From Mesoproterozoic magmatism to collisional Cretaceous anatexis: Tectono-magmatic history of the...

From Mesoproterozoic magmatism to collisionalCretaceous anatexis: Tectonomagmatic historyof the Pelagonian Zone, GreeceFilippo Luca Schenker1,2, Jean-Pierre Burg1, Dimitrios Kostopoulos3, Evangelos Moulas2,Alexander Larionov4, and Albrecht von Quadt1

1Department of Earth Sciences, ETH Zurich, Zurich, Switzerland, 2Institute of Earth Sciences, University of Lausanne, Lausanne,Switzerland, 3School of Science, Faculty of Geology and Geoenvironment, National and Kapodistrian University of Athens,Athens, Greece, 4Centre of Isotopic Research, All-Russian Geological Research Institute (VSEGEI), St. Petersburg, Russia

Abstract The magmatic history of the Pelagonian Zone, in northern Greece, is constrained with secondaryion mass spectrometer (SIMS) U-Pb dating on zircons of various granitoids whose structural positions weredefined with respect to the regional main foliation. Ages pertain to four groups: (i) Mesoproterozoic(circa 1430Ma) crystallization of granites inferred from inheritedmagmatic zircon cores that have been partiallymolten during the (ii) Neoproterozoic at circa 685Ma (metamorphic zircon rims) and subsequently intrudedby a Neoproterozoic leucogranite (circa 600Ma). (iii) Late- or post-Variscan calc-alkaline granitoids (315–301Ma)were in turn intruded by a subvolcanic dike at about 280Ma. In the Early Permian the εNd(t) in magmasdecreased from �7.3 to �1.3, hinting to mantle-derived melts produced during extension. Rifting is furtherheralded by two acidic and one mafic dike containing Lower-Middle Triassic zircons (246–242Ma). (iv) EarlyCretaceous anatectic melts at 117±8Ma formed during regional metamorphism. This age is the first report ofin situ anatexis in the Pelagonian Zone. Cretaceous anatexis developed during the Mesozoic collision ofPelagonia with the Eurasian margin. Major- and trace-element geochemistry of amphibolites further attestsfor the complex pre-Alpine tectonic history with Neoproterozoic calc-alkaline and back-arc geochemicalsignature and Triassic alkali-magmatism.

1. Introduction

The Pelagonian Zone of Greece (Figure 1) is a Permo-Carboniferous continental margin [Yarwood and Aftalion, 1976;Vavassis et al., 2000] built on Neoproterozoic and older continental basement that was separated from southernEurasia during the breakoff of Pangea [Anders et al., 2007] and subsequently reunited to Eurasia during Alpineaccretion [e.g., Ricou et al., 1998]. Recent geochronological work documented a polymagmatic history withPrecambrian [Anders et al., 2006; Zlatkin et al., 2014], Permo-Carboniferous [Yarwood and Aftalion, 1976; Vavassis et al.,2000; Anders et al., 2007], Triassic [Pe-Piper, 1998], and Early Cretaceous [Anders et al., 2007] magmatic events.However, detailed field relationships between the successive magmatic events were not specified. Documentingthe relationships betweenmagmatic events and deformation is fundamental in unraveling the tectonic setting intowhich the magmas were emplaced. The Cretaceous event, in particular, is inconsistent with current plate tectonicsinterpretations, which usually place Pelagonia in the footwall of the Tethyan suture zone, far away from the activeorogenic trench at that time (Figure 2) [Ricou et al., 1998; van Hinsbergen et al., 2005; Sharp and Robertson, 2006].According to these interpretations, Cretaceous magmas can neither source from arc magmatism nor from collisionof Pelagonia with Eurasia. In the Early Cretaceous reconstruction of Papanikolaou [1989, 2009], Pelagonia has beentectonically roofed from the northeast by a nonmetamorphic to low-metamorphic passivemargin (the Axios/VardarZone, Figure 2). Nevertheless, in this case also, Pelagonia in the footwall of the suture and the low-thermal conditionsof the accretion are inconsistent with arc magmatism or collisional anatexis. To integrate all magmatic ages intoa tectonic framework, we performed a systematic investigation by dating zircons of granitoids from differentstructural positions and in different textural relations with respect to themain, regional foliation of Pelagonian rocks.Zircon U-Pb dating was performed with sensitive high-resolution ion microprobe (SHRIMP) II. The ages obtainedspan fromMesoproterozoic to Cretaceous. Three fabric-forming events, now showing foliations subparallel to eachother, were distinguished: (i) a Neoproterozoic synmelt foliation, (ii) a subsolidus foliation in late- to post-Variscanmetagranitoids, and (iii) an Early Cretaceous synmelt to postmelt foliation. The Cretaceous in situ anatexis is linked tocollision of Pelagonia with the Eurasian margin along the main Tethyan suture now located east of the Vardar.

SCHENKER ET AL. ©2014. American Geophysical Union. All Rights Reserved. 1

PUBLICATIONSTectonics

RESEARCH ARTICLE10.1002/2014TC003563

Key Points:• Wedescribe a newPrecambrian terranein the Pelagonian Zone

• First dating of Cretaceous collisionalmigmatites in the Pelagonian Zone

• Early Permian asthenospheric upwellingduring Gondwana breakoff

Supporting Information:• Readme• Figure S1• Figure S2• Figure S3• Figure S4• Figure S5• Figure S6• Figure S7• Figure S8• Figure S9• Figure S10• Figure S11• Table S1• Table S2• Table S3• Table S4• Table S5• Table S6• Table S7• Table S8• Table S9• Table S10• Table S11• Table S12• Table S13• Table S14• Table S15• Table S16

Correspondence to:F. L. Schenker,[email protected]

Citation:Schenker, F. L., J.-P. Burg, D. Kostopoulos,E. Moulas, A. Larionov, and A. von Quadt(2014), From Mesoproterozoic magma-tism to collisional Cretaceous anatexis:Tectonomagmatic history of thePelagonian Zone, Greece, Tectonics, 33,doi:10.1002/2014TC003563.

Received 24 FEB 2014Accepted 9 JUL 2014Accepted article online 17 JUL 2014

http://publications.agu.org/journals/

http://onlinelibrary.wiley.com/journal/10.1002/(ISSN)1944-9194

http://dx.doi.org/10.1002/2014TC003563

http://dx.doi.org/10.1002/2014TC003563

Figure 1. (a) Map [Burg, 2012] and (b) synthetic cross section through the Balkan region. Moho below the Hellenides from Sachpazi et al. [2007], below the Vardar/Axios andsouthern Rhodope from Papazachos [1998], and below the Rhodope from Boykova [1999].

Tectonics 10.1002/2014TC003563


1.1. Tectonic Setting

In the Dinaride-Albanide-Hellenide mountain chain (Figure 1a), convergence between Europe and Africa ledto the piling of various tectonic units such as the Pelagonian Zone in the accretionary wedge of the EarlyJurassic-Paleocene Tethyan subduction zone [e.g., Ricou et al., 1998]. The stacked crustal and oceanic pileswere exhumed in metamorphic domes (Figure 1b) during protracted extensional tectonics between theEocene and late Neogene [e.g., Lister et al., 1984; Jolivet and Brun, 2010; Burg, 2012]. The Pelagonian continentalblock was involved in four Mesozoic-early Cenozoic tectonic events [Robertson and Dixon, 1984; Dercourtet al., 1986]: (i) Late Jurassic-Early Cretaceous, southwestward obduction of ophiolitic segments from theVardar passive margin (a branch of Neotethys) onto Pelagonia [Bernoulli and Laubscher, 1972; Mercier et al.,1975]; (ii) Early Cretaceous tectonometamorphic event whose significance is disputed [e.g.,Yarwood and Dixon,1977; Schermer et al., 1990]; (iii) post-Cretaceous southwestward thrusting of Pelagonia which evolved into theEocene fold-and-thrust accretionary system of the external Hellenides [Renz and Reichel, 1945; Jacobshagenet al., 1978]; and (iv) Miocene extension [Schermer et al., 1990] attributed to the ongoing rollback of the Hellenicslab [e.g., Lepichon and Angelier, 1979; Brun and Faccenna, 2008].

1.2. Geological Overview

The Pelagonian Zone in continental Greece is thewesternmost unit of the Internal Hellenides (Figure 1). The nappecomplex of northern Pelagonia is composed, (according to Kilias and Mountrakis [1989]) from bottom to top of(i) ancient crystalline basement comprising orthogneisses, amphibolites, and metapelites; (ii) Permo-Τriassicvolcano-sedimentary rocks and Triassic-Jurassic carbonates, nonmetamorphosed to weakly metamorphosed atthe western side and higher grade to the east; (iii) erratic ophiolites, deformed in Late Jurassic-Early Cretaceoustimes; and (iv) transgressive Cretaceous limestones that pass upward into Paleocene turbidites.

In the study area, mafic dikes intruded Paleozoic and Proterozoic orthogneisses [Anders et al., 2007]. Thebasement gneisses are topped by fragmented marbles considered to be Paleozoic to Mesozoic according tostratigraphic correlations with northwestern Pelagonia [Mercier, 1968;Mountrakis, 1986; Sharp and Robertson,2006]. Gneisses and marbles are duplicated, indicating intra-Pelagonian imbrication [Kilias et al., 2010].Pliocene to recent sediments unconformably cover the Pelagonian basement [e.g., Godfriaux, 1968].

Structurally, the Pelagonian Zone is amostly flat-lying complex that bends downward toward its eastern tectoniccontacts, forming a large-scale antiform to the east (Figure 1b). The eastern flank is covered by stacked sequencesof Cretaceous platform carbonates and hemipelagic marls, serpentinized ultramafic and metavolcanic rocksascribed to the Axios/Vardar passive margin [Sharp and Robertson, 2006]. The western Pelagonia dips eastward atthe structurally lower contact (e.g., Parnassos, Figure 1b) and westward at the structurally upper contact with theophiolitic units. Two dominant regionalmetamorphic events were identified in the gneisses [Godfriaux et al., 1988]:

Figure 2. Early Cretaceous paleogeographic reconstructions through the Balkan region after the cited authors.

Tectonics 10.1002/2014TC003563


an Early Cretaceous to 98Ma amphibolite- to lower greenschist-facies metamorphism [Yarwood and Dixon,1977; Schermer, 1990; Lips et al., 1998], more prominent in the northern part, and a Paleocene-Eoceneblueschist-facies metamorphism in southern Pelagonia [Schermer, 1990; Schermer et al., 1990; Lips et al., 1998].

2. Sampling and Dating Methodology

Eleven samples (Figure 3) were selected to fully represent the range of granitoids andmetagranitoids found in thePelagonian Zone. Their main characteristics (structural context, mineral assemblage, texture, zircon morphology,and cathodoluminescence (CL) zoning) and chemical analytical results are summarized in Table 1 anddiscussed insection 3. Sampling aimed at identifying the timing between successive episodes of plutonism as specified fromstructural relationships, taking the regional main foliation as reference. In particular, we were interested in thetiming between the formation ofmigmatites (synfoliation, almost in situ leucosomes), prefoliation and postfoliationplutons and veins. Granitoids having an internal foliation subparallel to the regional foliation were interpreted

Figure 3. Geological map of the eastern Pelagonia gneiss dome with location of samples analyzed in this study.

