The mechanical and thermal properties of the subduction interface controlthe internal structure, kinematics,
and dynamics of a subduction zone (e.g., Agard etal.,2018; Cloos,1982; Gerya & Stöckhert,2002; Behr &
Becker,2018). Along the shallow interface (≤20km), direct observations of the megathrust and accretionary
wedge are possible through high-resolution seismic reflection imaging, ocean bottom seismometers, and ocean
drilling projects (e.g., Fagereng etal.,2019; H. Kimura etal.,2010; Park etal.,2002). However, seismic tomogra-
phy and earthquake seismology have limited spatial and temporal resolution (e.g., Calvert etal.,2011; Rondenay
etal.,2008) so the geometry and internal structure of the deep interface (∼20–70+ km) remain poorly understood.
The deep interface can be studied through geologic observations of exhumed high-pressure/low-temperature (HP/
LT) metamorphic rocks. Some of the most spectacular examples – for example, the Franciscan Complex (e.g.,
Cloos,1986; Wakabayashi,1990), Japan (Aoki etal.,2008; G. Kimura etal.,2012), and the Mediterranean region
(e.g., Brun & Faccenna,2008; Jolivet etal.,2003) – have profoundly shaped our understanding of subduction and
exhumation processes. Specifically, field studies provide constraints on the structural and kinematic evolution,
interface geometry, metamorphic pressure-temperature (P-T) trajectories and thermal structure, and timing and
rates of subduction and exhumation (e.g., Agard etal.,2018; Angiboust etal.,2016; Behr & Platt,2012; Dragovic
etal.,2015; Kotowski etal.,2021; Tewksbury-Christle etal.,2021). Geologic observations can validate or chal-
lenge the results of geodynamic simulations that model the kinematics and dynamics of plate boundary shear
zones (e.g., Cloos,1982; Gerya etal.,2002; Gerya & Stöckhert,2002; Penniston-Dorland etal.,2015; Warren
Abstract Exhumed high-pressure/low-temperature (HP/LT) metamorphic rocks provide insights into deep
(∼20–70km) subduction interface dynamics. On Syros Island (Cyclades, Greece), the Cycladic Blueschist Unit
preserves blueschist-to-eclogite facies oceanic- and continental-affinity rocks that record the structural and
thermal evolution linked to Eocene subduction. Despite decades of research, the metamorphic and deformation
history (P-T-D) and timing of subduction and exhumation are matters of ongoing discussion. We suggest that
Syros comprises three coherent tectonic slices and that each slice underwent subduction, underplating, and
syn-subduction return flow along similar P-T trajectories, but at progressively younger times. Subduction and
exhumation are distinguished by lineations and ductile fold axis orientations, and are kinematically consistent
with previous studies that document top-to-the-S-SW shear (prograde-to-peak subduction), top-to-the-NE shear
(blueschist facies exhumation), and then E-W coaxial stretching (greenschist facies exhumation). Amphibole
zonations record cooling during decompression, indicating return flow above a cold slab. Multi-mineral Rb-Sr
isochrons and compiled metamorphic geochronology show that the three slices record distinct stages of peak
subduction (53–52, ∼50, and 45Ma) that young with structural depth. Retrograde blueschist and greenschist
facies fabrics span ∼50–40 and ∼43–20Ma, respectively, and also young with structural depth. Synthesized
data sets support a revised tectonic framework for Syros, involving subduction of structurally distinct coherent
slices and simultaneous return flow of previously accreted tectonic slices in the subduction channel shear zone.
Distributed, ductile, dominantly coaxial return flow in an Eocene-Oligocene subduction channel proceeded at
rates of ∼1.5–5mm/yr and accommodated ∼80% of the total exhumation of this HP/LT complex.
KOTOWSKI ET AL.
© Wiley Periodicals LLC. The Authors.
This is an open access article under
the terms of the Creative Commons
Attribution License, which permits use,
distribution and reproduction in any
medium, provided the original work is
Subduction, Underplating, and Return Flow Recorded in the
Cycladic Blueschist Unit Exposed on Syros, Greece
Alissa J. Kotowski1,2,3 , Miguel Cisneros1,4,5 , Whitney M. Behr1,4 , Daniel F. Stockli1 ,
Konstantinos Soukis6 , Jaime D. Barnes1 , and Daniel Ortega-Arroyo1,7
1Department of Geological Sciences, Jackson School of Geosciences, University of Texas at Austin, Austin, TX, USA,
2Department of Earth and Planetary Sciences, McGill University, Montreal, QC, Canada, 3Now at Department of Earth
Sciences, Utrecht University, Utrecht, The Netherlands, 4Department of Earth Sciences, Geological Institute, Swiss Federal
Institute of Technology (ETH), Zurich, Switzerland, 5Now at Lawrence Livermore National Laboratory, Livermore, CA,
USA, 6Geology and Geoenvironment, National and Kapodistrian University of Athens, Athens, Greece, 7Now at Department
of Earth, Atmospheric & Planetary Sciences, Massachusetts Institute of Technology, Cambridge, MA, USA
• Syros is a tectonic stack composed of
three slices constructed by subduction
and underplating; peak subduction
ages young with structural depth
• The subduction-to-exhumation
transition is marked by kinematic
rotation and cooling during
• Metamorphic geochronology indicates
syn-subduction exhumation occurred
continuously in an Eocene-Oligocene
Supporting Information may be found in
the online version of this article.
A. J. Kotowski,
Kotowski, A. J., Cisneros, M., Behr, W.
M., Stockli, D. F., Soukis, K., Barnes,
J. D., & Ortega-Arroyo, D. (2022).
Subduction, underplating, and return flow
recorded in the Cycladic Blueschist Unit
exposed on Syros, Greece. Tectonics,
41, e2020TC006528. https://doi.
Received 12 SEP 2020
Accepted 15 MAY 2022
1 of 41
KOTOWSKI ET AL.
2 of 41
Syros Island, located in the central Aegean Sea (Figure1), is an ideal locality to study deep subduction interface
processes due to its exceptional preservation and exposure of HP/LT blueschist-to-eclogite facies assemblages
(Dürr etal.,1978; Okrusch & Bröcker,1990; Ridley,1982,1984). Despite decades of research on Syros, numer-
ous disagreements persist regarding the structural evolution, metamorphic conditions, and timing and mecha-
nisms of subduction and exhumation on the island (e.g., Aravadinou & Xypolias, 2017; Bröcker etal., 2013;
Keiter etal., 2004; Laurent et al., 2018, 2016; Lister & Forster, 2016; Ridley,1982; Ring & Layer,2003;
Rosenbaum etal.,2002; Schumacher etal.,2008; Skelton etal.,2019; Soukis & Stockli,2013; Trotet, Jolivet,
& Vidal, 2001). Furthermore, crustal-scale extensional detachments that accommodated the latest stages of
post-orogenic exhumation are well-documented across the Cyclades (Avigad & Garfunkel,1989,1991; Gautier
etal.,1993; Grasemann etal.,2012; Jolivet, L., Brun, & J. P,2010; Jolivet & Brun,2010; Schneider etal.,2018;
Soukis & Stockli,2013), but workers continue to debate the relative importance of major detachments during
syn-orogenic exhumation from peak conditions, and whether the strain was distributed or highly localized on
Syros (Bond etal.,2007; Keiter etal.,2004; Laurent etal.,2016; Lister & Forster,2016; Rosenbaum etal.,2002).
In this work, we present new structural and petrologic data and Rb-Sr geochronology, and integrate our results
with a synthesis of previously published geochronology, to propose a new model for the evolution of the CBU on
Syros. Our results refine the island's deformation-metamorphism history, and shed light on the kinematics, meta-
morphic conditions, and timing of subduction and return flow in the Hellenic subduction zone. This work has
implications for rates and mechanisms of HP/LT rock exhumation and provides a broader framework for regional
construction of the Attic-Cycladic Complex.
Figure 1. Regional tectonic map of the Cyclades, modified from Grasemann etal.(2012). Syros is outlined by the yellow
box. North Cycladic (NCDS), West Cycladic (WCDS), Paros-Naxos (PNDS), and Santorini (SDS) Detachment Systems are
outlined in white. Kinematic directions are from Aravadinou etal.(2016), Forster etal.(2020), Grasemann etal.(2012), Huet
etal.(2009), and references therein.
KOTOWSKI ET AL.
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2. Regional Geologic Setting
The Cycladic Islands and parts of mainland Greece comprise the Attic-Cycladic Complex (ACC), which is
divided into three units according to depositional age and metamorphic history. From structural top to bottom,
the units are: (a) the Upper Cycladic Unit, (b) the Cycladic Blueschist Unit, and (c) the Basal Unit (e.g., Altherr
etal.,1994; Avigad & Garfunkel,1989; Dürr etal.,1978; Jacobshagen,1986; van der Maar & Jansen,1983)
(Figure1). The Upper Cycladic Unit is a suite of ophiolitic slivers, altered carbonates ± serpentinites, Late
Cretaceous (70–100Ma) amphibolite-facies orthogneisses, and Miocene greenschist-facies meta-basalts, and
correlates with the Pelagonian realm exposed on mainland Greece (Papanikolaou,1987). The Upper Unit was
the upper plate during Late Cretaceous-Paleogene subduction and crops out above the Cycladic Blueschist Unit
(CBU) in the hanging wall of crustal-scale, Miocene detachment faults on several Cycladic Islands (Jolivet
etal.,2010,2013; Soukis & Stockli,2013).
The majority of the ACC is composed of the Cycladic Blueschist Unit (CBU) (Figure1). The CBU comprises
poly-metamorphosed tectonic slices (Dürr etal.,1978; Forster & Lister,2005,2008; Jolivet & Brun, 2010)
of the following protoliths: (a) (Jurassic?-to-) Cretaceous (∼80Ma) mafic igneous crust with enriched-MORB
and back-arc geochemical signatures±serpentinized mantle (Bonneau, 1984; Bulle etal.,2010; Cooperdock
etal.,2018; Fu etal.,2015; Seck etal.,1996; Tomaschek etal.,2003), (b) Triassic (∼240Ma) bimodal rift volcan-
ics (Bolhar etal., 2017; Bröcker & Keasling,2006; Bröcker & Pidgeon,2007; Keay, 1998; Robertson,2007)
blanketed by Triassic-to-Cretaceous, locally-sourced (e.g., from Triassic volcanics), rifted and passive conti-
nental margin siliciclastic and carbonate rocks (Löwen etal.,2015; Papanikolaou,2013; Poulaki etal.,2019;
Seman,2016; Seman etal.,2017), and (c) peri-Gondwanan basement cross-cut by Carboniferous calc-alkaline
granitoids (Flansburg etal.,2019; Keay,1998; Keay & Lister,2002).
Regionally, CBU lithologies record evidence for HP/LT metamorphism under blueschist-to-eclogite facies (‘M1’)
conditions between ∼53–40Ma (Cliff etal.,2016; Dixon,1976; Gorce etal.,2021; Lagos etal.,2007; Laurent
etal.,2017; Okrusch & Bröcker,1990; Ring, Glodny, etal.,2007; Schliestedt,1986; Tomaschek etal.,2003;
Wijbrans etal.,1990). The CBU was exhumed first within the subduction channel, leading to blueschist and
greenschist facies overprinting (e.g., Cliff et al., 2016; Kotowski & Behr,2019; Laurent etal.,2018; Ring
etal.,2020), and then in the footwalls of crustal-scale, low-angle normal faults of the North, West, and South
Cycladic (Grasemann etal.,2012; Jolivet etal.,2003,2010; Jolivet & Brun,2010; Ring & Layer,2003,2011;
Roche etal.,2016; Soukis & Stockli,2013), the Paros-Naxos (Bargnesi etal.,2013; Gautier etal.,1993; Linnros
etal.,2019), and the Santorini Detachment Systems (Schneider etal.,2018). Exhumation beneath ductile and
semi-brittle detachments led to the development of Metamorphic Core Complexes (MCCs) that locally also
produced a greenschist-facies (‘M2’) overprint (Bröcker,1990; Bröcker etal.,1993). As slab rollback initiated
and the arc migrated southward through the former forearc, Miocene I-type and S-type plutons intruded the
exhuming CBU, and MCC formation led to a local high-temperature, amphibolite-facies (‘M3’) overprint on
some islands (e.g., Paros and Naxos, Mykonos, and Ikaria) between ∼21–17Ma (Andriessen etal.,1979; Brichau
etal., 2007; Lister etal.,1984; Pe-Piper etal.,2002; Rabillard et al.,2018; Vanderhaeghe & Whitney,2004;
Wijbrans & McDougall,1988).
3. The CBU on Syros Island
3.1. Rock Types and Tectonostratigraphy
Syros is a small island (∼84km
2) in the central Cyclades and is dominantly composed of CBU with a klippe of
UU in the southeast in the hanging wall of the Oligo-Miocene Vari Detachment (Keiter etal.,2011; Ridley,1984;
Ring etal.,2003; Soukis & Stockli,2013) (Figure1). In the context of the Cyclades, Syros preserves some of the
most pristine HP/LT metamorphic rocks, in some places even recording peak assemblages with little to no retro-
gression (Kotowski & Behr,2019; Okrusch & Bröcker,1990; Ridley,1982); similar assemblages are preserved
on the island of Sifnos (Aravadinou etal.,2016; Roche etal.,2016).
Within the CBU on Syros, mafic blueschists and eclogites crop out along three tectonostratigraphic hori-
zons: Kampos Belt, Kini-Vaporia-Kalamisia, and Galissas-Fabrikas (Figures 2 and 3). Each horizon exposes
∼300–500 m (structural thickness) of blueschist-to-eclogite facies meta-basalts and gabbros, serpentinites,
and interlayered, foliated felsic gneiss/schist and glaucophane schist sequences (metamorphosed bimodal
KOTOWSKI ET AL.
