The dynamic evolution of fault zones at the seismogenic
brittle–ductile transition zone (BDTZ) expresses the delicate interplay
between numerous physical and chemical processes. Deformation and fluid flow
at the BDTZ are closely related and mutually dependent during repeating and
transient cycles of frictional and viscous deformation. Despite numerous
studies documenting in detail seismogenic faults exhumed from the BDTZ,
uncertainties remain as to the exact role of fluids in facilitating broadly
coeval brittle and ductile deformation at that structural level. We combine
structural analysis, fluid inclusion, and mineral chemistry data from
synkinematic and authigenic minerals to reconstruct the temporal variations
in fluid pressure (Pf), temperature (T), and bulk composition (X) of
the fluids that mediated deformation and steered strain localization along
BFZ300, a strike–slip fault originally active at the BDTZ. BFZ300 deforms
the Paleoproterozoic migmatitic basement of southwestern Finland and hosts
in its core two laterally continuous quartz veins formed by two texturally
distinct types of quartz – Qtz I and Qtz II, with Qtz I older than Qtz II.
Veins within the damage zone are formed exclusively by Qtz I. Mesostructural and
microstructural analysis combined with fluid compositional data indicate
recurrent cycles of mutually overprinting brittle and ductile deformation
triggered by oscillations of fluid pressure peaking at 210 MPa. Fluid
inclusion microthermometry and mineral pair geothermometry indicate that the
two documented quartz types precipitated from different fluid batches, with
bulk salinities in the 1 wt % NaCleq–5 wt % NaCleq range for Qtz I and in the 6 wt % NaCleq–11 wt % NaCleq range for Qtz II.
The temperature of the fluids involved with
initial strain localization and later fault reactivation evolved through
time from > 350 ∘C during Qtz I precipitation to
< 300 ∘C at the time of Qtz II crystallization. The peak
fluid pressure estimates constrain pore pressure oscillations between 80 and
210 MPa during the recorded faulting episodes. Our results suggest
variability of the physico-chemical conditions of the fluids steering
deformation (Pf, T, X), reflecting the ingress and effects of
multiple batches of fluid in the fault zone. Initial fluid-mediated
embrittlement generated a diffuse network of joints and/or hybrid–shear
fractures in the damage zone; subsequent strain localization led to more
localized deformation within the fault core. Localization was guided by
cyclically increasing fluid pressure and transient embrittlement of a system
that was otherwise under overall ductile conditions.
Our analysis suggests that fluid overpressure at the BDTZ can play a key
role in the initial embrittlement of the deforming rock and steer subsequent
strain localization.
Introduction
The physical and chemical properties of fault systems play a fundamental
role in controlling the rheological behaviour of the Earth's crust and in
steering channelled fluid flow (e.g. Caine et al., 1996). Deformation and
fluid flow are closely related and mutually dependent via a number of
feedbacks, such as the control that fluids exert upon the effectiveness of
deformation processes and the development of fault systems at all scales,
and the control by rock heterogeneities and/or fracture system topology on
the net fault transmissivity (e.g. Crider and Peacock, 2004). The nucleation
and development of permeable fault systems and the mechanisms whereby
individual faults may weaken and eventually fail are therefore complex
functions of a number of processes. In this perspective, the interaction
between fluid and mineral phases within fault rocks needs to be studied with
a system approach in order to single out the role and importance of all
processes involved (Kaduri et al., 2017).
An obvious effect of fluid involvement, particularly in crustal volumes that
have experienced large deformation-controlled fluid fluxes, is the
precipitation of authigenic and hydrothermal minerals within faults (Oliver
and Bons, 2001; Viola et al., 2016) and their immediately adjacent host rock
(Mancktelow and Pennacchioni, 2005; Garofalo, 2004). In the seismogenic
region of the crust, where fluids may even be the primary driver of the
seismic cycle (e.g. Miller, 2013), faults have been shown to have the
potential to function like a “fluid-activated valve”, whereby they
experience transient and cyclic fluid pressure build-up before sudden fluid
venting, pore pressure and mechanical strength drop concomitant with
seismic failure (e.g. Sibson, 1989, 1992b, 1993; Cox, 1995; Viola et al.,
2006; De Paola et al., 2007; Wehrens et al., 2016). Hydrothermal ore
deposits, in which fault networks focus relatively large volumes of ore fluids
and precipitate economic minerals (Cox et al., 2001; Boiron et al.,
2003; Moritz et al., 2006; Scheffer et al., 2017a),
are also pertinent examples of significant deformation-controlled fluid
ingress.
The seismogenic depth down to 10–15 km (e.g. Kohlstedt et al., 1995) is thus
a key region of the crust in which to study the whole range of fluid–rock
interaction processes occurring within fault zones. Deformation at that
depth might be accommodated under overall brittle–ductile conditions along
fault systems crossing or rooting into the brittle–ductile transition zone
(BDTZ). In detail, the deformation style in the BDTZ is generally
characterized by the cyclicity, also at the short timescale, between
brittle and ductile behaviour (Famin et al., 2004, 2005;
Siebenaller et al., 2013). This
is induced and regulated by the complex and transient interplay of numerous
parameters, among which are the lithological composition and transient variation
of temperature, pore pressure, and strain rate within the deforming system.
Field studies have documented unequivocally that ductile and brittle
deformation may even be simultaneously active during deformation as a
function of the transient and spatially heterogeneous evolution of the
chemical and physical parameters steering deformation, leading to the broad
coexistence of geological features expressing frictional deformation and
viscous creep and to mutual cross-cutting relationships thereof (e.g.
Guermani and Pennacchioni, 1998; Kjøll et al., 2015; Pennacchioni et al.,
2006; Wehrens et al., 2016; Scheffer et al., 2017b).
Veins are particularly important in this context because they attest to the
relative abundance of aqueous fluids in the deformation history (e.g. Cox et
al., 2001). Portions of the seismogenic crust that experience large fluid
fluxes host pervasive and vertically extensive vein networks (Sibson et al.,
1988), within which up to several million cubic metres of
hydrothermal minerals may deposit from the flowing fluid (e.g. Heinrich et
al., 2000; Cox, 2005; Bons, 2001; Garofalo et al., 2002). In
contrast, portions of the crust deforming in the absence of significant
fluid flow would show little to no evidence of veining, with only
synkinematic H2O-rich minerals within the fault rock attesting to
hydrous conditions (Mancktelow and Pennacchioni, 2004;
Menegon et al., 2017).
The physical–chemical conditions of fluid–rock interaction in the BDTZ have
been studied within exhumed faults by applying a set of geochemical tools
that include fluid inclusion analysis (e.g. Morrison, 1994;
Morrison and Anderson, 1998; Mulch et al., 2004; Ault and
Selverstone, 2008; Garofalo et al., 2014; Siebenaller et al., 2016; Compton
et al., 2017), determination of the isotopic compositions of fault fluids,
and mass transfer calculations between host rock and fault rocks (e.g.
Goddard and Evans, 1995; Garofalo, 2004; Mittempergher et al., 2014;
Spruzeniece and Piazolo, 2015). This approach yields important constraints
on the P–T conditions of fluid–rock interaction within the BDTZ, on the
source region of the fluids reaching and flowing within the deformation
zones, and on element mobility during syn-tectonic fluid flow. These
studies, however, do not specifically address the role of fluids in the
mechanisms that trigger and permit the aforementioned cycles of
brittle–ductile deformation. Open questions thus remain, such as what pressure, temperature, and composition (P, T, X) conditions are
best for a fluid to trigger brittle–ductile deformation cycles in a fault
system within the BDTZ and which fluid property is specifically most
effective in controlling the cycles.
(a) Simplified geological map of southwestern Finland modified after
Mattila and Viola (2014). (b) Geological sketch of Olkiluoto Island. The
upper right inset shows the poles to foliation planes measured from all
available Olkiluoto drill cores (N=4479, equal area, lower hemisphere
projection; Mattila and Viola, 2014). The lower left inset is a panoramic
photograph with an overlay drawing of the underground infrastructure (photo
courtesy of Posiva Oy, Finland). The red circle shows the depth location of
BFZ300. Coordinates are given in the local KKJ1 coordinate system.
In this work, we follow a multidisciplinary approach by combining mesostructural and
microstructural observations with the geochemical analysis of fluids,
petrographic documentation of fault rocks and veins, microthermometric
properties of fluid inclusion assemblages, electron probe microanalysis
(EPMA) of fault minerals, Raman spectrometry of fluid inclusions, and
electron probe cathodoluminescence imaging to study the effects of numerous
cycles of fluid–rock interaction that have occurred in a vein-rich
deformation zone at the seismogenic BDTZ and now exhumed as part of the
Paleoproterozoic continental crust of southwestern Finland. The studied
deformation zone belongs to an exhumed conjugate fault system that
experienced a complex history of structural reactivation and fluid flow.
Deformation zone BFZ300, the target of our study, crops out at ca. 426 m
below sea level within the deep Onkalo nuclear waste repository that is
presently being built on the island of Olkiluoto (Fig. 1a).
Our results allow us to constrain and describe the progressive evolution of
the deformation processes and the role of fluids involved both at fault
initiation and during the subsequent reactivation phases. We propose that
fluid pressure fluctuation cycles within an overall ductile environment at
the BDTZ triggered brittle–ductile cyclicity encompassing fracturing, vein
precipitation, and crystal–plastic deformation before renewed and
fluid-induced embrittlement. Our multi-technique approach made it possible to
determine many of the actual chemical and physical properties of the fluids
involved in the deformation process, leading to a well-constrained
conceptual mechanical model for the fault nucleation and subsequent
development.
Geological setting
The study area is located in southwestern Finland on the island of
Olkiluoto (Fig. 1a) within the Paleoproterozoic Svecofennian orogenic
province, which is formed by supracrustal high-grade metamorphic sequences
and plutonic rocks. The most abundant lithologies in the study area are
variably migmatitic metasedimentary rocks interleaved with metavolcanic rocks up to several metres thick in addition to calc-alkaline
synorogenic tonalite-trondhjemite-granodiorite-type (TTG) granitoids, as well as late orogenic leucogranites
(Fig. 1a, b). For a detailed lithological characterization of the area, we
refer the reader to Hudson and Cosgrove (2006) and Aaltonen et al. (2016).
Numerous studies carried out on Olkiluoto have highlighted the long
geological evolution of the region, which is commonly summarized by tectonic
models proposing either an evolution during a single and semi-continuous
Svecofennian orogenic event (Gorbatschev and Bogdanova, 1993) or,
alternatively, a sequence of up to five distinct accretion events leading to
the amalgamation of several microcontinents and island arcs at the margin of
the Archean craton between 1.92 and 1.79 Ga (e.g. Lahtinen et al., 2005). In
this scenario, several subduction systems developed, and the collision of
the involved microcontinents and island arc complexes resulted in
conspicuous continental growth, forming the major part of the
Paleoproterozoic domain of the Fennoscandian Shield (1.89–1.87 Ga).