Tectonics 10.1002/2014TC003563


Table

1.Lo

catio

n,FieldAspects,M

ineralAssem

blag

es,Z

ircon

Morph

olog

y,an

dAge

sof

Selected

Samples

Rock

Type

Sample

Coo

rdinates

FieldCha

racteristic

MineralAssem

blag

eaZircon

sMorph

olog

yan

dCLZo

ning

Age

Core

Age

Rim

Gen

eses

Grano

dioritesan

dGrano

dioriticGneisses

09-055

N40

°19′11

.6″

E022

°05′53

8″coarse

graine

d,with

out

foliatio

n,with

maficen

claves

pl,q

tz,afs,b

t,zo,amph

,spn

,zrn,

opaq

ues,±chl,±gt

prismaticcrystals(lw

=1.8–

4),

oscillatory

zoning

,inh

erite

dcores

297–

613Ma

304–

308Ma

late-or

post-

Variscanmag

matism

09-121

N40

°13′37

.0″

E022

°06′34

.5″

foliated,

coarse

graine

dpl,afs,q

tz,zo,

amph

,bt,

spn,

wm,zrn,o

paqu

esprismaticcrystals(lw

=1.2–

3.3),

oscillatory,p

atchyan

dchao

ticcores,bright

rims

310–

317Ma

307–

320Ma

late-or

post-

Variscanmag

matism

MigmatiticOrtho

gneiss

P12-04

1N40

°21′58

.3″

E022

°06′43

.0″

coarse

graine

d,foliated,

alternation

ofmelan

ocratic

andleucocratic

band

sat

thecentim

eter

scale

qtz,pl,afs,b

t,wm,g

troun

ded-prismaticto

roun

ded

(lw=1.1–

2.1),h

omog

eneo

us,

sector,o

roscillatory

cores

with

grey

rim

1061

–161

2Ma

679–

1424

Ma

Precam

brian

anatexiseven

t

Granitic

Vein

10-050

N40

°21′56

.3″

E022

°06′39

.0″

2–10

cmvein

inmigmatitic

orthog

neiss

qtz,afs,pl,w

m,b

t,zo,zrn,o

paqu

esfrag

men

tedan

droun

ded-

prismatic(lw

=1.8–

3.3),sector,

andoscillatory

zoning

559–

1977

Ma

318–

618Ma

clastic

grains

with

Paleozoic-

Proterozoiccores

Leucog

ranitesan

dLeucog

neiss

09-122

N40

°13′37

.0″

E022

°06′34

.5″

light

color,med

ium

graine

d,foliated

qtz,pl,afs,w

m,amph

,zo,zrn,o

paqu

esprismatican

dxeno

morph

ic(lw

=2.3–

3.2),o

scillatoryor

patchy

,som

ebright

rims

304–

605Ma

309–

310Ma

late-or

post-

Variscanmag

matism

P12-02

6N40

°19′10

.8″

E022

°06′24

.7″

light

color,fine

graine

d,crosscuttin

gthegran

odiorite

(09-05

5),n

otfoliated

pl,q

tz,afs,b

t,zo

roun

ded-prismaticto

xeno

morph

(lw=1.2–

3.5),o

scillatory

zoning

with

rims

303–

570Ma

298–

322Ma

late-or

post-

Variscanmag

matism

P12-04

2N40

°21′56

.0″

E022

°06′42

.6″

light

color,fine

graine

d,crosscuttin

gthemigmatiticorthog

neiss

qtz,pl,afs,w

m,g

teu

hedral-prismaticto

roun

ded

(lw=1.1–

3.2)

1294

–191

1Ma

591–

952Ma

Flurinamag

matism

And

esiticSubvolcanicDike

10-079

N40

°22′45

.1″

E022

°07′10

.1″

slightlygreen,

holocrystalin

e-po

rphy

ric,crosscuttingthe

gran

odioriticorthog

neiss

pl,q

tz,amph

,bt,zo,al,

spn,

wm,zrn,ap

from

roun

ded-ellip

ticto

roun

ded-


=1.4–

2.2),

blurred,

sector,o

scillatory,

andconv

olutezoning

279–

1022

Ma

253–

307Ma

mostly

detrita

lgrains,

late-Variscanintrusion

Aug

engn

eissWith

inAmph

ibolites

P12-01

2N40

°24′05

.9″

E022

°10′31

.3″

coarse

graine

d,foliated,

intrud

ingtheam

phibolite

spl,afs,zo,

bt,w

m,

spn,

zrn,

chl

tabu

lar-prismaticcrystal

(lw=1.6–

2.6),o

scillatoryzoning

245Ma

Triassicintrusion

even

t

Leucosom

esWith

inAmph

ibolites

10-005

N40

°22′23

.2″

E022

°05′14

.6″

med

ium

graine

d,leucocratic

netw

orkin

amph

ibolite

scrosscuttin

gthemainfoliatio

n

pl,q

tz,afs,b

t,zo,

cpx,am

ph,zrn


=1.2–

2),

sector

zoning

with

bright

rims,inhe

rited

cores

255–

237Ma

130–

243Ma

intrusionat

130Mathat

remob

ilizedTriassic

andolde

rzircon

s10

-132

N40

°21′27

.1″

E022

°05′28

.5″

med

ium-grained

,leu

cocratic

patche

sin

amph

ibolite

scrosscuttin

gthemainfoliatio

n

pl,afs,amph

from

roun

ded-ellip

ticto

roun

ded-


=1.1–

3.5),

blurred,

sector,o

scillatory,

andconv

olutezoning

130–

1573

Ma

97–3

01Ma

anatexiseven

tat

116Ma

that

remob

ilizedTriassic

andolde

rzircon

s

a Mineralab

breviatio

ns:afs=alkalifeldspar,aln=allanite,amp=am

phibole,ap

=ap

atite,bt=

biotite,chl=chlorite,cpx=clinop

yroxen

e,ep

=ep

idote,grt=

garnet,hbl=ho

rnblen

de,m

s=muscovite,

pl=plag

ioclase,qz

=qu

artz,spn

=sphe

ne,w

m=white

mica,zo

=zoisite

,and

zrn=zircon

.Mineralab

breviatio

nsafterWhitney

andEvan

s[2010].

Tectonics 10.1002/2014TC003563


as prefoliation magmas. Granitoids cutting the regional foliation were defined as postfoliation intrusions.Leucosomeswrapped in the regional foliation and, in places, discordant to it, were defined as synfoliationmelts.

Zircons were extracted with acoustic shock waves produced in the Selfrag apparatus, ETH Zürich. Standardheavy-liquid and magnetic separation methods delivered mineral fractions enriched in zircons which werehandpicked under a binocular microscope. These zircons were mounted in U- and Pb-free resin and polished

Figure 4. Field pictures of sample outcrops. GPS coordinates and rock descriptions in Table 1. (a) Migmatitic orthogneissP12-041. (b) Leucogranite dike P12-042 cutting the foliation (S) of country migmatites (same locality as P12-041). (c) Graniticvein cutting orthogneiss and migmatites (same locality as P12-041). (d) Typical outcrop where crosscutting relationshipsdefine successive magmatic events in chronological order: (i) granodiorite, (ii) leucogranite, and (iii) mafic dike. (e) Andesiticdike (10-079) cut by a mafic dike, both intruding granodioritic gneiss and leucogneiss. (f) Mylonitic augengneiss (P12-012)in amphibolites with a top-to-the SW sense of shear (arrow). (g) Network of leucocratic veins (10-005) in foliated amphibolite.(h) leucosome vein (10-132) crosscutting the foliation (S) in amphibolite.

Tectonics 10.1002/2014TC003563


to expose the internal parts of the minerals. Suitable zircon domains identified in CL images were then analyzedby SHRIMP II. Zircon characteristics were interpreted following Corfu et al. [2003]. Concentric oscillatory zoningwas interpreted as magmatic growth irrespective of position in the zircon (core or rim). Metamorphic zirconswere essentially recognized by their rounded or ovoid shape and by a bright or dark rim overgrowth, whereasmagmatic zircons were recognized from their euhedral shape with sharp edges and high aspect ratios. Tableswith isotopic ratios and CL images of the analyzed zircons are given in the supporting information data.

In situ U-Pb analyses on zircons were performed with the SHRIMP II of the Center of Isotopic Research at VSEGEI,Saint Petersburg, Russia. We applied a secondary electron multiplier in peak-jumping mode according to theprocedure ofWilliams [1998] and Larionov et al. [2004]. The primaryO2

� beamwas approximately 27 × 20μm in size.The results have been postprocessed with the SQUID v1.12 [Ludwig, 2005b] and ISOPLOT/Ex 3.22 [Ludwig, 2005a]software using the decay constants of Steiger and Jager [1977]. The ages were presented in Concordia [Wetherill,1956] and Tera-Wasserburg [Tera and Wasserburg, 1972] diagrams. Uncertainties given for individual analyses(ratios and ages) are at the 1σ level; the uncertainties in calculated Concordia ages are reported at the 2σ level.

3. Structural Position, Sample Description, Zircon Cathodoluminescence (CL)Images, and U-Pb Ages

Migmatitic orthogneiss (P12-041) was found in the core of the antiform, in the deep levels of the exposedPelagonian gneisses (Figure 3). The rock displays millimetric to decimetric leucocratic veins (neosome) borderedby biotite-, amphibole-, and garnet-rich zones (restite) in amatrix with gneissic texture (paleosome ormesosome)(Figure 4a). The leucosomes are generally contained in the main foliation defined by biotite and amphibole,

Figure 5. Cathodoluminescence images of representative zircon grains. (a) Heterogeneous zircon population in migmatiteP12-041 with metamorphic rims. (b) Heterogeneous population of zircons in leucogranite P12-042. Youngest cluster (n=4)of concordant ages at 599 ± 5Ma. (c) Zircons from sample 10-050 showing various ages without any correlation with thezoning pattern. (d) Zircons from granodioritic gneiss 09-121 with concordant ages at 315.1 ± 1.4Ma. (e) Zircons fromleucogneiss 09-122 with well-developed oscillatory zoning occasionally overgrown by a bright rim (concordant age at308 ± 2Ma). (f ) Zircons from granodiorite 09-055 showing oscillatory zoning with Concordia ages of 304± 1.5Ma. Age ofinherited core (Spot 09-055_8.1) at 613Ma.

Tectonics 10.1002/2014TC003563


except for some coarse-grained veins that locally cut this fabric. Such leucosomes are interpreted asautochthonous/para-autochthonous melts. The analyzed sample is constituted of qz-pl-afs-bt-wm-grt(mineral abbreviations in bottom of Table 1); it included both leucocratic and melanocratic domains, whichwere too thin to be physically separated. The leucocratic domains contain quartz, alkali feldspar ± garnet.

The zircons are rounded-prismatic to rounded showing length-to-width ratios (l/w ratios) of 1.1–2.1. The CLimages (Figures 5a) show zircon cores with homogeneous, sector, or oscillatory zoning truncated by greyrims. The contact between rims and cores is bulbous and serrated. Three oscillatory domains yielded agesbetween 1429 and 1467Ma, one sector domain was dated at 1061Ma, whereas one dark core gave an ageof 1189Ma (Figures 7a). The rim ages vary between 679 and 1424Ma. Thorium/uranium (Th/U) ratios(Figure 9a) of cores range between 0.22 and 1.38, which is typical of magmatic zircons, and Th/U ratios ofrims, between 0.01 and 0.12, are typical of metamorphic zircons [Teipel et al., 2004]. Two concordantpopulations were calculated from these rims; one at 723 ± 3.1Ma, the other at 683 ± 4Ma. The migmatiticevent is related to the growth of the 683± 4Ma, low Th/U ratio metamorphic rims. The age of the magmaticprotolith is tentatively interpreted as that of two magmatic inherited cores (e.g., spot analysis P12-041_11.2,Figure 5a) and one rim with well-developed oscillatory zoning concordant at 1432.1 ± 5.7 (Figure 7a).

