-south shear (Van Baelen 2010). (a) foliation-parallel shear continuing across the vein. (b) foliation-parallel shear is transferred across the vein by vein-wall parallel shear. Dots represent ISPs. Grey zones indicate zones of shearing. Vein array 1 is exemplary of (b).


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dolomite (Radke & Mathis 1980). Its clear growth direction from the vein wall to the vein center, combined with its peculiar crystal structure is a strong argument for crystal growth in an open fluid-filled cavity. Blocky quartz further filled the cavity, but also underwent deformation as evidences by the microstructures. Quartz clearly shows evidence of recovery (undulose extinction, subgrains, deformation bands and lamellae) and dynamic recrystallization (bulging, subgrain rotation). Also pressure dissolution, shear bands and brittle fracturing of quartz has been observed.

During and after sealing of the vein, the infill thus underwent crystal-plastic deformation, being indicative of relative high temperature conditions (300 to 400°C) and relatively low differential stress conditions. The twins in dolomite indicate temperature conditions above 300°C (Passchier & Trouw 2005). In quartz bulging has been identified as the most important dynamic recrystallization process, resulting in highly serrated grain boundaries and locally core-and-mantle microstructures, which may also indicate subgrain rotation recrystallization. Both are indicative for high temperature conditions (Stipp et al. 2002). An evolution towards planar microstructures in quartz and brittle fracturing in ankerite/Fe-dolomite and feldspar indicates a transition towards lower temperature conditions and/or higher strain rates and differential stresses (Blenkinsop 2000). In general, the range of microstructures observed are all indicative of deformation at brittle-plastic conditions.

On figure E12 the grey path represents a possible evolution of the discordant veins (type (b) on figure E11). In the initial stages of vein development a path of combined opening and shearing is followed. Sealing is, though, incomplete. First, ankerite/Fe-dolomite is precipitated along the vein wall. Locally this rim is detached due to vein-parallel shearing. Subsequently blocky quartz precipitation becomes dominant in the fluid-filled cavity. At the moment that sealing is almost complete, planar microstructures may have developed in the completely sealed parts of the vein during progressive increments of deformation. Since planar microstructures are superimposed on crystals with a sweeping extinction, they are probably posterior to recovery and dynamic recrystallization. However, recrystallization along grain boundaries affected by planar

microstructures indicates that recrystallization went on after shear band development. Finally, vein-parallel cracks fill with euhedral quartz, and vein-fragments in fault breccias are overgrown with clear quartz. These quartz-filled cracks are mode I cracks indicating that the maximum principals stress was still oriented at small angles to the vein wall (figure E10). Contrary to the high differential stresses suggested by the planar microstructures, a return to lower differential stresses and higher fluid pressures may be expected to create these tension microcracks. The euhedral infill, completely free of microstructures, implies that the infill did no longer undergo high temperatures or pressures.

Planar microstructures can be found in all veins independently of their thickness. They seem preferentially present in the larger crystals, while they are lacking in the smallest crystals and the host-rock grains. Based on an orientation analysis of the planar microstructures in quartz a potential technique has been developed to measure local shortening directions in thin sections. An extensive mapping of the shortening directions has demonstrated that the discordant veins (type (b) on figure E11) usually displays a shortening direction oblique or at high angle to the vein walls, which is thus perpendicular to the resultant extension direction related to the combined opening and shearing (figure E10). A shortening of 7 to 10% is estimated based on the microstructural interpretation.

Figure E12 – Schematic representation of the evolution of opening, vein-parallel shearing and vein infill of both end-member vein types (see figure E11) (Van Baelen 2010). The example presented during the field trip (figures E9 & E10) is of type (b).


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Although this value may be an underestimation, it is still of a completely different order than the large strains suggested by the kinematic model (figure E11). Therefore, the planar microstructures are believed to be kinematically related to the development of the veins, but probably overprinted in a late stage of vein-shape modification. The higher strain rates suggested by the planar microstructures may potentially represent high peak stresses related to the seismic cycle (e.g. Birtel & Stöckhert 2008).

The high degree of deformation of the vein quartz indicates a low preservation potential of the fluid inclusions to allow a characterization of the mineralizing fluids and the P-T trapping conditions. Two distinct populations of fluid inclusions could be identified: re-equilibrated inclusions and inclusions related to deformation (Van Baelen et al. 2009). The explosion structures observed in the re-equilibrated zone, indicate that earlier trapped inclusions had higher densities than the external fluid present during re-equilibration. This change in fluid density is related to a drop from (sub-)lithostatic fluid pressures to (supra-)hydrostatic fluid pressures and is classically interpreted as caused by rapid unroofing. There are, though, no indications for such rapid uplift in the HASB. Therefore, this isothermal decompression (figure E14) is interpreted as caused by the evolution of a closed fluid system, with (sub-)lithostatic fluid pressures, to a more open fluid system, with (supra-)hydrostatic fluid pressures. It is unknown whether this transition is related to a short-lived, possibly cyclic process acting at a local scale or a long-lived transition acting on a regional scale. A population of rectangular inclusions can be related to the deformation of the vein quartz. The systematic orientation of the inclusion walls may possibly indicate that the inclusions formed as cavities during glide along the positive or negative rhomb planes in the quartz crystals. These inclusions are often aligned in trails, resembling fluid inclusion planes. These rectangular inclusions necessitate high resolved shear stresses along the crystallographic planes, which infers high differential stresses and low – (supra-)hydrostatic – fluid pressures. The latter has been confirmed by the microthermometric analysis of the rectangular inclusions. A final population of fluid inclusions, occurring in the quartz of the late cracks, shows evidence of necking down, occurring most probably during cooling after or during heterogenization of the initially trapped fluid.

