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Keywords:

  • armorica;
  • fluid flow;
  • quartz veins;
  • microthermometry;
  • oxygen isotopes;
  • Variscan

Abstract

  1. Top of page
  2. Abstract
  3. Introduction
  4. Geological setting
  5. Methods
  6. Veining during progressive deformation
  7. Microthermometry and fluid composition
  8. Oxygen isotope and total organic carbon analyses
  9. Interpretation and discussion
  10. Conclusions
  11. Acknowledgements
  12. References

Quartz veins in the early Variscan Monts d’Arrée slate belt (Central Armorican Terrane, Western France), have been used to determine fluid-flow characteristics. A combination of a detailed structural analysis, fluid inclusion microthermometry and stable isotope analyses provides insights in the scale of fluid flow and the water–rock interactions. This research suggests that fluids were expelled during progressive deformation and underwent an evolution in fluid chemistry because of changing redox conditions. Seven quartz-vein generations were identified in the metasedimentary multilayer sequence of the Upper Silurian to Lower Devonian Plougastel Formation, and placed within the time frame of the deformation history. Fluid inclusion data of primary inclusions in syn- to post-tectonic vein generations indicate a gradual increase in methane content of the aqueous–gaseous H2O–CO2–NaCl–CH4–N2 fluid during similar P–T conditions (350–400°C and 2–3.5 kbar). The heterogeneous centimetre- to metre-scale multilayer sequence of quartzites and phyllites has a range of oxygen-isotope values (8.0–14.1‰ Vienna Standard Mean Ocean Water), which is comparable with the range in the crosscutting quartz veins (10.5–14.7‰ V-SMOW). Significant differences between oxygen-isotope values of veins and adjacent host rock (Δ = −2.8‰ to +4.9‰ V-SMOW) suggest an absence of host-rock buffering on a centimetre scale, but based on the similar range of isotope values in the Plougastel Formation, an intraformational buffering and an intermediate-scale fluid-flow system could be inferred. The abundance of veins, their well-distributed and isolated occurrence, and their direct relationship with the progressive deformation suggests that the intermediate-scale fluid-flow system primarily occurred in a dynamically generated network of temporarily open fractures.


Introduction

  1. Top of page
  2. Abstract
  3. Introduction
  4. Geological setting
  5. Methods
  6. Veining during progressive deformation
  7. Microthermometry and fluid composition
  8. Oxygen isotope and total organic carbon analyses
  9. Interpretation and discussion
  10. Conclusions
  11. Acknowledgements
  12. References

In an upper- to mid-crustal setting, under low-grade metamorphic conditions, fluid flow may occur from the centimetre to the regional scale. Whilst in most metamorphic settings fluid flow is considered to occur dominant along faults and through fractures (cf. Verweij 1993), regional mass transfer can be achieved by diffuse upper- and mid-crustal fluid convection, or a more pervasive fluid flow (cf. Etheridge et al. 1983; Ferry 1992; Oliver 2001). Hydrodynamically linked networks of fractures can play an important role in this regional mass transfer. Both fluid flow and fluid pressure promotes growth of these hydrodynamically linked networks, primarily because fracture growth is favoured in high pore-fluid pressure regimes (Cox et al. 2001; Oliver 2001).

The scale dependence of fluid-flow mechanism also applies when considering the open or closed nature of a fluid-flow system. Moreover, during progressive deformation, the fluid-flow system can evolve. For example, Kenis et al. (2000) showed that, during the progressive development of a kilometre-scale syncline, fluids remained isotopically buffered and in thermal equilibrium with the host rock, but still underwent a pressure–temperature (P–T) evolution. Such a rock-buffered vein system is indicated when the isotope (Sr, O, C) and trace element compositions of the veins are similar to those of the immediate surrounding host rock (Yardley 1975; Cartwright et al. 1994; Muchez et al. 1995; Travéet al. 1997; Janssen et al. 2005; Kenis et al. 2005). The opposite, i.e. a dissimilarity between vein and host-rock isotopic and trace element composition, is true for an open fluid-flow system (Oliver et al. 1993). Finally, an evolution from a rock-buffered to an open fluid-flow system is often observed in foreland fold-and-thrust belts (Van Geet et al. 2002; Roure et al. 2005; Breesch et al. 2007).

In the case of a more pervasive fluid flow in a heterogeneous metasedimentary multilayer sequence, inferring whether the fluid-flow system is open or closed on an intermediate scale (decimetre to metre) becomes less obvious. Due to variation in isotopic composition between the different lithologies and horizons present in the multilayer sequence, and the variable contribution of each rock type in the buffering process, the true nature of the fluid-flow system is difficult to determine. Intermediate and regional scale mass transfer and isotopic changes can be caused by a pervasive fluid flow of metamorphic fluids through a grain-scale network of fluid channels (Dipple & Ferry 1992; Cartwright et al. 1995; Oliver et al. 1998; Oliver 2001). However, little is known of the contribution to the regional flow of intermediate-scale fluid flow through a transient network of open fractures (i.e. dynamic permeability; Cox et al. 2001). The intimate relationship between fluid expulsion, metamorphism and progressive deformation in such a fluid-flow system could provide insight in both the dynamic processes of progressive deformation and the evolution in fluid chemistry (cf. Crispini & Frezzotti 1998).

Numerous studies on fluid-flow reconstruction are based on structural, geochemical or mineralogical analyses. This study, however, will combine detailed structural, stable isotope analysis and microthermometry on fluid inclusions to provide better insights in the scale and evolution of fluid flow and water–rock interaction.

In the Monts d’Arrée slate belt (MASB) in Brittany (France), a heterogeneous, metasedimentary multilayer sequence, i.e. the quartzites and phyllites of the Lower Palaeozoic Plougastel Formation, is well exposed. These siliciclastic rocks underwent greenschist metamorphism throughout the progressive contraction-dominated ‘Bretonian’ deformation event during the late Famennian and Tournaisian (approximately 370–350 Ma; Darboux 1981; Rolet 1982; Le Gall et al. 1992; van Noorden et al. 2007). Intense veining occurred during the deformation history (van Noorden 2007; van Noorden et al. 2007). These rocks are therefore suited to help better constrain the intimate relationship between progressive deformation and fluid expulsion, and to study intermediate-scale fluid flow in a lithologically heterogeneous sequence.

This study will address the following questions about fluid flow in the MASB. First, what is the veining history and how does it relate to the main deformation in the slate belt? Secondly, what type of fluids migrated through the metasedimentary sequence? Thirdly, did these fluids evolve and, fourthly, what is the origin of fluids and, likewise, what is the scale of fluid flow?

Geological setting

  1. Top of page
  2. Abstract
  3. Introduction
  4. Geological setting
  5. Methods
  6. Veining during progressive deformation
  7. Microthermometry and fluid composition
  8. Oxygen isotope and total organic carbon analyses
  9. Interpretation and discussion
  10. Conclusions
  11. Acknowledgements
  12. References

The Central Armorican Terrane (CAT) is exposed in the Armorican Massif (Brittany, France) (Fig. 1A), and is part of a particular segment of the late Palaeozoic, Pan-European Variscan orogen that is dominated by regional wrench tectonics. The South Armorican Shear Zone (SASZ), acting as the southern border of the CAT, localizes this wrenching and forms a major tectonic boundary, reflecting the suturing between the Armorican microplate and Gondwana and the associated closure of the South Armorican Ocean (Autran et al. 1994). North of the SASZ the CAT forms a low-grade, middle- to upper-crustal domain, composed of a nearly continuous Lower Ordovician (480 Ma) to Middle Carboniferous (320 Ma) metasedimentary sequence (Fig. 1A; Guillocheau & Rolet 1982). This sequence was deposited on top of Neoproterozoic, Brioverian (670–640 Ma) metasediments, representing the erosion product of a Pan-African basement (Cogné 1974; Le Corre et al. 1991; Autran et al. 1994; Chantraine et al. 2001). This Pan-African basement primarily crops out north of the North Armorican Shear Zone (NASZ) in the North Armorican Composite Terrane (NACT). From the Lower Ordovician (Arenig) until the Upper Devonian (late Famennian), the sedimentary record in the CAT reflects the development of an unstable platform. Dynamic deposits (e.g. olistostromes, turbidites) mark the start of a second deposition period, which is associated with increased tectonic activity in the CAT during the early Variscan, ‘Bretonian event’ (late Famennian – Tournaisian; approximately 370–350 Ma) (Rolet 1982; Darboux 1991; Le Gall et al. 1992).

image

Figure 1.  (A) Overview of the geological setting of the Armorican Massif in western France. (B) Detailed geological map of the Monts d’Arrée with indication of the different sample locations (Sant Kadou, Croaz Mélar, Roc’h Trévézel, Roc’h Tredudon and Roc’h Ar Feunteun) (after Cabanis et al. 1981; Babin et al. 1982; Castaing et al. 1987). NASZ, North Armorican Shear Zone; SASZ, South Armorican Shear Zone.