Leucogranite (P12-042) is a 30 cm thick dike discordantly intruding the migmatitic foliation of sample P12-041(Figure 4b). The weak internal granulometric layering is parallel to the dike and inclined at 30° with respect tothemigmatite foliation. Sample P12-042 has a mineral assemblage of qz-pl-afs-wm±grt (in order of decreasing

Figure 6. Cathodoluminescence images of representative zircon grains. (a) Homogeneous population with oscillatory zoningfrom leucogranite (P12-026) with concordant age at 301 ± 2Ma. (b) Heterogeneous zircon population from the andesiticsubvolcanic dike (10-079). (c) Oscillatory zoning in zircons from augengneiss P12-012 with concordant age at 245 ± 1.3Ma.(d) Zircons from leucocratic vein 10-005 in amphibolite showing cores at 246 ± 2Ma and bright rims at 130 ± 4.4Ma. (e)Zircons from leucosome 10-132 (Figure 4h) with a metamorphic rim at 116.4 ± 7.6Ma.

Tectonics 10.1002/2014TC003563


abundance) imparting to the rock its light color (Figure 4b). The texture is fine grained, with oriented whitemicas defining a weakly developed foliation.

The zircon morphology is heterogeneous, varying from euhedral-prismatic to rounded prismatic or rounded.CL images show zirconswith different internal textures. Most zircons have oscillatory, sector, or homogeneous corestruncated by an either dark or grey oscillatory rim (Figure 5b). A few grains show a second, thin (5μm thick), brightrim. The analyses reveal Mesoproterozoic cores (1294–1911Ma) and Neoproterozoic dark (598–697Ma) and grey(591–952Ma) rims (Figure 7b). Only one of the thin light rims could be dated; its age is 473±3Ma. TheTh/U ratios ofdark, grey, and bright rims are consistently lower than those of the cores (Figure 9b), hinting to a metamorphicrather than intrusive event. Nevertheless, two grey and one dark rim analyses are concordant at 697±8Ma(Figure 7c), nearly coeval with the 683±4Ma age interpreted as melting of the country P12-041 migmatite. We,therefore, deduce that the 697±8Ma population is exotic and was derived from the country rock. The age of theleucogranite intrusion is inferred from the four concordant analyses (two dark and two grey rims) to be 599±5Ma.

Granitic vein (10-050) is 5–20 cm thick and cuts the foliation of both a gneissic dioritic dike andmigmatite P12-041. Rare white micas and biotite lie within the weakly developed foliation in this vein. In the migmatite, thevein branches into thinner veins both discordant to and concordant with the foliation of the host (Figure 4c).Where discordant, these veins are folded in open asymmetric folds with axial planes subparallel to the mainfoliation, suggesting that the vein was deformed after intrusion. This light-colored, middle-grained, andslightly foliated sample is constituted, in decreasing abundance, of qz-afs-pl-wm-bt-zo-zrc and opaques.

The zircons are euhedral, prismatic to anhedral, and prismatic with rounded edges (Figure 5c). The l/w ratio is1.8–3.3. Cathodoluminescence images show zircons with a euhedral core overgrown by a 10–20μm thick,dark, and occasionally oscillatory rim. The rim shows only one dark zone in all zircons, whereas the coresexhibit either oscillatory, sector, dark homogeneous, or chaotic zoning patterns (Figure 5c). Core and rim agesspan a wide spectrum (Figure 7d): 18 analyses identify three Proterozoic cores (1977Ma, 1705Ma, and1147Ma), a Late Proterozoic cluster (concordant ages at 580 ± 10Ma, n=11) of three cores and eight rims,one Devonian rim (400 ± 8Ma), and three Carboniferous rims (concordant at 326 ± 7Ma). The heterogeneityin zoning, brightness in CL images, and ages (Figure 5c and 7d) suggests that these zircons were taken fromdifferent magmatic sources. The youngest concordant age cluster at 326 ± 7Ma is interpreted as the intrusionage of the vein, but an alternative explanation with intrusion at 580Ma followed by two metamorphisms at400Ma and at 326Ma cannot be ruled out. We prefer the first option because the 326Ma event is notrecorded in the zircons of the host migmatite (sample P12-041).

Granodioritic gneiss (09-121) was collected from the western limb of the antiform (Figure 3). Oriented crystals ofbiotite, white mica, and amphibole define a foliation subparallel to the local trend and plunging gently to theSW. Asymmetric, xenomorphic feldspar augens (1–3mm in diameter) indicate a top-to-the SW sense of shear.No outcrop of the contact with the country rock was found. The sample contains, in decreasing abundance, pl-afs-qz-zo-amph-bt-spn-wm-zrc and opaques. It is coarse grained and has a protomylonitic texture.

The CL images show prismatic zircons with slightly rounded edges (Figure 5d). The shape ratio is variable(1.2< l/w< 3.3). Four grains show rather well-developed magmatic oscillatory zoning. Nine grains show acore rimmed by a 1 to 10μm thick, bright band. The cores show oscillatory, patchy, or chaotic zoning.Independent of zircon domain, the 15 spot analyses yielded a concordant age at 315.1 ± 1.4Ma (LateCarboniferous, Figure 7e). The Th/U ratios of all measurements vary between 0.28 and 1.34 withoutcorrelation with age. These Th/U ratios are in the range of magmatic zircons [Teipel et al., 2004]. ThemagmaticTh/U ratios and the overlap in core and rim ages suggest intrusion at 315.1 ± 1.4Ma.

Leucogneiss (09-122) was collected 20m to the west of sample 09-121; the contact with the neighboringgranodioritic gneiss is not exposed. Plagioclase-qz-afs-wm-amph-zo-zrc and opaques form the mineralassemblage of this light-colored, medium-grained, and foliated sample. The foliation is defined by fine-grained,elongated quartz and plagioclase crystals, and whitemicas. Grain-size gradients of quartz-plagioclase aggregatesdefine a protomylonitic texture. Xenomorphic alkali feldspar (1mm in diameter) and aggregates of zoisite makeasymmetric augen. The foliation is subparallel to the foliation of sample 09-121 and to the regional trend.

Sample 09-122 contains both euhedral-prismatic and xenomorphic zircons (l/w ratios =2.3–3.2).Cathodoluminescence zoning revealed two zircon types (Figure 5e): (i) zircons with well-developed magmaticoscillatory zoning and (ii) zircons with a bright or oscillatory rim truncating a core with either oscillatory or patchy

Tectonics 10.1002/2014TC003563


Figure 7. Ahrens-Wetherill and Tera-Wasserburg Concordia plots and weightedmean age diagrams from SHRIMP II results.Ages with 95% confidence interval errors. Ellipses plotted with a 2σ error. Solid ellipses = concordant ages; dashedellipses = either mean age calculations with outliers omitted or discordant ages. Different age populations shown in grey.Mean square weighted deviation (MSWD)= sum of squares of weighted residuals divided by the degrees of freedom foreach calculation.

Tectonics 10.1002/2014TC003563


zoning. Spot analyses from type (i) zircons yielded three spots Late Carboniferous (304, 308, and 309Ma), onespot Devonian (383Ma) and two spots Proterozoic (554 and 605Ma) ages. In type (ii) zircons, four bright andoscillatory rims yielded Late Carboniferous ages (309–310Ma), whereas five cores yielded ages varying between440 and 605Ma. The Th/U ratios of all measurements vary between 0.03 and 1.12 without correlation with age.The seven youngest results (three oscillatory cores and four rims) give a concordant age at 308±2Ma (Figure 7f).The rims cannot be related to metamorphism because zircons of granodioritic gneiss 09-121, which crops outsome 10m away and has a magmatic age of 315±1.4Ma, do not have rims at circa 309–310Ma. The spread ofmostly concordant ages in different grains (Figure 7e) suggests an inherited nature of most dated zircons.

Granodiorite (09-055) was sampled on the western flank of the regional antiform (Figure 3). The granodioriticrocks show gradients in foliation intensity at the hectometer scale from nondeformed granodiorites tomylonites along top to the SW shear zones parallel to the regional trend. The foliation is defined by biotite,amphibole, and zoisite. The gneiss contains approximately 10 cm wide mafic lenses. The collected samplecontains, in order of decreasing abundance, pl-qz-afs-bt-zo-amph-spn-zrc-opaques ± chl ± grt and belongs toa nondeformed, slightly foliated, and coarse-grained granodiorite mass.

The CL images show a morphologically homogeneous zircon population (Figure 5f). Crystals are prismatic,elongated (l/w ratios = 1.8 to 4), and euhedral with slightly rounded edges. The dominant population showsrather well-developed magmatic oscillatory zoning truncating a core with either homogeneous, sector,planar, or patchy zoning. Nine oscillatory zircon domains, two patchy cores, one banded core, and onehomogeneous domain were analyzed. Th/U ratios of all measurements vary between 0.31 and 0.96 with theexception of one homogeneous zircon domain yielding a Th/U ratio of 0.03. All single spot analyses on theoscillatory domains yielded ages between 297 and 311Ma (Figure 7g). One patchy core was dated at 613Ma.The homogeneous, low Th/U domain yielded 370Ma. Eleven spot analyses yielded concordant ages at304 ± 1.5Ma (Figure 7g), which is taken as the intrusion age.

Leucogranite (P12-026) is part of a dike network intruding into granodiorite 09-055 (like in Figure 4d). Thethickness of the veins varies from 0.01 to 5m. Like the country rock, the leucogranite developed a prominentfoliation of elongated plagioclase, quartz, alkali feldspar, and biotite along top-to-the SW shear zones. Thesampled rock is composed of pl-qz-afs-bt-zo and is a fine-grained, weakly- to nonfoliated rock.

The zircons are mostly rounded-prismatic (l/w ratios = 1.2–3.5) with the exception of one xenomorphic grain.The CL images show a dominant population with continuous magmatic oscillatory zoning from bright coresto dark rims (Figure 6a). Four zircons showed oscillatory or chaotic inherited cores. Eleven analyses yieldedLate Carboniferous ages regardless of zoning domain. Seven of these analyses are concordant at 301±2Ma(Figure 7h), which is interpreted as the crystallization age. Two inherited cores have a Proterozoic age (565 and570Ma), probably dating assimilated magmatic material. Th/U ratios vary between 0.13 and 1.77, in the range ofmagmatic zircons [Teipel et al., 2004]. Two outliers yielded lower Th/U values without correlation to ages.

Andesitic subvolcanic dike (10-079) was sampled in the core of the antiform, to the northeast of migmatiteP12-041 (Figure 3). The dike is oblique to the regional foliation (defined by biotite and white mica) in both adioritic orthogneiss and a leucogneiss; it has, in turn, been intruded by a fine-grained mafic dike displaced bytop-to-the SSW shear bands (Figure 4e). At the outcrop, the sampled green andesitic subvolcanic dike displays aholocrystalline-porphyric texture typical of subvolcanic intrusions. Microscopically, the analyzed specimenshows green amphibole phenocrysts (length 1–5mm) bearing lamellae of Ti-phases (rutile or sphene), whichare typical of exsolution during cooling and/or greenschist metamorphism [Ernst and Liu, 1998]. Aggregates ofzoisite are pseudomorphs after phenocrystic plagioclase. Small grains (<0.5mm) of plagioclase, quartz, biotite,and epidote (with locally allanitic cores) are the major constituent of the crystalline matrix.