According to Lahtinen et al. (2005), this “Fennian accretionary event”
ended with a phase of orogenic collapse associated with regional extension
and remarkable crustal thinning between ca. 1.86 and 1.84 Ga. Renewed
compression ensued during collision of the “Sarmatian Plate” with the
previously consolidated Svecofennian Shield, causing major crustal
shortening, high-temperature regional metamorphism (Kukkonen and Lauri,
2009), and the emplacement of S-type granites (e.g. Ehlers et al., 1993).
Tectonic activity ascribable to this orogenic phase ceased with a distinct
orogenic collapse phase at 1.79–1.77 Ga (Lahtinen et al., 2005).
Pervasive reworking of the Svecofennian domain took place in the
Mesoproterozoic when the crust underwent significant stretching and was
intruded by voluminous Rapakivi granites and diabase dykes resulting from
the widespread melting of the lower crust at ca. 1.65–1.50 Ga. This tectonic
phase was probably due to the development of a rift along the present Baltic
Sea (Korja et al., 2001). Crustal thinning also caused the formation of the
“Satakunta Graben”, a NW–SE-trending graben located ca. 50 km to the north
of Olkiluoto, which was later filled by Mesoproterozoic sandstone (Jotnian
sandstones, Fig. 1a). The latest stage of crustal evolution in southern
Finland is expressed by the intrusion of 1.27–1.25 Ga, N–S-striking olivine
diabase dikes (Fig. 1a; e.g. Suominen, 1991).
As to the structural evolution of the study area, the bedrock was affected
by complex, polyphase ductile deformation between 1.86 and 1.81 Ga.
According to the evolutionary deformation scheme by Aaltonen et al. (2010)
the results of up to five different phases, referred to as D1–D5,
are preserved in the local structural record, each characterized by
structures with distinctive mineral composition, metamorphic grade, geometry,
and kinematics. The most relevant phases to our study are D2 to
D4. During these ductile episodes, a regional and pervasive NE–SW-striking and moderately SE-dipping foliation developed, strain localized
along mesoscopic shear zones parallel to subparallel to the foliation, and
extensive migmatization occurred under amphibolite facies metamorphic
conditions. NNE–SSW- and N–S-striking mylonitic shear zones also formed under
those conditions, whereas later ductile events developed under progressively
lower-grade metamorphism until ca. 1.7 Ga, when brittle deformation
became the dominant deformation style in response to progressive regional
exhumation and cooling (Mattila and Viola, 2014; Aaltonen et al., 2016). The
penetrative, inherited ductile grain that by then characterized the
crystalline basement and that was suitably oriented with regard to the
prevailing stress field was invariably reactivated. This is the case for
several NNE–SSW-striking faults mapped underground in the Onkalo repository,
which clearly overprint earlier D4 shear zones and fully exploit the
pre-existing ductile precursors. Other faults, such as BFZ300, do not show
any clear genetic relation to the older ductile fabric and cut it
discordantly.
As shown in the following section, BFZ300 belongs to a set of subvertical,
conjugate brittle–ductile to fully brittle strike–slip faults characterized
by N–S-trending sinistral and NW–SE dextral faults. Both sets document a
complex history of reactivation and contain evidence for cyclic and
transient switches between brittle and ductile deformation at all scales.
Mesostructural and microstructural studies show that the sinistral faults overprint
and probably reactivate a dextral mylonitic precursor related to earlier
localized ductile deformation (Prando et al., 2019). These faults
locally contain pseudotachylyte injections, which suggests seismic behaviour
during deformation (Menegon et al., 2018). In contrast, dextral
faults cut across the foliation, do not exploit any ductile precursors, and
do not host pseudotachylytes. BFZ300 belongs to this second group of faults.
In the following, we describe its architecture, reconstruct its deformation
history, and constrain the deformation mechanisms and faulting conditions
that prevailed during its nucleation and subsequent development. The
architecture and deformation history of the remarkably different conjugate
structure to BFZ300, which is a sinistral brittle–ductile deformation zone,
is described in the Part II companion paper (Prando et al., 2019).
Methods: fluid inclusion, mineral chemistry, and EBSD analysis
Field documentation and sampling were carried out at the underground Onkalo
exposures of BFZ300 (Fig. 1b), which were necessarily limited in extent to
the actual excavated volume of rock at the time of our study but that,
together with the logged diamond drill cores from the underground
exploration, allow for a well-constrained 3-D reconstruction of the local
geology.
Fluid inclusion measurements were conducted on “fluid inclusion
assemblages” (FIAs), i.e. on petrographically discriminated, cogenetic
groups of fluid inclusions located along trails or (less commonly) within
clusters (Bodnar, 2003a; Goldstein, 2003). By definition, FIAs
are groups of inclusions that have been trapped together (i.e. they are
cogenetic) at a specific stage of mineral formation and, as such, give the
highest level of confidence when characterizing the properties of trapped
fluids and discriminating possible stages of post-entrapment
re-equilibration (Bodnar, 2003b, and references therein). We
applied Roedder's identification criteria of FIAs according to the timing of
entrapment (i.e. primary, secondary, pseudosecondary) in order to link
stages of fluid entrapment with stages of brittle and ductile deformation of
quartz. In this regard, we considered only FIAs that exhibited both similar orientation and
petrographic characteristics at the scale of the thin section to be cogenetic and therefore
representative of one specific stage of brittle deformation and fluid
circulation.
In the selected samples, we studied 28 FIAs entrapped within two
distinct generations of quartz (named Qtz I and Qtz II) forming two
different generations of veins and exhibiting the least petrographic
evidence of post-entrapment overprinting by later ductile and/or brittle
deformation, which provided ca. 800 microthermometric properties. Due to the
well-documented tendency of fluid inclusions to modify their shape, volume,
and composition after their initial entrapment even under low deviatoric stress
conditions (e.g. Diamond et al., 2010; Kerrich, 1976; Tarantola et al.,
2010; Wilkins and Barkas, 1978), working on FIAs that show the least
possible degree of textural re-equilibration is essential when aiming at
constraining the physical and chemical properties of the fluid(s) involved
in the fault activity.
Microthermometric properties of fluid inclusions were determined at the
Department of Biological, Geological and Environmental Sciences of the
University of Bologna using a Linkam THMSG 600 heating–freezing stage
coupled with an Olympus BX51 polarizing microscope. The microthermometry
stage was calibrated by using synthetic fluid inclusion samples at -56.6,
0.0, and 374 ∘C, which correspond to the melting of CO2, ice
melting, and final homogenization of H2O inclusions, respectively.
Obtained accuracies were ±0.3∘C for final ice melting
temperature (Tmice) and ±3∘C for final
homogenization temperature (Thtot). In order to produce an internally
consistent dataset, all phase transitions were exclusively collected for
individual FIAs and measured with the same standard procedure. Samples were
first rapidly cooled to ca. -180∘C and then slowly heated to
detect the potential formation of a solid carbonic phase, eutectic phases,
salt hydrates, ice, and clathrates. The Thtot values were later determined
in the FIAs by heating the samples from room temperature and recording the
mode of homogenization (i.e. by bubble or liquid disappearance). All phase
transitions were measured by using the cycling method described by Goldstein
and Reynolds (1994), and care was taken to also record the minimum and
maximum values for each assemblage. Volume fractions of individual fluid
inclusions, determined as a percentage of the ratio ϕ=Vv/Vtot
(Diamond, 2003), were estimated optically at room temperature
using calibrated charts. Salinity, bulk densities, and isochores were
computed from the measured Tmice values using the HokieFlincs Excel
spreadsheet (Steele-MacInnis et al., 2012, and references therein).
Fluid inclusions were also analysed using micro-Raman spectrometry. Analyses
were carried out at the Department of Mathematical, Physical and Computer
Sciences of the University of Parma (Italy) using a Jobin-Yvon Horiba LabRam
spectrometer equipped with an He–Ne laser (emission line 632.8 nm) and
motorized XY stage. The spectral resolution of the measurements was
determined as nearly 2 cm-1. The confocal hole was adjusted to obtain a
spatial (lateral and depth) resolution of 1–2 µm. Most spectra were
obtained with a 50× objective (N.A. 0.75), although for shallow
inclusions a 100× objective (N.A. 0.90) was also used. The
calibration was made using the 520.7 cm-1 Raman line of silicon. A wide
spectral range (100–3600 cm-1) was scanned for each inclusion for the
presence of CO2, N2, CH4, and H2S, but the final
acquisitions were made mainly between 1100 and 1800 cm-1 for the study
of CO2 spectra and between 2500 and 3300 cm-1 for CH4 and
H2S. The acquisition time for each spectral window was 120–240 s, with
two accumulations. The power on the sample surface is nearly 1 mW, but the
power on the analysed inclusions has to be considered lower due to
reflections and scattering. Analyses were carried out on the vapour bubbles
of the fluid inclusions.
After the calculation of representative fluid inclusion isochores for each
FIA, the pressure corrections were assessed by using the crystallization
temperatures of two mineral pairs – namely chlorite–quartz and
stannite–sphalerite – as independent input parameters. Chlorite–quartz
temperatures were calculated by using the method of Bourdelle and
Cathelineau (2015), which assumes quartz–chlorite equilibrium and uses
ratios of chlorite endmember activities to link the chlorite compositions
with the corresponding formation temperatures through the quartz–chlorite
equilibrium constants. This method is based on the measurements of the
concentrations of the major chlorite components (Si, Fe, Mg) and can only be
applied to chlorites with (K2O+Na2O+CaO) < 1 wt %, which is indeed the case of our chlorites. To estimate the formation
temperature of cogenetic sulfides associated with Qtz II we used the
stannite–sphalerite formation temperature following the method proposed by
Shimizu and Shikazono (1985). This geothermometer uses the temperature
dependency of iron and zinc partitioning between stannite and sphalerite
(Nekrasov et al., 1979) as a temperature indicator of the association of Qtz
II stannite and sphalerite.
Electron probe microanalysis (EPMA) of fault minerals was carried out with a
JEOL-8200 wavelength-dispersive electron microprobe housed at the Department
of Earth Sciences of the University of Milan, Italy. The instrument fits five
wavelength dispersive spectroscopy (WDS) spectrometers utilizing lithium fluoride (LiFH), pentaerythritol (PETJ
and PETH), and thallium acid phthalate (TAP) analysing crystals and an
optical microscope. Samples were probed with a beam size of ∼1µm at a 15 keV and 5 nA beam current. Synthetic and natural materials
were used as calibration standards at the beginning of each session.
