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DISS. ETH NO. 22888
Magmatic geochemistry and geochronology
in relation to the geodynamic and metallogenic evolution
of the Banat Region and the Apuseni Mountains of Romania
A thesis submitted to attain the degree of
DOCTOR OF SCIENCES of ETH ZURICH
(Dr. sc. ETH Zurich)
presented by
DANIELA GALLHOFER
MSc Angewandte Geowissenschaften, Montanuniversität Leoben
born on 12.01.1985
citizen of Austria
accepted on the recommendation of
Prof. Dr. Christoph A. Heinrich, examiner
Dr. Albrecht von Quadt, co-examiner
Prof. emeritus Dr. Stefan M. Schmid, co-examiner
Prof. Dr. Gerhard Wörner, co-examiner
2015
Abstract
Abstract
The Apuseni-Banat-Timok-Srednogorie (ABTS) magmatic arc in southeastern Europe
formed as a result of NE-dipping subduction of the Neotethys ocean beneath the European
continental margin during the Late Cretaceous. This magmatic arc is associated with some of
Europe’s largest porphyry Cu-Au and epithermal Cu-Au deposits. However, the ore deposits
are not evenly distributed within the arc and porphyry-type and epithermal deposits occur only
in the central segments of the arc. Subsequently, the arc experienced intense deformation on a
lithospheric scale, resulting in the present L-shape, which complicates the interpretation of the
original setting of the arc. The timing and geochemical evolution of the arc magmatism is well
studied in the central and eastern segments, but information on the northernmost Banat and
Apuseni segments, is still scarce. New whole rock major and trace element data, radiogenic
isotope (Sr-Nd) data and U-Pb zircon dates for the Banat and Apuseni segments complement
the existing dataset of the ABTS magmatic arc. Integration with earlier studies of other arc
segments allows an overall reconstruction of the magmatic, geodynamic and metallogenic
evolution of the ABTS arc.
In the Apuseni Mountains, two additional phases of calc-alkaline magmatism occur. A
younger phase of Miocene magmatism is superimposed on the Late Cretaceous arc magmatism,
and older Jurassic island-arc calc-alkaline granitoids are associated with obducted Jurassic
ophiolites. Although the Miocene magmatism has characteristics of subduction-related
magmatism and is associated with ore deposits commonly found in magmatic arcs, it did not
form in a classical subduction-related setting. Previous authors suggested that it formed due to
extension-induced re-melting of a previously subduction-modified source. The close spatial
relationship between the Late Cretaceous and Miocene magmatism makes a common source
region likely. The new findings and geochemical data for the Late Cretaceous magmatism
permit a test and provide at least a permissive confirmation of the genetic link between Late
Cretaceous mantle metasomatism and Miocene post-subduction extensional melting.
The Jurassic ophiolites and calc-alkaline series are petrographically and geochemically
well characterized. However, radiogenic isotope data (Sr-Nd) yield important additional
constraints on the crustal contamination of the calc-alkaline granitoids, and might influence or
strengthen the interpretation of the tectonic setting. By combining radiogenic isotope data and
U-Pb zircon ages, I reinvestigate the tectonic setting of the calc-alkaline granitoids and estimate
i
Abstract
the maximum age of obduction of the Jurassic ophiolites and calc-alkaline granitoids in the
Apuseni Mountains.
This thesis consists of three chapters, which were written as co-authored papers of which
I am the leading author. The first chapter was published in ‘Tectonics’. Included are also an
overall introduction and a general conclusions section with recommendations for future
research.
The first chapter aims at refining the reconstruction of the original arc geometry and
better constrain the tectonic evolution of the Late Cretaceous ABTS magmatic arc. The Late
Cretaceous magmatic arc can be divided into five geochemically distinct, partly mineralized
and partly barren segments: the (1) Apuseni, (2) Banat, (3) Timok, (4) Panagyurishte and (5)
Eastern Srednogorie segments. Trace elements and isotopic signatures of the arc magmas
indicate a subduction-enriched source in all segments and variable contamination by continental
crust. The continental arc was active for 25 Ma (92.2 to 66.8 Ma), and systematic across-arc
age and isotopic trends are observed in nearly all arc segments. Progressively younger ages
towards the paleo-trench indicate gradual steepening of the subducting Neotethys slab, away
from the upper plate European continental margin. Steepening of the slab enhances
asthenospheric corner flow in the overriding plate, which is detected by decreasing 87Sr/86Sr
(0.70577 to 0.70373) and increasing 143Nd/144Nd (0.51234 to 0.51264) ratios over time in some
segments. Large-scale shear zones and related strike-slip sedimentary basins formed
contemporaneously with arc magmatism in the Panagyurishte and Timok segments, indicate
mild transtension in these central arc segments, whereas the deep marine co-magmatic basin in
the Eastern Srednogorie segment records strong orthogonal extension. Porphyry-Cu and
epithermal deposits formed exclusively in the central arc segments that experienced only mild
transtension during contemporaneous shearing, which favored lower crustal high-pressure
fractionation resulting in ‘adakite-like’ signatures, and accumulation of volatiles and metals.