The rock contains zircons that differ in size, morphology, and zoning pattern (Figure 6b). The l/w ratio variesfrom 1.4 to 2.2, reflecting the diversity of shapes from elongated-prismatic with rounded edges to anhedral,subrounded, and multifaceted. The analyzed and dated zoning patterns (Figure 6b) are enumerated inchronological order: (i) core-to-rim patchy zoning at 999–1022Ma, (ii) core-to-rim sector zoning at 612Ma (ina soccer ball zircon), (iii) unzoned or patchy core at 447Ma with unzoned dark rim at 271–285Ma (SpotAnalysis (SA) 10-079_1.2 and 1.3, Figure 6b), (iv) oscillatory core at 306Ma with bright rim at 307Ma (core andrim ages are identical within error), (v) laminar core with a 307Ma oscillatory rim, (vi) core-to-rimhomogeneous zoning at 300Ma with some brighter patches, (vii) chaotic bright core with an oscillatorybright rim at 286Ma (SA10-079_8.1, Figure 6b), and (viii) oscillatory core at 279Ma (SA 10-079_3.1, Figure 6b)

Tectonics 10.1002/2014TC003563


with a 253 ± 7Ma bright rim (SA 10-079_3.2, Figure 6b). The Th/U ratios vary between 0.02 and 0.92 withoutage correlation. The spread of concordant ages (Figure 8a) suggests the exotic origin of the dated zircons. Aprecise intrusion age cannot be inferred from this 700Ma range in ages, which is too large to reflectcontinuous magmatism. The largest age population is Permo-Carboniferous. In this population, four analysesare concordant at 305.3 ± 4.6Ma, four at 279 ± 8.1Ma, and one zircon rim at 253 ± 7Ma (Figure 8b). Thezircons at 305Ma are possibly related, within error, to the granodioritic intrusion (sample 09-055). We suggestthat the andesitic intrusion is dated by the youngest, statistically strongest 279 ± 8.1Ma population. Intrusionmay be younger if the single zircon rim at 253 ± 7Ma is considered. Notably, the low U concentration(61 ppm) precludes rejuvenation due to severe Pb loss because of metamictization. Field relations supportthe post-Carboniferous age since the sampled dike intruded granodioritic gneisses similar in appearance andin mineral assemblage to the 304Ma old granodiorite (sample 09-055).

Figure 8. Ahrens-Wetherill and Tera-Wasserburg Concordia plots andweightedmean age diagrams from SHRIMP II results. Ages with 95% confidence interval errors.Ellipses plotted with a 2σ error. Solid ellipses = concordant ages; dashed ellipses = eithermean age calculations with outliers omitted or discordant ages. Different agepopulations shown in grey. MSWD= sum of squares of weighted residuals divided by the degrees of freedom for each calculation.

Tectonics 10.1002/2014TC003563


Augengneiss (P12-012; Figure 3) was taken from an approximately 20 cm thick layer occurring in anamphibolite on the eastern flank of the antiform (Figure 4f). Both contacts and biotite-bearing foliation of theaugengneiss are parallel to the regional trend. Asymmetric porphyroclasts of quartz and feldspar in the gneiss andleucocratic lenses in the amphibolites show a top-to-the SW sense of shear (Figure 4f). The host amphibolite is alsointruded by a leucogranite preserving fine-grained, chilled margins. The mineral association of the analyzedaugengneiss is pl-afs-zo-bt-wm-spn-zr-chl (listed in decreasing order of abundance).

The zircons are prismatic, mostly with a tabular morphology (1.6< l/w< 2.6). The dominant populationshowed a continuous magmatic oscillatory zoning with a grey to bright core rimmed by a dark zone(Figure 6c). Only one inherited core was observed. All 11 spot analyses are concordant at 245 ± 1.3Ma(Figure 8c), reflecting the magmatic intrusion age. Thorium/uranium ratios range between 0.49 and 0.77without correlation with age, but consistent with a magmatic character.

Leucocratic vein (10-005) was collected from the axis of the antiform (Figure 3), tectonostratigraphically abovemigmatite P12-041. It is part of a network that cuts the amphibolite country rock foliation defined by the peakmetamorphic assemblage of hornblende and biotite (Figure 4g); the amphibolitic foliation and theleucocratic veins have been folded together, suggesting ductile deformation of both rocks, in agreementwith the inferred high temperature of metamorphism. The vein sample is constituted by afs-pl-qz-bt-zo-cpx.It is medium grained with plagioclase and quartz defining an interlobate texture. Alkali feldspar is associatedwith clinopyroxene at the contact between the leucocratic veins and the host amphibolite.

Zircons from sample 10-005 have a euhedral, prismatic morphology with a moderate elongation (1.2< l/w< 2).The CL images show homogeneous grains with weak oscillatory and sector zoning. Only 2 out of the 14 analyzedzircons have xenocrystic cores with darker oscillatory zoning (zircon 10-005_6, Figure 6d). In most zircons,a≤10μm bright rim is separated from the homogeneous core by a bulbous surface (zircon 10-005_11 and10-005_13, Figure 6d). Fifteen analyses of zircon cores yielded a concordant age cluster at 246±2Ma (Figure 8d)whatever the zoning pattern (e.g., oscillatory, sector, or xenocrystic core). Bright rims sufficiently thick (i.e.,> 10μm)to bemeasured yielded Tera-Wasserburg ages of 130±4.4Ma (Figure 8d). The zircon cores display highTh/U ratios(0.63–1.33), with ratios> 1 being typical of magmatic zircons frommafic rocks [Heaman et al., 1990]. The Th/U ratioin the bright rims varies between 0.07 and 0.24 with an outlier at 0.65 (Figure 9c). These low values fall in the

Figure 9. Th/U ratios of core and rim analyses of (a) migmatitic orthogneiss, (b) leucogranite truncating the migmatitic foliation, (c) leucocratic vein, and(d) leucocratic patches in amphibolites.

Tectonics 10.1002/2014TC003563


domain of metamorphic Th/U ratios of zircons [Teipel et al., 2004]. The leucocratic veins intruded the hostamphibolite during or after the metamorphic peak conditions. Accordingly, the 130Ma rim growth is linked to thecrystallization of the leucocratic melt. The 246Ma cores reflect the age of the source protolith.

Leucocratic pocket (sample 10-132) occurs inside amphibolites in the western flank of the antiform, at the deepesttectonostratigraphic levels (Figure 3). The amphibolites lie within orthogneisses as lenses showing an internalhornblende-dominated foliation subparallel to the regional trend. On the roof of the amphibolites, shear zonesexhibit a top-to-the NE sense of shear. The rare and small (max 20cm long) leucocratic pockets show both diffuseand sharp contacts with the foliation in the amphibolite. Garnet in the amphibolite host (5m away from the samplepoint) indicates upper amphibolite-faciesmetamorphism. The diffusemargins (Figure 4h) and the spatial associationwith garnet-amphibolites suggest locally-produced, quasi in situ melt. The analyzed sample is constituted ofmedium-grained plagioclase and alkali feldspar and rare hornblende defining an equigranular texture.

The zircon morphology of sample 10-132 varies between prismatic with rounded edges to subrounded forms.This is reflected in the low l/w ratio (1.1–3.5) of some grains. Subrounded and resorbed grains are typical ofmetamorphism-modified zircons. The zircons exhibit two types of CL pattern: (i) inhomogeneous and patchy or(ii) oscillatory and chaotic xenocrystic core mantled by a bright rim with rarely preserved weakly pronouncedoscillatory rims (Figure 6e). The largest age population is 678± 15Ma on average (Figure 8e). While theoscillatory xenocrystic cores consistently yielded a Proterozoic age, the heterogeneous and patchy cores varied,with ages from 1573Ma to 130Ma. The bright rims around these cores yielded a population of 16 ages witha mean at 116.4 ± 7.6Ma (Figure 8f). Bright rims not belonging to this cluster yielded ages of 301Ma, 281Ma,and 242Ma. The Th/U ratios of all bright rims were< 0.04, plotting in the domain of metamorphic rims [Teipelet al., 2004]. The circa 130Ma cores have also low Th/U values. The other analyses yielded values ranging from0.21 to 1.1 (Figure 9d). The zircon morphology and the low Th/U ratios suggest that the bright rims at116.4± 7.6Ma record the metamorphic/melting event. The concordant age at 242Ma (Figure 8e) is interpretedas that of the crystallization of the paleosome (e.g., the host amphibolite) since it is the second youngest ageafter the 116Ma population. The older zircon cores are interpreted as exotic within the 242Ma mafic intrusion.

4. Geochemical Constraints

Major- and trace-element geochemistry and Sr and Nd isotope analyses were performed on the datedgranitoids and on amphibolites to further investigate the origin of themagmas and the tectonic environmentof their formation. Tables of major and trace elements, of Sr and Nd isotopes and sample coordinates aregiven in the supporting information data.

Whole-rock XRF analyses were performed on fused disks using a Panalytical Axios wavelength dispersivespectrometer (WDXRF, 2.4 KV) at ETH-Zürich. Samples were ground to fine powder in an agate mill and mixedwith lithium tetraborate at 1:5 ratio and molten to homogenous glass disks. The spectrometer is set up for 12major and minor elements (SiO2, TiO2, Fe2O3, MnO, MgO, CaO, Na2O, K2O, P2O3, Cr2O3, and NiO) and 19 traceelements (Rb, Ba, Sr, Nb, Zr, Hf, Y, Ga, Zn, Cu, Co, V, Sc, La, Ce, Nd, Pb, Th, and U).

Trace elements were measured at ETH Zürich with laser ablation-inductively coupled plasma-massspectrometry on XRF fused disks calibrated without matrix-matching standards [Günther et al., 2001]. Dataacquisition time per spot was about 1min, and energy density was 15 J/cm2 at a frequency of 12 Hz. For eachdisk we analyzed three spots (90–120μm diameter). CaO XRF values were used as internal standards andcalibrated against National Institute of Standards and Technology (NIST) 610 for data correction with SILLSsoftware (http://www.geopetro.ethz.ch/research/orefluids/software). The expected measuring error is ~2%,close to the detection limit, and even smaller at higher concentrations.

Thirteen representative samples were analyzed for Sr and Nd isotopes on powdered bulk samples. Samples weredissolved in HF (3mL) and HNO3 (1mL) for 5 days in Teflon beakers at 180°C. This process was repeated after evapo-ration. Solutions were analyzed with thermal ionization mass spectrometry using a Triton plus multicollector spec-trometer. Standards were National Bureau of Standards (NBS) 987 (0.710234±4) for Sr and Ndi (0.512100±3) for Nd.