Figure E13 – Mapping of the shortening bisector on the different vein arrays on outcrop [D]. The shortening direction are parallel to the vein wall in the case of thin veins without supposed vein-parallel displacements, and generally oblique or at high angle in the case of thick veins with supposed vein-parallel shear (e.g. vein array 1 – see figure E9) (Van Baelen 2010).


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Based on the isochores and a number of assumptions with respect to the overall geothermal gradient and the maximum temperature conditions, the P-T conditions of fluid trapping can be tentatively derived (figure E14). In the case of the re-equilibrated fluid inclusions (figure E14(a)), the delimited P-T field indicates that rather low pressures (60-170 MPa) existed at relative high temperatures (300-390°C). This is in agreement with the suggested re-equilibration due to an isothermal decompression of about 100MPa. The internal overpressure, resulting in the explosion structures, may result from a pressure drop of the external fluid pressure relative to inclusions initially trapped at lithostatic fluid pressures. For the rectangular inclusions, related to deformation, a similar P-T field can be determined (270-390°C and 60-220 MPa). These fluid inclusions, related to the planar microstructures, thus still indicate rather high temperatures. The homogenization temperatures suggest rather low, sub-lithostatic fluid pressures during trapping, especially in the lower temperature range. This is in agreement with the related microstructures (deformation lamellae and shear bands), typical for high differential stresses (Blenkinsop 2000), thus low fluid pressures (Sibson 2000; see also figure A13 – stop A).

In general, the temperature, pressure and composition of the fluid inclusions correspond to the overall fluid evolution within the HASB during the Variscan orogeny. An evolution from gaseous CO2-N2(-gas) and aqueous-gaseous

H2O-NaCl(-gas) to pure H2O-NaCl inclusions has indeed been observed. This evolution is interpreted as an evolution from peak metamorphic conditions, with the release of nitrogen from organic matter or from the maturation of NH4

+-bearing minerals (Kenis et al. 2005a) (see also figure D4 – stop D) to a pure, low-saline H2O-NaCl fluid system, marking the onset of retrograde conditions.

A KINEMATIC MODEL

All structures described can be fitted in one progressive deformation history which is governed by a subvertical compression and a subhorizontal, NS-trending extension. The presence of a south-dipping foliation played a crucial role in the eventual expression of the deformation in these brittle-plastic, mid-crustal deformation conditions, which systematically shows top-to-the-south kinematics. Geodynamic implications from kinematic models, the microstructural analysis and microthermometry all document a transition from low differential stresses accompanied by high fluid pressures to high differential stresses and low fluid pressures and ultimately a return to low differential stresses and relatively high fluid pressures (figure E15(g)). The deformation history most probably occurred at temperatures of 300 to 400°C and at depths of 6 to 10 km, with the exception of the local and final overgrowth of quartz, occurring at decreasing temperatures. The peak differential stresses, and lower fluid pressures, occurred during the development of the planar microstructures and possibly the development of the kink bands. These planar microstructures may very well be related to high differential stresses in a context of co-seismic loading and post-seismic creep in the brittle-plastic regime (e.g. Birtel & Stöckhert 2008). The faulting, interpreted as the last

Figure E14 – Representation of the isochores of the aqueous H2O-NaCl inclusions in a P-T diagram, with indication of the – lithostatic and hydrostatic – geothermal gradients of 30°C/km and 50°C/km and the maximum pressure and temperature condtions (Kenis et al. 2005). (a) re-equilibrated fluid inclusions, with arrow indicating isothermal decompression. (b) fluid inclusions related to deformation. Dark domains indicate possible P-T trapping conditions (Van Baelen 2010).


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important step in the deformation history, may thus be the last, brittle, expression of this extensional deformation.

Late-orogenic destabilization of a slate belt

The south-dipping foliation corresponds to the Variscan cleavage, representing the contraction-dominated main stage of the Variscan orogeny in the HASB. The reconstructed geodynamic model must therefore be placed relatively late with respect to the Variscan deformation history in this southern part of the Ardenne allochthon, or even post-date it, when the Variscan cleavage became unstable.