Post-emplacement deformation and associated extension concealed the rather simple geometry
of this continental arc, but allowed the preservation of near-surface ore deposits.
The second chapter explores a genetic link between Late Cretaceous and Miocene calcalkaline magmatism in the Apuseni Mountains. The Late Cretaceous arc magmatism in the
Apuseni Mountains is more silicic than in the other arc segments, and rhyolitic ignimbrites are
restricted to this arc segment. Mantle-derived melts presumably underwent a polybaric
evolution in the mid to upper crust. Low Sr/Y ratios, decreasing Sr contents and Dy/Yb ratios
indicate fractionation of a plagioclase- and amphibole-bearing assemblage, probably at mid
crustal levels (~20 km). High
87
Sr/86Sr80Ma ratios (up to 0.716278) and low
ii
143
Nd/144Nd80Ma
Abstract
ratios (as low as 0.512187) require the addition of a maximum of 60% partial crustal melts to
mantle derived magmas. The andesitic to dacitic melts then ascended to shallow crustal levels
(<8 km) and evolved to high silica rhyolitic melts. Explosive volcanism might have triggered a
rapid loss of volatiles, which prevented the formation of porphyry-style deposits. Miocene remelting of the subduction-modified mantle led to calc-alkaline magmatism, and associated AuTe epithermal and Cu-Au porphyry-deposits. Two groups of Miocene magmas differ in their
87
Sr/86Sr and trace element ratios, and presumably evolved via distinct pathways in the
continental crust. A slightly older high 87Sr/86Sr (0.706529-0.707596) group assimilated local
crust and presumably underwent fractionation at mid to upper crustal levels. A low 87Sr/86Sr
(0.703819-0.705431) group shows extreme enrichments in Sr, Ba and La and has ‘adakite-like’
trace element characteristics (high Sr/Y, La/Yb). These signatures were probably acquired by
the addition of small-degree partial melts of hydrous mafic cumulates which formed in the crust
during the Late Cretaceous arc magmatism. Both groups of Miocene magmas are unusually Aurich, which might be explained by re-melting of Au+(Cu, Te)-rich sulfides left in the mantle
after extraction of the Late Cretaceous arc magmas.
The third chapter aims to refine the timing and tectonic setting of the Jurassic ophiolites
and the calc-alkaline series in the Apuseni Mountains. The South Apuseni ophiolites are the
northernmost occurrence of ophiolites presently overlying the Dacia continental unit in
Romania, Serbia, Macedonia (FYROM) and Greece. The ophiolites show dominantly MORBtype affinities, but slight enrichments in Th and U and depletion in Nb probably point to their
formation in a marginal or back-arc basin. Four gabbros from the ophiolitic sequence yielded
U-Pb ages between 158.9 and 155.9 Ma (Late Jurassic). Calc-alkaline granitoids (158.6 to 152.9
Ma, Late Jurassic), which intruded the ophiolites, show subduction-related trace element
signatures (high LILE, low HFSE). Their low Sr and high Nd isotope ratios exclude their
formation due to obduction-induced melting of metasediments deposited on the continental
margin, or in a collisional to post-collisional setting. The Sr and Nd isotope ratios overlap with
that of the ophiolites and indicate that the calc-alkaline series was formed in an island arc setting
with none to only a very minor contribution of subducted sediment to the mantle source. Due
to a lack of substantial crustal input, the island arc series must have already been emplaced in
the ophiolites before the entire sequence was obducted onto the continental margin. Therefore,
the age of the youngest island-arc granitoid, ~153 Ma (Late Kimmeridgian), yields an estimate
for the maximum obduction age for the South Apuseni ophiolites.
iii
Abstract
iv
Zusammenfassung
Zusammenfassung
Der Apuseni-Banat-Timok-Srednogorie (ABTS) magmatische Gürtel in Südosteuropa
entstand in der späten Kreide in Folge von nordwärts gerichteter Subduktion des Neotethys
Ozeans unter den europäischen Kontinentalrand. Einige der größten porphyrischen und
epithermalen Kupfer-Gold Lagerstätten Europas sind mit den magmatischen Gesteinen dieses
Gürtels verbunden. Allerdings sind die Erzlagerstätten nicht gleichmäßig verteilt und
porphyrische und epithermale Lagerstätten treten nur in den zentralen Segmenten des
magmatischen Gürtels auf. Nach Intrusion der magmatischen Gesteine wurden die
lithosphärischen Platten des Kontinentalrands großräumig deformiert. Dies führte zu einer
Deformation des magmatischen Bogens in seine heutige L-Form und erschwert die
Interpretation der ursprünglichen Konfiguration. Die zeitliche und geochemische Entwicklung
dieses Magmatismus in den zentralen und westlichen Segmenten sind bereits gründlich
untersucht, aber für die nördlichsten Banat und Apuseni Segmente gibt es nur spärliche
Informationen.