Analytical 1σ errors are typically < 4 % for major
elements and for the minor elements.
(a) View to the north and interpretation of the structural elements
of BFZ300. (b) Lower hemisphere, equiangular projection of conjugate fault
segments (blue: dextral faults; red: sinistral
faults), cleavage (green), and Qtz I chlorite veins infilling
joints (black). (c) Slickensides (white dashed line) and
slickenlines (black dashed lines) on a chlorite-decorated, NW–SE-striking
fracture plane at the vein–host interface, indicating dextral strike–slip
kinematics.
Panchromatic cathodoluminescence (CL) imaging was also performed by
using the CL CCD detector adjacent to the optical microscope of the
JEOL-8200 on the sections used for microstructural work. The
electron beam was focused on the sections with an accelerating voltage of
15 kV and a 30 nA beam current. Black-and-white digital images were
collected with a 40× magnification by beam mapping with the CCD detector at
a spatial resolution of 1 µm (beam resolution), which
resulted in imaged areas of 27.8×22.2 mm. The exposure time for
image acquisition was 120 s.
Petrographic thin sections were later studied with the scanning electron
microprobe (SEM) to investigate the crystallographic preferred orientation
(CPO) of selected sites of the quartz veins from the fault core (sample name
TPH-120-4; see Fig. 2 for sample location). Samples were analysed with a
JEOL-6610 SEM equipped with a Nordlys Nano electron backscatter diffraction (EBSD) detector, hosted at the
Electron Microscopy Centre of the University of Plymouth, UK. EBSD detailed
results are reported in the Supplement.
BFZ300 architecture with examples of representative structural
features. The red rectangles locate the areas where detailed outcrop photos
were taken. Stars locate hand and drill core samples. Stars with a black
outline identify samples used for the microthermometric study. Note that the
fault is made of two main segments offset laterally at a sinistral
compressive step-over zone. Fault core quartz veins are shown by thicker
black lines in the schematic model (centre of figure), while blue and white
lines highlight the positions of the two types of quartz veins in the outcrop
pictures. (a) Damage zone made of millimetre-thick, en echelon veins connected by
conjugate shear segments. (b) Detail of (a) showing fractures filled by the
first quartz generation (Qtz I). (c) Two distinct generations of
quartz–chlorite veins recognized in the fault core (Qtz I and Qtz II). (d) Detail of the sinistral compressional step-over zone characterized by
multiple and parallel T fractures, filled by Qtz I. A brecciated body is
cross-cut by the Y planes. (e) Tensional fracture infilled by Qtz I. (f) Compressional structures (P shears) from the step-over zone and relationships
between Qtz I and Qtz II. The Riedel geometry suggests that the Qtz II vein
formed due to the reactivation of the internal principal slip zones
(YII). Note the Qtz II vein cutting the Qtz I vein. (g) Juxtaposed Qtz I
and Qtz II veins. Qtz I veins are thinner and made of a translucent, small-grained quartz. In contrast, Qtz II veins, which contain pockets of sulfide
aggregates, are thicker and made of larger and euhedral quartz. Chlorite
occurs as a minor phase in both veins. Notice the presence of a cataclastic
band between the two veins. (h) Spatial continuity of the chlorite aggregates
within the Qtz II veins, which always grow orthogonal to the vein boundaries.
This open-space filling texture suggests hybrid conditions of reactivation of
the older Qtz I veins. (i) Small quartz breccia formed between the two
generations of quartz veins.
ResultsBFZ300 fault architecture
The studied BFZ300 section is located at a depth of 426 m b.s.l. and is
about 8 m long (Fig. 2a). It strikes NNW–SSE and dips very steeply to
subvertically to the southwest (Fig. 2b). It cuts through high-grade veined
migmatite, interlayered with gneiss and pegmatitic granite. The fault is a
strike–slip fault system formed by two main subparallel fault segments
connected by a mesoscopic sinistral step-over zone. Subhorizontal striae
defined by elongated trails of chlorite grains and kinematic indicators, such
as chlorite slickensides (Fig. 2c) and R and R' planes, invariably indicate
dextral strike–slip kinematics. The most striking mesoscopic characteristic
of BFZ300 is the presence in the fault core of a composite set of almost
continuous quartz veins (between 1 and 20 cm in thickness) along the entire
exposed strike length. A schematic representation of the fault zone is shown
in Fig. 3.
The fault contains a 0.5–2 m thick damage zone separated from the undeformed
host rock by two discrete bounding surfaces (YI planes according to
Tchalenko, 1970; Fig. 2a). The damage zone can be defined in the field on the
basis of the presence of a fractured volume containing sets of conjugate
dextral and sinistral hybrid fractures (Fig. 3a) intersecting to form a
tight acute angle of ca. 38∘ (Figs. 2b, 3a). Laterally continuous,
NNW–SSE-striking Mode I fractures (joints) invariably bisect this angle
(Fig. 2b), helping to constrain the stress field orientation at the time of
fracture formation, with the greatest compressive stress axis σ1 parallel to the Mode I fracture strike and oriented ca. NNW–SSE.
Joints are sharp and have a regular spacing of ca. 10 cm. The joints and the
hybrid fractures of the damage zone contain quartz, referred to as Qtz I
hereinafter, forming veins up to 1–1.5 cm thick (Fig. 3a). Fractures and
faults decorated by Qtz I have a translucent look that reflects the generally
fine grain size of Qtz I (< 1 cm, Fig. 3b). Locally they are formed
by en echelon tensional segments connected by shear planes not decorated by
any quartz infill (Fig. 3b). Joints occur also as barren fractures defining
a penetrative sympathetic fracture cleavage (sensu Basson and Viola, 2004; green
lines in Fig. 2b). Field evidence also suggests that fracture density within
the damage zone tends to increase towards the fault core.
The fault core is bounded by two main discrete slip surfaces (YII;
Figs. 2a, 3d, f, h). It contains, and is defined by, two distinct
generations of quartz veins (Fig. 3c) that are interrupted and offset
laterally by a metric sinistral step-over zone (Fig. 3d–f). The main quartz
vein of the core is infilled by quartz exhibiting the same mesoscopic
appearance of Qtz I in the damage zone; we therefore refer to it as a Qtz I
vein. It is accompanied by a younger, subparallel vein formed by a
milky-white type of quartz with a significantly larger grain size than Qtz I
(> 1 cm) that we refer to as Qtz II (Fig. 3c). Locally, pockets
of cataclasite and breccia formed at the expense of the host migmatitic
gneiss are also observed along and between the two veins (Fig. 3g, i).
The Qtz II vein exhibits a quite irregular, curved geometry (Fig. 3c, h)
and a variable thickness up to a maximum of ca. 20 cm. The minimum Qtz II
vein thickness coincides spatially with an apparent lateral displacement of
the vein. The BFZ300 core varies in thickness between 20 and 30 cm along
most of the exposed fault length but becomes thicker (up to 50 cm) in the
compressional step-over zone that connects the two fault segments that are
offset laterally by ca. 1 m. The sinistral step-over zone is defined by
synthetic T fractures (Fig. 3d, e) and contains a decimetric brecciated
lens (Fig. 3d). T fractures are filled by Qtz I veins (Fig. 3e).
Chlorite is present as a secondary phase, with a modal abundance between 5 % vol
and 10 % vol in both Qtz I and Qtz II veins. In Qtz I veins it occurs as
euhedral–subhedral crystals up to 1–2 mm in size (Fig. 3g). Chlorite is
present mostly as a disseminated, interstitial phase concentrated mainly in
the internal part of the Qtz I veins (Fig. 3g). In the Qtz II vein, however,
it occurs as elongated crystals (5–8 mm in length) arranged perpendicularly
to the walls of the vein, which suggests orthogonal dilation at the time of
opening (Fig. 3h). The Qtz II vein also contains small (1–2 cm) aggregates
of sulfides (sphalerite, pyrite, galena, and chalcopyrite) mainly
concentrated in the central part of the vein (Fig. 3g).
As observed in the field, the presence of Qtz I veins along the joints in
the damage zone and the continuity of the fault core Qtz I vein suggest Mode
I fracturing during Qtz I emplacement (Figs. 2a, 3a, c). The semi-continuous
parallelism of Qtz I and Qtz II veins in the fault core, combined with the
location of the Qtz II vein along the walls of the Qtz I vein, suggests the
partial reactivation of the Qtz I vein during Qtz II emplacement. Dilation
leading to Qtz II emplacement exploited and further reworked the Qtz I host-rock contact, which seemingly had a lower tensile strength than the pristine
migmatite. The reconstructed time relationship between the two vein
generations is also consistent with local evidence of the Qtz II vein partly
cross-cutting the Qtz I vein (Fig. 3f).
BFZ300 microstructural analysis
To constrain the spatial and temporal association of fault rocks and the
type of fluid involved in the deformation, several outcrop samples, each
representative of a specific structural domain, were collected (TPH-120-2,
TPH-120-3, TPH-120-4, TPH-120-5, and TPH-120-6), in addition to samples PH21
and PH22 from diamond drill cores that intersect BFZ300 at the same depth in
an area that is currently not excavated. From these samples we prepared 10
petrographic thin sections (samples TPH-120-2, TPH-120-4, TPH-120-6, PH21,
and PH22) and 9 doubly-polished sections for fluid inclusion analysis
(thickness ∼150µm; samples TPH-120-2, TPH-120-4,
TPH-120-6, PH21, and PH22). Due to the extensive reactivation of the fault
zone and the consequent obliteration of the fluid inclusion (FI) record, the FI study was
carried out only in samples TPH-120-4, TPH-120-6, and PH21. Hand samples and
drill cores localities are shown in Fig. 3.
The microstructural work was carried out on oriented petrographic thin
sections cut orthogonally to the migmatitic foliation and parallel to the
slickenlines.
In the following we provide a description of the microstructural
characteristics of BFZ300 by detailing our findings and observations
separately for the main structural domains of the fault zone.
Microtextural characteristics of Qtz I from the damage zone of
BFZ300 (sample: TPH-120-2). (a) Stitched photomicrographs of a Qtz I vein
interconnecting with a sinistral shear band (crossed nicols). Faulting
kinematics is suggested by drag folds in the host rock. (b) Tip of Qtz I vein
hosted by a sericite-rich cataclastic band of the host rock. (c) Detail of
panel (a) showing open-space filling texture in the Qtz I vein. Notice the
sericite microfractures cross-cutting Qtz I. (d) Panchromatic
cathodoluminescence image of Qtz I showing healed microfractures cross-cutting
the crystal.