Figure E15 – Geodynamic overview of the deformation history (time, from left to right, is not to scale) (Van Baelen 2010). (a) kinematic model for the thin planar veins (end-member type (a) in figure E11). (b) kinematic model for the thick, discordant veins (end-member type (b) in figure E11). (c) development of planar microstructures in the vein quartz. (d) development of the kink bands in outcrop. (e) late activation of foliation as fault. (f) tentative evolution of deformation conditions (dots) plotted on evolving crustal strength curves (curves taken from Sibson (1983) for 30°C/km). (g) tentative evolution of differential stress and fluid pressure (axes not to scale).


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Despite the fact that the veins studied have complex geometries and vein wall-foliation relationships, their initial shape can always be restored to EW-trending, steeply south-dipping to nearly vertical ‘parental cracks’/’precursor veins’ with high aspect ratios, crosscutting the cleavage at an angle of 10 to 30° (figure E16(d)). The veins are locally arranged in south-dipping (and possibly north-dipping) en-enchelon vein arrays. They are interpreted as tension fractures, formed by hydraulic fracturing in a regime of low differential stress, with the maximum principal stress oriented sub-parallel to the veins (figure E16(d)-inset). Based on the drastic change in this specific attitude of the maximum principal stress to near vertical and the low differential stresses, it is postulated that the veins formed closely after the tectonic inversion from compression to extension (see figure A13 – stop A). When tectonic compression further decreased and became significantly smaller than the vertical overburden load, thus increasing differential stress, vein-shape modification was initiated due to the activation of the foliation as shear zones (figure E16(e)). Comparison of initial and final configuration in all kinematic models proposed (figure E11), suggests that the bulk strain accommodated by this event is a subvertical shortening and a NS-trending extension. The regional occurrence of comparable veins in the Eifel depression indicates that this destabilization of the Variscan slaty cleavage may have resulted in a regional NS-extension and thinning of the Pragian rocks in the core of the slate belt. This may have had a profound influence on the overlying rocks, which must have accommodated a comparable amount of strain. It is therefore speculated that the overlying rock pile may possibly have accommodated the strain by non-localized flexures at first (figure E16(e)) and upper-crustal normal faulting during protracted extension (figure E16(f)). These normal faults may be rooted in the brittle-plastic detachment zones, as materialized by the particular vein occurrences in the Eifel depression. Faulting along these upper-crustal normal faults may eventually have generated transient high differential stresses or high strain rates, recorded in the planar microstructures superimposed on the vein quartz during the late stages of vein development (cf. Birtel & Stöckert 2008). The kink bands also suggest a sequential process and possibly high strain rates and may also have formed at these conditions or later, but definitively prior to brittle activation of the foliation as discrete fault planes.

The regional occurrence of these discordant veins in the Eifel depression suggests that this extensional destabilization is at least regional, even reflecting a belt-wide tectonic inversion (see also stop C). The low-permeability of the slates and the absence of through-going high-permeability structures may have favored the development of en-echelon small-displacement, fracture networks. It is proposed that when extension progressed, foliation-parallel shear zones developed linking up the individual deforming veins. In this configuration, initial parental cracks evolve to dilation jogs (see also stop C). Such structures become extremely important as permeability paths, since they are intermittently opened, even in conditions in which tension fracturing is normally replaced by frictional sliding. The presence of the cleavage anisotropy, optimally oriented for activation, may thus been crucial for the continuous combined opening and shearing of the veins, even when differential stresses started to increase after the tectonic inversion. This may explain why the discordant veins in the Herbeumont area underwent a progressive deformation during a much longer period of time than the two major vein generations formed at the onset of Variscan orogeny (see stop A, B and D). The timing of the late-orogenic destabilization of the slate belt is difficult to assess. The depth obtained from microthermometry (6 to 10 km) does not significantly differ from the maximum burial depth, inferred for the cHASB (maximum 10 km; Kenis et al. 2005a). Uplift did not yet have a major influence. Deformation must therefore have occurred relatively short after the contraction-dominated stage of the Variscan orogeny in the HASB. The main contraction at the height of the HASB occurred in late Viséan-Namurian times (ca. 325 Ma), during the so-called Sudetic stage of the Variscan orogeny (Fielitz & Mansy 1999). A similar geodynamic evolution has been described in SW England by Shail & Leveridge (1997), also part of the Rhenohercynian zone (figure O3). Structures related to the post-convergence deformation, occurring from the latest Carboniferous till the earliest Permian (ca. 305-295 Ma) (Shail & Leveridge 2009), show a lot of similarities with the structures


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observed in the Herbeumont area. They are characterized by zones of distributed shear, detachments and later high-angle normal faults, all reflecting top-to-the-south kinematics. Therefore, we believe it is fair to assume that the destabilization of the HASB may have taken place around 300 Ma, possibly coinciding with the ongoing contractional deformation at the Variscan front (known as the Asturian stage).