4.1. Geochemistry of Major Elements

Most of the granitoids, metagranitoids, and the metaandesitic dikes classify as calc-alkaline rocks in theA = Na2O + K2O, F = FeO, M = MgO (AFM) and in the MgO/(MgO+ FeOtot) versus SiO2 diagrams

Tectonics 10.1002/2014TC003563


http://www.geopetro.ethz.ch/research/orefluids/software

(Figure 10); exceptions are samples 10-132, P12-012, and P12-042, which classify as tholeiitic(Figure 10). By contrast, most of the amphibolites classify as tholeiites (Figure 10), but samples 09-114and 10-004 plot in the calc-alkaline field (Figure 10). These classifications are interpreted with cautionand will not be further taken into account in this study because the mobility of Mg, Na, and K inmetamorphic fluids can lead to deceiving tectonogenetic significance.

4.2. Geochemistry of Trace Elements

The chondrite-normalized trace-element patterns of the granitoids display enrichment in light over heavyrare earth element (REE), typical for calc-alkaline melts. REEs decrease gradually from La to Gd with a negativeEu anomaly of different magnitude in different samples (Figure 11a). This Eu anomaly is attributed toplagioclase fractionation from melts with low fO2 [e.g., Philpotts, 1970]. Heavy rare earth element (HREE)patterns are flat at values between 9 and 20 (normalized); samples 10-005 and P12-012 show higher HREEconcentrations. Sample P12-042 has remarkably lower light REE (LREE) concentrations. Most analyzedgranitoids plot in the volcanic arc granite field of the Rb versus Y +Nb discrimination diagram (Figure 12a).Since other postcollisional granites (e.g., Grandes Rousses, Querigut, Adamello, and Oman) plot in the samerange as the Pelagonian granitoids (Figure 12a), discriminating between volcanic arc calc-alkaline granitesand postcollisional calc-alkaline granites is senseless. Sample P12-012 classifies as within-plate granite.Furthermore, Rb is mobile in metamorphic fluids; hence, the Cretaceous metamorphismmay have altered Rbconcentrations and shifted analyses to the fields of syncollision or volcanic arc granites without overlap withthe postcollisional granite field (Figure 12a). Most samples have also a volcanic arc signature in the Rb-Hf-Tadiagram [Harris et al., 1986] (Figure 12b), though they plot close to the within-plate and the late- orpostcollision calc-alkaline fields. Owing to the known Rb mobility, these rocks may have been in the within-plate field before Cretaceous metamorphism (Figure 12b). Exceptions are sample 09-122 characterized as awithin-plate granite and sample P12-042 classified as late- or postcollision calc-alkaline granite. Because ofhigh-grade metamorphism, the classification of leucosome 10-132 in the upper Rb apex (Figure 12b) has notectonic significance.

Three amphibolite groups were distinguished on the basis of the chondrite-normalized REE patterns(Figures 11b and 11c):

1. Group I: samples 10-004 and 09-114 resemble the granitoid patterns (Figure 11a) with a steep slope fromLa (Lan (normalized)> 100) to a negative anomaly in Eu, followed by a flatter trend toward the HREE. Traceelements normalized to normal mid-oceanic ridge basalt (N-MORB) show a hump in large-ion lithophileelement (LILE) + Th followed by troughs in Nb-Ta, P, and Ti (Figure 11c). Such patterns are similar to those ofGlobal Subducting Sediment Composition (GLOSS) [Plank and Langmuir, 1998] and/or lavas that haveassimilated sediment (including from continental extensional settings). Considering the similarity of LREEsegments with those in all granitoids and the Eu anomalies suggestive of low-pressure crystal fractionation,

Figure 10. Major element discrimination plots for magmatic rocks. Groups in ellipses defined according to the REE patterns (Figure 11). (a) Binary diagramMgO/(MgO+ FeOtot) versus SiO2 (wt %). (b) AFM diagram [Irvine and Baragar, 1971].

Tectonics 10.1002/2014TC003563


these samples may have been formed by crustal contamination involving assimilation with fractional crys-tallization processes. In the Nb/Y versus Zr/TiO2 discrimination diagram (Figure 12c), these amphibolites ploton the border between the andesite/basalt and the subalkaline basalt field. The Nb/Yb versus Th/Yb dia-gram (Figure 12d) shows that the mafic rocks have assimilated Th along a trend that originates from theenriched mid-ocean ridge basalt (E-MORB) composition suggesting crustal contamination of Th throughmagma-crust interaction.

2. Group II: samples 10-133, 10-134, 10-135a, and 10-135b show a progressive depletion from La (Lan ~120)to Lu (Lun ~50) with a positive Eu anomaly representing plagioclase accumulation from a melt. There is aremarkable enrichment in all elements from Th to Yb with positive Th and Zr-Hf anomalies and a distinctdepletion in Large-ion lithophile elements (LILE). Two hypotheses may explain their untypical pattern(different from GLOSS, ocean island basalt, and MORB): (a) the parental melt was initially enriched in traceelements and was depleted in LILEs during metamorphism. However, Nb and Ta values are too low toreflect alkali, oceanic island basalt melts. For trace elements normalized to N-MORB, the concentration fromCeto Yb is higher than in Group I, which is supposed to be enriched by a sediment signature. (b) The high Ce to Ybvalues (especially Zr and Hf) may relate to local melt percolation (see sample 10-132, Figure 4h), which mayhave enriched the rock in trace elements during partial melting. Group II belongs to the andesite/basalt field inthe Nb/Y versus Zr/TiO2 diagram (Figure 12c). The Nb/Yb versus Th/Yb diagram shows that the mafic rockshave assimilated Th along a trend that originates from the E-MORB composition (Figure 12d) corroboratinghypothesis (a) of an already enriched parental melt. In any case, hypothesis (b) can be discarded, since theanalyzed amphibolites were collected meters away from leucocratic patches and veins.

3. Group III: samples 09-112c, 09-037, P12-020, P12-028, and 10-012 have nearly flat patterns atREEn ~ 17 ppm, resembling continental tholeiites from extensional settings [e.g., Fodor and Vetter, 1984].

Figure 11. Chondrite-normalized [Sun and McDonough, 1989] REE trace-element diagram of the Pelagonian (a) granitoids and (b) amphibolites. Sample 10-132 notplotted because concentrations were too close to detection limit. (c) N-MORB normalized [Gale et al., 2013] trace-element patterns of the Pelagonian amphibolites.GLOSS=global subducting sediment composition [Plank and Langmuir, 1998]. BAB=back-arc basin basalt composition [Gale et al., 2013].

Tectonics 10.1002/2014TC003563


Trace elements normalized to N-MORB have a flat trend around 1 from Nb to Yb with a LILE humpcharacteristic of MORB basalts contaminated with crustal LILE. This trend is typical for back-arc basins(Figure 11c), unless the LILE hump is a contamination of metamorphic fluids. No or a faint negative Euanomaly is observed. Group III plots in the subalkaline basalt field of the Nb/Y versus Zr/TiO2 diagram. Inthe Nb/Yb versus Th/Yb diagram, Th enrichment from crustal rocks through magma-crust interaction isabsent to moderate (Figure 12d).

4.3. Sr and Nd Isotopes87Sr/86Sr values of granitoids range between 0.709279 and 0.737409 (Figure 13a). Sample P12-042 showshigh 87Sr/86Sr ratio (0.737409), which is attributed to radiogenic production of 87Sr since the Precambrian.The 87Sr/86Sr values of the Permo-Carboniferous samples P12-026 (0.706645), 09-055 (0.709506), 09-121(0.708019), and 09-122 (0.708933) are typical crustal values. The Triassic augengneiss P12-012 has anintermediate value (0.724015) between the Precambrian and the Permo-Carboniferous intrusions. The143Nd/144Nd values between 0.512083 and 0.512328 of all granitoids are also typical crustal signatures.However, the andesitic dike (10-079) has lower 87Sr/86Sr (0.703935) and higher 143Nd/144Nd (0.512328) valuesthan the granitoids. The 87Sr/86Sr ratio of sample 10-079 is lower than the mantle UR value (0.704171)

Figure 12. Trace-element discrimination plots for magmatic rocks. (a) Rb versus (Y +Nb) [Pearce et al., 1984]. Ellipses = compositional field of postcollisional granites.(b) Rb-Hf-Ta discrimination diagram for intermediate and acid intrusive rocks in continent-continent collision zones [Harris et al., 1986]. (c) Nb/Y versus Zr/TiO2 plotoriginally designed for volcanic rocks [Winchester and Floyd, 1977], yet commonly used to discriminate between metamorphic subalkaline and alkaline mafic rocks[Pearce, 1996]. (d) Th/Yb versus Th/Yb plot [Pearce, 1983] designed for volcanic rocks, but the immobility of the discriminating elements make it valid also tometamorphic rocks. Groups in ellipses defined according to the REE patterns (Figure 11).

Tectonics 10.1002/2014TC003563


calculated at the time of intrusion (279Ma). Similarly, the 143Nd/144Nd value shows primitive, mantle valuesclose to 143Nd/144Nd CHUR (279Ma) ratios (0.512278).87Sr/86Sr values of the amphibolites spread from 0.705712 to 0.720101 and 143Nd/144Nd values from0.511939 to 0.512907. The large spread is a consequence of the high 87Sr/86Sr (0.720101) and low143Nd/144Nd ratios (0.511939) of sample 09-114. This sample is incorporated into the Precambrianmigmatitesthat may have introduced a crustal component in its isotopic signature. Discarding this anomaly, theamphibolites have 143Nd/144Nd values reflecting a more primitive mantle signature than granitoids.Differently, the 87Sr/86Sr in amphibolites plots in the same range as the granitoids. The Sr isotopes were likelymobile during metamorphism, homogenizing the isotopic signal in the Pelagonian basement except for10-079, which kept the mantle signature.

The εNd(t) [Depaolo and Wasserburg, 1976] was calculated according to the inferred intrusion age t of thedated samples, (Figure 13b). The εNd(t) decreased from �3.5 to �5.2 between 315.1 ± 1.4Ma (granodiorite09-121) and 308± 2Ma (leucogneiss 09-122). Thereafter, εNd(t) increased to �4.3 at 304 ± 1.5Ma(granodiorite 09-055), to decrease again to �7.3 at 301 ± 2Ma (leucogranite P12-026). After then, the εNd(t)rose to �1.9 at 292 ± 5Ma (sample Pl61 of Anders et al. [2006]) and to �1.3 at 279 ± 8Ma (andesitic dike 10-079). Sample V1 of Anders et al. [2006] is not taken into account for interpretation since the age error overlapsthe Late Carboniferous-Early Permian (circa 305 to 275Ma) granitoids. In the Triassic, the εNd(t) keptincreasing in the mafic rocks to circa �0.8 (10-134 and 10-135, 242 ± 9Ma) and lessening in the augengneiss�3.7 (sample P21-012, 245 ± 1.3Ma). A post-Carboniferous age is inferred for amphibolites P12-020 andP12-028 (geochemical back-arc signature of Group III) since they crosscut gneisses similar in aspect and inmineralogy to Carboniferous granitoids 09-055 and P12-026. Assuming a Jurassic intrusion age, thecalculated εNd(t) increases again to 5.6 for sample P12-020 and to 3.8 for sample P12-028.

5. Discussion5.1. U-Pb Ages

The SIMS zircon U-Pb ages of granitoids and metagranitoids of the Pelagonian Zone offer a considerable agespectrum from 1977Ma to 97Ma. Migmatite P12-041, leucogranite P12-042, leucogneiss 09-122, granitic vein10-050, and subvolcanic dike10-079 have an up to 1300Ma internal age spread, with single grains (e.g.,sample P12-041) having up to 1000Ma of age difference between core and rim. This wide andheterogeneous age distribution (Table 2) expresses the Precambrian polymagmatic and polyorogenic historyof the Pelagonian continental block and illustrates the complexity of intracontinental magmatic systems.