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The early Variscan, ‘Bretonian’, orogenic event resulted from the late Devonian to early Carboniferous closure of the Rheic Ocean and the related oblique docking of the Léon Terrane (Fig. 1A). This ‘Bretonian event’ is characterized by major top-to-the-northwest overthrusting, nappe-stacking and a significant crustal thickening of the central Armorican crust (Rolet et al. 1994; Faure et al. 2005; van Noorden 2007; Sintubin et al. 2008). Subsequently, during middle to late Carboniferous, intracontinental deformation resulted in a significant wrenching, predominantly evidenced along the SASZ (e.g. Gapais & Le Corre 1980; Jégouzo 1980) and syn-orogenic collapse (Faure et al. 2005).

The MASB consists of a highly deformed, low-grade metamorphic metasedimentary multilayer sequence of the Upper Silurian (Pridoli) to Lower Devonian (Lochkovian) Plougastel Formation (Fig. 1B). This siliciclastic formation composed of a centimetre- to metre-scale alternation of dark phyllites and quartzites reflecting a proximal turbiditic facies (Fig. 2; J. T. Renouf, unpublished observation). Completely exposed single sections are absent, but the minimum thickness of the whole sequence reaches the order of 500 m (J. T. Renouf, unpublished observation). The phyllites consist of quartz, chlorite and muscovite, with subordinate quantities of feldspar. The quartzites are rich in quartz with lesser amount of chlorite, muscovite and feldspar (van Noorden 2007). The Plougastel Formation overlies massive quartzites of the Grès Armoricain Formation (Fig. 1B; Darboux 1981; Darboux & Le Gall 1988).

image

Figure 2.  Lithostratigraphy of the Upper Silurian to Lower Devonian Plougastel Formation, with (A) massive white to grey quartzitic sandstones in beds that can reach a thickness of approximately 2 m, (B) a dark-coloured homogeneous silty facies, (C) a multilayer facies of silty shales and quartzites that occur as beds of maximum 0.5 m and (D) a phyllitic facies with few interlayered quartzite beds. The minimum thickness of the whole sequence is 500 m (J. T. Renouf, unpublished observation).

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The MASB reflects a contraction-dominated deformation event caused by a single progressive north-northwest-verging thrusting, with an incipient east-northeast-west-southwest wrench-dominated strain during later stages. Strike-slip structures (e.g. extensional shear bands (ESBs), drag folds) are randomly distributed and are consistent evidence of a dextral, belt-parallel movement. These local strain heterogeneities represent incipient strain partitioning and the development of embryonic domains of wrench-dominated deformation in an overall pure-shear dominated transpressional regime (van Noorden 2007; van Noorden et al. 2007). The radiometric age of the Huelgoat granite (336 ± 13 Ma; Fig. 1B; Peucat et al. 1979) postdates most of the deformation (approximately 370–350 Ma) and furthermore indicates that the main deformation in the MASB can be fitted into the overall thrusting and nappe-stacking during the ‘Bretonian’ orogenic event (approximately 370–350 Ma; cf. Faure et al. 2005; Sintubin et al. 2008).

Methods

  1. Top of page
  2. Abstract
  3. Introduction
  4. Geological setting
  5. Methods
  6. Veining during progressive deformation
  7. Microthermometry and fluid composition
  8. Oxygen isotope and total organic carbon analyses
  9. Interpretation and discussion
  10. Conclusions
  11. Acknowledgements
  12. References

Fluid-inclusion techniques

Quartz-vein samples were studied with a conventional petrographic microscope, and different types of fluid inclusions were identified. Strongly deformed zones and intensely recrystallized quartz crystals yield clusters of fluid inclusions, but the intracrystalline deformation of these crystals may induce leakage and re-equilibration of fluid inclusions, which can result in changes of bulk density and composition (Roedder 1984; Boullier et al. 1989; Sterner & Bodnar 1989; Vityk & Bodnar 1995). Therefore, these crystals were avoided as much as possible.

Microthermometric analyses were performed on double-polished 150-μm thick wafers using a Linkam THMSG 600 heating-cooling stage and a USGS-modified fluid inclusion stage. Calibration was performed using synthetic Syn Flinc fluid inclusion standards and the stages were calibrated at −56.6, −21.2, 0.0 and 374.1°C. Reproducibility was within 0.2°C for melting temperatures (Tm) and <3°C for total homogenization temperatures (Th). Quantifications of the gaseous components in the fluid inclusions were performed by Laser Raman microspectrometry, using a LabRam (Dilor, Jobin-Yvon) Raman microspectrometer fitted with a CCD detector, at the Université H. Poincaré (Nancy). The 514.5-nm ionized Argon laser (Spectra Physics) with a spot size of 2–3 μ was used as a source of exciting radiation. The laser was focused on the sample using an Olympus optical microscope (objective 80×). Raman analyses data were obtained a few degrees above homogenization temperature in the gas phase, using a Linkam stage fixed on the microscope of the Raman microprobe.

All salinities were calculated by using the equation proposed by Bodnar (1993), based on Tm ice (Fig. 6C, D). Due to the presence of clathrate, most values indicate maximum salinities. Salinities could not be calculated by use of Tm clath, because of the presence of other gases in addition to CO2. Total homogenization always occurred into liquid state.

image

Figure 6.  Histograms of fluid-inclusion microthermometric data of different vein generations of the MASB (Brittany, France). Histograms show both aqueous–gaseous and aqueous fluid inclusions. Primary inclusions occur in growth zones and a cluster of inclusions in vein type 6, is interpreted as primary. Secondary inclusions occur both in trails and clusters. (A, B) Temperatures of last melting of CO2 (Tm CO2) in respectively primary and secondary inclusions. (C, D) Temperatures of last melting of ice (Tm ice) in respectively primary and secondary inclusions. (E, F) Temperatures of last melting of clathrates (Tm clath) in respectively primary and secondary inclusions. (G, H) Homogenization temperatures of CO2 (Th CO2) in respectively primary and secondary inclusions. (I, J) Total homogenization temperatures (Th tot) in respectively primary and secondary inclusions.

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Stable-isotope techniques

Twenty-six oxygen-isotope analyses were performed on samples of the seven vein generations and 12 analyses on their associated host rock. Samples were taken from different localities in the MASB over a distance of 20 km (Fig. 1B). Vein quartz and whole rock separation was achieved by microdrilling. Host-rock samples (both quartzites and phyllites of the Plougastel Formation) were drilled at a distance of approximately 5 cm away from the corresponding veins to avoid sampling of alteration zones, but still close enough to the vein. All separates were analysed using a laser fluorination procedure, involving total sample reaction with excess ClF3 using a CO2 laser as a heat source. Reproducibility is better than 0.3‰ (1σ). Results are reported in standard notation (δ18O) as per mil (‰) deviations from Vienna Standard Mean Ocean Water (V-SMOW).