Damage zone
Qtz I veins within the damage zone cut across the migmatitic host rock and
form the infill of conjugate sets of hybrid fractures, which, when studied
at the microscale, appear to be formed by dilatant segments joined by
cataclastic shear fractures (Fig. 4a). Shearing on the latter is well
documented by the asymptotic bending into the shear surfaces of foliation
planes formed by the alignment of chlorite and muscovite, both partly
altered to sericite and chlorite, respectively (Fig. 4a). Qtz I infilling
the tensional segments has an average grain size between 200 µm and 3 mm and exhibits a rather heterogeneous texture, from purely blocky to mixed
elongated–blocky (Fig. 4b, c). The largest crystals (800 µm to 1 mm)
are elongated and stretched from the vein walls towards the inner part of
the vein (Figs. 4c, 5a). At least two episodes of vein growth and renewed
dilation, as indicated by the presence of median lines (MLs), are visible
within one of the studied veins and confirm a syntaxial growth mechanism for
the vein (Fig. 5; e.g. Bons et al., 2012). Medial lines are defined by the
alignment of chlorite, sericite, and carbonate aggregates (Fig. 5a, b, d).
Blocky euhedral quartz crystals are also found varying in grain size between
300 and 600 µm. These crystals are juxtaposed against very-fine-grained quartz (< 200 µm) within sericite-rich cataclastic
bands (Fig. 4b). These cataclasites contain also hydrothermally altered
host-rock fragments including pervasively altered K-feldspar-bearing lithic
fragments and phyllosilicates.
Microtextural characteristics of Qtz I from the damage zone of
BFZ300 (samples PH21 and TPH-120-2). (a) Stitched microphotographs of a Qtz I
vein showing elongated–blocky texture with crystals growing obliquely with
respect to the vein boundaries, which suggests growth under oblique
dilatation. A series of median lines (MLs) are marked by (b) sericite crystals
suggesting repeated crack and seal. Quartz crystals show low-temperature
crystal–plastic deformation by undulose extinction and extinction bands. (c) Detail of plastic deformation in damage zone quartz veins: distorted crystals
showing incipient bulging and intracrystalline fracturing. (d) Detail (plane-polarized light) of an ML and secondary fractures both decorated by vermicular
chlorite and aggregates of rare-earth-element-bearing carbonate.
With the exception of the blocky variety, Qtz I crystals exhibit various
degrees of crystal–plastic deformation and recovery. They contain widespread
evidence of undulose extinction and extinction bands (Fig. 5b), and
incipient bulging along grain boundaries is also evident (Fig. 5c).
Millimetric intracrystalline barren fractures are also recognized (e.g. Fig. 5c). Cathodoluminescence imaging of Qtz I from the damage zone also shows
the presence of a dense network of healed quartz microfractures (Fig. 4d),
which demonstrates healing subsequent to brittle deformation and fracturing.
Chlorite occurs along the ML of the veins, secondary cracks, along grain
boundaries, and as inclusions within quartz crystals. It has a vermicular
texture (Fig. 5d) and crystal dimensions up to 50 µm.
Microtextural characteristics of Qtz I from the BFZ300 core (sample
TPH-120-4). (a) Stitched photomicrographs showing the typical heterogeneous
grain size of Qtz I (30–800 µm). (b) Evidence of plastic deformation
of Qtz I from the fault core given by bulging of the largest crystals, wide
extinction bands, and undulose extinction. Note the late brittle fractures
cross-cutting all the previously formed plastic features. (c) Intracrystalline
deformation bands within a large crystal. (d) Detail of (c) showing the
typical grain size of the band (50–250 µm). Intracrystalline deformation
bands are oriented at < 30∘ with respect to the BFZ300 vein
walls and can be up to 2 mm in length. (e) Intercrystalline deformation band
showing a thickening at the triple junction of larger grains. These
intercrystalline bands are parallel to the strike of BFZ300.
Fault core
In the fault core, Qtz I grain size reaches the smallest observed value
(range: 30–800 µm; Fig. 6a), although it is strongly variable within
the vein, suggesting the presence of heterogeneous and complex structural
sub-domains. The earliest post-vein emplacement deformation stage is
reflected by the low-temperature, intracrystalline deformation of the
largest crystals (400–800 µm in size). Undulose extinction, wide
extinction bands (WEBs; Derez et al., 2015), and bulging along grain
boundaries are the most common microstructures ascribable to this
deformation stage (Fig. 6a, b). A first brittle deformation event is
documented by narrow, intracrystalline fractures that cross-cut the largest
quartz crystals (Fig. 6c) and which locally contain new grains of quartz
ranging in size between 20 and 100 µm (Fig. 6d). In more detail, these
new grains form parallel bands that are oriented at a low angle (< 30∘) to the vein walls and that can be up to 2 mm in length and
200 µm in thickness. Plastically deformed Qtz I crystals hosting
these intracrystalline bands of new grains are cut across by another later
set of subparallel intercrystalline fractures, which are interpreted as the
expression of yet another deformation event that occurred under overall
brittle conditions. These fractures are parallel to the strike of BFZ300 and
are in turn sealed by partly recrystallized new quartz grains (grain size:
50–150 µm; Fig. 6e). The cathodoluminescence imaging of these
fractures shows that they are sealed, yielding a homogeneous dark signal
(Fig. S1 in the Supplement). They are locally decorated by
trails of fluid inclusions (Fig. S2a, d) and
can be up to 2.5 cm in length and up to 500 µm in width (Fig. 6a).
EBSD maps were acquired along some of these intercrystalline bands, and
results suggest that the new grains sealing the fractures reflect the
combined effect of initial cracking, grain nucleation, and subsequent partial
dynamic recrystallization (Fig. S2b, c).
Microstructural characteristics of Qtz II from BFZ300 (samples
TPH-120-6, PH22). (a) Stitched photomicrographs of Qtz II vein from the fault
core. Notice the coarse quartz crystals and their elongated–blocky texture.
Primary growth textures are sometimes visible and are marked by solid
inclusions and decrepitated FIAs. (b) Radiate chlorite crystals along a
prismatic Qtz II crystal boundary. Note that Qtz II is cross-cut by numerous
trails of FIs. (c) Panchromatic cathodoluminescence image of the same large
Qtz II crystal from panel (b), showing radiate chlorite along the crystal
boundary and a primary growth zone cut by a set of healed fractures. (d) Euhedral quartz crystals set within opaque phases and cross-cut by a network
of thin microfractures. (e) Reflected-light photomicrograph showing the
opaque mineral assemblage typically associated with Qtz II, i.e. subhedral to
anhedral sphalerite, pyrite, and galena. Chalcopyrite is a minor phase and
occurs as small round inclusions within sphalerite (chalcopyrite “disease”)
or as large subhedral–anhedral masses together with galena.
Qtz II within the fault core is typically coarse grained (individual
crystals: 300 µm–1 cm in size) and exhibits a regular blocky texture
devoid of any shape or crystal preferred orientation (Fig. 7a). Locally,
these large crystals display primary growth textures, such as primary FIAs
oriented parallel to specific crystallographic planes. With the exception of
undulose extinction, Qtz II does not show clear evidence of plastic
deformation, although cathodoluminescence imaging of optically continuous
Qtz II has also shown that a dense network of healed quartz microfractures
locally cross-cuts Qtz II crystals (Fig. 7c). These are relatively thin
(hundreds of micrometres thick) networks that are poorly visible to invisible by
standard petrographic analysis. The only petrographic evidence for these
healed microfractures within quartz is represented by well-defined trails of fluid
inclusions cross-cutting primary growth bands (Fig. 7d).
Chlorite is the second most abundant phase within the fault core Qtz I and
Qtz II veins and occurs with a variety of textures. Aggregates of vermicular
chlorite similar to that occurring in the damage zone are also present in
Qtz I from the core (Fig. 8e), although chlorite with flaky and radiate
textures (Fig. 8f) is also present. The latter is generally 100–300 µm in size and is in textural equilibrium with quartz and rare calcite.
Radiate chlorite overgrowing fractured Qtz II (Fig. 7b) suggests late Qtz II
precipitation.
Associated with Qtz II, a sulfide assemblage made of pyrite, sphalerite,
galena, and chalcopyrite (Fig. 7d, e; see also Fig. 3g) forms aggregates
that are commonly located along quartz grain boundaries. These aggregates
have dimensions between 10 and 600 µm.
Microstructures of the cataclasite juxtaposing Qtz I and Qtz II
veins (sample TPH-120-4). (a) Stitched photomicrographs covering the contact
between the two quartz veins and the intervening 5 mm thick cataclastic band.
(b) Cataclastic band containing large Qtz I fragments (8–12 mm) embedded
within a finer matrix (20–200 µm in size) of sericite and
recrystallized quartz. The largest crystals show lobate boundaries,
suggesting dissolution and local resorption along the clast–matrix interface.
(c) Stylolite seams at the boundary of the cataclasite that strike parallel
to the BFZ300. (d) Reflected-light photomicrograph showing anhedral to
subhedral pyrite, chalcopyrite, stannite, and sphalerite arranged along the
stylolite as residual products of pressure solution. (e) Vermicular and
radiate (f) chlorite aggregates associated with Qtz I close to the
cataclastic band.
Multiply reworked breccias and cataclasites occur within and cross-cut
BFZ300. In the studied sections, a cataclastic band between 5 and 8 mm thick
cross-cuts both Qtz I and Qtz II veins (Fig. 8a), but it is in turn cross-cut by
a different quartz–radiate–chlorite vein displaying evidence of syntaxial
growth. This cataclasite contains poorly sorted and angular quartz clasts
between 8 and 12 mm in size set in a finer (20–200 µm in size) white
mica–quartz matrix. The largest quartz fragments show irregular, lobate
grain boundaries and are affected by undulose extinction. We interpret these
textures as the product of dissolution and cataclastic reworking of Qtz I
Parallel sets of stylolitic seams trend ca. N–S, parallel to the strike of
BFZ300, and mark the two sides of the cataclastic band (Fig. 8a, c). They
host anhedral sphalerite, stannite, galena, pyrite, and chalcopyrite (Fig. 8d), which are coeval with the formation of the Qtz II vein. We interpret
the presence of these anhedral sulfide minerals along the stylolite as the
product of passive concentration by pressure solution.
Schematic summary of main microstructures, fluid properties, and P–T
deformation conditions in the quartz veins of the BFZ300 fault.
Note: microstructures are coupled with the corresponding FI types and P–T
constraints derived from the collected dataset. See text for a further
explanation. Note that we combine structural and geochemical data to
constrain the relationships between stages of mineral-scale deformation and
fluid circulation, which in turn defines the relative chronology of stages
of fluid flow during faulting.