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volcanics) in varying proportions (Bröcker & Keasling,2006; Dixon & Ridley,1987; Keiter etal.,2011). Along
Kampos Belt (i.e., Kampos mélange), eclogitic meta-gabbros, blueschist facies bimodal meta-volcanics, and
serpentinite/chlorite-talc schists are most abundant. Meta-gabbro pods (varying in size from ∼1 to at least
Figure 2. Geologic and structural map of Syros Island, modified from Keiter etal.(2004,2011). Structural elements and
locations of the Syringas Marker Horizon are from Keiter etal.(2011). Constraints on protolith ages are from the references
discussed in Section3.1. Protolith ages are color coded according to rock type. Localities discussed in this study are shown in
bold italics, new Rb-Sr sample names and locations are in italics. The thick dashed black line marks a hypothesized location
of the lower nappe-bounding shear zone; its location is primarily constrained by detrital zircon maximal depositional ages and
KOTOWSKI ET AL.
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3) are commonly veiled by metasomatic reaction rinds that developed
at contacts with thin serpentinite carapaces, but clastic meta-sediments and
bimodal meta-volcanics comprise the volumetric majority of the surrounding
matrix (Dixon & Ridley,1987; Keiter etal.,2011; Ridley,1982) (Figure2).
Kini, Vaporia (north of Ermoupoli), and Kalamisia are primarily composed
of fine-grained mafic blueschist, and contain pods and lenses of eclogite
(centimeters-to-decimeters in diameter) and meters-thick layers of serpen-
tinite/talc schist (Keiter et al., 2011; Kotowski & Behr, 2019). Fabrikas
comprises coarse-grained glaucophane-bearing eclogite pods (centimeters to
meters in diameter) within a fine-grained matrix of mafic blueschists and
quartz-mica schists, capped by meta-carbonate (Kotowski & Behr, 2019;
Ring etal.,2020; Skelton etal.,2019). Keiter etal.(2011) suggested that
mafic blueschists and eclogites are genetically related, and changes in volume
proportions of lithologies reflect primary lateral and/or vertical ‘facies
changes’ of an enriched-MORB or back-arc igneous suite. Throughout the
structural section on Syros, the CBU is internally coherent but commonly
contains outcrop-scale block-and-matrix structures reflecting competency
contrasts between different protoliths (e.g., basalts vs. gabbros) and meta-
morphic rock types (e.g., blueschists vs. eclogites) (Keiter et al., 2011;
Kotowski & Behr,2019).
The majority of the CBU comprises a ∼6–8 km section of interca-
lated meta-volcanic and meta-sedimentary schists, and calcite- and
dolomite-marbles with Jurassic-to-Cretaceous depositional ages (Keiter
etal.,2004; Löwen etal.,2015; Papanikolaou,2013; Seman et al., 2017)
(Figures2 and3). Keiter etal.(2004,2011) documented a series of boudi-
naged marbles, cherts, and albite-bearing quartzite, which they named the
Syringas Marker Horizon and interpreted as primary sedimentary layer-
ing(orange dots shown in Figure2). The sequence crops out at three or four
structural levels suggesting it marks several km-scale thrust sheets (Dixon
& Ridley,1987; Keiter etal.,2011; Ridley,1982). Repetition of the Syrin-
gas Marker Horizon by km-scale folding is unlikely because the largest observable upright folds within this
sequence have amplitudes of several hundreds of meters and the marker horizon never appears to be overturned
(Keiter etal.,2011). Furthermore, Keiter etal.(2011) documented the repetition of distinct packages of bimodal,
rift-related meta-volcanics (also mapped as “banded tuffitic schists”) that have Triassic magmatic protolith ages
(Bröcker & Keasling,2006; Keay,1998; Pe-Piper etal.,2002; Seman, 2016) (Figure 2), and Seman (2016)
presented detrital zircon (DZ) Maximum Depositional Ages (MDAs) for meta-sedimentary rocks that may point
to old-on-young tectonostratigraphic inversions. Both results appear to support imbrication (cf. Figure3).
3.2. Previously Proposed P-T-D-t Paths
Previously published P-T-D evolutions for Syros fall into two categories. Some workers have argued that the
majority of deformation and metamorphism on the island is exhumation-related, following peak pressure condi-
tions of ∼20–24 kbar(Laurent etal.,2016; Lister & Forster,2016; Trotet, Jolivet, & Vidal,2001) (Figure4a).
These studies interpret mafic blueschists and eclogites to occupy the top of the structural pile and separate
them from underlying meta-sedimentary rocks along extensional shear zones (Forster & Lister,2005; Laurent
etal.,2016,2018; Trotet, Vidal, & Jolivet,2001). This model implies that lithologically distinct rock pack-
ages were juxtaposed during syn-orogenic exhumation (Forster & Lister,2005; Laurent etal.,2016). Unaltered
and retrogressed eclogite has been documented throughout the structural section on Syros, which is considered
evidence that all rocks experienced high-pressure conditions during subduction. However, lithologic packages
that currently occupy different structural depths could have followed different P-T paths during exhumation (cf.
Laurent etal.,2018; Trotet, Jolivet, & Vidal,2001; Trotet, Vidal, & Jolivet, 2001), and/or could have been
subducted and accreted/underplated at different times (Laurent etal.,2017; Lister & Forster,2016). Such a
model could explain reported differences in P-T estimates across Syros; mafic blueschists and eclogites may have
Figure 3. Condensed lithostratigraphical column modified from Keiter
etal.(2011). Our structural, petrologic, and geochronologic synthesis supports
dividing the CBU into three sub-units comprising different proportions and
distributions of lithologies. Ages of peak subduction metamorphism young
with structural depth.
KOTOWSKI ET AL.
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Figure 4. (a) Compilation of proposed P-T-D histories for the Cycladic Blueschist Unit on Syros. (b) Closure temperature versus time for compiled metamorphic
geochronology listed in TableS2 in Supporting InformationS1. This dataset comprises 127 datapoints made up of 208 individual ages (some data points are weighted
means), from 18 studies and 5 chronometers, from work published between 1987-2022. The black arrows at the bottom labeled ‘structural development’ shows that the
timing and tectonic significance of progressive Eo-Oligocene deformation events and corresponding fabric development is contentious.
KOTOWSKI ET AL.
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been subducted slightly deeper, earlier, compared to meta-sedimentary lithologies (as discussed by Schumacher
Alternatively, other authors have suggested that prograde deformation and metamorphism reached ∼16–18
kbar and is locally preserved, but the exhumation-related strain was partitioned into weaker lithologies (Bond
etal.,2007; Cisneros etal.,2021; Keiter etal.,2004,2011; Ridley,1982; Rosenbaum etal.,2002) (Figure4a).
These studies interpret mafic blueschist and eclogites to record primary relationships with surrounding schists
and marbles or to have been juxtaposed with the schists and marbles during early underthrusting (Blake Jr
etal.,1981; Hecht,1985; Keiter etal.,2004; Ridley,1982). For either of those cases, map-scale lenses of mafic
blueschists and eclogites at Vaporia, Kalamisia, and Fabrikas (Figure2) need not be separated from surround-
ing CBU by faults or shear zones (i.e., the structurally highest Kampos sub-unit of Laurent etal.(2016)), but
instead could occupy a range of structural depths throughout the tectonostratigraphic pile (Keiter etal.,
Figure3). This model implies that meta-mafic and meta-sedimentary rocks that occupy similar structural levels
were subducted together and experienced similar P-T histories during subduction and subsequent exhumation
(Cisneros etal.,2021; Keiter etal.,2011; Schumacher etal.,2008).
Although existing metamorphic ages help to roughly distinguish prograde from retrograde fabrics and the timing
of subduction versus exhumation, differentiating between these P-T-D models has been challenging because of
the difficulty in assigning geologic significance to ages (Figure4b). Two age clusters are most commonly cited
for the timing of peak subduction on Syros: ∼53–50Ma (U-Pb zircon, Ar/Ar and Rb-Sr white mica, Lu-Hf garnet;
Cliff etal.; Lagos etal.; Lister & Forster; Tomaschek etal.), and both ∼52Ma and
∼45Ma for different underplated slices (Ar/Ar white mica; Forster & Lister; Glodny & Ring;
Laurent etal.; Lister & Forster). Recently, Uunk etal.(2022) suggested that Syros is composed of
three lithologically distinct sub-units that reached similar peak P-T conditions, but were underplated at progres-
sively younger times, as recorded by garnet-whole rock Lu-Hf isochrons (cf. Figure3). However, the timing of
retrogression recorded by Ar/Ar and Rb-Sr ages span the entire Eocene. Maximum CBU temperatures do not
appear to have exceeded those required for diffusional resetting of the Ar/Ar and Rb-Sr systems, therefore it is
unclear whether retrograde blueschist-to-greenschist facies white mica ages record incomplete isotopic mixing,
and/or partial or continuous recrystallization during exhumation, beneath the isotopic closure temperature of the
Ar/Ar and Rb-Sr systems (Figure4b) (e.g., Bröcker etal.,2013; Cliff etal.,2016; Laurent etal.,2017; Rogowitz
etal.,2014; Uunk etal.,2018). An additional challenge is that many geochronologic data points in Figure4b were
collected without a clear framework for linking the ages to specific deformation fabrics.
4. Approach and Methodology
4.1. Structural and Microstructural Analysis
Following detailed mapping by Keiter etal.(2004) and Keiter etal.(2011) (map in Figures2,5–7), we collected
new structural data at several localities from Northern Syros (Figure5), Central Syros (Figure6), and Southeast-
ern Syros (Figure7). We measured planar and linear structural elements, including foliations and cleavages, axial
planes to folds, fold axes, and mineral growth, crenulation, and stretching lineations. We constructed π circle
diagrams to constrain fold orientations by plotting poles to metamorphic foliation planes. Each color on a given
stereonet in Figures5–7 corresponds to poles to foliations of a specific rock type, or poles to foliations defining
single outcrop-scale folds (e.g., Figures8i and 8k). Our measurements used to produce π diagrams were all
derived from cylindrical structures; even if folds have curved hinge lines on larger scales, we measured folds in
locations where hinge lines are locally straight. We calculated poles to mean circles to determine fold axis orien-
tations (bold circles) and compared them to fold axes that could be directly measured (diamonds) and mineral
lineations (open circles). We documented minerals defining lineations and fold axes, porphyroblast stability
and kinematic context (i.e., pre-, syn, post-kinematic with respect to surrounding fabric), and break-down and
replacement textures (Figures8–12) to constrain metamorphic conditions of deformation. Fold axis orientations
and mineral lineations do not provide a sense of shear, but rather stretching and/or transport directions during
shearing. However, we also documented many outcrop-to micro-scale shear sense indicators and supplemented
our kinematic interpretations with literature constraints. Microstructural analysis (Figure9) (139 total samples,
21 studied in detail) and quantitative EMPA analyses of zoned minerals (6 samples) refined our interpreted P-T-D
history (Figures10 and12).
KOTOWSKI ET AL.
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4.2. Rb-Sr Geochronology
We selected five samples for multi-mineral Rb-Sr geochronology. This technique has been applied to exhumed
HP/LT metamorphic rocks to date deformation and metamorphism with considerable success (Angiboust
etal.,2016; Cliff etal.,2016; Freeman etal.,1997; Glodny etal.,2008,2005; Kirchner etal.,2016; Ring, Will,
etal.,2007). The primary assumption required to construct a multi-mineral isochron is that the phases defining
the isochron were co-genetic, and thus share the same initial Sr composition. We separated and selected miner-
als that we hypothesized were co-genetic based on our structural and microstructural results, and quantitatively
tested this hypothesis by identifying phases that were in isotopic disequilibrium (i.e., fall off the isochron) (Cliff
Figure 5. Geology and structural elements of Northern Syros. Base map and foliation orientations are from Keiter etal.(2011). Black fold axes are Keiter
etal.(2011)'s intrafolial ‘F2 shear folds.’ The spread of orientations is the result of superposed folding, as older folds were progressively reoriented by S-vergent simple
shear during subduction (cf. Keiter etal.,2004). Average lineation orientations are shown in white arrows outlined in blue and greenfor glaucophane and actinolite
measurements, respectively; all measured lineations are plotted on the stereonets. Colored planes are best-fit π circles fitting poles to foliation planes defining single
outcrop-scale folds. Cross section A1-A2-A3 is modified from Keiter etal.(2011). Topographic contours are 20m.
KOTOWSKI ET AL.
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& Meffan-Main,2003). Penetrative foliations lend support to the assumption of syn-kinematic recrystallization
of selected minerals, which can reset the Sr isotopic signature between mica and co-genetic phases to tempera-
tures as low as 300°C (Müller etal.,2000). Furthermore, diffusional resetting of the Rb-Sr system is thought to
occur at temperatures >550–600°C (Glodny etal.,2008; Inger & Cliff,1994), which exceeds maximum tempera-
tures in the CBU. Therefore, we interpret our Rb-Sr ages as (re-)crystallization ages associated with deformation.
Following Glodny etal.(2003,2008), we cut out ∼5cm
3 cubes of rock from hand samples to isolate specific
fabrics corresponding to different stages of the deformation history. Samples were crushed with a small hammer
Figure 6. Geology and structural elements of Central Syros. Average lineation orientations are shown in white arrows outlined in blue and greenfor glaucophane and
actinolite measurements, respectively; all measured lineations are plotted on the stereonets. Topographic contours are 20m. See Figure5 caption for further details.
KOTOWSKI ET AL.
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Figure 7. Geology and structural elements of Southeast Syros. Black arrows with the circles are upright fold axes in the Vari Unit. Average lineation orientations are
shown in white arrows outlined in blue and greenfor glaucophane and actinolite measurements, respectively; all measured lineations are plotted on the stereonets. Cross
section A6-B2-B3 is modified from Keiter etal.(2011); compare with Laurent etal.(2016). Topographic contours are 20m. See Figure5 caption for further details.
KOTOWSKI ET AL.