Figure E16 – Integrated geodynamic model and deformation history as recorded by the quartz veins in the Herbeumont area (Van Baelen 2010). (a) Sedimentation in the Ardenne-Eifel rift basin (ca. 400 Ma) (E: Emsian; Sg: Siegenian). (b) veining during compressional tectonic inversion at the onset of Variscan orogeny (see stops A, B & D). (c) main stage of Variscan contraction in the HASB (ca. 325 Ma). (d) veining associated with the extensional tectonic inversion during the late stages of Variscan orogeny (ca. 300 Ma). (e) increasing differential stress, vein-parallel shearing and vein-shape modifications; resulting in bulk horizontal extension and vertical shortening. (f) normal faulting, rooted in brittle-plastic detachment zones (circles: earthquake hypocenter).


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References Adams, R. & Vandenberghe, N. 1999. The Meuse valley section across the Condroz-Ardennes (Belgium) based on a

predeformational sediment wedge. In: Palaeozoic to Recent tectonics in the NW European Variscan Front Zone (edited by Sintubin, M., Vandycke, S. & Camelbeeck, T.). Tectonophysics 309, 179-195.

Asselberghs, E. 1946. L'Eodévonien de l'Ardenne et des régions voisines. Mémoires de l'Institut géologique de l'Université de Louvain 14, 1-598.

Birtel, S. & Stöckhert, B. 2008. Quartz veins record earthquake-related brittle failure and short term ductile flow in the deep crust. Tectonophysics 457, 53-63.

Beugnies, A. 1986. Le métamorphisme de l'aire anticlinale de l'Ardenne. Hercynica 2(1), 17-33.

Blenkinsop, T. 2000. Deformation Microstructures and Mechanisms in Minerals and Rocks. Kluwer Academic Publishers, Dordrecht.

Bos, A., De Haas, G. J. L. M., Voncken, J. H. L., Van der Eerden, A. M. J. & Jansen, J. B. H. 1987. Hydrothermal synthesis of ammonioum-phlogopite. Geologie en Mijnbouw 66, 251-258.

Brühl, H. 1969. Boudinage in den Ardennen und in der Nordeifel als Ergebnis der inneren Deformation. Geologische Mitteilungen 8, 263-308.

Bultynck, P. & Dejonghe, L. 2001. Devonian lithostratigraphic units (Belgium). Geologica Belgica 4(1-2), 39-69.

Childs, C., Nicol, A., Walsh, J. J. & Watterson, J. 1996. Growth of vertically segmented normal faults. Journal of Structural Geology 18(12), 1389-1397.

Cloos, E. 1947. Boudinage. Transactions of the American Geophysical Union 28(4), 626-632.

Corin, F. 1932. A propos du boudinage en Ardenne. Bulletin de la Société belge de Géologie 42, 101-117.

Cosgrove, J. W. 1997. Hydraulic fractures and their implications regarding the state of stress in a sedimentary sequence during burial. In: Evolution of Geological Structures in Micro- to Macro-scales (edited by Sengupta, S.). Chapman & Hall, London, 11-25.

Cox, S. F., Knackstedt, M. A. & Braun, J. 2001. Principles of Structural Control on Permeability and Fluid Flow in Hydrothermal Systems. In: Structural Controls on Ore Genesis. Reviews in Economic Geology 14. Society of Economic Geology, Boulder, 1-24.

Darimont, A. 1986. Les inclusions fluides de quartz filoniens d'Ardenne. Annales de la Société géologique de Belgique 109, 587-601.

Darimont, A., Burke, E. A. J. & Touret, J. L. R. 1988. Nitrogen-rich metamorphic fluids in devonian metasediments from Bastogne, Belgium. Bulletin de Minéralogie 111, 321-330.

Debacker, T. N. 2001. Palaeozoic deformation of the Brabant Massif within eastern Avalonia: how, when and why? Unpublished Ph. D. thesis, Universiteit Gent.

Debacker, T. N., Dewaele, S., Sintubin, M., Verniers, J., Muchez, P. & Boven, A. 2005. Timing and duration of the progressive deformation of the Brabant Massif, Belgium. Geologica Belgica 8(4), 20-34.

Debacker, T. N., Seghedi, A., Belmans, M. & Sintubin, M. 2008. Contractional kink bands formed by stress deflection along pre-existing anisotropies? Examples from the Anglo-Brabant Deformation Belt (Belgium) and the North Dobrogea Orogen (Romania). Journal of Structural Geology 30(8), 1047-1059.

Dejonghe, L. 2008. Le coulour de décrochement dextre de l'Ourthe dans l'axe Erezée - Saint-Hubert (Haute Ardenne, Belgique) et son implication sur le tracé des failles longitudinales. Geologica Belgica 11(3-4), 151-165.

Dejonghe, L. & Hance, L. 2001. Carte géologique de Wallonie, feuille Champlon - La Roche-en-Ardenne (60/1-2) & Notice explicative. Ministère de la Région Wallonne, Namur.