ML: median line; Blg: bulging.
Fluid inclusion dataFluid inclusion petrography
The studied FIAs invariably contain a two-phase fluid (liquid–vapour) and
are mainly arranged in secondary trails within Qtz I crystals in the damage
zone (Type S1) and also within Qtz I fault core, where they form dismembered
(Type S2) trails and also appear as individual clusters inside the crystals
affected by crystal–plastic deformation (Type S3). Within Qtz II, FIAs are
arranged as pseudosecondary (Type PS) and secondary (Type S4) trails.
Representative examples of FI petrographic features are shown for each
structural domain in Fig. 9. Table 1 provides a schematic representation of
the location of the FI types, in addition to their location within the fault
architecture and their fluid properties.
Characteristic textures of FIAs hosted within the BFZ300 quartz
(samples PH21, TPH-120-4, TPH-120-6). (a) Secondary trails cross-cutting large
Qtz I crystals of the damage zone. (b) Detail of (a) showing the phase ratios
of one of the studied secondary assemblages (FIA 3), most representative of
Type S1 FIA. (c) Long secondary transgranular trails cross-cutting Qtz I of
the fault core, dismembered by intercrystalline fractures, infilled by quartz
new grains. Qtz I fault core also hosts a set of short sub-trails developed at
a high angle with respect to the long trails. (d) Detail of Type S2 FIA
entrapped along a preserved secondary fracture trail. (e) Small inclusions
(< 1 µm) arranged along the boundaries of new polygonal quartz.
(f) Example of Type S3 FIA arranged as isolated clusters inside ductile-deformed fault core Qtz I. These trails formed during a brittle deformation
stage that predates ductile re-crystallization. (g) Pseudosecondary FIA
associated with Qtz II chlorite (FIA 11). The enlargement shows the phase
ratio details. (h) Small-scale view of secondary FIAs cross-cutting Qtz II.
(i) Detail of secondary trails cross-cutting euhedral Qtz II (FIA 13). In all
photographs north points up.
Damage zone. Within Qtz I grains (Fig. 9a, b), secondary FIAs are found as trails (Fig. 9a) that parallel what we interpret as healed, old microfractures. In these
assemblages, FIs have a maximum size between 2 and 20 µm, a regular
equi-dimensional shape (i.e. negative crystal morphology), and a volume
fraction, ϕ (ϕ=Vv/Vtot, see Sect. 3), ranging
between 5 and 15 % (Fig. 9b).
Fault core. Qtz I grains host secondary FIAs (Type S2), which are transgranular trails
(Fig. 9c) along healed joints and hybrid fractures. These trails are locally
interrupted and dismembered by aggregates of new, fine-grained quartz grains
(Fig. 9c) and generate a texture that is indeed typical of Qtz I from the
fault core (Fig. 6a). Fluid inclusions vary in size between 1 and 10 µm, have a ϕ of 10 %–20 %, and show a negative crystal morphology
(Fig. 9d). Fluid inclusions are also found as isolated clusters inside
intensely recrystallized quartz domains (Fig. 9c). FIAs inside these
recrystallized quartz domains were pervasively obliterated during later
episodes of ductile deformation. The development of WEBs, intercrystalline
bands, and bulging resulted in the remobilization (i.e. “transposition”
sensu Anderson et al., 1990) of these assemblages. This is regularly observed and
is documented, for instance, by the presence of short, secondary trails of
regularly shaped inclusion oriented at a high angle with respect to a
longer parent trail (Fig. 9c). Morphologically, these trails resemble the
transposed trails documented in high-grade metamorphic rocks (Andersen et
al., 1990; Van den Kerkhof et al., 2014). Different types of fluid inclusion
morphologies are found within the intensely recrystallized quartz domains
(Fig. 9f). Negative crystal morphology is observed in some areas of the
selected samples, but it is uncommon. The
“dismembered” morphology is instead more typical (Vityk and Bodnar, 1995; Tarantola et al.,
2010), and is observed in the relatively large inclusions (> 20 µm). This morphology is made of a central (often empty) inclusion,
showing several tails and re-entrants, surrounded by three-dimensional
clusters of small “satellite” inclusions. These clusters might be arranged
with a quasi-planar geometry inside the host (i.e. in a trail-like fashion).
Another typical texture found in most assemblages is the “scalloped”
morphology of small- to medium-sized inclusions (< 10–15 µm),
which is defined by the presence of indentations, embayments,
irregularities, and sharp tips of the inclusion walls (Fig. 9f). Small
inclusions (< 1 µm) are also found at the edge of the straight,
regular boundaries of new quartz grains; they are mostly dark, i.e. they are
vapour-rich or empty, and are equant in shape (Fig. 9e). Although small
inclusions do not allow for a microthermometric study of the fluid-phase
behaviour in this structural domain, they confirm the complex reactivation
history of BFZ300.
Qtz II contains both pseudosecondary (Type PS) and secondary (Type S4)
assemblages (Fig. 9g, i, h). The first type is arranged in trails that cut
at a low angle the hosting quartz but not the neighbouring phases (e.g.
chlorite). In these assemblages, FIs are relatively large (2–45 µm)
and exhibit an elongated shape and ϕ values between 15 % and 30 % (Fig. 9g). Type S4 FIAs (Fig. 9h) host two-phase inclusions whose size (5–35 µm) is similar to that of PS trails but show a ϕ between 30 %
and 40 % (Fig. 9i).
Primary FIAs are also present along growth planes of Qtz II, where they have
a relatively large size (20–50 µm; Fig. S3a, b, c). They present irregular and dismembered textures,
which suggest intense post-entrapment re-equilibration.
In summary, our microtextural study shows that the FIAs to be selected for
the microthermometric study are only those hosted within Qtz I and Qtz II
crystals with little to no recrystallization and whose inclusions have
textures corresponding to the least intense post-entrapment re-equilibration
(Bodnar, 2003b, and references therein; Tarantola et al., 2010). These are
the pseudosecondary and secondary FIAs in which dendritic or transposed
inclusions are absent and in which the host quartz exhibits only undulose
extinction (S1, S2, S4, and PS).
Microthermometric data of the studied FIAs. Panels (a)–(d) show the bulk
salinities of individual FIAs calculated from the Tmice data, while panels
(e)–(h) refer to the temperatures of final homogenization in the same
assemblages. Notice that the data report the properties of individual FIAs
according to their occurrence within Qtz I of the damage zone, Qtz I from the
fault core, and Qtz II from the fault core. Notice that pseudosecondary (PS)
and secondary (S) FIAs identify progressive later stages of fluid entrapment
and can be used to constrain the fluid properties in the fault zone. Notice
also that the measured ranges of Thtot spread across T intervals that
are too large to represent entrapment at equilibrium (e.g. FIA 7 of Qtz I
from fault core: 130–320 ∘C), which suggests post-entrapment
re-equilibration of the inclusions. Fluid bulk composition is expressed as
salinity, which is conventionally reported as weight percent of NaCl
equivalents (wt % NaCleq; Roedder, 1984).
Microthermometry
Damage zone. Secondary FIAs hosted within Qtz I from the damage zone (Type S1) show a
range of Tmice between -5.9 and -0.1∘C, which corresponds
to a salinity of 0 wt % NaCleq–9 wt % NaCleq (Fig. 10a). In these FIAs, final
homogenization (Thtot) occurs into the liquid phase (i.e. by
disappearance of the vapour bubble) and mainly between 150 and 400 ∘C
(Fig. 10e).
Fault core. The secondary FIAs hosted within Qtz I in the fault core (Type S2) show a
range of Tmice between -8.2 and -0.4∘C, which corresponds
to salinities between 0 wt % NaCleq and 14 wt % NaCleq (Fig. 10b), and final
homogenization occurs into the liquid phase between 130 and 410 ∘C (Fig. 10f).
Pseudosecondary FIAs entrapped within Qtz II (Type PS) show a range of
Tmice between -13.6 and -0.1∘C, which corresponds to a
salinity range between 0 wt % NaCleq and 18 wt % NaCleq (Fig. 10c); final
homogenization occurs into the liquid phase and between 150 and
440 ∘C (Fig. 10g). Secondary FIAs in Qtz II (Type S4) show a
range of Tmice between -11 and 0 ∘C, which corresponds to
a 0 wt % NaCleq–15 wt % NaCleq range of salinity (Fig. 10d), while final
homogenization into the liquid phase is between 130 and 430 ∘C (Fig. 10h).
As no gases were determined during microthermometric analysis (i.e. melting
of carbonic-phase or clathrate hydrates was not detected during the
freezing experiments), additional micro-Raman analysis was performed on a
set of representative FIAs (samples: TPH-120-4; TPH-120-6; PH21; PH22).
Aqueous fluid inclusions hosted by both Qtz I and Qtz II show peaks at
the characteristic wavenumbers of CH4 (2917 cm-1) and CO2 (1388 cm-1). These peaks were determined as weak in all spectra, and CO2 detection was sporadic in a few inclusions of only one sample of the fault
core (TPH-120-4). Although spectroscopic detections, the CO2- and
CH4-bearing inclusions are not systematically associated with specific
quartz vein generations or microstructures (i.e. intracrystalline healed
cracks, WEB planes, intercrystalline fractures). We therefore cannot
associate the presence of CO2 and/or CH4 to any specific
deformation stage of the fault.
Such spectroscopic determinations are consistent with the lack of
microthermometric evidence of carbonic-phase or clathrate hydrates during
the freezing experiments (Rosso and Bodnar, 1995; Dubessy et al., 2001).
The impossibility of detecting CO2- and CH4-bearing fluids during the
freezing experiments indicate a gas pressure that is systematically lower
than that required to observe clathrate dissociation (e.g. 1.4 MPa in
CO2–H2O fluids; Rosso and Bodnar, 1995); i.e. it shows low gas
concentrations. As a consequence, we have modelled the fluid phases as
simple H2O–NaCl systems.
Representative chlorite EPMA from various structural zones of BFZ300.
Chlorite composition has been determined for several generations of
chlorites in association with Qtz I and II, namely vermicular chlorite
associated with Qtz I from the damage zone, vermicular and radiate chlorite
associated with Qtz I from the fault core, and radiate chlorite associated
with Qtz II from the fault core (Table 2). Chlorite compositions are shown
in Fig. 11a and are expressed as a function of the Fe/(Fe+Mg) ratio.
Chlorite compositional data are presented according to the structural domain
of the fault they are associated with and to the corresponding texture.