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between sheets of paper and sieved and separated by grain size. Grain size fractions 125–250μm and 250–500μm
were separated based on magnetic susceptibility using a Frantz magnetic separator. Mineral separates were picked
by hand under a microscope, and white mica separates were cleaned of inclusions by gently smearing them in a
mortar and pestle and washing them through a sieve with ethanol. All Rb and Sr isotopic separation and analyses
were conducted at the University of Texas at Austin in the Radiogenic Isotopic Clean Lab. All separates (except
apatite) were cleaned in 2N HCl to remove surficial contamination and spiked with mixed high Rb/Sr and low
Rb/Sr spikes. We followed methodology for mineral dissolution, isotope column chemistry, Thermal Ionization
Mass Spectrometry (Sr analyses), Solution Inductively Coupled Plasma Mass Spectrometry (Rb analyses), and
estimating uncertainties in isotopic ratios as described in Kirchner etal.(2016). Reproducibility on replicate
USGS Standard Hawaiian Basalt (BHVO) Rb measurements determines the uncertainty of the Rb-Sr ratio, and
long-term reproducibility on the NBS987 Sr standard determines the uncertainty of the Sr ratio (Table2). Ages
were calculated using the IsoplotR toolbox (Vermeesch,2018) with the
87Rb decay constant of 1.3972±0.0045
−1 (Villa etal.,2015).
Figure 8. Selected field photos showing prograde (a)–(d) and retrograde (e–l) deformation and metamorphism. (a) Preservation of primary igneous breccias at Grizzas.
(b) Right-side-up sequence of oceanic lithosphere at Kini. (c), (d) SS at Kini contains lawsonite pseudomorphs and omphacite with glaucophane- and garnet-filled
pressure shadows. Black arrows in the close-up photo of (c) point to pseudomorphs with garnet inclusions. (e) DT1 retrogression under blueschist-facies conditions
produced local static glaucophane coronas formed around pinched eclogite lenses at Vaporia. (f) Coaxial E-W stretching of calcite clasts in meta-conglomerate at
Delfini during DT1−2. (g), (h) SS is cut by ST1 crenulation cleavage at Azolimnos. (h) Two glaucophane lineations record transposition of SS (black arrow, parallel to the
pen) into alignment with crenulation hinges (white arrow) during DT1. (i)–(k) DT2 greenschist facies retrogression and upright folding at Delfini (i), Fabrikas (j), and
Lotos (k). (i) White arrows point to FS folds along the limbs of FT fold. Dashed white lines mark the axial planar ST cleavage. (j) Non-coaxial, top-to-the-E extensional
shear under retrograde blueschist-to-greenschist facies conditions. (k) SS cross-cut by DT folding; fold axes trend E-W. (L) Coaxial, lineation-parallel DT2 brittle
boudinage of epidote-rich lenses in greenschists.
KOTOWSKI ET AL.
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5. Structures and Deformation Fabrics
The CBU on Syros records evidence for three main phases of deformation and metamorphism herein referred
to as DR, DS, and DT1−2 (Table1). Subscripts follow an alphabetical order according to the relative age of defor-
mation, that is, DR is the oldest observed deformation, and DT1−2 is the youngest. Each phase led to spaced to
penetrative foliation development, and/or ductile folding of older fabrics. Kinematic indicators, metamorphic
mineral assemblages, and porphyroblast zonations described herein demonstrate that DR and DS developed on
the prograde path and are best preserved in mafic blueschists and eclogites (but are locally preserved as textural
relicts in bimodal meta-volcanics and meta-sediments), and DT developed on the retrograde path and is best
recorded by meta-volcanic and meta-sedimentary schists.
5.1. DR — Prograde Fabric Development During Subduction Under Blueschist Facies Conditions
DR is the earliest recognizable prograde event but it is not visible at the outcrop-scale. DR likely formed a strong,
penetrative SR foliation that is locally recorded as inclusion trails in garnet porphyroblasts at Kampos (Figures9a
and9b) and is tightly folded during DS. DR inclusion trails are commonly oblique to the external foliation and are
defined by glaucophane, omphacite, and white mica. However, some garnets contain an internal foliation that is
Figure 9. Selected photomicrographs showing prograde (a)–(f) and retrograde (e, g–k) deformation and metamorphism. (a),(b) Internal SR inclusion trails from Lia
Beach (A, PPL; B, XPL). (c) SS contains syn-kinematic garnet porphyroblasts with foamy quartz inclusion trails that are rotated but continuous with respect to the
dominant external foliation. (d) DS garnets include pseudomorphs after lawsonite (comprising epidote and white mica). (e), (f) SS in mafic blueschists from Kini. (e) SS
is cut by DT1 crenulation under glaucophane-stable conditions in mafic blueschists from Lia Beach. (f) Omphacite and garnet in DS Kini blueschists have asymmetric
pressure shadows filled with high-pressure minerals. (g)–(i) DT1 retrogression in bimodal meta-volcanics at Kampos (g), Azolimnos (h) and Kalamisia (i). (h) DT2
crenulation transposes SS and strengthens as albite, chlorite, and actinolite form. (i) Omphacite and paragonite break down to epidote, blue amphibole, and albite. (j) DT2
brittle micro-boudinage of epidote porphyroblasts in Lotos greenschists. (k) Final stages of DT2 are characterized by post-tectonic albite growth at Lotos.
KOTOWSKI ET AL.
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Figure 10. False-colored X-Ray maps and representative BSE images of DS in Kini blueschists (a),(b), a composite DS-DT1
fabric in Azolimnos bimodal meta-volcanics (c),(d), DT1 in Kalamisia blueschists (e),(f), and DT2 in Fabrikas quartz-mica
schists (g),(h). Quantitative analyses of sodic amphiboles in (b) KCS53 and (f) KCS12B are shown in Figure12; white mica
analyses from (h) KCS65 are shown in Figure11 and FigureS2in Supporting InformationS1.
KOTOWSKI ET AL.
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continuous with the external foliation, and others preserve no internal foliation. This suggests that garnet growth
occurred prior to, during, and after DR (during DS) in different lithologies.
5.2. DS — Prograde-to-Peak Fabric Development During Subduction Under Blueschist to Eclogite Facies
5.2.1. DS Structures
DS is best recorded at Grizzas and Kini (Figure6b), with relicts preserved on Kampos Belt (Figure5c), at Lia
Beach, and at Azolimnos (Figure7b). DS produced a dominant SS foliation in mafic blueschists, meta-cherts,
and bimodal meta-volcanics at Grizzas that is parallel to the axial planes of intrafolial folds (FS), and rotated and
boudinaged quartz veins. This folding event is characterized by shallowly to moderately plunging SW-trending
fold axes clustering around 205–251°/15–35°; glaucophane mineral lineations are similarly oriented (Figure5b).
In rare cases, outcrop-scale prograde metamorphism was not associated with penetrative deformation, indicated
by the preservation of igneous protolith features such as pillow lavas (Grizzas, cf. Keiter etal.), intrusive
relationships (Kini, cf. Kotowski and Behr and Laurent etal.), and magmatic breccias (e.g., at
Grizzas, Episkopi, Figure8a).
Kini dominantly records DS deformation-metamorphism; it is bounded by high-angle normal faults and is structur-
ally discordant with respect to the surrounding CBU (Figure6b; cf. Keiter etal.). In one location, serpen-
tinite wraps around the base of massive meta-gabbros, which transitions upward into fine-grained blueschists,
suggesting local preservation of an attenuated section of metamorphosed oceanic lithosphere (Figure8b). Similar
Figure 11. White mica chemistry and micro-textures. Samples from Azolimnos, Lotos, Kini, and Delfini were targeted
for Rb-Sr geochronology. (a) Quantitative phengite and paragonite EMPA analyses. (b) Composite DS-DT1 texture in
Azolimnos retrograde blueschist, characterized by intergrown phengite and paragonite. (c) DT2 texture in a greenschist
facies meta-volcanic from Lotos. Well-aligned white mica defines the foliation, containing relict amphibole porphryoblasts
breaking down to quartz, albite, and mica. Phengite is enriched in Fe near reaction zones, but exhibits only slight reduction
in Si a.p.f.u. relative to darker cores. (d) DT2 texture in Delfini meta-volcanics. Phengite-paragonite pairs are in textural
equilibrium. Phengite locally occurs as lineation-parallel overgrowths on paragonite cores.
KOTOWSKI ET AL.
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to Grizzas, the DS fabric in Kini blueschists contains isoclinal folds (FS) with shallowly south-plunging fold axes.
This fold generation is recorded by a 182°/33° fold axis in Kini schists (Figure6b; Figure8d). The SS axial planar
cleavage seen in Kini mafic blueschists (e.g., Figures8c and8d) is also seen as textural relicts in quartz-mica rich
lithologies, as at Azolimnos (Figure8g). In some localities, blue amphibole lineations define great circles, likely
reflecting folding of earlier (DR) fabric during DS (Figures5c and6b; relicts at Azolimnos in Figure7b). In other
localities, blue amphibole lineations appear to be reoriented into moderately S- or SW-plunging clusters (e.g.,
Grizzas and Kini, Figures5b and6b). Similarly, Keiter etal.(2004,2011) documented a significant spread of fold
axis orientations which they attributed to superposed folding that systematically reoriented older fold hinges via
S-vergent simple shear during prograde-to-peak subduction (i.e., their D2, black fold axes in Figure5).
Locally, centimeter-sized, prismatic pseudomorphs after lawsonite indicate that lawsonite grew at the culmination
of DS but did not survive peak conditions. Syn-to-post-kinematic porphyroblasts overgrow the mafic blueschist
foliation at Grizzas and Lia, decorate foliation-parallel compositional layers at Kini (Figure8c), and commonly
contain inclusions of garnet, and are included by garnet (Figure8c, closeup). Pseudomorphs are weakly attenu-
ated along the limbs of folds, but preserve their diamond-like shapes in fold hinges (Figure8c).
5.2.2. DS Microstructures and Mineral Chemistry
DS micro-textures in meta-sedimentary rocks are characterized by strong quartz-mica cleavage-microlithon SS
fabrics and locally record rotated inclusion trails in garnets that are mostly continuous with external foliations
(Figure9c). Quartz-rich microlithons have strong lineation-parallel shape-preferred orientations, and mica-rich
cleavages comprise intergrown phengite and paragonite (Figure9c and Figure10c). Lawsonite pseudomorphs
preserved as inclusions in garnet comprise intergrown epidote and white mica (Figure9d). Garnet compositional
Figure 12. Amphibole mineral chemistry and micro-textures. (a) Quantitative amphibole EMPA analyses (Leake
etal. classification scheme). All analyses have NaB>1.5 apfu except for those indicated with an asterisk. (b) DT1
static growth zonations in glaucophane contained in retrogressed eclogite pod. (c) DT1 lineation-parallel zonations developed
in glaucophane-filled strain shadow fringing garnet porphyroclasts. (d) Greenschists preserve relict DT1 sodic amphibole as
inclusions in epidote, and matrix amphibole records lineation-parallel compositional changes during DT2 retrogression.
KOTOWSKI ET AL.
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zoning varies between samples; some record complex pulses of Ca-enrichment (e.g., Figures10c and10d) while
others record Mn-rich cores and Fe-rich rims (Figure10g).
DS micro-textures in mafic blueschists are characterized by compositional segregation defined by glaucophane-rich
and epidote-rich layering alternating on the mm-scale (∼50–200μm grain size) (Figure9e). The SS foliation
contains syn-kinematic porphyroblasts of garnet and omphacite (∼300 μm-5 mm), and contains rutile with minor
titanite overgrowths (Figures9f and10a). Syn-kinematic phengitic white mica is chemically homogeneous and
has 3.35–3.45 Si atoms p.f.u (Figures11a and Figure S2 in Supporting InformationS1). Omphacite and garnet
deflect local foliations, and have pressure shadows and strain caps composed of glaucophane, phengite, and para-
gonite, and/or more omphacite (Figures8d and10a). Omphacite porphyroblasts in Kini blueschists have cores
of low-Na, high-Mg omphacite, fringed by asymmetric, syn-kinematic pressure shadows of high-Na, low-Mg
omphacite (Figure10a). DS amphibole is glaucophane (Figures10a and12a). Rare examples reveal glaucophane
cores with thin, patchy rims (Figure10b) that trend toward lower Al
iv+Fetot) values and higher (Na+K)A
(Figures12a and Figure S2 in Supporting InformationS1). In the samples studied, garnet compositional zoning
is less pronounced than in meta-sedimentary rocks; we observed weak Mn-enrichment in some garnet cores. See
Laurent etal.(2018) for more details on garnet zoning.
5.3. DT — Retrograde Fabric Development, Crenulation, and Re-Folding Through Blueschist-to-
Greenschist Facies Conditions
5.3.1. DT Structures
DT1 is best recorded at Kampos Belt and Palos (Figures5a and5c), Azolimnos (Figure7b), Kalamisia (Figure7a),
and locally at Kini (Figure6b). DT1 structures refold older SS foliations into inclined-to-upright, open-to-tight,
shallowly to moderately N- and NE-plunging folds (Figures5c and6d,6e,7a,7b; Figure S4 in Supporting Infor-
mationS1). Glaucophane, calcite, and quartz mineral and stretching lineations are oriented parallel to FT fold
hinge lines (Figures5c and7a,7b). Along Kampos Belt, DT1 fold axes span ∼335–055°/15–45°, with a cluster of
moderately N-plunging folds (e.g., Figure5c). At Azolimnos, DT1 folding locally develops an upright crenulation
cleavage (ST) that cuts the SS foliation (Figures7a and7b,8g). Cm-scale spaced cleavages are parasitic to larger
open folds with 045°/5–10° fold axes and steep axial planes. At Azolimnos, glaucophane lineations define a great
circle and swing from N to NE into alignment with FT1 crenulation hinge lines (Figure8h). Crenulation of Kini
rocks is defined by a vertical, NE-striking ST1 cleavage that cross-cuts mafic blueschists (Figure6b).