At least four episodes characterize the pre-Alpine gneiss. The oldest is deciphered from the inheritedmagmatic zircon cores of P12-041 migmatite at ca. 1430Ma. Zircon cores may date Mesoproterozoic

Figure 13. Sr and Nd isotopes plots. (a) 143Nd/144Nd and 87Sr/86Sr ratios of granitoids and amphibolites. Groups defined according to REE patterns (Figure 11). (b)Evolution of the εNd at the intrusion time (t) of each sample. εNd for amphibolites P12-020 and P12-028 calculated for an age t of 171Ma. Top right: εNd (0Ma) valuesfor modern back-arc and MORB magmatism from PetDB Database (www.earthchem.org/petdb) and from Depaolo [1988], respectively.

Tectonics 10.1002/2014TC003563


www.earthchem.org/petdb

gran

odio

rites

and

gr

anod

iorit

ic g

neis

ses

ande

sitic

sub

-vo

lcan

ic d

ike

leuc

ogra

nite

s an

d le

ucog

neis

s

gran

itic

vein

mig

mat

itic

orth

ogne

iss

leuc

osom

es w

ithin

am

phib

olite

s

auge

ngne

iss

with

in

amph

ibol

ites

09-0

55

Pl3

7

10-0

79

P12

-042

09-1

22

P12

-026

10-0

50

P12

-041

09-1

21

P12

-012

10-0

05

10-1

32

Pro

tero

zoic

Cre

tace

ous

Jura

ssic

Tria

ssic

Per

mia

nD

evon

ian

Cam

bO

rdv

Silu

Car

b

(And

ers

et

a

l. 20

07)

Roc

k T

ype

Sam

ple

304

± 1

.561

3

315.

1 ±

1.4

683

± 4

723

± 3

.116

12

1061

323

± 7

.258

0 ±

10

?40

119

77-1

705

1174

245

± 1

.3

307

± 1

1

301

± 2

570-

565

308

± 2

500

610-

550

599

± 5

697

± 8

1911

-160

812

94

279

± 8

.130

5.3

± 4

.6

253

± 7

?10

00

246

± 2

130

± 4

.4

116

± 7

.624

2 ±

9.1

1573

678

± 1

5

mos

t rec

ent m

agm

atic

age

mos

t rec

ent m

etam

orph

ic a

ge

mag

mat

ic o

r m

etam

orph

ic

mag

mat

ic a

ge (

mos

tly in

herit

ed c

ores

)

met

amor

phic

age

1432

± 6

Table

2.Su

mmaryof

Zircon

U-PbAge

sin

MaObtaine

din

ThisStud

y

Tectonics 10.1002/2014TC003563


intrusions that were subsequently molten in the Neoproterozoic, as recorded by the metamorphic zircon rims at683±4Ma (P12-041). These migmatites (P12-041) were intruded by leucogranites at 599±5Ma (P12-042). Thecirca 80Ma gap between the 683Ma migmatites and the 599Ma leucogranites suggests that the latter intrusionspertain to another event, possibly a postcollision calc-alkaline intrusion as suggested by the Rb-Hf-Ta plot(Figure 12b). Other witnesses of these circa 700–600Ma events were reported in the Pelagonian Zone both to thenorth and to the south of the study area andwere collectively attributed to the Florina Terrane [Anders et al., 2006].The magmatism of the Florina Terrane coincides with the Pan-African or Cadomian orogenies.

The fourth episode includes Late Carboniferous-Early Permian granitoids with the calc-alkaline signature(Figures 12a and 12b): the granodioritic gneiss (09-121) at 315.1 ± 1.4 truncated by the leucogneiss (09-122) at308 ± 2Ma and the granodiorite (09-055) at 304 ± 1.5Ma intruded by the leucogranite (P12-026) at301 ± 2Ma. Permo-Carboniferous intrusions are found also in the Moesian Platform, in the lower unit of theRhodope and in the Pontides [Okay et al., 2006], and they were intruded into the Variscan Eurasia [e.g.,Stampfli and Borel, 2002; Anders et al., 2007]. These Late Paleozoic plutons, also identified in previous works[Yarwood and Aftalion, 1976; Vavassis et al., 2000; Reischmann et al., 2001; Anders et al., 2007], were interpretedas part of the Paleozoic to Early Mesozoic active continental margin during Paleotethys subduction belowPelagonia [e.g., Stampfli and Borel, 2002]. Nevertheless, the absence of Late Carboniferous accretionarycomplexes, blueschists or eclogites, and deep forearc basins in the Balkans does not lend support forsubduction at that time. Calc-alkaline melts may form in postcollisional settings, deriving from a molten calc-alkaline basement. Distinguishing magmatic arc and postcollisional settings with geochemistry only is athorny problem since postcollisional granites cannot be explained in terms of a single crustal or mantlesource [e.g., Pearce et al., 1984]. In fact, thermomechanical numerical models of collapsed or extendedorogens show contemporaneous melting in the upwelling asthenosphere and in the lower crust [Schenkeret al., 2012]. The granitoids produced in such cases may follow an isotopic evolution that fudges the“classical” end-members (e.g., syncollisional, volcanic arc, and within-plate magmatism) because ofnonintuitive isotopic mixing. Furthermore, the waning stages of an orogen are often linked to the breakoff ofthe subducting plate [Davies and Von Blanckenburg, 1995]. The melts produced during this mechanical andthermal reequilibration are indistinguishable from subduction volcanic arc magmatism, even thoughsubduction in a strict sense is no longer active [Sizova et al., 2014].

This line of thoughts and knowledge excludes the calc-alkaline signature as final criterion to recognize activesubduction. The Early Permian increase of the εNd(t) (301-279Ma, Figure 13b) and the 87Sr/86Sr ratio(0.703935) of the 279Ma andesitic subvolcanic dike may instead indicate mantle contribution fromasthenospheric melts during lithospheric extension. Asthenospheric upwelling triggers high thermalconditions that melted the lower crust. The 500–1600Ma inherited zircon cores (Table 2) stem from a nearlycomplete assimilation of the Meso-Neoproterozoic basement. Similarly, the calc-alkaline signatureof< 301Ma melts is also inherited from the assimilated rocks. This interpretation fits the Permian extensionalhistory of the lower unit of Rhodope [Baziotis et al., 2014] and the late Variscan history of northern andwestern Europe [e.g., Menard and Molnar, 1988; Burg et al., 1994; Wilson et al., 2004] where postorogenicprocesses contributed to the growth and differentiation of continental lithosphere in the collapsed orogen[Costa and Rey, 1995].

Acidic (sample P12-012) and basic (amphibolites of Group II) Triassic intrusions are attributed to rifting of theTethys Ocean(s). The three Lower-Middle Triassic (245–242Ma) samples fit with sedimentary [e.g., Stais andFerrière, 1991; Brown and Robertson, 2003] and magmatic [Pe-Piper, 1998; Himmerkus et al., 2009] constraintssignaling Triassic rifting in the Hellenides. The bimodality of εNd(t) between augengneiss P12-012 andamphibolites 10-134 and 10-135 (Figure 13b) indicate coeval intrusions of crustal- and mantle-derived meltsduring extension [Schenker et al., 2012]. Cretaceous zircon rims at 130-116Ma and foliation in samples 10-005and 10-132 corroborate a tectonometamorphic event in the Early Cretaceous. This event was previouslyhinted from Rb/Sr [Yarwood and Dixon, 1977] and from 40Ar/39Ar (at 100–130Ma) geochronology (in thesouthern Pelagonian Zone) [Schermer, 1990; Schermer et al., 1990; Lips et al., 1998, 1999]. Sample 10-132 yieldsthe first direct record of Cretaceous melting in the Pelagonian Zone. Mapping indicates that anatexis ispunctuated, not as regional as the amphibolite-facies metamorphism accompanying anatexis. The heat forthis regional metamorphism cannot have originated from the advected Cretaceous magmatism, which wasvolumetrically too small; it had to source from in situ heat production during collision [e.g., Burg and Gerya, 2005].Since the age of migmatization at 116±8Ma is not distinguishable from the regional 40Ar/39Ar cooling ages, the

Tectonics 10.1002/2014TC003563


inferred cooling rates suggest that Cretaceous metamorphism occurred at the end of the collision-related burialand during cooling/exhumation of Pelagonia. Collision involved the Eurasianmargin and promontories (Pelagoniaand lower unit of Rhodope) detached/rifted away from the Eurasianmargin in the Triassic (Figure 14). This tectonichistory implies that (i) the Tethys suture is nowadays located to the east of the Axios/Vardar Zone, (ii) the erraticophiolitic fragments mapped onto and southwest of Pelagonia (e.g., Vourinos) are allochthonous from thenortheast, and (iii) Pelagonia was unified to the lower unit of Rhodope before Alpine collision (Figure 14).

The tectonic reconstruction in Figure 14 is based on (i) the Early Jurassic paleogeographic configuration(Figure 14a) showing from the southwest to the northeast the external Hellenides, the Pindos basin, thecontinental block formed by the Pelagonia and lower unit of the Rhodope, the passivemargin of the Axios/Vardar,and the main mature Tethyan lithosphere; and (ii) the northeast dipping subduction of a unique lithospheric slabas seen on seismic tomography [Bijwaard et al., 1998; van Hinsbergen et al., 2005].

Three convergence events can be suggested (Figure 14a): (i) In the Late Jurassic-Early Cretaceous, theVardar units were imbricated [Bernoulli and Laubscher, 1972; Baumgartner, 1985; Chiari et al., 2013] belowthe main obduction from the northeast. (ii) In the late Early Cretaceous, collision or arrival of Pelagoniawithin the Rhodopian nappe pile produced regional metamorphism that culminated with themigmatization dated in this work at 117 ± 8Ma. And (iii) in the Paleocene-Eocene, Pelagonia was thrustsouthwestward forming the fold-and-thrust accretionary system of the external Hellenides [Renz andReichel, 1945; Jacobshagen et al., 1978].

Figure 14. (a) Paleogeographic reconstruction at the maximal Tethyan extension during the Early Jurassic with timing of oceanic obduction and Axios/Vardar andPelagonia thrusting. (b) Position of the Pelagonian Block during peak metamorphism at circa 116Ma. Erratic ophiolites are fragments of ophiolites resting onto thePelagonia (e.g., Vourinos ophiolites).

Tectonics 10.1002/2014TC003563


5.2. Timing of Deformation: Ages of Foliations

On the large-scale, the orientation of the main foliation outlines the regional architecture of Pelagonia as anantiform. The ages of pre-, syn-, and post-foliation granitoids suggest three fabric-forming events whoserelationships are sketched in Figure 15:

1. A Neoproterozoic, synmelt foliation in the migmatitic orthogneiss (P12-041) at 683±4Ma cut by aNeoproterozoic leucogranite at 599±5Ma (the postfoliation dyke P12-042, Figure 4b) and by a Carboniferousgranitic vein at 326±7Ma (sample 10-050).

2. A late- to post-Variscan subsolidus foliation in orthogneisses and leucogneisses, assumed to be of the samegeneration as the 304Ma granodiorite (samples 09-055 and P12-026), truncated by the subvolcanic dike at279±8Ma (sample 10-079). In places, the alignment of euhedral alkali feldspar suggests that the foliation in theorthogneisses is also magmatic.