Veining during progressive deformation

  1. Top of page
  2. Abstract
  3. Introduction
  4. Geological setting
  5. Methods
  6. Veining during progressive deformation
  7. Microthermometry and fluid composition
  8. Oxygen isotope and total organic carbon analyses
  9. Interpretation and discussion
  10. Conclusions
  11. Acknowledgements
  12. References

The highly deformed multilayer sequence of the MASB is characterized by the abundance of veins, which predominantly consist of quartz, but chlorite and muscovite have been recognized as well. During detailed field-based and microstructural analysis, seven main generations were identified (van Noorden 2007; van Noorden et al. 2007). This division is based on crosscutting relationships of the veins with other structural features (e.g. bedding planes, cleavage planes and folds). Each generation can thus be placed within the relative time frame of the burial (approximately 420–370 Ma) and the progressive deformation (approximately 370–350 Ma) of the MASB (Fig. 3).

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Figure 3.  Diagram showing the different types of quartz veins and their chemical composition. Time indications are relative and based on the age of the sedimentary sequence and the timing of the granitic intrusions. On the right side, an overview is shown of the chronological evolution of fluids that circulated through the rocks of the MASB, as evidenced by microthermometry and Raman analysis of fluid inclusions. The first fluids were composed of equal amounts of CO2 and CH4 (type 3). With time, percentage of CH4 increased slightly (types 4 and 6). KB, kink bands. Secondary fluid inclusions are enriched in CO2 (type 7). Finally, a late H2O–NaCl fluid circulated through the rocks of the MASB.

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Regular bedding-perpendicular quartz veins (type 1) are omnipresent and are clearly pre-cleavage, evidenced by the presence of cleavage developed in these veins (Fig. 4A). They occur in quartzitic layers and display a fibrous crystal-growth habit. The veins are typically crosscut by bedding-parallel stylolites. Furthermore, these veins are organized in regularly spaced arrays. Type-2 quartz veins are related to the fold-hinge line and are rather rare (Fig. 4B). These syn-folding, syn-cleavage veins are either symmetrical or asymmetrical and are also observed between boudinage features. The most common type of quartz veins (type 3) (Fig. 4C) are represented by rod-type quartz veins, associated with phyllitic sequences. They are typical, syn-kinematic and are subparallel to cleavage. The veins locally display asymmetric folding on vertical sections and extension features on horizontal sections. Characteristically, this is expressed by ESBs, which predominantly display localized dextral shearing. Type-4 veins are expressed by regular quartz veins in quartzitic layers, and are likely to be associated with cleavage refraction (Fig. 4D). Typically, they display fibrous growth. These veins have a syn- to postcleavage chronology, but occur prior to massive quartz veining (types 6 and 7). Type-5 quartz veins crosscut the phyllitic sequence and are parallel to kink-band boundaries (Fig. 4E). Therefore, these quartz veins form a late vein generation and have a postfolding, postcleavage appearance. Massive quartz veins of type 6, associated with quartzitic layers, have a very irregular, blocky character (Fig. 4F). It is possible that this type of massive quartz vein is related to the veins of type 7. The latter is typically NS-trending and contains discrete veins of postfolding and postcleavage origin, with blocky character and very irregular shapes (Fig. 4G). Type-7 veins represent the latest veining event. Generally, they are regular spaced and are related to small-scale quartz veins, which diminish away from the massive quartz veins of type 7.

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Figure 4.  Overview of the seven different quartz-vein generations from the MASB, based on crosscutting relations of the veins with other structural features. (A) Type-1 veins are regular bedding-perpendicular quartz veins in quartzitic layers. They are pre-cleavage and crosscut by bedding-parallel stylolites. (B) Type-2 quartz veins are related to fold hinge line parallel deformation and therefore supposed to be syn-folding and syn-cleavage. They are also observed between boudin structures. (C) The most common type of quartz veins is type 3, which is associated with phyllitic beds. They are folded, show a dextral sense of shear and lie subparallel to the cleavage. This type is syn-folding and syn-cleavage. (D) Type-4 quartz veins are associated with cleavage refraction in quartzitic layers. These veins have fibrous crystals and are syn- to postcleavage. (E) Type-5 veins are fine regular quartz veins crosscutting phyllites and parallel to kink-band boundaries. They are supposed to be postfolding and postcleavage. (F) Massive quartz veins of type 6 are associated with quartzitic layers, have irregular, blocky crystals and are postfolding and postcleavage. (G) Type-7 veins are also massive quartz veins associated with quartzitic layers, but show a clear north-south strike. These veins are postfolding and postcleavage.

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Microthermometry and fluid composition

  1. Top of page
  2. Abstract
  3. Introduction
  4. Geological setting
  5. Methods
  6. Veining during progressive deformation
  7. Microthermometry and fluid composition
  8. Oxygen isotope and total organic carbon analyses
  9. Interpretation and discussion
  10. Conclusions
  11. Acknowledgements
  12. References

Samples of the different vein generations were taken from locations throughout the MASB and examined in order to identify the types of fluids present during the progressive deformation in the slate belt.

According to the chemical composition and the observed phase transitions during heating, different types of fluid inclusions were distinguished. A first subdivision was made based on their composition; they can be divided into two-phase aqueous and three-phase mixed aqueous–gaseous fluid inclusions. Based on their behaviour during heating, mixed aqueous–gaseous fluid inclusions can be further subdivided into A-type (Tm clath < Th CO2) and B-type (Tm clath > Th CO2) fluid inclusions (Fig. 5; cf. Kenis 2004). This classification is based on the relative position of the melting temperature of clathrate with respect to the homogenization temperature of CO2. Concerning the ice-melting temperature, B-type fluid inclusions can further be divided into type B1 (Th CO2 < Tm ice) and B2 (Th CO2 > Tm ice) (Fig. 5).

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Figure 5.  Classification of the mixed aqueous–gaseous fluid inclusions of the MASB, based on the differences between clathrate melting and homogenization temperatures (after Kenis 2004).

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Type-1 quartz veins

In type-1 quartz veins, clusters of two-phase aqueous (aqueous liquid and aqueous vapour) fluid inclusions were identified (Fig. 6, Table 1). They have a H2O–NaCl composition with salinities varying between 0.2 and 4.8 eq.wt% NaCl. Homogenization temperatures range from 154 to 333°C (Fig. 6J).