EventContext Diagnostic Structures MetamorphismExample Localities
DRSubduction • Only preserved as inclusion trails in garnets and as early
fabric (SR) that is tightly folded during DSlawsonite-blueschist N/A
• Axial plane schistosity (SS) associated with tight to isoclinal
folds (FS) that transpose the SR foliation, with S-SW-
plunging fold axes
Grizzas and Kini
• S-SW mineral and stretching lineations
• Dominantly non-coaxial with top-S-SW sense of shear,
locally non-penetrative in mafic lenses (e.g. Grizzas)
• Crenulation cleavage (ST) associated with upright, open-to-
tight folds (FT) that fold SS
• SS foliation continuously reworked and retrogressed
• Fold axes and mineral lineations rotate from N-NE (DT1) to
E-W (DT2) as a function of strain
• Dominantly coaxial, but locally non-coaxial (e.g. near the
Vari Detachment at Fabrikas, Kalamisia)
• Ductile to semi-brittle boudinage in later stages
progressing to greenschist
Note. Where DR is preserved, it is recorded as inclusion trails within garnet; however, garnet growth occurred at different times in different lithologies (i.e., before,
during, and after DR). Some garnets lack an internal foliation, and others contain inclusion trails that are oblique to (pre-DS) or continuous with SS (syn-DS).
Summary of Interpreted Deformation-Metamorphism Events in the CBU on Syros
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DT2 is characterized by E-W oriented mineral and stretching lineations that are primarily indicative of green-
schist facies conditions (e.g., Lotos, Delfini; Figures6a and6c) but locally preserve blueschist facies conditions
where strain was highly non-coaxial (i.e., Fabrikas; Figures7c and8j), and can be seen in a wide range of rock
types throughout central and southern Syros. At Vaporia, the mafic blueschists and eclogites and the surround-
ing meta-sedimentary rocks develop identical DT2 structures (Figures6d and6e). Single greenschist facies FT2
folds range in geometry from open to tight and have near-vertical, E-NE-to E-W striking axial planes. FT2 fold
axes cluster strongly around ∼070–110°/5–30° (Figures6a,6c and8i,8k), and mineral and stretching lineations
defined by actinolite, quartz, calcite, and relict glaucophane are oriented parallel to FT2 hinge lines (Figures6a, 6c
and6e). Older SS foliations are progressively reworked during DT2 creating a composite retrogressed foliation that
is visible as S- and Z-folds (e.g., Figures8i and8k) with hinge-limb layer thickness variations locally exceeding
20:1 (FigureS4 in Supporting InformationS1). FT2 folds have axial planar cleavages decorated with actinolite,
epidote, and chlorite. Coaxial stretching parallel to FT2 fold hinges is common, resulting in semi-brittle to brittle
boudinage of epidote-rich lenses visible from the meso-to the micro-scale, as competent lithologies become brit-
tle during exhumation (Figure8l). At Delfini, shear sense clast counting of a carbonate meta-conglomerate (GPS:
37°27’36” N/024°53’46” E) reveals conflicting and/or ambiguous shear sense. This is indicative of dominantly
coaxial strain during reworking of a composite foliation that develops syn-kinematically with respect to upright
folding (cf. Figure S4 in Supporting InformationS1). Although DT strain is primarily coaxial, strongly asym-
metric strain occurs locally on the E-SE side of the island. Non-coaxial DT1−2 is best preserved at Kalamisia and
Fabrikas, respectively. At Fabrikas for example, outcrop-scale extensional top-to-the-E shear bands and boudi-
nage cross-cut eclogite pods are decorated by glaucophane (partially replaced by actinolite) and quartz (Kotowski
& Behr,2019; Laurent etal.,2016).
5.3.2. DT Microstructures and Mineral Chemistry
DT1 microstructures transpose and retrogress older SS foliations (creating a composite, or reworked, SS folia-
tion), record geochemical evidence for retrogression through primarily blueschist facies conditions, and are
primarily coaxial. Crenulation hinges that record DT1 in mafic blueschists are defined by high-Si white mica and
glaucophane that has an identical composition to glaucophane defining the SS foliation (Lia Beach, Figure9e;
FigureS2 in Supporting InformationS1). Coaxial DT1 deformation in mafic blueschists is evidenced by symmet-
ric strain shadows around partially chloritized garnets. During DT1, SS-defining blue amphibole grows in
the symmetric strain shadows and records lineation-parallel growth zonations trending from glaucophane to
magnesio-riebeckite (Vaporia, Figures12a and12c) and locally becomes actinolitic (e.g., Kampos, Figure9g).
Some static textures record the same compositional trend (e.g., Figure12a and12b). At Kalamisia, extensional
C-C’ fabrics are well-developed in thin sections, and C’ top-to-the-ENE shear bands are decorated with albite,
paragonite, and phengite (Figures10e and10f). C’ cleavages are also defined by finely recrystallized blue amphi-
bole that records lineation-parallel core-to-rim zonations from high-Al riebeckite to low-Al (and lower (Na+K)
A) riebeckite (Figures9i,10f, and12a). Omphacite and paragonite porphyroblasts record the breakdown reaction
omphacite+ paragonite+ H2O= sodic amphibole+ epidote+ albite (Figure9i), and rutile is overgrown by
syn-kinematic titanite (Figure 10e). In quartz-mica schists, the retrogressed SS foliation comprises alternat-
ing glaucophane-rich and quartz-mica±albite-calcite layering. Bimodal meta-volcanics at Azolimnos exhibit
strong foliations defined by intergrown phengite and paragonite that are mostly in textural equilibrium. Phengite
is compositionally homogeneous with consistent Si values of 3.33± 0.01 atoms p.f.u. (n = 24). Paragonite
grains locally contain flakes of phengite that may indicate partial or incomplete recrystallization during shearing
(Figure11b). Locally, a syn-DT1 axial planar cleavage, ST1, is defined by actinolite, albite, phengite, and parago-
nite in the cores of upright FT1 folds (Figure9h).
DT2 microstructures transpose and retrogress older SS foliations, and are primarily coaxial and record geochem-
ical evidence for retrogression under greenschist facies conditions (e.g., Delfini and Lotos). Locally DT2 was
non-coaxial and developed under blueschist facies conditions (e.g., Fabrikas). Mafic greenschists that record DT2
comprise strongly retrogressed SS foliations that are defined by fine-grained white mica, albite, epidote, actino-
lite, chlorite, calcite, and titanite (∼50–500μm grain size), and contain lineation-parallel epidote porphyroblasts
(∼2–5mm) and unoriented, mat-like albite porphyroblasts (∼1–5mm) (Figures9j and9k). Amphibole occurs in
two distinct contexts: as inclusions in epidote and albite porphyroblasts, and as a dominant SS foliation-forming
phase. Amphibole inclusions record core-rim zonations evolving from magnesio-riebeckite to winchite, and
matrix amphibole record core-rim zonations evolving from ferro-winchite to actinolite (Figures12a and12d).
KOTOWSKI ET AL.
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SS-defining, syn-DT2 epidote porphyroblasts have pressure shadows filled with white mica, calcite, and albite, and
are boudinaged with necks filled with quartz and calcite (Figures9j and12d).
Foliation-forming white mica in meta-volcanics at Lotos is phengitic, with average Si atoms p.f.u. of 3.41±0.04
(n=63). Locally, white mica exhibits core-rim zonations characterized by a subtle increase in Fe, decrease in Al,
and decrease in Si; Si atoms p.f.u. decrease slightly from core to rim but still overlap with cores within error (dark
cores=3.42±0.04, bright rims=3.37±0.02; Figures11a and11c). Bright zones are also concentrated around
reaction pockets where relict amphibole breaks down to albite, quartz, and white mica (Figure11c). At Delfini,
greenschist facies meta-volcanics exhibit strong foliations defined by intergrown, texturally-equilibrated phengite
and paragonite. Rare phengites preserve core-rim zonations characterized by decreasing Si from 3.45±0.05
(n=7) to 3.33±0.2 (n=31). More commonly, paragonite grains develop syn-kinematic, lineation-parallel phen-
gitic tails that have Si values similar to foliation-forming phengites (3.35±0.04, n=11) (Figures11a and11d).
In blueschist facies fabrics at Fabrikas, the retrogressed SS foliation comprises syn-DT2 epidote porphyroblasts
that contain rotated inclusion trails of quartz and glaucophane and inclusions of garnet that preserve syn-DS
spessartine-to-almandine zonations (Figures10g and10h). Phengite and paragonite define C- and C’-planes of
an extensional, top-to-the-E shear fabric. Phengitic white mica reveals a tight range of Si atoms p.f.u (∼3.33–3.39
a.p.f.u, FigureS2 in Supporting InformationS1), and Si content of C- and C’-defining phengite is identical
(Figures10g and Figure S2 in Supporting InformationS1). Lineation-parallel brittle micro-boudinage of epidote
and amphibole porphyroblasts is common; epidote boudin necks are filled with quartz, and blue amphibole
boudin necks contain green amphibole needles. A planar ST2 fabric that cuts SS is only found in the core of FT2
folds (i.e., ST2 crenulation cleavage at Delfini, Cisneros etal.).
6. Multi-Mineral Rb-Sr Isochron Petrochronology
Results from the five samples selected for Rb-Sr geochronology are shown in Figure13 and Table2. The samples
include a meta-basalt, two meta-volcanosedimentary rocks, a greenschist facies reaction rind around an epidote
pod, and greenschist mineralization in an epidote lens boudin neck, and record distinct stages of the structural
evolution as outlined in Section5 (i.e., DS, DT1, and DT2). All of the isochrons described herein have Mean Stand-
ard Weighted Deviations (MSWDs) greater than 1, which suggests that the data dispersion exceeds that predicted
by analytical uncertainties (Wendt & Carl,1991). However, we consider our Rb-Sr ages reliable records of true
deformation-metamorphism events (see TableS1 in Supporting InformationS1). This is because the isochrons
were constructed from mineral suites that our structural and petrographic observations suggest are co-genetic,
the co-linearity of the data is striking (with some justifiable exceptions discussed below), and in constructing an
isochron we are directly testing which minerals are cogenetic and which are not. The high MSWD values may
reflect an underestimation of our analytical uncertainties, or minor Rb-Sr disequilibrium during progressive
Figure 13. Multi-mineral Rb-Sr isochrons from a Kini omphacite-epidote-glaucophane schist (SY1616), Azolimnos glaucophane-mica schist (KCS1617), and
Delfini actinolite-mica schist (KCS1621). Gray insets are zoom-ins of low Rb/Sr separates outlined in black boxes. Faded gray symbols were excluded from isochron
calculations. Multiple white mica separates for each isochron are shown in black symbols. Sample SY1616 records DS in the northern nappe, KCS1617 records DT1
in the central nappe, and KCS1621 records DT2 in the central nappe. DT retrogression pre-dates the onset of regional core complex capture. Mineral abbreviations:
ep=epidote, glc=glaucophane, om=omphacite, grt=garnet, wm=white mica, chl=chlorite.
KOTOWSKI ET AL.
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Sample ID and summary Mineral Grain size (μm)
(ppm) Sr (ppm) 87Rb/86Sr 87Sr/86Sr +2σ
SY1616: Kini omphacite-epidote-glaucophane schist epidote (L18-001) 75-125 0.12 1008 0.00035 0.703224 0.000014
Solution on 10 points: 53.5+0.6Ma glaucophane (L18-010) 75-125 0.15 143 0.00301 0.703225 0.000014
Initial 87/86 Sr: 0.703209+0.000012 omphacite (L19-099) 75-125 0.28 31.3 0.02571 0.703235 0.000014
MSWD=4 white mica (L19-097) 125-250 (0.5A) 0.31 18.8 0.04765 0.703244 0.000014
white mica (L19-093) 250-500 (0.2A) 0.31 15.5 0.05829 0.703234 0.000014
white mica (L18-009) 125-250 (0.5A) 0.37 16.6 0.06439 0.703284 0.000020
garnet #1 (L18-011) 250-500 0.29 13.0 0.06562 0.703261 0.000014
white mica (L19-094) 250-500 (0.3A) 1.16 43.8 0.07644 0.703248 0.000014
white mica (L19-096) 250-500 (0.4M) 6.05 142 0.12305 0.703289 0.000014
white mica (L19-095) 250-500 (0.5M) 34.9 23.7 4.26296 0.706398 0.000014
removed from isochron
garnet #2 (L19-098) 250-500 0.35 12.2 0.08338 0.703353 0.000014
KCS1617: Azolimnos glaucophane-mica schist white mica (L19-103) 250-500 (0.6M) 9.15 199 0.13309 0.706681 0.000014
Solution on 7 points: 45.5+0.3Ma white mica (L19-102) 250-500 (0.5M) 14.7 179 0.23810 0.706776 0.000014
Initial 87/86 Sr: 0.706592+0.000022 glaucophane (L18-002) 250-500 0.31 2.68 0.33478 0.706783 0.000014
MSWD=8 white mica (L19-100) 250-500 (0.3M) 33.8 178 0.55161 0.706927 0.000014
white mica (L18-007) 125-250 (0.8M) 4.66 13.9 0.97194 0.707200 0.000014
white mica (L19-101) 250-500 (0.4M) 112 112 2.87985 0.708433 0.000014
white mica (L18-005) 250-500 (0.7M) 219 42.5 14.89794 0.716067 0.000014
removed from isochron
epidote (L18-003) 0.70 1486 0.00136 0.706668 0.000014
garnet #1 (L19-004) 0.69 7.53 0.26530 0.706583 0.000014
garnet #2 (L19-104) 0.87 3.97 0.63469 0.706733 0.000014
KCS1621: Delfini actinolite-mica schist epidote 125-250 0.97 1961 0.00143 0.706597 0.000014
Solution on 7 points: 37.1+0.1Ma white mica (L19-225) 250-500 (0.6M) 54.0 258 0.60594 0.706951 0.000014
Initial 87/86 Sr: 0.706626+0.000033 white mica (L19-222) 125-250 (0.5M) 326 38.6 2.37806 0.707878 0.000014
MSWD=13 chlorite (L19-226) 250-500 (0.25M) 9.40 11.3 2.41488 0.707852 0.000014
white mica (L19-224) 125-250 (0.6M) 150 155 2.79655 0.708052 0.000014
white mica (L19-223) 125-250 (0.5M) 142 173 24.45665 0.719330 0.000014
white mica (L19-221) 250-500 (0.4M) 369 20.7 51.64803 0.733354 0.000015
removed from isochron
phengite (L19-220) 125-250 (0.4M) 343 46.8 21.16233 0.717173 0.000014
SY1644: Delfini mineralization in epidote boudin neck epidote (L19-041) >2000 0.23 2170 0.00031 0.706608 0.000014
Solution on 3 points: 36.1+2.6Ma actinolite (L19-042) ∼250-1000 16.9 123 0.39700 0.706899 0.000014
Initial 87/86 Sr: 0.706655+0.00058 white mica (L19-040) >1000 303 29.9 29.24565 0.721388 0.000014
SY1402: Lotos reaction rim around epidote pod apatite <100 2.49 726 0.00992 0.704969 0.000008
Solution on 5 points: 34.9+5.8Ma white mica (L19-029) <125 204 12.5 47.20997 0.724376 0.000014
Initial 87/86 Sr: 0.70455+0.00363 white mica (L19-030) 125-250 227 9.54 68.82184 0.739526 0.000015
MSWD=76,000 white mica (L19-031) 250-500 234 7.13 95.01809 0.753424 0.000015
white mica (L19-032) 250-500 203 8.65 67.84958 0.736426 0.000015
Summary of Rb and Sr Concentrations and Measured Ratios From Analyzed Samples
KOTOWSKI ET AL.