3. An Early Cretaceous foliation with synanatectic to postanatectic melts at 116±8Ma (sample 10-132). The samefabric-forming event was responsible for the top-to-the SW shearing that deformed the 245±1.3Maaugengneiss (sample P12-012, Figure 3f).

No structural and petrographical criterion could be used to distinguish the three foliations depicted bythe preferred orientation of biotite, white mica, and hornblende in the orthogneisses and hornblendein the amphibolites. The mineral assemblages hint to similar P and T conditions, whether foliation ismetamorphic or magmatic. On the regional scale, all foliations are parallel, within errors of compassmeasurements. Therefore, all foliations have either become parallel to each other during the lastductile event, or they may all have been originally subparallel (shear zones developing along oldfoliations), before the development of the antiform. Our structural observation showed a pattern ofheterogeneous shear deformation, with strain localization into nearly foliation-parallel shear zones(Figure 15). Hence, shear-induced rotation during the Early Cretaceous has most likely parallelized allfoliations, yet dissecting patches of Precambrian rocks in which primary and ancient structuralrelationships, like the 599 ± 5Ma dike (sample P12-042) cutting the 683 ± 4Ma migmatitic orthogneiss(sample P12-041), are preserved.

Figure 15. Schematic field relationships with magmatic ages of the Pelagonian basement. Not to scale.

Tectonics 10.1002/2014TC003563


6. Conclusions

1. New U-Pb results reveal the polymagmatic and polyorogenic history of the Pelagonian Zone withinherited zircon ages up to Paleoproterozoic (1977Ma). The circa 1430Ma granitic precursor of the circa683Ma migmatites may represent the oldest rock found in Greece. These basement rocks were cut by acirca 600Ma leucogranite. Such Neoproterozoic thermal/magmatic events are consistent with thesyncollisional to postcollisional evolution of the Pan-African or Cadomian active margin. The discoveredPrecambrian rocks are attributed to the Florina Terrane, on the basis of similar ages further north andsouth in the Pelagonian Zone [Anders et al., 2006].

2. The Neoproterozoic rocks were assimilated into the 315-279Ma magmatism with a calc-alkaline, volcanicarc signature typical for the Eurasianmargin. In the Early Permian, the rise of εNd(t) suggests an increase ofasthenospheric melt tentatively attributed to asthenospheric upwelling during lithospheric extension.

3. Middle-Late Triassic alkaline acidic and mafic magmatism (246-242Ma, sample P12-012 and the cores ofzircons of samples 10-005 and 10-132) is coeval with the opening of the Tethys Ocean(s) in the north-eastern margin of the developing Vardar basin [Stais and Ferrière, 1991; Brown and Robertson, 2003]. Thisrifting event separated the Pelagonian from Eurasia. Early Jurassic inversion of plate motions led to (a)Late Jurassic-Early Cretaceous southwestward obduction of Tethyan ophiolites over Pelagonia and (b) lateEarly Cretaceous collision between Pelagonia and the European (Rhodopian) margin.

4. The role of Pelagonia during late Early Cretaceous collision was thought to be unimportant due to itsconjectured distal location from the orogenic trench [Ricou et al., 1998; van Hinsbergen et al., 2005;Sharp and Robertson, 2006]. The first dating of Cretaceous migmatites (sample 10-132 at 116 ± 8Ma) in thePelagonian Zone represents incipient in situ melts during regional metamorphism of Pelagonia collidedwith the Eurasian margin. In this tectonic reconstruction, the “root” suture is located at the easternmostparts of the Vardar/Axios Zone.

ReferencesAnders, B., T. Reischmann, D. Kostopoulos, and U. Poller (2006), The oldest rocks of Greece: First evidence for a Precambrian terrane within

the Pelagonian Zone, Geol. Mag., 143(1), 41–58.Anders, B., T. Reischmann, and D. Kostopoulos (2007), Zircon geochronology of basement rocks from the Pelagonian Zone, Greece:

Constraints on the pre-Alpine evolution of the westernmost internal Hellenides, Int. J. Earth Sci., 96(4), 639–661.Baumgartner, P. O. (1985), Jurassic sedimentary evolution and nappe emplacement in the Argolis Peninsula (Peloponnesus; Greece), Mém.

Soc. Helv. Sci. Nat., 99, 1–111.Baziotis, I., E. Mposkos, and P. D. Asimow (2014), Continental rift and oceanic protoliths of mafic-ultramafic rocks from the Kechros Complex,

NE Rhodope (Greece): Implications from petrography, major and trace-element systematics, and MELTS modeling, Int. J. Earth Sci., 103,981–1003.

Bernoulli, D., and H. Laubscher (1972), The palinspastic problem of the Hellenides, Eclogae Geol. Helv., 65(1), 1007–1118.Bijwaard, H., W. Spakman, and E. R. Engdahl (1998), Closing the gap between regional and global travel time tomography, J. Geophys. Res.,

103(B12), 30,055–30,078, doi:10.1029/98JB02467.Boykova, A. (1999), Moho discontinuity in central Balkan Peninsula in the light of the geostatistical structural analysis, Phys. Earth Planet.

Inter., 114(1-2), 49–58.Brown, S. A. M., and A. H. F. Robertson (2003), Sedimentary geology as a key to understanding the tectonic evolution of the Mesozoic-Early

Tertiary Paikon Massif, Vardar suture zone, N Greece, Sediment. Geol., 160(1-3), 179–212.Brun, J. P., and C. Faccenna (2008), Exhumation of high-pressure rocks driven by slab rollback, Earth Planet. Sci. Lett., 272(1-2), 1–7.Burg, J.-P. (2012), Rhodope: From Mesozoic convergence to Cenozoic extension. A petro-structural review, J. Virtual Explor., 42, 1–44.Burg, J. P., and T. V. Gerya (2005), The role of viscous heating in Barrovian metamorphism of collisional orogens: Thermomechanical models

and application to the Lepontine Dome in the Central Alps, J. Metamorph. Geol., 23(2), 75–95.Burg, J.-P., J. Van Den Driessche, and J. P. Brun (1994), Syn- to post-thickening extension in the Variscan belt of Western Europe: Modes and

structural consequences, Géol. Fr., 3, 18.Chiari, M., P. O. Baumgartner, D. Bernoulli, V. Bertolotti, M. Marcucci, A. Phodiades, and G. Principi (2013), Late Triassic, Early and Middle

Jurassic Radiolaria from ferromanganese-chert ‘nodules’ (Angelokastron, Argolis, Greece): Evidence for prolonged radiolarite sedimentation in theMaliac-Vardar Ocean, Facies, 59, 391–424.

Corfu, F., J. M. Hanchar, P. W. O. Hoskin, and P. Kinny (2003), Atlas of zircon textures, Rev. Mineral. Geochem., 53, 469–500.Costa, S., and P. Rey (1995), Lower crustal rejuvenation and growth during post-thickening collapse—Insights from a crustal cross-section

through a Variscan metamorphic core complex, Geology, 23(10), 905–908.Davies, J. H., and F. Von Blanckenburg (1995), Slab breakoff—A model of lithosphere detachment and its test in the magmatism and

deformation of collisional orogens, Earth Planet. Sci. Lett., 129(1-4), 85–102.Depaolo, D. J. (1988), Nd Isotope Geochemistry: An Introduction, 187 pp., Springer, Berlin, Heidelberg, New York.Depaolo, D. J., and G. J. Wasserburg (1976), Interferences about magma sources and mantle structures from variations of

143Nd/

144N,

Geophys. Res. Lett., 3(12), 743–746, doi:10.1029/GL003i012p00743.Dercourt, J., et al. (1986), Geological evolution of the Tethys Belt from the Atlantic to the Pamirs since the Lias, Tectonophysics, 123(1-4),

241–315.Ernst, W. G., and J. Liu (1998), Experimental phase-equilibrium study of Al- and Ti-contents of calcic amphibole in MORB—A semiquantitative

thermobarometer, Am. Mineral., 83(9-10), 952–969.

Tectonics 10.1002/2014TC003563


AcknowledgmentsThis work was supported by ETHResearch grant ETH-09 09-3. All the dataused in this work are provided in thesupporting information. We thank JanPleuger and Laurent Jolivet for theirthorough reviews, which helped clarify-ing our results and interpretations.

http://dx.doi.org/10.1029/98JB02467

http://dx.doi.org/10.1029/GL003i012p00743

Fodor, R. V., and S. K. Vetter (1984), Rift-zone magmatism—Petrology of basaltic rocks transitional from CFB to MORB, southeastern Brazilmargin, Contrib. Mineral Petrol., 88(4), 307–321.

Gale, A., C. A. Dalton, C. H. Langmuir, Y. Su, and J.-P. Schilling (2013), The mean composition of ocean ridge basalts, Geochem. Geophys.Geosyst., 489–518, doi:10.1029/2012GC004334.

Godfriaux, I. (1968), Etude géologique de la région de l’Olympe (Gréce), Ann. Géol. Pays Hell., 19, 1–271.Godfriaux, I., J. Ferriere, and A. Schmitt (1988), Le développement en contexte continental d’un métamorphisme HP/BT: Les “schistes bleus”

tertiaires Thessaliens, Bull. Geol. Soc. Greece, 20, 175–192.Günther, D., A. von Quadt, R. Wirz, H. Cousin, and F. J. Dietrich (2001), Elemental analyses using laser ablation-inductively coupled plasma-mass

spectrometry (LA-ICP-MS) of geological samples fused with Li2B4O7 and calibrated without matrix-matched standards, Mikrochim. Acta,136, 101–107.

Harris, N. B. W., J. A. Pearce, and A. G. Tindle (1986), Geochemical characteristics of collision-zone magmatism, Geol. Soc. London Spec. Publ.,19, 67–81.

Heaman, L. M., R. Bowins, and J. Crocket (1990), The chemical-composition of igneous zircon suites—Implications for geochemical tracerstudies, Geochim. Cosmochim. Acta, 54(6), 1597–1607.

Himmerkus, F., T. Reischmann, and D. Kostopoulos (2009), Triassic rift-related meta-granites in the Internal Hellenides, Greece, Geol. Mag.,146(2), 252–265.

Irvine, T. N., and W. R. A. Baragar (1971), Guide to chemical classification of common volcanic rocks, Can. J. Earth Sci., 8(5), 523–548.Jacobshagen, V., S. Dürr, F. Kockel, K. O. Kopp, G. Kowalczyk, H. Berckhemer, and D. Büttner (1978), Structure and Geodynamic Evolution of the

Aegean Region, pp. 537–564, E. Schweizerbart’sche Verlagsbuchhandlung, Stuttgart.Jolivet, L., and J. P. Brun (2010), Cenozoic geodynamic evolution of the Aegean, Int. J. Earth Sci., 99(1), 109–138.Kilias, A., and D. Mountrakis (1989), The Pelagonian nappe: Tectonics, metamorphism and magmatism, Bull. Geol. Soc. Greece, 23(1), 29–46.Kilias, A., W. Frisch, A. Avgerinas, I. Dunkl, G. Falalakis, and H. J. Gawlick (2010), Alpine architecture and kinematics of deformation of the

northern Pelagonian nappe pile in the Hellenides, Aust. J. Earth Sci., 103(1), 4–28.Larionov, A. N., V. A. Andreichev, and D. G. Gee (Eds.) (2004), The Vendian Alkaline Igneous Suite of Northern Timan: Ion Microprobe U–Pb Zircon

Ages of Gabbros and Syenite, Mem. Geol. Soc. London, 30, 69–74.Lepichon, X., and J. Angelier (1979), Hellenic arc and trench system—Key to the neotectonic evolution of the eastern Mediterranean area,

Tectonophysics, 60(1-2), 1–42.Lips, A. L. W., S. H. White, and J. R. Wijbrans (1998),

40Ar/

39Ar laserprobe direct dating of discrete deformational events: A continuous record of

early Alpine tectonics in the Pelagonian Zone, NW Aegean area, Greece, Tectonophysics, 298(1-3), 133–153.Lips, A. L. W., J. R. Wijbrans, and S. H. White (1999), New Insights From

40Ar/

39Ar Laserprobe Dating of White Mica Fabrics From the Pelion Massif,

Pelagonian Zone, Internal Hellenides, Greece: Implications for the Timing of Metamorphic Episodes and Tectonic Events in the Aegean Region,pp. 457–474, Geological Society, London.