Table 1.   Vein type, sample, sample location, microthermometry and Raman microspectrometry data of the primary and secondary fluid inclusions of quartz veins in the Plougastel Formation.
TypeSampleLocalityPrimarySecondary
1MA03MN86Sant KadouAbsentClusters
    Tmice: –2.9 to –0.1 °C (20)
    Thtot: 154 to 333 °C (25)
    Salinity: 0.2 to 4.8 eq.wt% NaCl
2MA04MN23Roc’h ar FeunteunGrowth zoneAbsent
   Tmice: –5.1 to –2.3 °C (12) 
   Thtot: 208 to 303.6 °C (11) 
   Salinity: 3.9 to 8.0 eq.wt% NaCl 
3MA04MN10Roc’h Tredudon Growth zoneTrails
 MA03MN75Croaz Mélar TmCO2: –61.1 to –57.4 °C (18)Clusters
 MA03MN64 Tmice: –4.9 to –0.2 °C (21)TmCO2: –59.5 to –57.0 °C (40)
 MA03MN67 Tmclath: 5.1 to 14.3 °C (24)Tfm: –20.9 to –15.8 °C (3)
   ThCO2: –26.7 to –19.1 °C (11)Tmice: –3.0 to 0.0 °C (46)
   Thtot: 282 to 363 °C (13)Tmclath: 6.1 to 13.8 °C (48)
   Salinity: 0.3 to 7.7 eq.wt% NaClThCO2: –20.9 to 18.5 °C (19)
   CO2: 48 to 67.5 mol%Thtot: 280 to 363 °C (25)
   CH4: 27.2 to 35.5 mol%Salinity: 0.0 to 4.9 eq.wt% NaCl
   N2: 5.0 to 17.7 mol%CO2: 60.2 to 93.7 mol%
   H2S: 0.0 to 0.3 mol%CH4: 6.0 to 36.8 mol%
    N2: <0.5 to 6.0 mol%
    H2S: 0.0 to 0.3 mol%
4MA03MN52Roc’h TrévézelGrowth zoneAbsent
   TmCO2: –61.3 to –57.7 °C (7) 
   Tmice: –1.2 to –1.6 °C (10) 
   Tmclath: 12.6 to 15.0 °C (6) 
   ThCO2: –14.4 to –12.8 °C (4) 
   Thtot: 202 to 360 °C (6) 
   Salinity: 2.1 to 2.9 eq.wt% NaCl 
   CO2: 55.2 to 56.5 mol% 
   CH4: 39.5 to 40.2 mol% 
   N2: 4.0 to 4.5 mol% 
5MA03MN51Roc’h TrévézelAbsentTrails
    Tmice: –2.6 to –1.0 °C (18)
    Thtot: 64 to 214 °C (18)
    Salinity: 1.7 to 4.3 eq.wt% NaCl
6MA03MN73bCroaz MélarGrowth zoneTrails
   ClusterTmCO2: –61.9 to –57.8 °C (5)
   TmCO2: –68.0 to –57.0 °C (23)Tfm: –10.1 °C (1)
   Tmice: –1.5 to –1.1 °C (4)Tmice: –0.1 °C (1)
   Tmclath: 8.0 to 15.2 °C (19)Tmclath: 11.9 to 13.3 °C (5)
   ThCO2: 9.8 to 25.3 °C (14)ThCO2: 20.4 to 24.5 °C (4)
   Thtot: 337 to 382 °C (16)Thtot: 255 to 384 °C (24)
   Salinity: 1.9 to 2.6 eq.wt% NaClSalinity: 0.2 eq.wt% NaCl
   CO2: 48 to 57.5 mol%CO2: 73.2 to 90.3 mol%
   CH4: 40.0 to 49.4 mol%CH4: 6.5 to 26.3 mol%
   N2: 1.0 to 3.3 mol%N2: 0.4 to 3.5 mol%
   H2S: 0.0 to 0.2 mol% 
7MA03MN69Croaz MélarAbsentCluster
    TmCO2: –62.9 to –57.4 °C (24)
    Tfm: –36.2 to –37.7 °C (4)
    Tmice: –3.1 to –1.3 °C (17)
    Tmclath: 9.8 to 14.1 °C (20)
    ThCO2: 5.7 to 20.2 °C (23)
    Thtot: 315 to 330 °C (7)
    Salinity: 2.2 to 5.1 eq.wt% NaCl
    CO2: 67.0 to 80.0 mol%
    CH4: 13.5 to 22.2 mol%
    N2: 6.3 to 10.5 mol%
    H2S: 0.1 to 0.3 mol%

Type-2 quartz veins

Type-2 quartz veins contain inclusions in growth zones, which could be interpreted to be of primary origin (Roedder 1984), and clusters (Fig. 6, Table 1). All inclusions are aqueous and contain two phases at room temperature. Maximum salinities vary between 3.9 and 8.0 eq.wt% NaCl. Homogenization temperatures (Th tot) of the primary inclusions range between 208 and 304°C (Fig. 6I, Table 1).

Type-3 quartz veins

Samples of asymmetric quartz veins contain primary fluid inclusions, occurring in growth zones. Clusters of fluid inclusions have also been identified (Fig. 6, Table 1). The inclusions are mixed aqueous–gaseous fluid inclusions (H2O–CO2–X). At room temperature, two (aqueous and gaseous liquid) or three phases (aqueous liquid, gaseous liquid and vapour) could be distinguished. Fluid inclusions behave according to the characteristics of a B2 fluid (Fig. 5). Total homogenization temperatures of the mixed aqueous–gaseous fluid inclusions vary between 282 and 363°C (Fig. 6I, Table 1), and maximum salinities range from 0.3 to 7.7 eq.wt% NaCl. Raman analysis showed the presence of 27.2–35.5% CH4, 5–17.7% N2 and traces of H2S in addition to 48–67.5% CO2 (Fig. 7, Table 1).

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Figure 7.  Diagram showing Raman microspectrometry values of fluid inclusions from vein types 3, 4, 6 and 7. Secondary inclusions and clusters have lower CO2 and CH4 values according to the primary inclusions. Increasing methane content from vein types 3–7 is annotated with an arrow.

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Non-asymmetric quartz veins show evidence for re-equilibration. Clusters of fluid inclusions in these samples were therefore considered to be of secondary origin. Two types of fluid inclusions could be distinguished. The first type contains mixed aqueous–gaseous inclusions, characterized by an H2O–CO2–X composition and represents A1-type fluid inclusions (Fig. 5). At room temperature, two (aqueous and gaseous liquid) or three phases (aqueous liquid, gaseous liquid and vapour) could be distinguished. Maximum salinities vary between 0.0 and 4.9 eq.wt% NaCl and Th tot ranges between 280 and 363°C (Table 1, Fig. 6J). In addition to 60.2–93.7% CO2, there is a presence of 6.0–36.8% CH4, <0.5–6.0% N2 and traces of H2S as demonstrated by Raman microspectrometry (Fig. 7, Table 1). The second type of inclusions consists of two-phase (liquid and vapour) aqueous fluid inclusions with an H2O–NaCl composition. Maximum salinities of this type vary between 1.7 and 2.9 eq.wt% NaCl.

Type-4 quartz veins

Primary fluid inclusions occurring in growth zones in type-4 quartz veins consist of aqueous–gaseous fluid inclusions (H2O–CO2–X; Table 1). At room temperature, two (aqueous and gaseous liquid) or three phases (aqueous liquid and gaseous vapour and liquid) could be distinguished. Fluid inclusions in type-4 quartz veins are B1 inclusions and have Th tot varying between 202 and 360°C (Fig. 6I). Calculated maximum salinities range from 2.1 to 2.7 eq.wt% NaCl. Raman analyses demonstrated the presence of 55.2–56.5% CO2, 39.5–40.0% CH4 and 4.0–4.5% N2 (Fig. 7, Table 1).

Type-5 quartz veins

Fluid inclusions in type-5 quartz veins only occur in trails, interpreted as secondary inclusions (Fig. 6, Table 1). These two-phase aqueous inclusions have an H2O–NaCl–X composition with salinities between 1.7 and 4.3 eq.wt% NaCl. Total homogenization temperatures are between 166 and 214°C, with one outlier at 64°C (Fig. 6J).

Type-6 quartz veins

Primary fluid inclusions of type-6 quartz veins, occurring in growth zones and clusters, consist of two- and three-phase mixed aqueous–gaseous fluid inclusions (H2O–CO2–X; Table 1). These primary inclusions are A1-type inclusions (Fig. 5). Th tot values range from 337 to 382°C and calculated maximum salinities from 1.9 to 2.6 eq.wt% NaCl. Raman analysis indicated the presence of 48.0–57.5% CO2, 40–49.4% CH4, 1.0–3.3 mol% N2 and traces of H2S (Fig. 7, Table 1).

Secondary trails contain mixed aqueous–gaseous inclusions with Th tot values between 255 and 384°C (H2O–CO2–X; Fig. 6J, Table 1). In addition to 73.2–90.3% CO2, there is a presence of 6.5–26.3% CH4 and 0.4–3.5% N2 as demonstrated by Raman microspectrometry (Fig. 7, Table 1).