Delvaux de Fenffe, D. & Laduron, D. 1991. Caledonian and Variscan structures in the Rocroi-Ardenne Lower Paleozoic basement (Belgium and adjacent countries). Annales de la Société géologique de Belgique 114, 141-162.

de Ruette, A. 1972. La légende des Cailloux de Mousny. Ardenne et Famenne 46, 79-129.


70

Fielitz, W. & Mansy, J.-L. 1999. Pre- and synorogenic burial metamorphism in the Ardenne and neighbouring areas (Rhenohercynian zone, central European Variscides). In: Palaeozoic to Recent tectonics in the NW European Variscan Front Zone (edited by Sintubin, M., Vandycke, S. & Camelbeeck, T.). Tectonophysics 309, 227-256.

Fourmarier, P. 1914. Observations sur le versant Nord et al partie centrale du synclinal de l'Eifel dans la région de Bertrix-Herbeumont. Annales de la Société géologique de Belgique 41, B323-B330.

Goscombe, B. D., Passchier, C. W. & Hand, M. 2004. Boudinage classification: end-member boudin types and modified boudin structures. Journal of Structural Geology 26(4), 739-763.

Gosselet, J. 1888. L'Ardenne. Mém. Expl. Carte géol. France, 1-889.

Hance, L., Ghysel, P., Laloux, M., Dejonghe, L. & Mansy, J.-L. 1999. Influence of a heterogeneous lithostructural layering on orogenic deformation of the Variscan Front Zone (eastern Belgium). In: Palaeozoic to Recent tectonics in the NW European Variscan Front Zone (edited by Sintubin, M., Vandycke, S. & Camelbeeck, T.). Tectonophysics 309, 161-177.

Harker, A. 1889. On local thickening of dykes and beds by folding. Geological Magazine 6(69-70).

Hatert, F., Deliens, M., Houssa, M. & Coune, F. 2000. Native gold, native silver, and secondary minerals in the quartz veins from Bastogne, Belgium. Bulletin de l'Institut Royal des Sciences Naturelles de Belgique, Sciences de la Terre 70, 223-229.

Helsen, S. 1995. Burial history of Palaeozoic strata in Belgium as revealed by conodont colour alteration data and thickness distributions. Geologische Rundschau 84, 738-747.

Hilgers, C., Bücker, C. & Urai, J. L. 2006. Fossil overpressures compartments? A case study from the Eifel area and some general aspects. In: 11. Symposium "Tektonik, Struktur- und Kristallingeologie" (edited by Phillipp, S., Leiss, B., Vollbrecht, A., Tanner, D. & Gudmundsson, A.). Universitätsdrucke Göttingen, Göttingen, Germany, 87-89.

Holmquist, P. J. 1931. On the relations of the "boudinage-structure". Geol. Fären. Färhandl. 53, 193-208.

Hugon, H. 1983. Structures et déformation du Massif de Rocroi (Ardennes). Bulletin de la Société géologique et minéralogique de Bretagne 15, 109-143.

Hugon, H. & Le Corre, C. 1979. Mise en évidence d'une déformation hercynienne en régime cisaillant progressif dans le Massif de Rocroi (Ardenne). Comptes Rendus de l'Académie des Sciences, Paris 289, 615-618.

Jongmans, D. & Cosgrove, J. W. 1993. Observations structurales dans la région de Bastogne. Annales de la Société géologique de Belgique 116, 129-136.

Kenis, I. 2004. Brittle-Ductile Deformation Behaviour in the Middle Crust as Exemplified by Mullions (Former "Boudins") in the High-Ardenne Slate Belt, Belgium. Aardkundige Mededelingen 14, 1-127.

Kenis, I., Muchez, P., Sintubin, M., Mansy, J.-L. & Lacquement, F. 2000. The use of a combined structural, stable isotopic and fluid inclusion study to constrain the kinematic history at the northern Variscan front zone (Bettrechies, France). Journal of Structural Geology 22(5), 598-602.

Kenis, I., Muchez, P., Verhaert, G., Boyce, A. J. & Sintubin, M. 2005a. Fluid evolution during burial and Variscan deformation in the Lower Devonian rocks of the High-Ardenne slate belt (Belgium): sources and causes of high-salinity and C-O-H-N fluids. Contribution to Mineralogy and Petrology 150, 102-118.

Kenis, I. & Sintubin, M. 2007. About boudins and mullions in the Ardenne-Eifel area (Belgium, Germany). Geologica Belgica 10(1-2), 79-91.

Kenis, I., Sintubin, M., Muchez, P. & Burke, E. A. J. 2002. The "boudinage" question in the High-Ardenne slate belt (Belgium): a combined structural and fluid inclusions approach. In: Tectonophysics (edited by Labaume, P., Craw, D., Lespinasse, M. & Muchez, P.). Tectonophysics 348, 93-110.

Kenis, I., Urai, J. L., van der Zee, W., Hilgers, C. & Sintubin, M. 2005b. Rheology of fine-grained siliciclastic rocks in the middle crust - evidence from structural and numerical analysis. Earth and Planetary Science Letters 233, 351-360.