Lister, G. S., G. Banga, and A. Feenstra (1984), Metamorphic core complexes of cordilleran type in the cyclades, Aegean Sea, Greece, Geology,12, 221–225.

Ludwig, K. R. (2005a), User’s manual for ISOPLOT/Ex 3.22. A geochronological toolkit for Microsoft excel, Berkeley Geochronology CenterSpec Pub, 71.

Ludwig, K. R. (2005b), SQUID 1.12 a user’s manual. A geochronological toolkit for Microsoft excel, Berkeley Geochronology Center Spec Pub, 22.Menard, G., and P. Molnar (1988), Collapse of a Hercynian Tibetan Plateau into a Late Paleozoic European Basin and Range Province, Nature,

334(6179), 235–237.Mercier, J. (1968), Etude géologique des zones Hellénides en Macèdoine centrale (Grece), Ann. Géol. Pays Hell., 20, 17–92.Mercier, J., J. P. Vergely, and J. Bebien (1975), Les ophiolites helléniques “Obductées” an Jurassique supérieur sontelles les vestiges d’ un

ocean téthysien ou d’une mer marginale peri-europeenne?, C. R. Somm. Séances Soc. Géol. Fr., 17, 108–112.Mountrakis, D. (1986), The Pelagonia zone in Greece: A polyphase deformed fragment of the Cimmerian continent and its role in the

geotectonic evolution of the eastern Mediterranean, J. Geol., 94, 335–347.Okay, A. I., M. Satur, and W. Siebel (Eds.) (2006), Pre-Alpide Palaeozoic and Mesozoic Orogenic Events in the Eastern Mediterranean Region,

pp. 389–405, Geological Society, London.Papanikolaou, D. (1989), Are the medial crystalline massifs of the eastern Mediterranean drifted Gondwanan fragments?, in GCP No:276,

Newsletter No:1, edited by D. J. Papanikolaou and F. P. Sassi, Geol. Soc. Greece Spec. Publ., 1, 63–90.Papanikolaou, D. (2009), Timing of tectonic emplacement of the ophiolites and terrane paleogeography in the Hellenides, Lithos, 108(1–4),

262–280.Papazachos, C. B. (1998), Crustal P- and S-velocity structure of the Serbomacedonian Massif (Northern Greece) obtained by non-linear

inversion of traveltimes, Geophys. J. Int., 134(1), 25–39.Pearce, J. A. (1983), Role of the sub-continental lithosphere in magma genesis at active continental margins, in Continental Basalts and

Mantle Xenoliths, edited by C. J. Wawkesworth and M. J. Norry, pp. 230–249, Shiva, Nantwich, U. K.Pearce, J. A. (Ed.) (1996), A User’s Guide to Basalt Discrimination Diagrams, Geological Association of Canada.Pearce, J. A., N. B. W. Harris, and A. G. Tindle (1984), Trace-element discrimination diagrams for the tectonic interpretation of granitic-rocks,

J. Petrol., 25(4), 956–983.Pe-Piper, G. (1998), The nature of Triassic extension-related magmatism in Greece: Evidence from Nd and Pb isotope geochemistry, Geol. Mag.,

3, 331–348.Philpotts, J. A. (1970), Redox estimation from a calculation of Eu2+ and Eu3+ concentrations in natural phases, Earth Planet. Sci. Lett., 9, 257–268.Plank, T., and C. H. Langmuir (1998), The chemical composition of subducting sediment and its consequences for the crust and mantle,

Chem. Geol., 145(3-4), 325–394.Reischmann, T., and D. Kostopoulos (2007), Terrane accretion in the internal Hellenides, Geophys. Res. Abstr., 9, 05337.Reischmann, T., D. K. Kostopoulos, S. Loos, B. Anders, A. Avgerinas, and S. A. Sklavounos (2001), Late Paleozoic magmatism in the basement

rocks southwest of the Mt. Olympos, Central Pelagonian Zone, Greece: Remnants of a Permo-Carboniferous magmatic arc, Bull. Geol. Soc. Greece,25, 985–993.

Renz, C., and M. Reichel (1945), Beiträge zur stratigraphie und paläontologie des ostmediterranen jungpaläozoikums und dessen einordnung imgriechischen gebirgsystem, Eclogae Geol. Helv., 38(2), 221–313.

Ricou, L. E., J. P. Burg, I. Godfriaux, and Z. Ivanov (1998), Rhodope and Vardar: The metamorphic and the olistostromic paired belts related tothe Cretaceous subduction under Europe, Geodin. Acta, 11(6), 285–309.

Robertson, A. H. F., and J. E. Dixon (Eds.) (1984), Introduction: Aspects of the Geological Evolution of the Eastern Mediterranean, pp. 1–74,Geological Society, London.

Tectonics 10.1002/2014TC003563


http://dx.doi.org/10.1029/2012GC004334

Sachpazi, M., A. Galve, M. Laigle, A. Him, E. Sokos, A. Serpetsidaki, J. M. Marthelot, J. M. P. Alperin, B. Zelt, and B. Taylor (2007), Mohotopography under central Greece and its compensation by Pn time-terms for the accurate location of hypocenters: The example of theGulf of Corinth 1995 Aigion earthquake, Tectonophysics, 440(1-4), 53–65.

Schenker, F. L., T. Gerya, and J. P. Burg (2012), Bimodal behavior of extended continental lithosphere: Modeling insight and application tothermal history of migmatitic core complexes, Tectonophysics, 579, 88–103.

Schermer, E. R. (1990), Mechanisms of blueschist creation and preservation in an a-type subduction zone, Mount-Olympus Region, Greece,Geology, 18(11), 1130–1133.

Schermer, E. R., D. R. Lux, and B. C. Burchfiel (1990), Temperature-time history of subducted continental-crust, Mount Olympos Region, Greece,Tectonics, 9(5), 1165–1195, doi:10.1029/TC009i005p01165.

Sharp, I. R., and A. H. F. Robertson (Eds.) (2006), Tectonicsedimentary Evolution of the Western Margin of the Mesozoic Vardar Ocean: EvidenceFrom the Pelagonian and Almopias Zones, Northern Greece, pp. 373–412, Geological Society, London.

Sizova, E., T. Gerya, and M. Brown (2014), Contrasting styles of Phanerozoic and Precambrian continental collision, Gondwana Res., 25,522–545.

Stais, A., and J. Ferrière (1991), Nouvelles données sur la paléogéographie mésozoique du domaine Vardarien: Les bassins d’Almopias et dePéonias (Macédoine, Helleniques Internes septentrionales), Bull. Geol. Soc. Greece, 25, 491–507.

Stampfli, G. M., and G. D. Borel (2002), A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries andrestored synthetic oceanic isochrons, Earth Planet. Sci. Lett., 196(1-2), 17–33.

Steiger, R. H., and E. Jager (1977), Subcommission on geochronology—Convention on use of decay constants in geochronology and cos-mochronology, Earth Planet. Sci. Lett., 36(3), 359–362.

Sun, S. S., andW. F. McDonough (1989), Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes, inMagmatism in the Ocean Basins, edited by A. D. Saunders and M. J. Norry, Geol. Soc. London Spec. Publ., 42, 313–345.

Teipel, U., R. Eichhorn, G. Loth, J. Rohrmuller, R. Holl, and A. Kennedy (2004), U-Pb SHRIMP and Nd isotopic data from the western BohemianMassif (Bayerischer Wald, Germany): Implications for upper Vendian and lower Ordovician magmatism, Int. J. Earth Sci., 93(5), 782–801.

Tera, F., and G. J. Wasserburg (1972), U-Th-Pb systematics in 3 Apollo 14 basalts and problem of initial Pb in lunar rocks, Earth Planet. Sci. Lett.,14(3), 281.

van Hinsbergen, D. J. J., E. Hafkenscheid, W. Spakman, J. E. Meulenkamp, and R. Wortel (2005), Nappe stacking resulting from continentallithosphere below subduction of oceanic and Greece, Geology, 33(4), 325–328.

Vavassis, I., A. De Bono, G. M. Stampfli, D. Giorgis, A. Valloton, and Y. Amelin (2000), U-Pb and Ar-Ar geochronological data from thePelagonian basement in Evia (Greece): Geodynamic implications for the evolution of Paleotethys, Schweiz. Miner. Petrog., 80(1), 21–43.

Wetherill, G. V. (1956), Discordant uranium-lead ages, I, Trans. AGU, 37, 320–326.Whitney, D. L., and B. W. Evans (2010), Abbreviations for names of rock-forming minerals, Am. Mineral., 95, 185–187.Williams, I. S. (1998), U-Th-Pb geochronology by ion microprobe, in Applications of Microanalytical Techniques to Understanding Mineralizing

Processes, edited by M. A. McKibben, W. C. Shanks III, and W. I. Ridley, pp. 1–35, Reviews Economic Geology, Ann Arbor.Wilson, M., E.-R. Neumann, G. R. Davies, M. J. Timmerman, M. Heeremans, and B. T. Larsen (2004), Permo-Carboniferous Magmatism and Rifting

in Europe, Geological Society, London.Winchester, J. A., and P. A. Floyd (1977), Geochemical discrimination of different magma series and their differentiation products using

immobile elements, Chem. Geol., 20(4), 325–343.Yarwood, G. A., and M. Aftalion (1976), Field relations and U–Pb geochronology of a granite from the Pelagonian Zone of the Hellenides

(High Pieria, Greece), Bull. Soc. Geol. Fr, 18(2), 259–264.Yarwood, G. A., and J. E. Dixon (1977), Lower Cretaceous and younger thrusting in the Pelagonian Rocks of High Piera, Greece, Colloq. Aegean

Region, Athens 6th, 1, 269–280.Zlatkin, O., D. Avigad, and A. Gerdes (2014), Peri-Amazonian provenance of the Proto-Pelagonian basement (Greece), from zircon U-Pb

geochronology and Lu-Hf isotopic geochemistry, Lithos, 184-187, 379–392.

Tectonics 10.1002/2014TC003563


http://dx.doi.org/10.1029/TC009i005p01165

From Mesoproterozoic magmatism to collisional Cretaceous anatexis: Tectono-magmatic history of the...

Documents

Transcript of From Mesoproterozoic magmatism to collisional Cretaceous anatexis: Tectono-magmatic history of the...