Type-7 quartz veins

Type-7 quartz veins contain clusters of fluid inclusions and only very few isolated fluid inclusions (Fig. 6, Table 1). Both consist of two- and three-phase mixed aqueous–gaseous fluids (H2O–CO2; Table 1). Almost all fluid inclusions in type-7 veins behave according to the characteristics of B2-type inclusions (Fig. 5). Th tot values range from 315 to 330°C and maximum salinities vary between 2.2 and 5.1 eq.wt% NaCl (Fig. 6I, Table 1). Raman microspectrometry showed the presence of 13.5–22.2% CH4, 6.3–10.5% N2 and traces of H2S in addition to 67.0–80.0% CO2 (Fig. 7, Table 1).

Oxygen isotope and total organic carbon analyses

  1. Top of page
  2. Abstract
  3. Introduction
  4. Geological setting
  5. Methods
  6. Veining during progressive deformation
  7. Microthermometry and fluid composition
  8. Oxygen isotope and total organic carbon analyses
  9. Interpretation and discussion
  10. Conclusions
  11. Acknowledgements
  12. References

Stable-isotope data potentially provide information on the transport processes, the source and the volume of fluids associated with vein formation (Gray et al. 1991; Kenis et al. 2005). An oxygen-isotope study of the host rock and veins was performed, in order to determine if the mineralizing fluid was buffered by its surrounding host rock and if there was a vein evolution through time (e.g. Marquer & Burkhard 1992).

The δ18O values of vein quartz and host-rock samples are plotted for every outcrop area and for every vein generation in Fig. 8A, B. Vein quartz δ18O values vary between 10.5‰ and 14.7‰, with coexisting host-rock samples, ranging from 8.0‰ to 14.1‰. Most of the phyllites, which represent a combination of mainly quartz, muscovite and chlorite, have a lower δ18O value than the quartzite samples. The different vein generations show a mean value of 13.1‰ and an SD of 1.0‰. The variation of oxygen-isotope values between the different sample localities is rather limited, with a broader range for host rocks, than for veins (Fig. 8A). Differences in δ18O between vein quartz and corresponding host rock range from −2.8‰ to +4.9‰ (Fig. 9).

image

Figure 8.  δ18O data from vein quartz and associated host rock (phyllite, silicified phyllite and quartzite) in the MASB. (A) δ18O versus sample localities and (B) the range of δ18O of the different vein generations. Vertical line at 13.1‰ indicates the mean value of δ18O in the quartz-vein samples. SD of these values is 1.0‰.

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image

Figure 9.  Diagram showing δ18O of the host rock versus δ18O of the quartz veins in the MASB. Δ = 0 marks the line where vein quartz and associated host rock would have the same values. Numbers 1 and 2 show values close to the Δ = 0 line, which can be interpreted as host-rock buffering. In contrast, similar host-rock samples (number 3 and 4) do not support this model of host-rock buffering.

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In order to know whether or not the Plougastel Formation is a possible source for the methane, carbon dioxide and nitrogen content in the fluid inclusions, total organic carbon (TOC) was measured on 20 powdered phyllites and silicified phyllites. Measurements were carried out by the Walkley and Black dichromate method (Nelson & Sommers 1982). Because of the absence of an experimentally determined correction factor for phyllites, a factor of 1.3 was utilized as proposed by Nelson & Sommers (1982). Measured values range between 0.12% and 0.37% TOC, with a mean value of 0.24 and an SD of 0.07% TOC.

Interpretation and discussion

  1. Top of page
  2. Abstract
  3. Introduction
  4. Geological setting
  5. Methods
  6. Veining during progressive deformation
  7. Microthermometry and fluid composition
  8. Oxygen isotope and total organic carbon analyses
  9. Interpretation and discussion
  10. Conclusions
  11. Acknowledgements
  12. References

The progressive deformation in the MASB is recorded by seven different quartz-vein generations, occurring in the highly deformed, metasedimentary multilayer sequence of the Plougastel Formation. Syn-cleavage vein types 2–4 can be considered as veins formed during the main ‘Bretonian’ deformation event (370–350 Ma) affecting the MASB, while the postcleavage vein generations 5–7 were late- to postorogenic.

Composition of fluids in the MASB

The aqueous inclusions mainly have an H2O–NaCl composition. This type of fluid occurs in primary and secondary inclusions. Raman microspectrometry showed that the aqueous–gaseous fluid inclusions in the MASB contain CO2, CH4, N2 and H2S (Table 1). Primary and isolated aqueous–gaseous fluid inclusions have a bulk composition of about 55 mol% CO2, 40 mol% CH4 and 5 mol% N2, while secondary aqueous–gaseous fluid inclusions contain about 80–90 mol% CO2, 6–15 mol% CH4 and 1–5 mol% N2 (Fig. 7, Table 1). Raman analyses of other aqueous–gaseous fluid inclusions in some clusters show a similar composition and equal CO2- and CH4-content as inclusions that were identified as primary. In contrast, Raman analysis showed that other clusters of aqueous–gaseous fluid inclusions were enriched in CO2 and had the same bulk composition as secondary inclusions.

The compositional range of these fluid inclusions is compatible with metamorphic fluids produced by dehydration and decarbonization under low-grade metamorphic conditions in sedimentary sequences (Shepherd et al. 1991; Yardley 1997). CO2 is the most common C-rich component and can originate from decarbonization (thermal maturation) processes of sedimentary sequences. In general, considerable amounts of CH4 can form in rocks, which have been chemically re-equilibrated to produce CO2 during retrograde conditions. In that case, the oxygen fugacity being buffered by the rock minerals is reduced from high-grade metamorphic conditions towards lower temperatures and may cause a change of fluid composition (van den Kerkhof & Thiéry 2001). Because of this low-oxygen fugacity, CH4 is generally a common gas compound in diagenetic and low-grade metamorphic rocks (van den Kerkhof & Thiéry 2001). N2 is released during maturation of NH4-substituting minerals in wall rock phyllosilicates or feldspars due to prograde metamorphism (Shepherd et al. 1991; Dee & Roberts 1993; van den Kerkhof & Thiéry 2001; Kenis et al. 2005), but note that N2 can also be bound in organic matter (Shepherd et al. 1991).

Evolution of fluids and P–T conditions

An evolution of fluid composition from an aqueous–gaseous H2O–CO2–NaCl–CH4–N2 fluid, with an equal CO2- and CH4-content to a more CH4-enriched composition, is observed within the different vein generations (Fig. 3). The composition of the C-rich phases in fluid inclusions and thus also the CH4/CO2 ratio depends on the redox and P–T conditions (Ohmoto & Kerrick 1977; Holloway 1984; Huizenga 2001). A change from CO2- to CH4-dominated inclusions can occur during retrograde metamorphism (e.g. the Isua Greenstone Belt, SW Greenland; Heijlen et al. 2006), reflecting more reducing conditions after peak-metamorphic conditions (approximately 460°C and 4 kbar). The composition of fluid inclusions is indicative of interaction between rock and fluid during low-grade metamorphism, suggesting a close relationship between regional metamorphism and formation of the veins (cf. Shepherd et al. 1991; Dee & Roberts 1993).

The pressure and temperature during metamorphism and deformation of the Plougastel Formation in the MASB varied between 350–400°C and 2–3.5 kbar (Darboux 1991). This temperature range is confirmed by microthermometric analysis of fluid inclusions in vein types 2, 3, 4, 6 and 7 showing homogenization temperatures mainly between 270 and 370°C (Fig. 6I). In addition, metamorphic and structural analyses do not indicate changes of the metamorphic grade during progressive deformation (Darboux 1991; van Noorden 2007). Therefore, the varying CH4/CO2 ratio is most likely due to changing redox conditions. At 350°C and 2 kbar, thermodynamic calculations indicated that small changes in the oxygen fugacity induce significant changes in the CH4/CO2 ratio (see also Dewaele et al. 2004).