20 of 41
metamorphism (perhaps due to incomplete recrystallization, e.g., Halama etal.) that does not significantly
affect our Rb-Sr ages (TableS1 in Supporting InformationS1).
Sample SY1616 is an omphacite-epidote-glaucophane schist collected at Kini Beach and records DS. The miner-
als selected for isotopic analysis define the fabric shown in Figures8d, 9f and10a. This sample yielded an age of
53.5±0.6Ma (MSWD=4) based on a 10-point isochron defined by epidote, glaucophane, omphacite, six white
mica separates, and garnet (Table2, Figure13). To test the robustness of the isochron, several two-to five-point
isochrons were calculated from combinations of the co-genetic phases; the age does not change but the MSWD is
reduced (=1 for 2-pt isochrons by definition; <1 for 3- and 4-pt, and 1.4–1.7 for 5-pt).
Sample KCS1617 is a glaucophane-mica schist collected at Azolimnos and records DT1. The minerals selected
for geochronology define a fabric identical to that shown in Figures 10c and 11b, corresponding to blue
sub-horizontal layering as seen in Figure8g (i.e., a compositeSS fabric reworked and recrystallized during DT1
retrogression and crenulation). Similar rocks in the CBU on Syros that comprise cm-to-dm-scale intercalations
of glaucophane-epidote and quartz-mica-rich layers have been interpreted as bimodal meta-volcanics (Bröcker
& Keasling,2006; Keiter et al.,2011). This sample was collected from a glaucophane-rich layer and yielded
an age of 45.5±0.3Ma (MSWD=8) based on a 7-point isochron defined by glaucophane and six white mica
separates (Table2, Figure13). Our microstructural and petrologic characterization of phengite-paragonite inter-
growths from this sample are consistent with the colinearity of white mica separates defining the KCS1617
isochron; phengite compositions are homogeneous, and phengite-paragonite pairs are mostly in textural equilib-
rium. However, two garnet separates fell off of the isochron and are discarded in the age calculation. We justify
this based on microstructural observations shown in Figure10d; garnets preserve complex Ca-zonation patterns
and may record pulsed growth. Furthermore, garnets are probably DS porphyroblasts and are not expected to be
in isotopic equilibrium with the DT1 fabric during crenulation cleavage development and incipient retrogression.
Previous work suggests that Sr isotopic zoning in garnets (Sousa etal.,2013) and/or isotopic inheritance from
earlier stages of metamorphism (Romer & Rotzler,2011) tend to make garnets poor candidates for Rb-Sr isoch-
rons. Recent Lu-Hf garnet geochronology from Azolimnos confirms that garnet growth is older than the fabric
we dated (Uunk etal.,2022). Adding epidote to the isochron does not change the age (45.43±0.46Ma, n=8),
but increases the MSWD to 23. Epidote is stable throughout subduction and exhumation and could record subtle
zonations that grew during subsequent deformation events and therefore may not be co-genetic (see Cisneros
Sample KCS1621 is an actinolite-mica schist collected from Delfini and records DT2. It was collected from a fold
limb of a structure like the one in Figures8i and is interlayered on the decimeter-scale with actinolite-epidote-chlo-
rite schists, meta-cherts, and mica-schists that locally preserve blue amphibole lineations (see FigureS4 in Support-
ing InformationS1, photos R-T for examplesof such structures). This sample yielded an age of 37.1±0.1Ma
(MSWD=13) based on a 7-point isochron defined by epidote, chlorite, and 5 white mica separates (Table2,
Figure13). For this sample, various combinations of 2- to 6-pt isochrons all yield ages of ∼35–37Ma with
MSWD varying from ≪1 (e.g., 3-pt epidote-chlorite-white mica), to 1 (e.g., 2-pt white mica-chlorite) to 21 (e.g.,
4-pt epidote-chlorite-white mica-white mica). Even isochrons that are not well-defined in high-Rb space (i.e.,
do not contain phengitic white mica) yield nearly identical ages to the 7-point isochron (TableS1 in Supporting
InformationS1). Our microstructural observations demonstrate that core-rim zonations in paragonite and phen-
gite are syn-kinematic and cogenetic, and that the dominant fabric is retrograde. Rare, relict higher-Si phengite
cores testify to an earlier higher-pressure history that was pervasively recrystallized under lower-P conditions.
Relict high-Si cores do not appear to impact our age calculations, but it is possible that the one mica separate we
discarded from the KCS1621 isochron could reflect incomplete recrystallization.
Sample SY1644 is a collection of minerals precipitated in the neck of a boudinaged brittle epidote-rich lens from
Delfini, and sample SY1402 is a greenschist facies reaction rind at the margin of an epidote-rich lens from Lotos.
These samples are representative of semi-brittle boudinage associated with DT2 stretching (e.g., Figure8l), which
are genetically related to the white mica-rich fabric shown in Figure11c. These samples yield ages with reason-
able uncertainties, but extremely high MSWDs. Sample SY1644 yielded an age of 36.1±2.6Ma (MSWD=82)
based on a 3-point isochron defined by epidote, actinolite, and phengite, and sample SY1402 yielded an age of
34.9±5.8Ma (MSWD=76,000) based on a 5-point isochron defined by apatite and 4 phengites (Table2). For
both samples, 2-pt isochrons yield ∼36Ma and ∼29–36Ma, respectively (MSWD=1; TableS1 in Supporting
InformationS1). Based on our microstructural and petrologic observations of metamorphic replacement reactions,
KOTOWSKI ET AL.
21 of 41
Si- and Fe-zoning in phengites, and pervasive chloritization, it is unsurpris-
ing that these samples yielded poorer isochrons. Therefore, we consider these
data qualitative. However, these ages are similar to and trend slightly younger
than KCS1621, which is consistent with our structural observations.
7.1. Synthesis of Structural and Petrologic Data and Interpreted
7.1.1. DR Deformation and P-T Conditions
We interpret the DR fabric as the oldest recognizable in the CBU, formed under
lawsonite-blueschist facies conditions based on several lines of evidence: (a)
DR inclusion trail mineralogy (e.g., glaucophane, omphacite, and phengite),
(b) pseudomorphs of DR−S lawsonite included in DS garnets from metabasites
on Syros (also seen on Sifnos) (Okrusch & Bröcker,1990; Ridley,1982),
and (c) syn-kinematic DR−S omphacite blasts recording up-pressure, core-to-
rim zonations marked by increasing jadeite component (Figure 10a) (cf.
Thompson,1974). Lawsonite and epidote appear to have both been stable
in mafic bulk compositions during DR, with lawsonite growing later on the
prograde path under higher-pressure conditions (cf. Ballevre etal., 2003).
This is consistent with textural observations of lawsonite growing both syn-
and post-tectonic with respect to the SR foliation, incorporating inclusions of
garnet (which also grows near peak pressures, cf. Baxter & Caddick;
Dragovic etal.[2012,2015]), and being included by garnet.
7.1.2. DS Deformation and P-T Conditions
Deformation stage DS captures peak metamorphic conditions, and produced:
(a) an axial plane schistosity, SS, associated with tight to isoclinal folds
(FS) that fold SR and have SSW-plunging fold axes, (b) SSW-to-S-plung-
ing mineral lineations, (c) a blueschist-to-eclogite facies fabric containing
syn-kinematic garnet, omphacite, and (now pseudomorphed) lawsonite
porphyroblasts, and (d) chemical zonations in glaucophane and omph-
acite that record syn-kinematic increase in pressure and temperature. Our
structural observations are consistent with previous studies that attribute top-SSW thrust-sense kinematics to
prograde-to-peak subduction (Keiter etal.,2004; Laurent etal.,2016; Philippon etal.,2011). New and compiled
metamorphic geochronology demonstrates that different structural levels of the CBU on Syros experienced peak
conditions and DS deformation at different times (i.e., younging with structural depth, cf. Figure15; discussed
further below). However, it appears that each tectonic slice experienced similar P-T trajectories, including peak
P-T, despite subducting at different times. Uunk etal.(2022) came to a similar conclusion by combining garnet
Lu-Hf geochronology with thermodynamic modeling; they suggested that garnets throughout the Syros CBU
grew at similar pressures (19–21kbar) but at different times.
We do not provide new quantitative constraints on DS metamorphic conditions, but peak P-T for the DS fabric
shown in Figure14 are consistent with our petrologic observations and previous studies. Peak temperatures have
been calculated from garnet-omphacite major element exchange for mafic blueschists and eclogites (450–550°C)
(Laurent etal.,2018; Okrusch & Bröcker,1990; Rosenbaum etal.,2002; Schliestedt,1986); the upper limit of
glaucophane stability in marble (∼500°C at ∼15–16kbar; Schumacher etal.); and calculated lawsonite-out
reactions that predict up-temperature, prograde dehydration according to the reaction lawsonite=epidote+para-
gonite+H2O) at ∼400–500°C over ∼12–20kbar (depending on bulk rock and fluid composition) (Evans,1990;
Liou,1971; Philippon etal.,2011; Schumacher etal.,2008) (Figure14). Raman Spectroscopy of Carbona-
ceous Material from graphite schists suggests slightly higher temperatures of ∼540–560°C (Laurent etal.,2018).
Observed porphyroblast stability (e.g., lawsonite pseudomorph inclusions in garnets and vice versa), amphibole
chemistry, and the volumetric dominance of glaucophane-bearing marbles throughout the CBU on Syros are
generally consistent with peak T of ∼500–550°C.
Figure 14. Preferred, schematic P-T-D-t path for the Cycladic Blueshist Unit,
consistent with observations and analytical results from this study. The shape
of the path is modified from Schumacher etal.(2008). Amphibole stability
fields constraining DT2 temperatures are from Otsuki and Banno(1990).
The timing of metamorphism labeled along the P-T loop corresponds to
progressive subduction and exhumation of three distinct tectonic slices (the
northern, central, and southern nappes; see Sections6, 7.1 and Figure15).
Mineral abbreviations: crs=crossite (sodic amphibole), mg-rieb=magnesio-
riebeckite, wch=winchite, act=actinolite. Facies fields defined in Figure4.
KOTOWSKI ET AL.
22 of 41
Reported peak pressures for DS are variable in the literature and challenging to reconcile. Early conventional
thermobarometry suggested peak P of ∼12–18kbar in mafic blueschists and eclogites (Dixon,1976; Okrusch
etal.,1978; Okrusch & Bröcker,1990; Schliestedt,1986). These pressures are supported by solid inclusion
quartz-in-garnet barometry constraining garnet growth conditions at Kini, Kalamisia, Delfini, and Lotos
to ∼13–17kbar (Behr & Becker,2018; Cisneros et al., 2021). Gorce et al. (2021) recently investigated one
sample from Fabrikas, and demonstrated that the majority of pressures derived from solid inclusion barometry
(∼17–19kbar) overlap within the error of pressures derived from thermodynamic modeling that account for
garnet fractionation, although the P-T models trend slightly higher (∼20–22kbar). Other thermodynamic models
suggested rocks reached ∼19–21kbar (Uunk etal.,2022) or even as high as ∼22–24kbar (Laurent etal.,2018;
Skelton etal.,2019; Trotet, Jolivet, & Vidal,2001). We consider this unlikely based on the abundance of SS para-
gonite and the absence of kyanite in meta-mafic rocks, which suggests that the upper stability limit of paragonite
at ∼20–23kbar was not reached (Okrusch & Bröcker,1990; Schliestedt,1986; Skelton etal.,2019) (Figure14),
although we acknowledge that the kyanite-in reaction is strongly dependent on bulk rock composition (cf. Laurent
etal.,2018). Large differences in P-T estimates between traditional phase equilibria and recent thermodynamic
modeling may reflect arbitrary choices of thin section domains selected as representative bulk compositions (e.g.,
Lanari & Engi,2017). This effect has been demonstrated for CBU lithologies on Syros (see Figure 15 in Laurent
etal.) and is especially likely in garnet-bearing rocks, due to the strong disequilibrium effect that garnet
exerts on local bulk composition (Lanari, & Engi,2017; Lanari & Engi,2017; Lanari & Duesterhoeft,2018). It
is also possible that higher-P conditions are real, but have not yet been sampled by solid inclusion techniques.
Figure 15. Metamorphic age versus structural depth along the cross-section line A1-A6 as shown in Figure2 (modified from Keiter etal.,2011). Locations that crop
out on the cross section line are labeled; in the upper panel, other locations are indicated that project into or out of the page at the structural level shown (e.g., Kini is
a normal fault-bounded block on the west side of the island and therefore is not shown on the cross section). Only ages that were confidently linked to our proposed
deformation scheme are included. Clusters of ages outlined in black boxes are derived from the same locality and collapse onto a single point on the cross section.