Kenis, I., Urai, J. L., van der Zee, W. & Sintubin, M. 2004. Mullions in the High-Ardenne Slate Belt (Belgium). Numerical model and Parameter Sensitivity Analysis. Journal of Structural Geology 26(9), 1677-1692.

Klement, C. 1888. Analyse chimique de quelques minéraux et roches de Belgique et de l'Ardenne française. Bulletin du Musée royale de l'Histoire naturelle de Belgique 5, 159.

Lacquement, F. 2001. L'Ardenne Varisque. Déformation progressive d'un prisme sédimentaire pré-structuré, de l'affleurement au modèle de chaîne. Société Géologique du Nord, Lille.

Lacquement, F., Averbuch, O., Mansy, J.-L., Szaniawski, R. & Lewandowski, M. 2005. Transpressional deformations at lateral boundaries of propagating thrust-sheets: the example of the Meuse Valley Recess within the Ardennes Variscan fold-and-thrust belt (N France - S Belgium). Journal of Structural Geology 27, 1788-1802.


71

Lambert, A. & Bellière, J. 1976. Caractères structuraux de l'éodévonien aux environs de Bastogne. Annales de la Société géologique de Belgique 99, 283-297.

Le Gall, B. 1992. The deep structure of the Ardennes Variscan thrust belt from structural and ECORS seismic data. Journal of Structural Geology 14(5), 531-546.

Lohest, M., Stainier, X. & Fourmarier, P. 1908. Compte rendu de la session extraordinaire de la Société Géologique de Belgique, tenue à Eupen et à Bastogne les 29, 30 et 31 août et le 1, 2 et 3 septembre 1908. Annales de la Société géologique de Belgique 35, B351-B434.

Mansy, J.-L. & Lacquement, F. 2003. Le Paléozoïque du Nord de la France et de la Belgique. Géologues 133-134, 7-24.

Mansy, J.-L., Everaerts, M. & De Vos, W. 1999. Structural analysis of the adjacent Acadian and Variscan fold belts in Belgium and northern France from geophysical and geological evidence. In: Palaeozoic to Recent tectonics in the NW European Variscan Front Zone (edited by Sintubin, M., Vandycke, S. & Camelbeeck, T.). Tectonophysics 309, 99-116.

Meilliez, F., André, L., Blieck, A., Fielitz, W., Goffette, O., Hance, L., Khatir, A., Mansy, J.-L., Overlau, P. & Verniers, J. 1991. Ardenne - Brabant. Sciences Géologiques, Bulletin 44(1-2), 3-29.

Meilliez, F. & Mansy, J.-L. 1990. Déformation pelliculaire différenciée dans une série lithologique hétérogène: le Dévono-Carbonifère de l'Ardenne. Bulletin de la Société géologique de France (8), 6(1), 177-188.

Michot, P. 1980. Le segment tectogénique calédonien belge. Mémoire de l'Académie royal de Belgique, classe des Sciences. Collection in 8°, 2ème série 43, 1-61.

Micklethwaite, S., Sheldon, H. A. & Baker, T. 2010. Active fault and shear processes and their implications for mineral deposit formation and discovery. Journal of Structural Geology 32(2), 151-164.

Muchez, P. & Sintubin, M. 1998. Contrasting origin of palaeofluids in a strike-slip fault system. Chemical Geology 145, 105-114.

Muchez, P., Slobodnik, M., Sintubin, M., Viaene, W. & Keppens, E. 1997. Origin and migration of palaeofluids in the Lower Carboniferous of Southern and Eastern Belgium. Zentralblat für Geologie und Paläontologie, Teil I 1995 11/12, 1107-1112.

Muchez, P., Sintubin, M. & Swennen, R. 2000. Origin and migration pattern of palaeofluids during orogeny: discussion on the Variscides of Belgium and northern France. Journal of Geochemical Exploration 69-70, 47-51.

Mukhopadhyay, D. 1972. A note on the mullion structures from the Ardennes and North Eifel. Geologische Rundschau 61, 1037-1049.

Nüchter, J.-A. & Stöckhert, B. 2006. Fracturing and transformation into veins beneath the crustal scale brittle ductile transition - a record of co-seismic loading and post-seismic relaxation. Geophysical Research Abstracts 8, 05890.

Nüchter, J.-A. & Stöckhert, B. 2007. Vein quartz microfabrics indicating progressive evolution of fractures into cavities during postseismic creep in the middle crust. Journal of Structural Geology 29(9), 1445-1462.

Oncken, O., Plesch, A., Weber, J., Ricken, W. & Schrader, S. 2000. Passive margin detachment during arc-continent collision (Central European Variscides). In: Orogenic Processes: Quantification and Modelling in the Variscan Belt (edited by Franke, W., Haak, V., Oncken, O. & Tanner, D.). Special Publications 179. Geological Society, London, 199-216.