Based on the precipitation temperature and the quartz-vein δ18O-value, the oxygen-isotope composition of the ambient fluid could be calculated (Matsuhisa et al. 1979; Zhang et al. 1989). Calculated δ18O-values of water in vein generation 2, 3, 4 and 6, based on average homogenization temperatures in primary fluid inclusions, are respectively 6.0‰, 7.0‰, 6.6‰ and 8.8‰ V-SMOW. These values are within the range, typical for metamorphic fluids (Rollinson 1993).

Microthermometric data indicate an increase in total homogenization temperatures with time (Fig. 10). There are three possible explanations for this apparent increase in temperature. First, pressure dropped during trapping, while temperature remained constant (Fig. 11A), which is a common occurrence during rapid exhumation. Secondly, pressure remained constant and temperature increased (Fig. 11B). Finally, a combination of both, i.e. there was an increase in temperature and a change in pressure. A temperature increase could be due to an increase in geothermal gradient. The absence of an independent geobarometer or geothermometer makes it not possible to exclude one of the three interpretations.

image

Figure 10.  Diagram showing average total homogenization temperatures (Th tot) versus average salinity values of vein types 2–4 and 6. Arrow marks the decrease in salinity and increase in Th tot from vein types 2 to 3.

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image

Figure 11.  P–T diagrams showing possible mechanisms to interpret the increased homogenization temperatures (Th) in successive vein generations. (A) Variation in Th due to a pressure drop while trapping temperature (Tt) is constant. (B) The increase of Th reflects an increase in Tt in the successive vein generations.

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In addition, a decrease in maximum salinity is present in the successive vein generations. This decrease could be due to metamorphism. Decreasing or increasing salinity during metamorphism depends on the original mineralogy of the sedimentary sequence and the different chemical reactions that occur during changes in pressure and temperature (Smith & Yardley 1999). Dehydration and transformation of phyllosilicates in the Plougastel Formation is a likely source for water during prograde metamorphism (Burst 1969; Hower et al. 1976). This release of water causes a decrease in salinity of the fluids. This effect could explain the evolution in salinity between vein types 2, 3, 4 and 6. However, salinities reported from vein types 3, 4 and 6, in contrast to vein type 2, represent maximum salinities and are influenced by the amount and type of the gaseous components. Therefore, the evolution in maximum salinity from vein types 3–6 cannot be unequivocally interpreted.

Buffering and source of fluids

By analysing the isotopic composition of the different vein generations, the type of buffering, scale and evolution of the fluid-flow system can be retrieved. Host-rock buffering, as described by Gray et al. (1991) and Oliver (1996), refer to the isotope values of veins and associated host rock that lie in a restricted range. A fluid-dominated system, in contrast, would reflect isotopic variation of the fluid system, through time and space, within the different vein generations (Gray et al. 1991). In the Plougastel Formation, however, rock-buffering and evolution through time is not visible in the oxygen isotope set. The Δ-values of the measured oxygen isotopes (i.e. difference between vein and host-rock) range from −2.8‰ to +4.9‰ (Fig. 9). Two samples (number 1 and 2 on Fig. 9) have values close to Δ = 0, which could be interpreted as veins being buffered by quartz-rich rocks. However, the values of another quartzite, silicified phyllite (number 3 and 4) and of all phyllites, do not support this model. In general, there seems to be a larger contribution from fluids in isotopic equilibrium with quartz than with phyllosilicates.

The Plougastel Formation consists of a centimetre-scale alternation of phyllite and quartzite beds, resulting in an isotopic variability of the rocks. The isotopic composition of the veins could reflect an average composition of the surrounding rock and exhibits a disequilibrium in the isotopic relationship with their immediate host rock (cf. Richards et al. 2002).

Bons (2001), Richards et al. (2002) and Knoop et al. (2002) described quartz vein–host rock fractionation values up to at least 5‰ on a formation scale and relate this to intraformational heterogeneities and temperature differences during deformation. A variation between 8.0‰ and 14.7‰δ18O in the case of both veins and rocks within the Plougastel Formation does not exclude buffering on the formation scale. Variations through P–T conditions cannot be deduced from the oxygen-isotope values, due to the large range of host-rock values (8.0‰ to 14.1‰ V-SMOW).

The source of the CO2, CH4 and N2 can be diverse in sedimentary sequences affected by low-grade metamorphism. Total remaining organic carbon values in the pelitic beds of the Plougastel Formation is between 0.12% and 0.37% TOC. These relatively high values for a 415-Ma-old metamorphosed rock reflect the sedimentary environment. The Plougastel Formation was deposited as turbidites in the neighbourhood of an unstable platform at the Northern Gondwana Margin (J. T. Renouf, unpublished observation). This environment could accumulate much primary organic matter, which could be the source of methane, carbon dioxide and nitrogen during low-grade metamorphism.

The apparent contrast between an evolution in gas chemistry of the fluid inclusions and the absence of an evolution in the stable-isotope composition is interpreted to be due to the fact that these characteristics result from different processes. The evolution in gas chemistry is related to changes in redox conditions, while the oxygen-isotope composition reflects mixing of fluids from different layers in the metasedimentary sequence.

Conclusions

  1. Top of page
  2. Abstract
  3. Introduction
  4. Geological setting
  5. Methods
  6. Veining during progressive deformation
  7. Microthermometry and fluid composition
  8. Oxygen isotope and total organic carbon analyses
  9. Interpretation and discussion
  10. Conclusions
  11. Acknowledgements
  12. References

The fluid-flow system in the MASB, visualized by a distributed, but discrete system of veins, evolved during the progressive deformation of a siliciclastic, organic carbon-rich multilayer sequence in low-grade metamorphic conditions. Seven vein generations were identified and placed within the ‘Bretonian’ deformation history of the MASB, ranging from pre- to postcleavage. The abundance of the veins, their well-distributed and isolated occurrence and the direct relationship with the progressive deformation suggest that fluid flow primarily occurred in a dynamic, transiently generated network of open fractures. Primary fluids, generated during the deformation history, range in composition from a H2O–CO2–NaCl–CH4–N2 fluid, with equal amounts of CO2 and CH4, to a more CH4-enriched composition. This evolution of fluid composition is most probably determined by redox and metamorphic conditions in the metasediments. The oxygen isotopic composition of the quartz veins was not buffered by the immediate host rock. Fluids expelled from the multilayer sequence of the Plougastel Formation contributed to the average oxygen isotopic composition of the ambient fluids.

Although the chemical evolution and scale of fluid flow in the Plougastel Formation has been established here, new investigations in a broader area should reveal a better understanding of intermediate-scale migration within a regional fluid-flow system.

Acknowledgements

  1. Top of page
  2. Abstract
  3. Introduction
  4. Geological setting
  5. Methods
  6. Veining during progressive deformation
  7. Microthermometry and fluid composition
  8. Oxygen isotope and total organic carbon analyses
  9. Interpretation and discussion
  10. Conclusions
  11. Acknowledgements
  12. References

We thank Herman Nijs for the carefully prepared double-polished wafers and Thérèse Lhomme of the University H. Poincaré at Nancy (France) for help with the Raman analyses. Thanks to editor R. H. Worden, reviewers N. Oliver, A. McCaig and D. Quinn, and two anonymous reviewers for their helpful and constructive comments. M. Sintubin is a Research Professor and M. van Noorden and I. Berwouts are Research Assistants of the ‘Onderzoeksfonds K. U. Leuven’. This work forms part of the research project G.0159.03 of the F.W.O.-Vlaanderen, and of the research projects OT/02/31 and OT/06/33 of the ‘Onderzoeksfonds K. U. Leuven’.