Delfini symbols marked with stars were reported as blueschist-facies fabrics by Cliff etal.(2016); however, local preservation of glaucophane under greenschist facies
conditions can be due to CO2-bearing fluids. Bold half arrows outlined in red indicate interpreted nappe-delimiting ductile shear zones, which were likely reworked as
extensional structures during exhumation.
KOTOWSKI ET AL.
23 of 41
7.1.3. DT Deformation and P-T Conditions
DT represents retrograde deformation under blueschist-to-greenschist facies conditions during exhumation. DT is
distinguished by: (a) transposition of the SS foliation during the formation of upright, open to tight FT folds and
progressive new (ST) fabric development, (b) lineation orientations that rotate from N-NE (DT1) to E-W (DT2)
with progressive strain and (in general) increasing greenschist facies retrogression, (c) dominantly coaxial, but
locally non-coaxial deformation, and (d) chemical zonations in amphibole tracking syn-kinematic decrease in
pressure and temperature during the development of a composite, reworked foliation (e.g., SS is locally deformed
and metamorphosed during DT). Our structural data are consistent with previous studies that identified top-NE
and top-ENE extensional shear as well as E-W coaxial stretching (at different times and locations, as discussed
further below) during exhumation (e.g., Bond etal.,2007; Keiter etal.,2004; Laurent etal.,2016; Rosenbaum
etal.,2002; Trotet, Vidal, & Jolivet,2001).
During DT, foliation-forming amphiboles transition from glaucophane to (magnesio) riebeckite, to winchite, to
actinolite. The progressive decrease of total Al, NaB, and (Na+K)A in amphibole indicates that P and T decreased
as DT evolved. Glaucophane coronas that develop around eclogite pods during DT1 are chemically similar to
syn-DS glaucophane, and retrogressed glaucophane records decreasing Al
vi (KCS53, KCS52B) and NaB (KCS53)
from core to rim, and a minor increase in (Na+K)A (Figure12, FigureS2 in Supporting InformationS1). These
signatures indicate decompression and potentially slight warming (Ernst & Liu,1998; Laird & Albee,1981;
Moody etal.,1983; Raase,1974; Robinson,1982), at the subduction-to-exhumation transition.
DT2 is characterized by foliation-forming sodic-calcic amphiboles, and local relicts of sodic amphiboles are found
as inclusions in porphyroblasts. The chemical transition from sodic-to-calcic amphibole as recorded in Syros
CBU rocksindicates cooling during decompression (Brown,1977; Ernst & Liu,1998; Laird & Albee,1981;
Maruyama etal.,1983; Moody etal.,1983; Otsuki & Banno,1990; Schmidt,1992; Thompson,1974) through
albite-epidote blueschist facies and eventually greenschist facies conditions (Figure 14). This P-T trend is
supported by the abundance of albite and titanite overgrowths on rutile (this study), boudin neck quartz-calcite
oxygen isotope temperatures and quartz-in-epidote inclusion barometry (Cisneros etal.,2021), and decreases
from core-to-rim in celadonite component of foliation-forming white micas (Laurent etal. and this study).
While we cannot rule out an initial phase of isothermal decompression at high pressures, our documented amphi-
bole geochemical zonations are better explained by cooling during decompression at moderate pressures and do
not support a positive thermal excursion into the epidote-amphibolite facies field (e.g., edenite, pargasite, cros-
site), as suggested by Laurent etal.(2018), Lister and Forster(2016), and Trotet, Jolivet, and Vidal(2001) P-T-D
paths. Notably, Aravadinou etal.(2016) reported syn-kinematic amphibole zonations from retrograde fabrics in
the CBU on Sifnos that also support exhumation along a cooling during the decompression pathway (see also
Schmädicke & Will).
7.2. Synthesis of Previously Published Metamorphic Geochronology
We compiled published metamorphic geochronology from 1987 through 2022 and took inventory of the descrip-
tions of deformation fabrics and metamorphic textures provided by the authors to re-evaluate the significance of
Eocene and Oligocene ages in the context of subduction versus exhumation. A full compilation can be found in
Supplementary FigureS1 and TableS2. We applied several qualitative filters to the dataset to derive a subset of
ages that we can confidently attribute to fabric-forming events. The filters are justified as follows:
Zircon U-Pb ages are robust records of igneous crystallization, but as metamorphic ages they can be difficult to
place in pro- or retrograde context (Liu etal.,2006; Poulaki etal.,2021; Tomaschek etal.,2003; Yakymchuk
etal.,2017). We include U-Pb ages from Tomaschek etal.(2003) for comparison with other ages, but we do not
rely on them for island-scale interpretations.
Garnet Lu-Hf and Sm-Nd are commonly considered reliable indicators of ‘peak’ subduction ages (i.e., maximum
depths) (Gorce etal.,2021; Lagos etal.,2007; Uunk etal.,2022), because HP/LT garnets tend to grow rapidly
following reaction overstepping (Baxter & Caddick,2013; Dragovic etal., 2012,2015). For example, Lagos
etal.(2007) reported evidence for rapid, pulsed garnet growth near peak conditions from tight clustering of Lu-Hf
ages even though samples exhibited different Lu zoning profiles and distributions between their cores and rims
(see also Skora etal.). In this case, this refutes the possibility that garnet cores grew significantly earlier
than their rims somewhere along the prograde path. Uunk etal.(2022) employed a stepwise dissolution technique
KOTOWSKI ET AL.
24 of 41
that may have preferentially removed younger rim ages, and they did not provide constraints on Lu zoning in their
samples, but nonetheless succeeded in differentiating several statistically distinct ‘bulk’ Lu-Hf garnet-whole rock
isochron ages for rocks occupying distinct structural levels on the island. However, evidence for protracted garnet
growth is also locally present. Gorce etal.(2021) reported Sm-Nd ages derived from garnet cores and rims that
were statistically distinguisable, and concluded that garnets nucleated near peak conditions and continued to grow
during decompression over a span of ∼5 Myr. In this case, coupling geochronology with thermodynamic mode-
ling provided a clear tectonic context for multi-stage or continuously evolving metamorphism.
White mica Ar/Ar has the potential to capture the timing of metamorphism during fabric development. However,
this system is highly susceptible to disequilibrium, partial (re-) crystallization and mixed ages, and/or unpre-
dictable loss or gain of radiogenic products, making it difficult to interpret the geological significance of an age
(Bröcker etal.,2013; Laurent etal.,2016; Lister & Forster,2016; Maluski etal.,1987). For the final dataset,
we only included five Ar/Ar step-heating ages with strong plateaus from micro-drilled grains which all had clear
micro-textural context (Laurent etal.,2017), and one strong plateau age from a well-characterized marble shear
zone (Rogowitz etal.,2014). We acknowledge that in other HP terranes, even strong plateau ages have been previ-
ously attributed to excess Ar (Sherlock & Arnaud,1999). However, the Ar ages included in this study overlap
within reported error of independent Rb-Sr isochrons from rocks at the same locality and/or similar structural
levels, which suggests that at least locally, excess Ar is absent (or apparently absent; cf. Ruffet etal.).
Rb-Sr isochrons are typically considered reliable constraints on fabric ages when the selected fabrics, and
minerals defining them, are well-characterized (Bröcker & Enders, 2001; Bröcker et al., 2013; Glodny &
Ring,2022; Skelton etal.,2019). Furthermore, constructing a Rb-Sr isochron directly tests the assumption that
selected minerals were in isotopic equilibrium during metamorphism. This validates interpretation of Rb-Sr
ages as deformation-metamorphism events even if whole rock powders serve asSr anchors (e.g., Bröcker &
Enders,2001). Micro-drilling of white micas and co-genetic Sr-rich phases (epidote or calcite) also provide
strong textural context for regressed ages (Cliff etal.,2016).
In some cases, we propose alternative interpretations of published data based on our own structural observations.
Skelton etal.(2019), for example, interpreted three of their Rb-Sr isochrons from Fabrikas as peak metamor-
phic ages (i.e., DS), but we interpret Fabrikas fabrics as DT1−2, associated with early exhumation (cf. Figure7c).
This revised interpretation is supported by previous petrologic observations of eclogite breakdown to blueschist,
replacement of glaucophane by actinolite, and core-to-rim decrease in celadonite component of foliation-forming
phengitic white mica (Kotowski & Behr,2019; Laurent etal.,2018) (see also Figure10), and structural studies
that report dominantly extensional top-to-the-E, exhumation-related deformation immediately beneath the Vari
Detachment (Kotowski & Behr,2019; Laurent etal.,2016). Top-E extensional kinematics at Fabrikas clearly
contrasts with other localities where prograde, top-SSW deformation is preserved. Sm-Nd garnet geochronology
coupled with thermodynamic modeling of a sample from Fabrikas further support this interpretation; Gorce
etal.(2021) estimated peak conditions were reached at ∼45Ma (garnet cores) and blueschist facies retrogression
occurred through ∼40Ma (garnet rims). The ∼40Ma age for blueschist retrogression overlaps with Rb-Sr isoch-
rons for blueschist fabrics presented by Skelton etal.(2019).
In another case, Cliff etal.(2016) analyzed micro-drilled phengites from blueschist-to-greenschist facies (i.e.,
DT1 to DT2) extensional fabrics in calc-schists and quartz-mica schists. Four of their samples from Delfini
were described as blueschist-facies (data points marked with black stars in Figure15); however, we observed
penetrative greenschist facies deformation at Delfini (DT2). Glaucophane is locally preserved in abundance in
calc-schists at Delfini, and elsewhere on Syros. Rather than reflecting blueschist facies conditions during defor-
mation, this could be due to a glaucophane-stabilizing, CO2-bearing fluid under greenschist facies P-T conditions
(Kleine etal.,2014), or channelized fluid flow at lithological boundaries leading to heterogeneous retrogression
Finally, Rogowitz etal.(2014) dated phengites from a top-E extensional greenschist facies marble shear zone,
and hypothesized the ages would be Miocene in accordance with the regional ‘M2’. They interpreted their Eocene
ages as evidence that Miocene deformation did not reset the isotopic signature. However, these authors concluded
that the microstructures in their marble mylonite sample were consistent with calcite deformation at ∼300°C. This
suggests their ages capture a true Eocene recrystallization event (e.g., strong E-W stretching during greenschist
KOTOWSKI ET AL.
25 of 41
facies DT2) below the Ar/Ar closure temperature in white micas (∼400–450°C, cf. Hames & Bowring;
In Figure15, the refined compilation (n= 47) and new Rb-Sr geochronology (n=5) are projected onto the
cross-section line drawn in Figure2. Ages are labeled according to fabric generation. Faded data points were
assigned textural identities but do not record penetrative strain (e.g., randomly oriented, radiating cluster). Key
observations from new and compiled geochronology include:
1. DS, blueschist-to-eclogite facies deformation-metamorphism spans ∼53 to ∼45 Ma, and is captured by a
multi-mineral Rb-Sr isochron (this study), and Lu-Hf and Sm-Nd garnet ages.
2. DS ages are oldest and well-clustered at Grizzas and Kini (∼53–52Ma), younger at Azolimnos (∼49Ma), and
youngest at Fabrikas (∼45Ma).
3. DT1, retrograde blueschist facies deformation-metamorphism spans ∼50–40 Ma (Rb-Sr isochrons and Ar/
Ar single grain analyses) and youngs with structural depth, that is, from Kampos, to Azolimnos, to Fabrikas.
4. DT2, retrograde greenschist facies deformation-metamorphism spans ∼42–20Ma (Rb-Sr isochrons and one
Ar/Ar age) and youngs with structural depth, that is, from Palos (∼43–35Ma), to Delfini (∼35–28Ma), to
5. Rocks that presently occupy different structural levels developed distinct fabric generations contempo-
raneously. Examples include: Azolimnos DS and Kampos DT1 (∼49Ma), Fabrikas DS and Kampos DT1−2
(∼45Ma), Fabrikas DT1 and Palos DT2 (∼43–38Ma), and Posidonia DT2 and non-penetrative greenschist meta-
morphism in the north (faded symbols, ∼25–20Ma). In other words, retrograde blueschist and greenschist
facies deformation-metamorphism occurred first in the structurally highest units and progressed structurally
downwards with time.
8. A Revised Tectonic Model for the CBU on Syros
Here we synthesize protolith age constraints and our structural, petrologic, and geochronologic data to propose
a revised tectonic model for the CBU on Syros. First, we present a pre-subduction configuration, then discuss
a stepwise reconstruction of progressive subduction, underplating, and exhumation, leading to the proposed
three-part tectonic stack exposed on Syros today
8.1. Pre-Subduction Configuration
Figure16 builds on previous work (e.g., van Hinsbergen etal.,2020; Papanikolaou,1987,2013; Ring etal.,2010)
and illustrates a highly schematic paleogeographic setting for the protoliths of the CBU on Syros and South-
ern Cyclades immediately prior to subduction at ∼60Ma. Peri-Gondwanan Cycladic Basement, cross-cut by
Carboniferous magmatism (∼315Ma on Syros, Tomaschek etal. ; 330–305Ma in Southern Cyclades,
Flansburg etal. ), was rifted in the Triassic (∼245–237Ma, Bröcker & Pidgeon; Keay).