Vermicular chlorite associated with Qtz I in the damage zone has an XFe range
between 0.50 and 0.55, while vermicular chlorite associated with Qtz I from
the fault core has an XFe of 0.53. Radiate chlorite associated with Qtz I
from the fault core has an XFe range between 0.71 and 0.81, while radiate
chlorite associated with Qtz II from the fault core is between 0.65 and
0.80, constraining compositions within the ripidolite and aphrosiderite
endmembers.
Chlorite chemical composition diagram and mineral pair
geothermometry applied to the assemblages of the Qtz I and Qtz II veins. (a) Chlorite compositional diagram based on Hey (1954). Green, red, pink, and
light blue symbols indicate distinct chlorite textures in association with
Qtz I and Qtz II veins. (b) Chlorite–quartz formation temperature estimated
using the method of Bourdelle and Cathelineau (2015). The maximum temperature
is from the Qtz I chlorite pair from the fault core. (c) Estimated
temperature of formation of sphalerite–stannite in association with the Qtz II
vein (based on Shimizu and Shikazono, 1985). The region of the plot that was
calibrated with this geothermometer lies between the 250 and 450 ∘C isotherms.
Temperature–composition relationships for the quartz–chlorite pair portrayed
in the T–R2+–Si diagram of Bourdelle and Cathelineau (2015) show that,
in the hypothesis of quartz–chlorite equilibrium, the precipitation of
vermicular chlorite within the Qtz I of the damage zone took place in the
175–240 ∘C range (green diamonds in Fig. 11a). This range is
distinct from that of the vermicular and radiate chlorite from Qtz I of the
fault core, which is probably ca. 350 ∘C because the measured
R2+–Si compositional parameters (R2+= Mg+Fe) plot at the
edge of, or slightly outside, the calibrated region of the Bourdelle and
Cathelineau plot (red diamonds in Fig. 11a). We stress that the high-T
chlorite plots parallel to the 350 ∘C isotherm, suggesting that
it most probably precipitated at the same or at a similar temperature.
Radiate chlorite associated with Qtz II from the fault core is instead
compatible with an equilibrium precipitation at 160–220 ∘C
(light blue diamonds in Fig. 11a).
Representative EPMA of sulfides associated with Qtz II.
a Located within cataclastic band and close to stylolite.
b Located along stylolite.
Sphalerite and stannite compositions from locations indicated by b have
been used to calculate the temperatures of sphalerite–stannite equilibrium
following the geothermometer of Shimizu and Shikazono (1985). See text for a
further explanation.
The collected EPMA data show that the sulfides associated with Qtz II have
compositions that approach those of pure phases (Table 3). Pyrite has trace
element concentrations (Cu, As, Pb, Ni, Zn) that are in general below the
EPMA detection limit, while galena, sphalerite, and chalcopyrite show only
some significant trace contents of Fe and Zn (e.g. Fe: 0.22 wt %–1.00 wt % in
galena; Zn: 0.11 wt %–3.95 wt % in chalcopyrite). Pyrite and sphalerite from
the Qtz II veins (Fig. 7e) have trace element concentrations that are,
again, mostly below detection limits.
The stylolites bordering the cataclasite bands described above and formed at
the contact between the Qtz I and Qtz II vein contain pyrite, galena, and
the sphalerite–stannite pair (Fig. 8a, c, d), with the latter showing the
largest compositional variation. This pair represents a mineral
geothermometer because the partitioning of Zn and Fe between sphalerite and
stannite was demonstrated to be temperature dependent but pressure
independent (Nekrasov et al., 1979; Shimizu and Shikazono, 1985). In the
14 analysed pairs, stannite shows a range of Zn concentrations varying
between 0.48 wt % and 3.25 wt %, while those of Fe, Cu, and Sn vary
within narrow ranges (Fe: 12.74±0.56 wt %; Cu: 28.30±0.33 wt %; Sn: 27.65±0.71 wt %). Sphalerite in the pair has
concentrations of Fe and Zn of 7.63±0.87 wt % and 56.68±1.17 wt %, respectively. These ranges allow for the calculation of the
partition coefficient (KD) of the following reaction: Cu2FeSnS4 (in
stannite) + ZnS (in sphalerite) =Cu2ZnSnS4 (in stannite) +
FeS (in sphalerite). We have used the logkD–T relationship of Shimizu
and Shikazono (1985) to calculate the formation temperature of the
pair, which is portrayed in the
(CCu2FeSnS4/XCu2ZnSnS4)–(XFeS/XZnS) plot of Shimizu and
Shikazono (Fig. 11b). The resulting 220–305 ∘C interval lies at
the low end of, or slightly outside, the 250–350 ∘C interval of
the geothermometer.
Therefore, we consider the 250–305 ∘C interval as the best
estimation of the formation T of sphalerite–stannite in the stylolite.
Discussion
Our work constrains the architecture and the environmental conditions at
which BFZ300 deformation took place. Field and petrographic observations
support the idea of transiently elevated fluid pressures, cyclic
frictional–viscous deformation and progressive, and discrete strain
localization (Figs. 2, 3). Analytical data suggest that these deformation
cycles took place at the BDTZ. In the following, we discuss these
constraints by systematically considering our different analytical results.
Fluid inclusion data and mineral pair geothermometry
Field evidence combined with microstructural observations, fluid inclusion
analyses, and the documented distinct generations of synkinematic chlorites
confirm that Qtz I and Qtz II veins precipitated from distinct batches of
aqueous fluid (i.e. H2O–NaCl) that infiltrated the fault zone during
different stages of its evolution.
We documented a wide range of bulk salinity for each FIA entrapped within
the quartz veins in each structural domain (Fig. 10a–d). This suggests
post-entrapment re-equilibration of fluid inclusions (Bakker and Jansen,
1990; Diamond et al., 2010). The Thtot varies between ca. 130 and 440 ∘C without a clear mode or a skew (Fig. 10e, h), indicating that no
common range of entrapment temperature can be identified in the dataset.
Therefore, we conclude that even the properties of petrographically intact
FIAs do not correspond to chemically well-preserved assemblages. Indeed, the
ranges of Thtot in individual FIAs are typically of the order of
150–200 ∘C (Fig. 10e–h), i.e. a value that is much higher than
the ∼10∘C range expected for homogeneous FIAs
entrapped isochorically and isoplethically (Fall et al., 2009; Vityk and
Bodnar, 1998) and that demonstrates post-entrapment re-equilibration
(Vitik and Bodnar, 1998; Bodnar, 2003b; Sterner
and Bodnar, 1989; Invernizzi et al., 1998). A major implication of fluid
inclusion re-equilibration in our study is that the calculated fluid
properties do not rigorously reflect those of the pristine fluid originally
entrapped within BFZ300, but rather those of a fluid that modified its
properties during the fault activity.
Then, a possible approach to interpret our FI dataset is a comparison with
experimental work on synthetic fluid inclusions subjected to a range of
post-entrapment re-equilibration conditions (Bakker, 2017; Bakker and
Jansen, 1990, 1991, 1994; Vityk and Bodnar, 1995, 1998; Vityk et al., 1994;
Invernizzi et al., 1998). A straight comparison to the
experiments is in our case difficult because most experimental work was
carried out at high P–T conditions (500–900 ∘C; 90–300 MPa), and
also only few experiments were carried out under deviatoric stress
conditions that approach those of natural rocks (Diamond et al., 2010;
Tarantola et al., 2010). Despite these limitations, however, some key
experimental results provide fundamental constraints on our dataset. First,
both hydrostatic and uniaxial compression experiments showed that in each
re-equilibrated FIA a number of inclusions survive the
modified post-entrapment P–T conditions virtually intact, showing that only severe deformation
leads to total re-equilibration and complete obliteration of pristine
inclusions (i.e. Δσ > 100 MPa in uniaxial
compression experiments; > 400 MPa change of confining P in
hydrostatic experiments). Second, under conditions leading to only low to
moderate re-equilibration, the bulk chemical composition of the fluid
inclusions does not change significantly from that of the pristine
inclusions.
All of this implies that natural quartz samples with microstructures typical
of moderate T deformation, such as deformation lamellae, deformation bands,
undulose extinction and bulging, and hosting FIAs with moderately
re-equilibrated textures, should still contain a number of inclusions whose
properties resemble those of the pristine fluid. In this scenario, our
microthermometric dataset can be used to constrain the more probable
salinity ranges of the fluid batches which triggered the BZ300 reactivation
stages. Two possible interpretations of the microthermometric dataset can
follow and we can give the corresponding different salinity ranges for the fluids.
One possibility is that the different quartz veins and the fluids trapped
within fluid inclusions originated from multiple pulses of a single low-salinity fluid, with a salinity between 0 wt % NaCleq and 7 wt % NaCleq, as shown by
the distribution of > 70 % of the bulk salinities skewed
towards values of 7 wt % NaCleq or less (Fig. 10a–d). Thus, it is possible
that aliquots of the 0 wt % NaCleq–7 wt % NaCleq FIAs from Qtz I and II crystals from
both the damage zone and fault core are still representative of the pristine
sampled fluid. These inclusions would be those that survived or were
relatively less affected by deformation events post-dating their entrapment.
Inclusions falling outside the most typical 0 wt % NaCleq–7wt % NaCleq salinity range
would instead correspond to those which progressively modified their
properties as a consequence of fluid–rock interaction during faulting and
that experienced significant H2O loss and consequent salinity increase
during the successive stages of fault deformation (Bakker
and Jansen, 1990; Diamond et al., 2010). The large documented range of
Thtot lacking a specific mode observed in individual FIAs is the
product of fluid density changes caused by fluid inclusion re-equilibration
during post-entrapment deformation. This would have happened repeatedly and
cyclically within the host quartz during all ductile and brittle stages of
deformation of the multistage deformation history of BFZ300.
Alternatively, multiple batches of fluids with different salinities (from
low to intermediate salinity) may have infiltrated and evolved within BFZ300
during its activity. In fact, considering the salinity dataset presented for
each structural domain, fluid salinity can be seen clustering in restricted
ranges typical for each domain: (1) the salinity of 60 % of secondary fluid
inclusions in Qtz I from the damage zone is between 0 wt % NaCleq and 1 wt % NaCleq; (2) > 80 % of the secondary inclusions in Qtz I from the fault core
preserve a salinity in the 1 to 5 wt % NaCleq range; (3) 75 % of
pseudosecondary inclusions in Qtz II show salinity values between 6 wt % NaCleq and 11 wt % NaCleq; and (4) ∼70 % of the secondary inclusions
trapped within Qtz II show salinity values between 0 wt % NaCleq and 3 wt % NaCleq.