Oncken, O., von Winterfeld, C. & Dittmar, U. 1999. Accretion of a rifted passive margin: The Late Paleozoic Rhenohercynian fold and thrust belt (Middle European Variscides). Tectonics 18(1), 75-91.

Passchier, C. W. & Trouw, R. A. J. 2005. Microtectonics. Springer-Verlag, Berlin.

Pilgers, A. & Schmidt, W. 1957. Die Mullion-Structuren in der Nord-Eifel. Abh. Hess. Landesamtes Bodenforsch. 20, 1-53.

Piqué, A., Huon, S. & Clauer, N. 1984. La schistosité hercynienne et le métamorphisme associé dans la vallée de la Meuse, entre Charleville-Mézières et Namur (Ardennes franco-belges). Bulletin de la Société belge de Géologie 93(1-2), 55-70.

Plesch, A. & Oncken, O. 1999. Orogenic wedge growth during collision - constraints on mechanics of a fossil wedge from its kinematic record (Rhenohercynian FTB, Central Europe). In: Palaeozoic to Recent tectonics in the NW European Variscan Front Zone (edited by Sintubin, M., Vandycke, S. & Camelbeeck, T.). Tectonophysics 309, 117-139.

Quirke, T. T. 1923. Boudinage, an unusual structural phenomenon. Geological Society of America Bulletin 34, 649-660.

Radke, B. M. & Mathis, R. L. 1980. On the formation and occurrence of sadle dolomite. Journal of Sedimentary Petrology 50, 1149-1168.

Ramsay, A. C. 1866. The Geology of North Wales. Memoirs of the Geological Survey of Great Britain 3, 1-75.

Ribbert, K.-H. 1972. Erläuterungen zu Blatt C5502 Aachen. Geologisches Landesamt Northrhein-Westfalen, Krefeld.


72

Rondeel, H. E. & Voermans, F. M. 1975. Data pertinent to the phenomenon of boudinage at Bastogne in the Ardennes. Geologische Rundschau 64, 807-818.

Schavemaker, Y. A., de Bresser, J. H. P., Van Baelen, H. & Sintubin, M. 2012. Geometry and kinematics of the low-grade metamorphic 'Herbeumont shear zone' in the High-Ardenne slate belt (Belgium). Geologica Belgica 15(3), 126-136.

Schroyen, K. 2000. Caledonische en Varistische Fluïda en de metamorfe evolutie van het Stavelot-Vennmassief. Unpublished Ph. D. thesis, Katholieke Universiteit Leuven.

Schroyen, K. & Muchez, P. 2000. Evolution of metamorphic fluids at the Variscan fold-and-thrust belt in eastern Belgium. Sedimentary Geology 131, 163-180.

Shail, R. K. & Leveridge, B. E. 1997. Late Carboniferous to Triassic reactivation of Variscan basement in the western English Channel: evidence from onshore exposures in south Cornwall. Journal of the Geological Society, London 154, 163-168.

Shail, R. K. & Leveridge, B. E. 2009. The Rhenohercynian passive margin of SW England: Development, inversion and extensional reactivation. Comptes Rendus Geoscience 341, 140-155.

Sibson, R. H. 1983. Continental fault structure and the shallow earthquake source. Journal of the Geological Society, London 140, 741-767.

Sibson, R. H. 2000. Tectonic controls on maximum sustainable overpressure: fluid redistribution from stress transitions. Journal of Geochemical Exploration 69-70, 471-475.

Sibson, R. H. 2001. Seismogenic Framework for Hydrothermal Transport and Ore Deposition. In: Structural Controls on Ore Genesis. Reviews in Economic Geology 14. Society of Economic Geology, Boulder, 25-50.

Sibson, R. H. 2005. Hinge-parallel fluid flow in fold-thrust belts: how widespread?. Proceedings of the Geologists’ Association 116(3-4), 301-309.

Sintubin, M. 2008. L'Ardenne varisque: un témoin de la marge septentrionale de l'océan Rhéique. In: Géochronique 105, 26-29.

Sintubin, M., Debacker, T. N. & Van Baelen, H. 2009. Early Palaeozoic orogenic events north of the Rheic suture (Brabant, Ardenne): A review. Comptes Rendus Geoscience 341, 156-173.

Spaeth, G. 1986. Boudins and Mullions in devonischen Schichtfolgen der Nordeifel und Ardennen - Erscheinungsbild, Entstehung und Umformung. Nachrichten Deutsche Geologische Gesellschaft 35, 74-75.

Stainier, X. 1907. Sur le mode de gisement et l'origine des roches métamorphiques de la région de Bastogne. Mémoire de l'Académie royal de Belgique, classe des Sciences. Collection in 4°, 2ème série.

Stipp, M., Holger, S., Heilbronner, R. & Schmid, S. M. 2002. Dynamic recrystallization of quartz: correlation between natural and experimental conditions. In: Deformation mechanisms, rheology and tectonics: current status and future perspectives (edited by de Meer, S., Drury, M. R., de Bresser, J. H. P. & Pennock, G. M.). Special Publications 200. Geological Society, London, 171-190.