Syn-rift bimodal volcanic and sediments intruded and blanketed the hyper-extended margin; these will become
the diagnostic marker horizons referred to as banded tuffitic schists and bimodal meta-volcanics mapped by
Keiter etal.(2011) (orange and dark gray in Figure16; cf. Figure2). Rifting was followed by passive margin
sedimentation of psammites, debris flows, and carbonates from the Triassic (∼230Ma) through the Cretaceous
(∼75Ma) (Löwen etal.,2015; Poulaki etal.,2019; Seman,2016; Seman etal.,2017). Carbonates interbedded
with clastic sediments may be the protolith for the Syringas Marker Horizon (Keiter etal.,2011). Cretaceous
igneous rocks (∼80Ma, Tomaschek etal.) dissected the potentially hyper-extended basement and passive
margin sedimentary sequence, forming a small oceanic-affinity (backarc?) basin (Bonneau,1984; Fu etal.,2012;
Keiter etal.,2011). Bröcker and Keasling(2006) also reported ∼80Ma crystallization ages for zircons in black-
wall alteration zones around a jadeitite block in the Kampos Belt and interpreted the ages to record hydrothermal
and/or metasomatic processes in a Cretaceous subduction zone, but this has not yet been confirmed (e.g., Bulle
etal.,2010). Major and trace elements, REE patterns, and stable oxygen isotopes in serpentinites on Syros are
consistent with formation in either an abyssal hyper-extended margin or mid-ocean ridge/fracture zone environ-
ment, rather than a mantle wedge source (Cooperdock etal.,2018).
Some previous studies propose that the ophiolite-affinity sequences within the CBU are meta-olistostromes or
meta-debris flows, due to the juxtaposition of various rock types and depositional/igneous crystallization ages
(Bröcker & Enders,1999; Dixon & Ridley,1987; Hecht,1985). However, we argue that the spatial distributions
KOTOWSKI ET AL.
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and contact relationships between different rock types do not require substantial mechanical mixing or subma-
rine landslides; they can instead be explained by the progressive evolution of the hyper-extended margin as
described above. Furthermore, the presence of several in-tact young-on-old stratigraphic relationships in the
detrital zircon record throughout the meta-sedimentary section (e.g., northern Syros) suggests that the strati-
graphic pile was not homogenized by submarine landslide or mélange mixing during subduction (cf. Poulaki
etal.,2019; Seman,2016; Seman etal.,2017).
The most interpretive parts of Figure16 are the locations of mafic intrusive igneous rocks. These rocks could
reflect off-axis, shallow intrusions related to Cretaceous seafloor spreading, or older mafic igneous rocks related
to Triassic rifting and represent the protoliths for mappable exposures of blueschist-eclogite lithologies (dark blue
in Figure2, commonly producing block-in-matrix shear zones). Protolith ages have not been determined for Kini,
Vaporia, Kalamisia, or Fabrikas mafic rocks on Syros, but zircons in meta-gabbroic pods in other block-in-matrix
exposures on Tinos and Samos, and all studied blocks on Kampos Belt on Syros, yield Cretaceous ages (Bröcker
& Enders,1999; Bröcker etal.,2014; Bröcker & Keasling,2006; Bulle etal.,2010). Regardless of their origin,
the key point is that protoliths for mafic blueschists and eclogites were distributed throughout the CBU before
subduction, rather than only coming from the small ocean basin in the north. This interpretation is supported by
metamorphic geochronology that demonstrates Kampos/Kini and Fabrikas experienced peak metamorphism at
different times that do not overlap within statistical uncertainty (see Figure15 and discussion below).
This paleogeographic interpretation allows us to split the CBU on Syros into three sub-domains characterized by
distinct, but related, protolith assemblages (dashed boxes in Figure16; see also Figure3). These sub-domains are
the precursors to each of the three main tectonic slices that comprise the structural pile on Syros.
8.2. Peak Subduction of the Palos-Gramatta-Kampos Nappe (∼53Ma)
The Palos-Gramatta-Kampos nappe (northern nappe) comprises remnants of the Cretaceous oceanic lithosphere
and associated syn-to-post-magmatic sedimentation over a Triassic-Jurassic rift margin (Figure16). Seman(2016)
estimated a Maximum Depositional Age of 80±6Ma for the protolith of the Gramatta schists. The zircons were
interpreted to be sourced from a Cretaceous volcanic center and deposited locally in a pelagic environment.
Figure 16. Schematic paleogeographic reconstruction of the protoliths for the Cycladic Blueschist Unit (CBU) at ∼60Ma. The zoomed-in cross section is modified
from Seman(2016). Stepwise evolution of the CBU during subduction is shown in the next figure.
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Seman(2016) suggested that overlapping depositional and crystallization ages between the Gramatta schists and
Grizzas/Kampos meta-igneous rocks indicate the two were genetically related. Alternatively, since MDAs do
not unequivocally constrain depositional ages, the Gramatta schist protoliths could be significantly younger than
80Ma. Additionally, coeval ages of metabasic blocks and detrital zircons could suggest both protoliths have the
same provenance but the sedimentary material was mixed during transport. The latter scenario requires that there
was no active volcanism in the region after 80Ma. Regardless, these observations suggest that the pre-subduction
relationship of sedimentary rocks deposited nearby an active or recently extinct volcanic center was preserved
throughout subduction and exhumation.
Our view of this structurally highest subunit differs from that of Laurent etal.(2016)'s ‘Kampos subunit’ in that
it does not solely comprise meta-mafic lithologies, and it does not include the map-scale meta-mafic lenses at
Vaporia, Kalamisia, and Fabrikas (see also Section9). Garnet Lu-Hf from Grizzas and Kampos Belt, and new
Rb-Sr isochrons from Kini yield identical ages of DS fabric development within error, suggesting that Kini was
originally subducted as part of the northern nappe (Figure15), and was down-dropped by late-stage, high-angle
normal faults to its present position (cf. Keiter etal.,2011; Ridley,1984). Garnet Lu-Hf geochronology from
Uunk etal.(2022) supports our interpretation that Fabrikas and Katerghaki were subducted later than Kampos
Prograde-to-peak subduction was characterized by extremely high top-to-the-SSW asymmetric shear strain and at
least two stages of foliation development under blueschist-to-eclogite-facies conditions (DR and DS; Figure17a)
(Keiter etal.,2004; Laurent etal.,2016; Philippon etal.,2011). Our observations of early prograde SW-plunging
fold axes and mineral lineations preserved at Grizzas, Kini, and locally along Kampos Belt, are consistent with
previous observations of top-SW prograde shear sense. Girdled glaucophane lineations (e.g., Kini, Kampos)
record continuous kinematic rotation from SW to N-S during subduction. Metamorphism led to the formation of
blueschists and eclogites under identical P-T conditions, reflecting differences in bulk composition and/or proto-
lith texture (Kotowski & Behr,2019; Skelton etal.,2019), creating outcrop-scale block-and-matrix structures. As
such, subduction-related strain was very heterogeneous. This is evidenced by rheologically strong meta-gabbros
at Grizzas and Kini that preserve primary igneous features (Keiter etal.,2004,2011; Kotowski & Behr,2019;
The northern nappe was underplated after DS and before DT exhumation, thus removing it from the active
subduction interface. Detrital zircon U-Pb data may support independent structural observations that suggest a
large thrust-sense shear zone separates the northern nappe from the central nappe beneath it (Keiter etal.,2004;
Laurent etal.,2016; Seman,2016) (Figure17a; drawn as the structurally highest ‘nappe-bounding’ thrust in cross
section in Figure15). These contacts are not discrete structures, but rather distributed shear zones that ‘smeared’
and locally mixed lithologies along unit contact zones. This thrust-sense shear zone placed Triassic and Creta-
ceous igneous rocks (Kampos) atop Cretaceous (Syringas) sediments. After it was removed from the interface,
the underplated nappe started to exhume, while subduction of the intermediate nappe continued beneath it.
8.3. Subduction-and-Imbrication of the Syringas-Azolimnos Nappe and Blueschist Facies Exhumation of
The Syringas-Vaporia-Azolimnos nappe (central nappe) occupies the central portion of the island and comprises
interbedded Triassic-to-Cretaceous meta-sedimentary schists, meta-volcanic schists, and meta-carbonates
(Figure16). In contrast to Laurent etal.(2016)'s central Chroussa subunit, we suggest that Vaporia and Kalamisia
meta-mafic lenses belong to this central slice and record primary intrusive and/or depositional relationships with
surrounding CBU meta-sediments that were sheared during subduction (cf. Keiter etal.,2011).
Until recently the timing of peak DS during subduction of the central nappe was unknown, but we hypothesized
that it must have reached peak conditions sometime between 52 and 45Ma based on well-constrained ages of
peak subduction in the northern and southern nappes. A weighted average of four garnet Lu-Hf isochrons from
Myttakas, Delfini, Azolimnos, and Chroussa of 48.6±0.6Ma supports our hypothesis (i.e., the ‘passive margin
domain’ of Uunk etal.). For Azolimnos specifically, Uunk etal.(2022) reported a garnet Lu-Hf isoch-
ron age of 49.0±0.5Ma. This is consistent with our Rb-Sr isochron from the same outcrop, which constrains
the timing of incipient blueschist facies retrogression at 45.5± 0.3 Ma, thus bracketing the timing of the
subduction-to-exhumation transition in an imbricated slice of the central nappe.
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Figure 17. Block diagrams illustrating the structural evolution and timing of subduction and exhumation recorded by the three tectonic slices in the Syros nappe stack.
Compare stepwise subduction of sub-units to the paleogeography in Figure16. Horizontal scaling is equivalent to subduction rates of ∼2–3cm/yr and diagrams are
roughly 2× vertically exaggerated. The thickness of the interface is exaggerated for clarity. Contacts between sheets are illustrated as thick black lines; we interpret these
structures to be distributed shear zones (∼10–100s m), rather than discrete localized structures.
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DS in the central nappe is largely overprinted during subsequent exhumation-related deformation, but early fabrics
are reminiscent of DS in the northern nappe and similarly consist of isoclinal folds and strong cleavage devel-
opment (e.g., textural relicts at Azolimnos). While DS developed in the central nappe, DT1 exhumation-related
blueschist facies fabrics formed at the same time in the northern nappe (Figures15 and17b).
Detrital zircon U-Pb geochronology and Maximum Depositional Ages (MDAs) of meta-sedimentary rocks in the
central nappe suggest that several old-on-young stratigraphic inversions may exist, which would imply that imbri-
cation occurred along cryptic ductile thrusts during subduction (Seman,2016) (thin black thrusts in Figure15,
pink stars in Figure17b). For example, Seman(2016) documented an old-on-young stratigraphic inversion where
Triassic meta-volcanics at Delfini are thrust atop Cretaceous meta-sediments east of Kini (Figure2). Even though
these structures cannot be seen in the field, the presence and locations of inferred thrusts are further supported
by the repeated Syringas Marker Horizon, which never appears overturned (orange circles and pink stars in
Figures15 and17b, respectively). Thus, we propose that the central nappe is bounded by larger nappe-delimiting
shear zones to its north and south, and also comprises smaller-scale thrusts accommodating internal imbrication
of CBU meta-sedimentary rocks, shown in the cross section in Figures2 and15. Imbrication is further supported
by garnet Lu-Hf ages which reveal at leasttwo distinct peak subduction events within the meta-volcano-sedi-
mentary rocks that occupy the central nappe of Syros and comparable lithologies in the CBU on Sifnos (Uunk
While we acknowledge that zircon U-Pb MDAs do not provide unequivocal constraints on protolith depositional
ages, and additional geochronologic and structural evidence is needed, they do provide intriguing insights into
the chronostratigraphy of the tectonic packages corroborated by differences in zircon provenance signatures.
Even though imbrication has interesting implications for the CBU's structural history, the present tectonic model
of progressive subduction, underplating, and return flow is supported by our data (particularly by the compiled
metamorphic geochronology) whether or not the central section is imbricated. Evidence for syn-subduction
imbrication has been documented elsewhere in the Cyclades, for example on Evia where lower marbles were
thrust over the upper marble sequence during prograde, top-to-the-SE shearing (Gerogiannis etal.,2021); on NE
Sikinos and Ios islands where several old-on-young inversions occur throughout the meta-sedimentary strata near
the CBU-Cycladic Basement contact (Poulaki etal.,2019); and on Milos (Grasemann etal.,2018). Therefore,
if syn-subduction tectonostratigraphic repetition is confirmed on Syros, imbrication may be a common and/or
recurring process during Eo-Oligocene subduction across the Cyclades.
During peak subduction of the central nappe (DS), DT1 deformation occurred in the northern nappe and was
characterized by upright folding, crenulation cleavage development, and NE-trending fold axes and mineral line-
ations. This structural transition is marked by ∼120–180° rotation in dominant mineral lineations and fold axis
orientations from the S-SW to the N-NE. We interpret the crenulation cleavage formed during DT1 to be a signa-
ture of the ‘subduction-to-exhumation transition,’ when rocks ‘turn the corner’ in the subduction channel, based
on the observation that crenulation lineations are decorated by high-pressure phases with compositions similar to
peak DS blueschist-to-eclogite facies conditions (Kini, Figure9e). Xypolias etal.(2012) also documented upright
ductile folding at the subduction-to-exhumation transition in the CBU in Evia. They interpreted the structures as
cross-folds that formed via constrictional flow under net compression and retrograde blueschist facies conditions.
DT1 and subsequent strain localized in weaker CBU meta-sediments during exhumation (e.g., Palos, Gramatta),
whereas prograde subduction-related fabrics are locally preserved in rheologically strong meta-gabbros at Griz-
zas and Kini. These observations support previous structural studies that suggest exhumation-related deformation
progressively localized toward the bottom of the structural pile, leading to more pervasive greenschist facies
overprints in the south of the island (Laurent etal.,2016; Lister & Forster,2016; Ring etal.,2020).
The structural base of the central nappe is difficult to pinpoint. However, metamorphic geochronology suggests
that it is somewhere below Azolimnos and is likely above the Fabrikas tectonostratigraphic horizon, which we
argue comprises the third and lowermost nappe. The presence of a nappe-bounding shear zone is also consistent
with progressive southward facies changes in the rock types, as carbonate horizons thin substantially and parag-
neissic material crops out at the island's southern tip, as well as the presence of thrust fault-bounded marble
klippe exposed locally on the southern portion of the island (Figures2 and15). Laurent etal.(2016) proposed a
nappe-bounding extensional shear zone across the island based on the observed intensity of greenschist overprint-
ing and the disappearance of marbles and suggested its western extent crops out as splaying shear zones above and
below the Delfini peninsula (i.e., their ‘Achaldi-Delfini shear zone’). If this is the case, then new and compiled
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geochronology suggests that greenschist facies overprinting in the southern slice spanned ∼36-20Ma. Alterna-
tively, if the nappe-bounding shear zone occupies a slightly deeper structural level (i.e., right beneath Delfini
peninsula, such that Delfini represents the lowermost portion of the central nappe and is heavily retrogressed
under greenschist facies conditions), then DT1−2 development in the central slice is slightly older (∼35-30Ma)
than in the southern slice (∼30–25 Ma). Our new Rb-Sr isochrons and recent garnet Lu-Hf ages from Uunk
etal.(2022) support the latter interpretation.