These clusters may best represent the original compositional ranges of
different batches of fluids, each involved during a different faulting
stage. Salinities outside these ranges may instead be explained again as
resulting from the post-entrapment re-equilibration of those fluids with
different salinities. This hypothetical scenario, in which chemically
distinct fluids entered into the fault and interacted with the rock at
different times (e.g. Selverstone et al., 1992; Boiron et al., 2003; Famin
et al., 2005), is also reinforced by several lines of observations, such as
the variation of chlorite composition, the slight change in
paragenesis–redox state within Qtz II and Qtz I veins (i.e. the absence of massive
sulfides in Qtz I), and the prolonged history of faulting (see below).
P–T diagrams showing the ranges of P–T trapping conditions of the
analysed fluid inclusions: (a) secondary inclusions in Qtz I from the fault
damage zone; (b) secondary inclusions from Qtz I in the fault core; (c) pseudosecondary inclusions trapped in Qtz II in the fault core; and (d) secondary inclusions in the Qtz II. Thin dashed lines indicate maximum and
minimum isochores of FIAs in each structural domain. The coloured areas
identify the probable P–T trapping ranges defined by (i) the slope and
position of the fluid inclusion isochores as determined by the most
representative salinity and homogenization temperature range (see the Supplement for details); (ii) mineral pair geothermometry and (iii) hydrostatic
and lithostatic fluid pressures computed assuming a regional geothermal
gradient of ca. 40 ∘C km-1 (assuming retrograde conditions of 4 kbar
and 650 ∘C; from Kärki and Paulamäki, 2006). The
liquid–vapour equilibrium curve for the H2O–NaCl modelled fluid is also
indicated.
Fully aware of the interpretative uncertainties of our dataset, we have
combined the microthermometric data of the studied FIAs with the independent
quartz–chlorite and sphalerite–stannite geothermometers to constrain the
most probable fluid pressure during the faulting events. With this approach,
we use the formation temperatures of the mineral pairs as independent
geothermometers and consider the intersection between these values and the
FIA isochores to derive the ranges of trapping pressure (Roedder and
Bodnar, 1980). In Fig. 12, we present the ranges of the possible pressure
(Pf) of the fluids involved during faulting as calculated by combining
the fluid inclusion data with the constraints provided by the mineral pair
geothermometry, the hydrostatic and lithostatic pressure gradients, and a
possible geothermal gradient (e.g. Van Noten et al., 2011;
Selverstone et al., 1995; Jaques and Pascal,
2017). The reconstructed regional gradients at the time of vein emplacement
are derived from peak metamorphic conditions (4–5 kbar; 650–700 ∘C, leading to ca. 40 ∘C km-1; from Kärki and
Paulamäki, 2006). Hydrostatic and lithostatic pressures are then
calculated by using pure water density and assuming a rock density of 2700 kg m-3, respectively. These gradients are used to constrain the upper
and lower bounds to physically possible fluid pressures. We computed the
maximum and minimum isochores by using the entire salinity and Thtot ranges obtained from the FIAs in each structural domain (Fig. 10). We also
computed the isochores of the inclusions with the most representative
salinity estimates for each structural domain obtained by comparing the
frequency diagrams (Fig. 10) with the Thtot vs. salinity plots
(Fig. S4). Considering the peak temperature of each
structural zone from the geothermometric estimations in combination with the
computed isocores, the estimated peak conditions of the fluid pressure are
(1) 80 MPa for Qtz I from the damage zone, (2) 210 MPa for Qtz I from the
fault core, (3) 140 MPa from pseudosecondary inclusions in Qtz II from the
core, and (4) 180 MPa from secondary inclusions in Qtz II, still from the core
(Fig. 12; Table 1).
In addition to the Pf peak conditions we can also constrain the
physically possible fluid pressure ranges for each stage of fluid ingress,
which are derived by considering the temperature range estimated for each
structural domain. Thus, for the damage zone, a Pf interval of 50–80 MPa (Fig. 12a) can be derived by intersecting the range of T obtained from
the chlorite–quartz pair in the Qtz I from the damage zone with the range of
isochores from the same quartz. As to the fault core, we combine the 350 ∘C constraint from the chlorite–quartz pair from Qtz I in the
fault core with the isochores from the same quartz, which yields Pf
between ca. 30 and 210 MPa (Fig. 12b). Similarly, the intersection between
the equilibrium T of the sphalerite–stannite pair in the Qtz II from the
fault core and the range of isochores of the pseudosecondary FIAs of Qtz II
defines Pf values between 50 and 140 MPa (Fig. 12c). Estimations from
secondary FIAs in Qtz II constrain a range between 40 and 180 MPa (Fig. 12d).
As also supported by the microstructures described above, we propose that
these values are sufficiently accurate to constrain at least four stages of
fault reactivation, each triggered by a fluid with distinct physical and
compositional properties.
As shown by the T vs. P plots in Fig. 12, the secondary FIAs entrapped in
Qtz I from the damage zone constrain the lowest value of Pf (i.e. 50–80 MPa) of the entire dataset. We interpret this not as representative of the
early BFZ300 localization, but rather as possibly resulting from fluid
entrapment during a later stage of fault reactivation at T∼200∘C. This is also consistent with the calculated temperature
of crystallization of the vermicular chlorite associated with Qtz I from the
damage zone (175–240 ∘C, Fig. 11b) and with the secondary nature
of the entrapped FIAs. Also, the most abundant salinities observed in the
Qtz I from the damage zone (0 wt % NaCleq–1 wt % NaCleq) coincide with the lowest
Thtot measured in the same structural domain. Later fracturing of Qtz I
in the damage zone may thus have been coeval with the formation of
vermicular chlorite preserved therein, which is found along secondary cracks
and median lines (Fig. 5d).
In light of these considerations, we propose that initial BFZ300
localization occurred in the presence of a fluid with T and P of at least
350 ∘C and 210 MPa, respectively. Later faulting continued by
cyclic brittle–ductile switches induced and assisted by fluid batches at
progressively lower temperature and pressure.
Conceptual model of the temporal and mechanical evolution of the
BFZ300 fault zone (see text for more details). Grey lines: traces of
metamorphic foliation. Black lines: fractures related to the BFZ300
structural development. (a) Initial embrittlement of the migmatitic basement
occurred by fracture coalescence (red line) under (b) initial lower
differential stress conditions and high fluid pressure, followed by a
transient increase in differential stress. A first generation of quartz veins
(Qtz I) precipitated inside the diffuse network of joints and hybrid–shear
fractures which formed during this first deformation stage. (c) Progressive
strain localization and fluid channelling within the fault core occurred by
(d) episodically renewed fluid pressure build-up driven by cycles of brittle
and ductile deformation. (e–g) Progressive exhumation and cooling of the
fault system occurred concomitant with several brittle reactivation episodes
of the fault zone under hybrid conditions and fluid pressure lower than
during the previous deformational stages. Lastly, a second generation of
quartz veins (Qtz II) was emplaced, mainly along the principal slip
boundaries of the fault core, following the Qtz I vein as shown by (h) the
strength profile across the fault architecture, which suggests lower tensile
strength values (and hence higher reactivation potential) along the Qtz I
vein and host-rock walls.
Structural evolution and fluid flow: a conceptual model
Based on the integration of field, microstructural, thermometric, and fluid
inclusion constraints (Table 1), we propose a conceptual model for the
structural evolution of BFZ300 (Fig. 13). The fault's finite strain results
from several slip episodes mediated by multiple events of fluid ingress and
fluid–rock interaction. A first constraint provided by our study is that the
analysis of the bulk chemical composition of the fluids that cyclically
ingressed the fault suggests the likely presence of several batches of
fluids of varying salinity and composition.
The embrittlement of the Olkiluoto metamorphic basement (time t1 in
Fig. 13a, b) represents the initial stage of the deformational history of
BFZ300, when conditions for brittle dilation and fracturing of the
Paleoproterozoic basement were first met in a transient fashion. We propose
that brittle failure under still ductile environmental conditions was caused
by transiently elevated Pf (> 210 MPa), as also demonstrated
by field evidence of hydrofracturing (pure tensional en echelon veins at
the BDTZ depth). Hydrofracturing of the host basement is also indicated by
the emplacement of Qtz I veins within the diffuse network of joints and
conjugate hybrid–shear fractures of the damage zone (Figs. 13a, 3a, b).
These brittle features are quite broadly distributed, suggesting an initial
volumetrically diffuse strain distribution. Their formation caused the
overall mechanical weakening of the host-rock volume, which in turn
facilitated later strain localization. Brittle structures formed during this
stage are discordant to the ENE–WSW-striking metamorphic foliation (Fig. 1b), which they cut at a high angle (Fig. 13a). Conditions for tensional and
hybrid failure require low differential stress, i.e. σ1–σ3≦4T, where T is the tensional strength of the
rock. Opening of fractures caused a stress drop, sudden increase in
permeability, fluid venting, and inhibited further build-up of Pf.
Dilatant fractures were partially infilled by Qtz I, which precipitated from
a first fluid with inferred low salinity (in the range between 1 wt % NaCleq and 5 wt % NaCleq). Precipitation of Qtz I and formation of veins within these
fractures caused hardening of the system. The progressive recovery of shear
stresses altered the overall background stress conditions such that failure,
after causing initial pure dilation, was later accommodated by hybrid
extensional and, eventually, by shear fracturing (Fig. 13b), thus forming
laterally continuous and interconnected shear fractures associated with
breccia pockets and cataclasites (Fig. 3d, g, i). Conjugate shear fractures
connected the previously formed extensional fractures through fracture
coalescence (e.g. Griffith, 1920; Sibson, 1996; Fig. 13a). At the
microscale this is demonstrated by the elongated–blocky texture of Qtz I
crystals from the damage zone (Figs. 4c and 5a), where crystals grew at a high
angle to the vein boundaries (thus suggesting initial near-orthogonal
dilation) that are physically connected by cataclastic shear bands to form a
fault-fracture mesh (e.g. Sibson, 1996). Cataclastic bands formed at the
expense of the migmatitic host rock are enriched in authigenic, synkinematic
sericite, likely due to the interaction between K feldspar and fluids
circulating in the dilatant fault zone (Fig. 4b). Shear fractures thus
deformed the migmatitic host rock to connect dilatant and mostly Qtz-I-filled tension gashes during a continuum of deformation. The conjugate
shear fractures ascribable to this stage invariably define tight acute
angles (Figs. 2b, 3a), which we take as further evidence of overall low
differential stress conditions at the time of failure (Fig. 13b).