Urai, J. L., Spaeth, G., van der Zee, W. & Hilgers, C. 2001. Evolution of mullion (formerly boudin) structures in the Variscan of the Ardennes and Eifel. Journal of the Virtual Explorer 3, 1-15.

Van Baelen, H. 2010. Dynamics of a progressive vein development during the late-orogenic mixed brittle-ductile destabilisation of a slate belt. Examples of the High-Ardenne slate belt (Herbeumont, Belgium). Aardkundige Mededelingen 24, 1-221.

Van Baelen, H. & Sintubin, M. 2008. Kinematic consequences of an angular unconformity in simple shear. An example from the southern border of the Lower Palaeozoic Rocroi inlier (Naux, France). Bulletin de la Société géologique de France 179(1), 73-87.

Van Baelen, H., Berwouts, I., Muchez, P. & Sintubin, M. 2009. Reequilibrated fluid inclusions as proxy for PTX-conditions of deformation and recrystallisation in vein quartz. Journal of Geochemical Exploration 101, 105.

van der Pluijm, B. A. & Marshak, S. 2004. Earth Structure. An introduction to structural geology and tectonics. W.W. Norton & Company, New York.

Vanbrabant, Y. & Dejonghe, L. 2006. Structural analysis of narrow reworked boudins and influence of sedimentary successions during a two-stage deformation sequence (Ardenne-Eifel region, Belgium-Germany). Memoirs of the Geological Survey of Belgium 53, 1-43.

Van Noten, K. 2011. Stress-State Evolution of the Brittle Upper Crust during the Early Variscan Tectonic Inversion as Defined by Successive Quartz Vein-Types in the High-Ardenne Slate Belt, Germany. Aardkundige Mededelingen 28, 1-241.


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Van Noten, K. & Sintubin, M. 2010. Linear to non-linear relationship between vein spacing and layer thickness in centimetre- to decimetre-scale siliciclastic multilayers from the High-Ardenne slate belt (Belgium, Germany). Journal of Structural Geology 32(3), 377-391.

Van Noten, K., Hilgers, C., Urai, J. L. & Sintubin, M. 2008. Late burial to early tectonic quartz veins in the periphery of the High-Ardenne slate belt (Rursee, North Eifel, Germany). Geologica Belgica 11(3-4), 179-198.

Van Noten, K., Van Baelen, H. & Sintubin, M. 2012. The complexity of 3D stress-state changes during compressional tectonic inversion at the onset of orogeny. In: Faulting, Fracturing, and Igneous Intrusion in the Earth’s crust (edited by Healy, D., Butler, R. W. H., Shipton, Z. K. & Sibson, R. H.). Special Publications 367. Geological Society, London, 51-69.

Verniers, J., Pharaoh, T. C., André, L., Debacker, T. N., De Vos, W., Everaerts, M., Herbosch, A., Samuelsson, J., Sintubin, M. & Vecoli, M. 2002. The Cambrian to mid Devonian basin development and deformation history of eastern Avalonia, east of the Midlands Microcraton: new data and a review. In: Palaeozoic Amalgamation of Central Europe (edited by Winchester, J. A., Pharaoh, T. C. & Verniers, J.). Special Publications 201. Geological Society, London, 49-93.

Voisin, L. 1987. Les Ardoisières de l'Ardenne. Editions Terres Ardennaises, Charleville-Mézières.

Walls, R. 1937. A new record of boudinage-structure from Scotland. Geological Magazine 74, 325-332.

Wegmann, C. E. 1932. Note sur le boudinage. Bulletin de la Société géologique de France 52, 477-491.

Winchester, J. A. & Team, T. P. T. N. 2002. Palaeozoic Amalgamation of Central Europe: new results from recent geological and geophysical investigations. Tectonophysics 360, 5-21.


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List of field trip participants

Field trip leaders

Manuel Sintubin KU Leuven Belgium [email protected] Koen Torremans KU Leuven Belgium [email protected] Tom Haerinck KU Leuven Belgium [email protected] Koen Van Noten Royal Observatory of Belgium Belgium [email protected] Hervé Van Baelen NIRAS/ONDRAF Belgium [email protected]

Field trip participants

Max Arndt RWTH Aachen Germany [email protected] Eric de Kemp Geological Survey of Canada Canada [email protected] Laurent Montesi University of Maryland U.S.A. [email protected] Sven Morgan Central Michigan University U.S.A. [email protected] Jens Walter Universität Göttingen Germany [email protected] Rudy Wenk UC Berkeley U.S.A. [email protected]













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DRT2013 19th International Conference on Deformation mechanisms, Rheology and Tectonics

http://ees.kuleuven.be/drt2013

Department of Earth and Environmental Sciences Celestijnenlaan 200E

B-3001 Leuven Belgium

http://ees.kuleuven.be/drt2013�

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