We suggest a slightly modified position of the lower-bound of the central nappe, which is shown in Figure2.
The location of this structure is primarily constrained by metamorphic geochronology at Fabrikas, Chroussa, and
Delfini; detrital zircon MDAs near Kini; the locations of the Syringas Marker Horizon; and locations of previ-
ously mapped marble klippe. We emphasize that the location of this structure is a hypothesis (hence the dashed
line and question marks). Further structural observations are needed to refine its location, for example, evidence
for a poly-deformed shear zone with an early thrust-sense history overprinted by younger extensional shear (cf.
Laurent etal.,2016). Both stages of deformation were likely accommodated across distributed shear zones rather
than discrete fault planes. As such, ductile shear zones (or fault planes) may be continuously transposed during
exhumation, erasing outcrop evidence of prior thrusting of the CBU. We propose that this nappe-bounding shear
zone accommodated the underplating of the central nappe around ∼49Ma while the southern nappe was still
subducting, and was subsequently reworked during exhumation.
8.4. Peak Subduction of the Fabrikas Nappe and Blueschist Facies Exhumation of the Central Nappe
The Fabrikas nappe (southern nappe) comprises Triassic meta-sedimentary schists, meta-volcanic schists, thinner
meta-carbonate horizons compared to the central nappe, and continental-affinity material of Posidonia (cf. Keiter
etal.,2011). This meta-sedimentary sequence was spatially associated with mafic igneous rocks with unknown
crystallization ages (Figure16). The primary difference between our southern slice and Laurent etal.(2016)'s
Posidonia subunit is that it contains the Fabrikas meta-mafic lens, which they placed in the structurally highest
Kampos subunit. Otherwise, our structural measurements and metamorphic observations are similar.
A Lu-Hf garnet-whole rock isochron from Katerghaki indicates that subduction of the Fabrikas nappe had already
begun by ∼48Ma (Uunk etal.,2022). The timing of peak subduction is constrained at ∼45Ma by garnet
core Sm-Nd crystallization ages (Gorce etal.,2021). This age is distinctly younger than peak subduction at
∼53Ma and ∼49Ma of the northern and central nappes, respectively. The subduction-to-exhumation transition,
or underplating event, of the southern nappe, is bracketed by peak subduction as recorded by garnet Sm-Nd ages,
blueschist facies metamorphism, and garnet rim growth during decompression at ∼40Ma (Gorce etal.,2021),
and blueschist facies retrogression recorded by Rb-Sr multi mineral isochrons between ∼42–39Ma (Skelton
Between ∼48–45Ma, rocks of the central nappe exhumed in the subduction channel under blueschist facies
conditions. Retrograde DT1 blueschist fabrics at Azolimnos, well-constrained at ∼45Ma (this study), overlap with
garnet Sm-Nd core ages at Fabrikas and therefore trend older than retrograde DT1 blueschist fabrics at Fabrikas,
which supports the separation of the central and southern tectonic slices. At this time, mafic blueschists and
eclogites and surrounding meta-sedimentary schists in the central nappe developed identical DT1−2 structures
(e.g., Vaporia and overlying meta-sedimentary rocks, and Kalamisia and Azolimnos; Figures6 and7). This indi-
cates that during DT1−2, mafic blueschists and eclogites, and surrounding meta-sedimentary rocks were exhumed
together, and in some places, the strain was partitioned between them. Therefore, even if mafic blueschists and
eclogites reached higher pressures on their prograde path, they must have been partially exhumed and juxtaposed
with CBU meta-sediments by ∼45Ma to explain concordant exhumation-related structures.
8.5. Exhumation of the Syros Nappe-Stack in the Subduction Channel Under Greenschist Facies
Conditions (Through ∼20Ma)
Between ∼44–20Ma, greenschist facies DT2 fabrics continuously developed throughout the accreted CBU stack,
younging systematically with structural depth, as each underplated nappe was exhumed in series from north
to south. Retrograde greenschist facies deformation-metamorphism occurred first in the structurally highest
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northern nappe and migrated structurally downward through time (cf. Glodny & Ring,2022; Ring etal.,2020;
Exhumation imparted penetrative deformation that progressively transposed older fabrics under blueschist facies
(DT1) and eventually greenschist facies (DT2) conditions. Upright folding that initiated under blueschist facies
conditions continued through greenschist facies conditions in the subduction channel. Kinematic rotation culmi-
nated in strongly E-W oriented fold axes and structures that formed under net compression, but were associated
with fold axis-parallel elongation and stretching (cf. Xypolias etal.,2012). Our new Rb-Sr isochron age from
Delfini provides precise constraints that exhumation was dominated by net compression under greenschist facies
conditions until at least ∼37Ma. Furthermore, exhumation-related DT1 and DT2 strains were dominantly coaxial
and well-distributed. This is evident from symmetric strain shadows on garnets, ductile pinching of partially
retrogressed eclogites at Agios Dimitrios, and outcrop-scale greenschist facies folds with sub-horizontal E-W
trending hinge lines with hinge-parallel symmetric boudinage of competent blueschist and epidote-rich lenses
(e.g., Delfini and Lotos; Figure8, Figure S4 in Supporting InformationS1).
The youngest dynamic DT2 greenschist facies fabrics associated with subduction channel exhumation are
∼25–20Ma and are recorded in the southern slice (Figure15). At this time, greenschist facies metamorphism
continued in the northern and central nappes, but was not associated with penetrative strain (e.g., random grains,
radiating clusters, decussate textures; Cliff etal.). These observations indicate strain progressively local-
ized toward the base of the stack through time. Patchy, static metamorphism in the northern and central nappes
may reflect local fluid availability as deformation migrated structurally downward.
8.6. Upper Plate Extension and Core Complex Capture
Slab rollback accelerated by ∼23–21Ma, which is constrained by dating of detachment faults and supra-detachment
sedimentary basins that developed in response to initial upper crustal extension (Gessner etal.,2013; Ring
etal.,2010). Rollback led to core complex capture, the southward migration of the volcanic arc through the
former forearc (e.g., the Tinos granite, 14.6±0.2Ma, Bolhar etal.), and continuous supradetachment
basin development through the late Miocene (e.g., Paros, ∼14–7Ma; Bargnesi etal.). CBU rocks were
exhumed in the footwall of the North and West Cycladic Detachment Systems and related smaller-scale structures
(Brichau etal.,2007; Grasemann etal.,2012; Jolivet etal.,2010; Soukis & Stockli,2013). On Syros, the Vari
Detachment operated as a semi-brittle to brittle extensional structure and accommodated late stages of exhuma-
Soukis and Stockli (2013) presented low-temperature zircon and apatite (U-Th)/He thermochronology and
concluded that the southern Syros CBU was juxtaposed with two structurally higher upper-plate units, the Upper
Unit (intermediate structural level) and Vari Unit (structurally highest), along at least two semi-brittle detachment
faults. While the Tinos Detachment exhumed CBU rocks between ∼22–19 on what would become neighboring
Tinos Island, low-angle normal faults juxtaposed the Vari and Upper Units on Syros. Exhumation of the Vari and
Upper Units at ∼13–15Ma was roughly coeval with magmatism on Tinos but the Syros CBU was exhumed later,
∼8–10Ma, beneath the Vari Detachment (Soukis & Stockli,2013).
Previous meso- and microscale observations demonstrate those kinematics along the Vari Detachment are
top-to-the-ENE (Soukis & Stockli,2013), consistent with our structural observations. High-angle normal faults
dip both SW and NE, indicating overall NE-SW extension (Soukis & Stockli,2013). The final exhumation of the
CBU on Syros occurred in multiple, temporally distinct, rapid episodes of unroofing. Exhumation beneath the
Vari Detachment was rapid, but only accommodated the final ∼6–9km of vertical exhumation (Ring etal.,2003).
9.1. Comparisons With Previous Tectonic Models
The tectonic model described above has several similarities and differences compared to previous models. First,
our results agree with previous studies suggesting that Syros is composed of distinct tectonic slices that reached
peak conditions at different times (Glodny & Ring,2022; Laurent et al.,2017; Lister & Forster,2016; Ring
etal.,2020; Skelton etal.,2019; Uunk etal.,2018,2022). Most of these studies propose two distinct slices in the
north and south, but our data suggest the likely presence of an additional third slice in between.
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The timing of peak subduction of Kampos/Kini versus Fabrikas compared to early retrograde blueschist defor-
mation at Azolimnos constrains the number of tectonic slices. Our new Rb-Sr isochron from Kini demonstrates
the DS fabric developed at 53.5±0.6Ma. This overlaps with independent constraints on the timing of garnet and
zircon metamorphic rim growth at Grizzas and Kini. Together these ages bracket the timing of peak subduction
of the structurally highest slice. This is statistically distinguishable from the peak subduction in Fabrikas rocks
at ∼45.3±1.0Ma (Gorce etal.,2021). Thus, to a first degree, these data lend further support to the presence of
two distinct structural slices in the north and south (Glodny & Ring,2022; Ring etal.,2020; Uunk etal.,2018).
Skelton etal.(2019) interpreted their Rb-Sr isochrons as records of peak subduction at ∼40Ma at Fabrikas, which
is younger than garnet core crystallization ages inferred to record peak metamorphism at ∼45Ma in the same
outcrop.As discussed above, the dominant outcrop-scale structures, metamorphic mineral chemistry, and mineral
replacement textures at Fabrikas are consistent with retrograde blueschist facies deformation. However, if Fabri-
kas did reach peak conditions at ∼40Ma, this supports the inference of a central slice above Fabrikas, since recent
garnet Lu-Hf and our new Rb-Sr isochron from Azolimnos indicate that rocks occupying a higher structural level
above Fabrikas experienced peak subduction at ∼49Ma (Uunk etal.,2022) and blueschist facies retrogression at
45.5±0.3Ma (this study). Our Azolimnos DT1 Rb-Sr isochron is statistically distinguishable (i.e., 2σ errors do
not overlap) from Fabrikas DT1 inferred from retrograde garnet growth (∼40.5±1.9Ma; Gorce etal.) and
blueschist fabric development (∼41.4±0.5, 41.6±1.5, and 39.6±1.2Ma; Skelton etal.). Therefore, if
the garnet core crystallization ages presented by Gorce etal.(2021) are accurate records of peak subduction of
the southern Fabrikas nappe, this also supports the interpretation of an intermediate slice, which was exhuming
at the same time as the Fabrikas nappe reached peak conditions.
We argue that mafic blueschists and eclogites do not exclusively occupy the structurally highest tectonic slice, in
contrast to Laurent etal.(2016) and Trotet, Jolivet, and Vidal(2001). Rather, our interpretation is that protoliths
for mafic blueschists and eclogites were present throughout the CBU before subduction and therefore appear
to record primary relationships (cf. Keiter etal.,2011). This implies that the mafic blueschists and eclogites at
Vaporia, Kalamisia, and Fabrikas are not separated from surrounding schists and marbles by shear zones and/or
detachments, as shown for the ‘Kampos subunit’ of Laurent etal.(2016). The primary observations supporting
that Fabrikas cannot belong to the same subducting unit as Kampos are that Fabrikas meta-volcanic record peak
metamorphism that is distinctly younger than that of Kampos and Kini, Fabrikas crops out toward the southern
end (i.e., bottom) of the dominantly north-northeast-dipping structural pile, and Fabrikas meta-volcanic are asso-
ciated with more meta-carbonate and meta-clastic sedimentary lithologies than Kampos and Kini suggesting they
represent subduction of different protolith assemblages. Moreover, the fact that Fabrikas occupies the immedi-
ate footwall of the Vari Detachment does not necessarily imply that it belongs to the structurally highest unit.
Even though the Vari Detachment has been interpreted as the paleo-subduction channel roof, continuous ductile
extension along a shallowly to moderately dipping structure throughout the Eo(?)-Oligocene, in addition to the
proposed 6–9km of semi-brittle exhumation accommodated by ∼20km of localized slip in the Miocene (Ring
etal.,2003), can easily explain tectonic removal of the uppermost units. (U-Th)/He ages and strong cataclastic
reworking at the base of the Vari Unit further attest to this (Soukis & Stockli,2013). These processes would
juxtapose structurally deeper CBU units with the Upper Unit in the hanging wall.
Our observations indicate that prograde textures are locally preserved in mafic blueschists and eclogites (cf.
Keiter etal.,2004), but the majority of the Syros CBU has been overprinted during subduction channel exhu-
mation (cf. Bond etal.,2007; Rosenbaum etal.,2002; Trotet, Vidal, & Jolivet,2001). Heterogeneous rock types
that occupy a given nappe were subducted and exhumed together, and therefore experienced identical P-T paths
(in contrast to Trotet, Vidal, & Jolivet; Trotet, Jolivet, & Vidal). Therefore, differences in strain,
metamorphic mineral assemblages, and/or preserved kinematics between mafic blueschists and eclogites and
meta-sedimentary rocks can be attributed to relative strengths, bulk composition, and fluid availability (and
composition) during metamorphism (see Schmädicke & Will for a similar discussion of P-T paths and
retrogression of the CBU on Sifnos).
9.2. Subduction Kinematics and Exhumation Processes
Kinematics of prograde subduction and the subduction-to-exhumation transition in the CBU are debated but have
direct implications for the geometry and mechanics of the subduction channel shear zone. Here, we summarize
the shear sense implications of our preferred tectonic and kinematic model, and then we discuss alternative