In synthesis, Qtz I veins from the damage zone are interpreted as the
expression of the earliest stage of fault nucleation, before strain
localization affected a progressively narrower rock volume to eventually
form the main fault core. Indeed, the mesoscale and microscale features observed
in the damage zone Qtz I, lacking the pervasive crystal–plastic deformation that
otherwise occurred in the fault core Qtz I, are used to document the initial
stage of embrittlement. Based on geometric, kinematic, and deformation style
characteristics, we tentatively assign this deformation episode to Stage 1
by Mattila and Viola (2014, their Fig. 18), i.e. to a discrete brittle
episode considered the expression of the earliest onset of brittle
conditions in southwestern Finland ca. 1.75 Ga, under overall NW–SE to
NNW–SSE transpressive conditions.
Further deformation of the BFZ300 (time t2 in Fig. 13c) occurred by
progressive inward strain localization and narrowing of the actively
deforming volume of the deformation zone (from a wide damage zone to a
narrow fault core). The early BFZ300 core, consisting of the main Qtz I vein,
is interpreted as having formed at this stage within an overall dextral
strike–slip kinematic framework. Emplacement of the Qtz I vein in the core
represents the last pulse of this brittle deformational episode (Fig. 13b).
Major fluid venting was likely associated with it such that the system,
once brittle failure in the core had occurred by hydrofracturing, moved back
to a more diffuse deformation style typical of the still prevailing ductile
conditions. Microscopic evidence of crystal–plastic deformation and dynamic
recrystallization (Fig. 6a, b; Table 1) overprinting the early brittle
structures of Qtz I in the fault core supports slow strain rate conditions
during deformation. However, this ductile background deformation was
punctuated by renewed and cyclically transient embrittlement as documented
by healed fractures shown by trails of secondary fluid inclusions cutting
across both the ductile fabrics and the earlier brittle deformational
features (Figs. 6c, S2a, b). EBSD analysis of the new grains documented
along healed microcracks also suggests that they likely nucleated from
fluids circulating in the early fractures before being later deformed in the
low-temperature plasticity regime. Thus, we show that at the BDTZ
“neocrystallization” by nucleation and growth in fractured fragments and
dynamic recrystallization (typically by bulging and subgrain rotation)
coexist and compete in the overall microstructural evolution of quartz (e.g.
Kjøll et al., 2015). Repeated fluid ingress and related deformation
would, in addition, have also caused some of the post-entrapment
equilibration of the FI, as discussed above.
Mattila and Viola (2014) describe a second brittle stage (referred to as
Stage 2, their Fig. 18) during which a ca. N–S- to NNE–SSW-oriented episode of
transpressional deformation affected southwestern Finland. Geometric and
temporal relationships between structures of Stage 1 and 2 (see also Viola
et al., 2009) were used to infer a clockwise rotation of the horizontal
compression direction from NW–SE (Stage 1) to NNE–SSW (Stage 2). Consistent
with the kinematic framework of Stage 2, we propose here that during
progressive regional exhumation and cooling to entirely brittle conditions,
the BFZ300 deformation continued through a further, distinct deformation
phase (t3 in Fig. 13e). This stage accommodated the selective
reactivation of the BFZ300 core, with renewed dilation due to the rotated
σ1 during Stage 2 acting subparallel to the strike of the Qtz
I vein in the BFZ300 core. Localized dilation in a still fluid-rich system
allowed for the emplacement of the Qtz II vein (Fig. 13e). Our estimations
indicate peak conditions of Pf 140 MPa and T≈305∘C. The BFZ300 core was reactivated by an intermediate-salinity fluid (in the
range between 6 wt % NaCleq and 11 wt % NaCleq) under overall hybrid conditions (Fig. 13f), as suggested by the irregular thickness and curved geometry of the Qtz
II vein therein and by the synkinematic chlorite crystals that are
stretched orthogonally to the vein boundaries (Fig. 3h). The Qtz II vein
invariably localized along at the contact between Qtz I and the host rock
(Figs. 3f, 13e), suggesting selective reactivation along the pre-existing
principal slip zones, which represented the weakest part of the fault (Fig. 13h). Evidence for mesoscale hybrid fracturing and our Pf estimates
(Fig. 12) suggest that Pf was lower than that of the earlier
deformation stages during Qtz I emplacement.
BFZ300 underwent one or more events of brittle fracturing and induration
(Fig. 13g), as suggested by the CL imaging of Qtz II crystals (Fig. 7c). The
fluid pressure peak value for this structural stage is ca. 180 MPa.
A possible, very late reactivation stage of unknown age is also documented
by the secondary chlorite associated with Qtz I in the damage zone (Fig. 5a, d). Also, the stylolitic seams striking parallel to the BFZ300 fault zone
suggest compression oriented ca. E–W, i.e. subparallel to the inferred
Sveconorwegian main shortening direction (e.g. Viola et al., 2011). The
sphalerite–stannite mineral pairs arranged along these structures were
possibly concentrated through pressure solution during this deformational
stage.
Skyttä and Torvela (2018) proposed that the BFZ300 is a brittle
structure localized onto a zone of incomplete structural transposition
inherited from the earlier ductile history of the Olkiluoto basement.
However, in our mesoscale and microstructural analysis we did not find
evidence of any ductile precursor, and we note that BFZ300 cuts the ductile
structural grain at a high angle, which excludes any reactivation of precursor
ductile fabrics.
Implications for seismic deformation at the base of the BDTZ
This study demonstrates the role of overpressured fluids in strain
localization during the incipient stages of fault nucleation and subsequent
reactivation(s) at the BDTZ. The maximum estimated fluid pressure and fluid
temperature conditions derived in this study (peak conditions of 210 MPa and
350 ∘C) are indeed realistic for the base of the seismogenic zone
in the continental lithosphere (e.g. Scholz, 1990, and references therein)
where the brittle–ductile transition for quartz occurs.
Mechanical models of long-term deformation (Rolandone and Jaupart, 2002)
propose that deformation at the brittle–ductile transition can be reasonably
described as mostly accommodated by intermittent and concomitant
coseismic slip and ductile flow. Hydrofracturing, as documented in this
study by the Qtz I and II veins, is possibly related in that context to
seismic failure. Faults accommodating hydrofracturing are indeed commonly
interpreted as seismogenic (e.g. Sibson, 1992a; Cox, 1995), particularly at
depth.
Our study confirms this view because BFZ300 contains not only brittle fault
rocks overprinting and overprinted by veins, but also clear-cut evidence of
mutually overprinting brittle and ductile deformation (Fig. 6). In light of
the field observations discussed and the constraints derived, we suggest
that BFZ300 behaved in a seismic way at least during the
emplacement of the principal Qtz I and Qtz II veins. In this perspective,
two possible scenarios can be considered to explain the genetic
relationships between BFZ300 and a possible seismic behaviour of the crust
during the Svecofennian orogeny. In a first scenario, the quartz veins of
the fault core would represent the result of coseismic rupture during the
mainshocks of a fully developed seismic cycle. Pore pressure fluctuations
caused the repeated transient embrittlement of the rock mass, which was
otherwise under overall ductile conditions. The documented brittle–ductile
cycles are thus the expression of coseismic fracturing and aseismic creep
between the individual shocks, as shown by viscous deformation overprinting
the brittle features, guided by the residual differential stress.
A second possibility is that faulting occurred in the absence of a
well-defined sequence of mainshocks and aftershocks. As in the case of
man-induced earthquakes triggered by high-pressure fluids during injection
of fluids (e.g. Healy et al., 1968), whereby deformation is typically
accommodated by diffuse swarms of low-magnitude seismicity rather than
well-defined mainshock–aftershock sequences (Cox, 2016), we propose that
BFZ300 might have localized strain by diffuse veining with crack-and-seal
textures (Cox, 2016). Breccias and cataclasites (Figs. 3, 8) mutually
overprinting with veins show that failure and veining were indeed broadly
coeval (e.g. Cox, 1995, 2016). Healing in fluid-rich environments can
occur over short periods of time (days to months) when compared with the recurrence
time of large earthquakes (10–100 years) (Olsen et al., 1998; Tenthorey and
Cox, 2006). Therefore, the documented repeated switches between brittle and
ductile deformations would then be steered by transient episodes of
fluid overpressuring but in this case would express the accommodation of
swarms of minor background earthquakes within overall ductile conditions.
Microstructures of fault rocks exhumed from the brittle–ductile transition
in other geological settings are mostly in agreement with our hypotheses of
seismic deformation. Transient and short-term high-stress deformation
followed by phases of stress relaxation, which is prevalently characterized
by recovery and recrystallization processes, has been documented by several
authors in deformed quartz (Trepmann and Stöckhert, 2003, 2013; Trepmann et
al., 2007, 2017; Bestmann et al., 2012).
Conclusions
This study shows that a multi-scale and multi-technique approach leading to
the generation of independent constraints offers the potential to
reconstruct in detail the evolutionary history of fault zones that have
experienced multiple events of reactivation triggered by fluid overpressure
and in which intense fluid–rock re-equilibration processes have taken place.
We document the localized, initial embrittlement of the Paleoproterozoic
basement of southwestern Finland at the BDTZ, which occurred by brittle
failure under overall ductile conditions in response to transiently high
fluid pressure and temperature (peak conditions: Pf > 210 MPa; T∼350∘C). Our results
further constrain the importance of cyclic seismicity and fluids in the
fragmentation of Precambrian cratons when deformed at the BDTZ, something
that is not yet that well understood for the Fennoscandian Shield. Our
study, moreover, provides potentially important inputs to many modern
geological applications, including site characterization of deep geological
disposal facilities for spent nuclear fuel. Results from the detailed
geological characterization of faults at the Olkiluoto site can thus be used
for continuously updating the geological site description and yield
further constraints on the mechanics of faulting at the BDTZ.
Data availability
All the data produced and used to write the paper are contained in it and in the corresponding Supplement.
The supplement related to this article is available online at: https://doi.org/10.5194/se-10-809-2019-supplement.
Author contributions
GV and JM conceptualized and designed the project and performed the fieldwork. BM and PSG performed the sulfides analysis and geothermometry. BM performed the petrography, microthermometry, and Raman analysis of fluid inclusions, chlorite compositions, and geothermometry. BM and LM selected microstructures to analyse with EBSD and LM acquired EBSD maps. BM, PSG, and GV wrote the paper with contributions by JM and LM.
Competing interests
The authors declare that they have no conflict of
interest.
Acknowledgements
We thank Oliver Vanderhaeghe and two anonymous reviewers for their constructive reviews, which led to a greatly improved paper.
Stephen F. Cox, Michael Stipp, and Alfons M. Van den Kerkhof are all warmly thanked for fruitful discussions during the early stages of
this work. Danilo Bersani and Andrea Risplendente are also thanked for their help with the Raman and SEM analyses. Financial support from Posiva Oy (no. 2105178) is acknowledged.
Review statement
This paper was edited by Bernhard Grasemann and reviewed by Olivier Vanderhaeghe and two anonymous referees.
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