References

  1. Top of page
  2. Abstract
  3. Introduction
  4. Geological setting
  5. Methods
  6. Veining during progressive deformation
  7. Microthermometry and fluid composition
  8. Oxygen isotope and total organic carbon analyses
  9. Interpretation and discussion
  10. Conclusions
  11. Acknowledgements
  12. References
  • Autran A, Lefort J-P, Debeglia N, Edel J-B, Vigneresse J-L (1994) Gravity and magnetic expression of terranes in France and their correlation beneath overstep sequences. In: Pre-Mesozoic Geology in France and Related Areas (ed. KeppieJD), pp. 4972. Springer-Verlag, Berlin.
  • Babin C, Darboux J-R, Hallegouet B, Garreau J, Melou M, Plusquellec Y, Morzadec P, Pelhate A, Thonon P (1982) Carte Géologique de la France (1/50.000) – Le Faou (275), BRGM, Orléans.
  • Bodnar RJ (1993) Revised equation and table for determining the freezing point depression of H2O–NaCl solutions. Geochimica et Cosmochimica Acta, 57, 6834.
  • Bons PD (2001) The formation of large quartz veins by rapid ascent of fluids in mobile hydrofractures. Tectonophysics, 336, 117.
  • Boullier A-M, Michot G, Pêcher A, Barrès O (1989) Diffusion and/or plastic deformation around fluids inclusions in synthetic quartz: new investigations. In: Fluid Movements – Element Transport and Composition of the Deep Crust (ed. BridgewaterD), pp. 34560. Kluwer Academic Publishing, Norwell, MA.
  • Breesch L, Swennen R, Dewever B, Mezini A (2007) Deposition and diagenesis of carbonate conglomerates in the Kremenara anticline, Albania: a paragenetic time marker in the Albanian foreland fold-and-thrust belt. Sedimentology, 54, 48396.
  • Burst JF (1969) Diagenesis of Gulf Coast clayey sediments and its possible relation to petroleum migration. American Association of Petroleum Geology Bulletin, 53, 7393.
  • Cabanis B, Chantraine J, Jebrak M, Dadet P, Herrouin Y, Huber P, Flageollet J-C, Chauris L, Garreau J, Beaujour A, Duhamel M (1981) Carte Géologique de la France (1/50.000) – Morlaix (240), BRGM, Orléans.
  • Cartwright I, Power WL, Oliver NHS, Valenta RK, McLatchie GS (1994) Fluid migration and vein formation during deformation and greenschist facies metamorphism at Ormiston Gorge, central Australia. Journal of Metamorphic Geology, 12, 37386.
  • Cartwright I, Vry J, Sandiford M (1995) Changes in stable isotope ratios of metapelites and marbles during regional metamorphism, Mount Lofty Ranges, South Australia: implications for crustal scale fluid flow. Contributions to Mineralogy and Petrology, 120, 292310.
  • Castaing C, Beurrier M, Herrouin Y, Rolet J, Thonon P, Clozier L, Garreau J (1987) Carte Géologique de la France (1/50.000) – Huelgoat (276), BRGM, Orléans.
  • Chantraine J, Egal E, Thiéblemont D, Le Goff E, Guerrot C, Ballèvre M, Guennoc P (2001) The Cadomian active margin (North Armorican Massif, France): a segment of the North Atlantic Panafrican belt. Tectonophysics, 331, 118.
  • Cogné J (1974) Le Massif Armoricain. In: Géologie de la France (ed. DebelmasJ), pp. 10561. Doin, Paris.
  • Cox SF, Knackstedt MA, Braun J (2001) Principles of structural control on permeability and fluid flow in hydrothermal systems. In: Deformation, Fluid flow and Ore Deposits (eds RichardsJ, TosdalR). Reviews in Economic Geology, 14, 124.
  • Crispini L, Frezzotti M (1998) Fluid inclusion evidence for progressive folding during decompression in metasediments of the Voltri Group (Western Alps, Italy). Journal of Structural Geology, 20, 173346.
  • Darboux J-R (1981) Caractérisation du régime cisaillant de la déformation hercynienne dans les Monts d’Arrée (Massif Armoricain, France). Comptes Rendus de l’Académie des Sciences, Paris, 292, 1497500.
  • Darboux J-R (1991) Evolution tectonosédimentaire et structuration synmétamorphe des zones externes du segment Hercynien ouest-européen: Le modèle du Domaine Centre Armoricain occidental. Thèse de doctorat d’état, Université de Bretagne Occidentale.
  • Darboux J-R, Le Gall B (1988) Les Montagnes Noires: cisaillement bordier méridional du bassin carbonifère de Châteaulin (Massif Armoricain, France). Caractéristiques structurales et métamorphiques. Geodinamica Acta, 2, 12133.
  • Dee SJ, Roberts S (1993) Late-kinematic gold mineralisation during regional uplift and the role of nitrogen: an example from the La Codosera area, W. Spain. Mineralogical Magazine, 57, 43750.
  • Dewaele S, Muchez Ph, Banks DA (2004) Fluid evolution along multistage composite fault systems at the southern margin of the Lower Palaeozoic Anglo-Brabant fold belt, Belgium. Geofluids, 4, 34156.
  • Dipple GM, Ferry JM (1992) Fluid flow and stable isotopic alteration in rocks at elevated temperatures with applications to metamorphism. Geochimica et Cosmochimica Acta, 56, 353950.
  • Etheridge MA, Wall VJ, Vernon RH (1983) The role of the fluid phase during regional metamorphism and deformation. Journal of Metamorphic Geology, 1, 20526.
  • Faure M, Mézème EB, Duguet M, Cartier C, Talbot J-Y (2005) Paleozoic tectonic evolution of medio-europa from the example of the French Massif Central and Massif Armoricain. Journal of the Virtual Explorer, 19, 126.
  • Ferry JM (1992) Regional metamorphism of the Waits River Formation, Eastern Vermont: delineation if a new type of giant hydrothermal system. Journal of Petrology, 33, 4594.
  • Gapais D, Le Corre C (1980) Is the Hercynian belt of Brittany a major shear zone? Nature, 288, 5746.
  • Gray DR, Gregory RT, Durney DW (1991) Rock-buffered fluid rock interaction in deformed quartz-rich turbidite sequences, eastern Australia. Journal of Geophysical Research, 96, 19681704.
  • Guillocheau F, Rolet J (1982) La sédimentation paléozoïque ouest-armoricaine. Histoire sédimentaire; relations tectonique-sédimentation. Bulletin de la Société Géologique et Minéralogique de Bretagne, 14, 4562.
  • Heijlen W, Appel PWU, Frezzotti M-L, Horsewell A, Touret JLR (2006) Metamorphic fluid flow in the northeastern part of the 3.8-3.7 Ga Isua Greenstone Belt (SW Greenland): a re-evaluation of fluid inclusion evidence for early Archean seafloor-hydrothermal systems. Geochimica et Cosmochimica Acta, 70, 307595.
  • Holloway JR (1984) Graphite–CH4–H2O–CO2 equilibria at low-grade metamorphic conditions. Geology, 12, 455.
  • Hower J, Eslinger EV, Hower ME, Perry EA (1976) Mechanisms of burial metamorphism of argillaceous sediment: 1. Mineralogical and chemical evidence. Geological Society of America Bulletin, 87, 72537.
  • Huizenga J-M (2001) Thermodynamic modelling of COH fluids. Lithos, 55, 10114.
  • Janssen C, Romer R, Hoffmann-Rothe A, Mingram B, Dulski P, Möller P, Al-Zubi H (2005) The role of fluids in faulting deformation: a case study from the Dead Sea transform (Jordan). International Journal of Earth Sciences, 94, 24355.
  • Jégouzo P (1980) The South Armorican Shear Zone. Journal of Structural Geology, 2, 3947.
  • Kenis I (2004) Brittle-Ductile deformation behaviour in the Middle Crust as exemplified by Mullions (Former “Boudins”) in the high Ardenne Slate Belt. Aardkundige Mededelingen, 14, 1127.
  • Kenis I, Muchez Ph, 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, 598602.
  • Kenis I, Muchez Ph, Verhaert G, Boyce AJ, 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. Contributions to Mineralogy and Petrology, 150, 10218.
  • Van Den Kerkhof AM, Thiéry R (2001) Carbonic inclusions. Lithos, 55, 4968.
  • Knoop SR, Kennedy LA, Dipple GM (2002) New evidence for syntectonic fluid flow across the hinterland-foreland transition of the Canadian Cordillera. Journal of Geophysical Research, 107, 207198.
  • Le Corre C, Auvray B, Ballèvre M, Robardet M (1991) Le Massif Armoricain. In: Massifs Anciens de France (ed. PiquéA), pp. 31103. Louis Pasteur University, Strasbourg.
  • Le Gall B, Loboziak S, Le Herissé A (1992) Le flanc sud du synclinorium carbonifère de Châteaulin (Massif armoricain, France): une bordure de bassin réactivée en contexte décro-chevauchant. Bulletin de la Société Géologique de France, 163, 1326.
  • Marquer D, Burkhard M (1992) Fluid circulation, progressive deformation and mass-transfer processes in the upper crust: the example of basement-cover relationships in the External Crystalline Massifs, Switserland. Journal of Structural Geology, 14, 104757.
  • Matsuhisa Y, Goldsmith JR, Clayton RN (1979) Oxygen isotopic fractionation in the system quartz-albite-anorthite-water. Geochimica et Cosmochimica Acta, 36, 113140.
  • Muchez Ph, Slobodnik M, Viaene W, Keppens E (1995) Geochemical constraints on the origin and migration of palaeofluids at the northern margin of the Variscan foreland, southern Belgium. Sedimentary Geology, 96, 191200.
  • Nelson DW, Sommers LE (1982) Total carbon, organic carbon, and organic matter. In: Methods of Soil Analysis. Part 2: Chemical and Microbiological Properties, 2nd edn (eds PageAL, MillerRH, KeeneyDR, MadisonWI), pp. 53980. Agronomy Series No. 9, American Society of Agronomy, Inc., Madison, WI.
  • Van Noorden M (2007) New insights into early Variscan geodynamics in the Armorican massif. Evidences of contraction-dominated deformation and incipient strain partitioning in the Palaeozoic of the Monts d’Arrée, Brittany, France. PhD Thesis, Departement Geografie-Geologie, K.U. Leuven, Leuven.
  • Van Noorden M, Sintubin M, Darboux J-R (2007) Incipient strain partitioning in a slate belt: evidence from the early Variscan Monts d’Arrée slate belt (Brittany, France). Journal of Structural Geology, 29, 83749.
  • Ohmoto H, Kerrick DM (1977) Devolatilization equilibria in graphitic systems. American Journal of Science, 277, 101344.
  • Oliver NHS (1996) Review and classification of structural controls on fluid flow during regional metamorphism. Journal of Metamorphic Geology, 14, 47792.
  • Oliver NHS (2001) Linking of regional and local hydrothermal systems in the mid-crust by shearing and faulting. Tectonophysics, 335, 14761.
  • Oliver NHS, Cartwright I, Wall VJ, Golding SD (1993) The stable isotope signature of kilometre-scale fracturedominated metamorphic fluid pathways, Mary Kathleen, Australia. Journal of Metamorphic Geology, 11, 70520.
  • Oliver NHS, Dipple GM, Cartwright I, Schiller J (1998) Fluid flow and metasomatism in the genesis of the amphibolite-facies, pelite-hosted Kanmantoo Copper Deposit, South Australia. American Journal of Science, 298, 181218.
  • Peucat J-J, Charlot R, Mifdal A, Chantraine J, Autran A (1979) Définition géochronologique de la phase bretonne en Bretagne centrale. Etude Rb/Sr de granites du domaine centre armoricain. Bulletin du B.R.G.M., 1, 34956.
  • Richards IJ, Connelly JB, Gregory RT, Gray DR (2002) The importance of diffusion, advection, and host-rock lithology on vein formation: a stable isotope study from the Paleozoic Ouachita orogenic belt, Arkansas and Oklahoma. Bulletin of the Geological Society of America, 114, 134355.
  • Roedder E (1984) Fluid Inclusions. Reviews in Mineralogy and Geochemistry, 12th edn. The Mineralogical Society of America, Washington, DC.
  • Rolet J (1982) La “phase bretonne en Bretagne”: état des connaissances. Bulletin de la Société Géologique et Minéralogique de Bretagne, 14, 6371.
  • Rolet J, Gresselin F, Jégouzo P, Ledru P, Wyns R (1994) Intracontinental Hercynian events in the Armorican Massif. In: Pre-Mesozoic Geology in France and Related Areas (ed. KeppieJD), pp. 195219. Springer-Verlag, Berlin.
  • Rollinson HR (1993) Using Geochemical Data: Evaluation, Presentation, Interpretation. Longman, Harlow.
  • Roure F, Swennen R, Schneider F, Faure J-L, Guilhaumou N, Osadetz K, Robion Ph (2005) Incidence of tectonics and natural fluid transfers on reservoir evolution in foreland Fold-and-Thrust Belts. Oil & Gas Science and Technology, Rev. IFP, 60, 67106.
  • Shepherd TJ, Bottrell SH, Miller MF (1991) Fluid inclusion volatiles as an exploration guide to black shale-hosted gold deposits, Dolgellau gold belt, North Wales, UK. Journal of Geochemical Exploration, 42, 524.
  • Sintubin M, Van Noorden M, Berwouts I (2008) Late Devonian – early Carboniferous contraction-dominated deformation in Central Armorica (Monts d’Arrée, Brittany, France) and its relationship with the closure of the Rheic Ocean. Tectonophysics. doi: DOI: 10.1016/j.tecto.2008.05.023.
  • Smith MP, Yardley BWD (1999) Fluid evolution during metamorphism of the Otago Shist, New Zealand: (I) Evidence from fluid inclusions. Journal of Metamorphic Geology, 17, 17386.
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  • Sterner SM, Bodnar RJ (1989) Synthetic fluid inclusions: VII. Re-equilibration of fluid inclusions of quartz during laboratory-simulated metamorphic burial and uplift. Journal of Metamorphic Geology, 7, 24360.
  • Travé A, Labaume P, Calvet F, Soler A (1997) Sediment dewatering and pore fluid flow along thrust faults in a foreland basin inferred from isotopic and elemental geochemical analyses (Eocene southern Pyrenees, Spain). Tectonophysics, 282, 37598.
  • Van Geet M, Swennen R, Durmishi C, Roure F, Muchez Ph (2002) Paragenesis of Cretaceous to Eocene carbonate reservoirs in the Ionian fold and thrust belt (Albania): relation between tectonism and fluid flow. Sedimentology, 49, 697718.
  • Verweij JM (1993) Hydrocarbon Migration Systems Analysis. Elsevier, Amsterdam.
  • Vityk MO, Bodnar RJ (1995) Do fluid inclusions in high-grade metamorphic terranes preserve peak metamorphic density during retrograde decompression? American Mineralogist, 80, 6414.
  • Yardley BWD (1975) On some quartz-plagioclase veins in the Connemara shists, Ireland. Geological Magazine, 112, 18390.
  • Yardley BWD (1997) The evolution of fluids through the metamorphic cycle. In: Fluid Flow and Transport in Rocks: Mechanisms and Effects (ed. YardleyBWD), pp. 99121. Chapman & Hall, London.
  • Zhang L, Liu J, Zhou H, Chen Z (1989) Oxygen isotope fractionation in the quartz-water-salt system. Economic Geology, 84, 164350.