Neue
geochemische
Haupt-und
Spurenelementanalysen,
radiogene
Isotopendaten (Sr, Nd) von Gesamtgesteinen, sowie U-Pb Alter von Zirkonen von
magmatischen Gesteinen aus den Banat und Apuseni Segmenten vervollständigen den
bestehenden Datensatz des ABTS Gürtels. Kombination der neuen Daten mit den bereits
publizierten Daten ermöglicht eine Rekonstruktion der magmatischen, geodynamischen und
metallogenetischen Entwicklung des gesamten ABTS Gürtels.
Im Apuseni Gebirge treten zwei weitere Phasen von kalk-alkalinem Magmatismus auf.
Eine Phase von miozänem Magmatismus überlagert den in der Kreide gebildeten
Kontinentalrand-Magmatismus und jurassische kalk-alkaline Granitoide sind mit obduzierten
jurassischen Ophioliten assoziiert. Obwohl der miozäne Magmatismus Ähnlichkeiten mit
Subduktions-bezogenem Magmatismus zeigt und Erzlagerstätten beinhaltet, die man
normalerweise an Kontinentalrändern oder in Inselbögen findet, wurde er dennoch nicht in
einem typischen Subduktionszonen-Szenario gebildet. Einige Autoren vermuteten, dass der
Miozäne Magmatismus in Folge von extensionsbedingtem Wiederaufschmelzen einer zuvor
durch Subduktion modifizierten Quellregion (‚source‘) gebildet wurde. Die räumliche
Überlappung des Magmatismus der späten Kreide und des Miozän legt eine gemeinsame
Quellregion nahe. Die neuen Erkenntnisse und geochemischen Daten des Kreide Magmatismus
v
Zusammenfassung
ermöglichen nun einen möglichen Zusammenhang zwischen diesen beiden Phasen zu
untersuchen.
Die jurassischen Ophiolite und kalk-alkaline Magmenserie sind petrographisch und
geochemisch gut charakterisiert. Radiogene Isotopendaten (Sr, Nd) können aber wichtige
zusätzliche Informationen über krustale Kontamination der kalk-alkalinen Granitoide liefern
und damit die Interpretation des tektonischen Szenarios beeinflussen. Radiogene Isotopendaten
wurden mit U-Pb Zirkon Altern kombiniert, um das tektonische Szenario der kalk-alkalinen
Granitoide erneut zu untersuchen und ein maximales Alter für die Obduktion der jurassischen
Ophiolite und kalk-alkalinen Granitoide im Apuseni Gebirge abzuschätzen.
Diese Arbeit besteht aus drei Kapiteln, die in Form von wissenschaftlichen
Manuskripten zur Veröffentlichung in Fachzeitschriften geschrieben wurden. Das erste Kapitel
wurde in ‚Tectonics‘ publiziert. Die Arbeit enthält auch eine Einleitung in die breitere Thematik
und ein abschließendes Kapitel mit Vorschlägen für zukünftige Forschung.
Das erste Kapitel zielt darauf ab, Rekonstruktionen der ursprünglichen Geometrie des
ABTS magmatischen Gürtels zu verbessern und seine tektonische Entwicklung in der späten
Kreide besser einzugrenzen. Der magmatische Gürtel kann in fünf geochemisch
unterschiedliche, teils mineralisierte, teils nicht mineralisierte Segmente eingeteilt werden: (1)
Apuseni, (2) Banat, (3) Timok, (4) Panagyurishte und (5) Ost Srednogorie. Spurenelemente und
Isotopensignaturen der Magmen deuten auf eine durch Subduktion angereicherte Quelle in allen
Segmenten und variable Kontamination der Magmen durch kontinentale Kruste. Der
magmatische Gürtel war 25 Ma (92.2 – 66.8 Ma) lang aktiv. Eine systematische Variation der
Alter und Isotopendaten quer zur Längserstreckung des magmatischen Gürtels tritt in fast allen
Segmenten auf. Zunehmend jüngere Alter in Richtung der Tiefseerinne deuten darauf hin, dass
die subduzierende ozeanische Platte der Neotethys kontinuierlich steiler wurde und sich von
der darüber liegenden europäischen Platte fort bewegte. Das Versteilen der ozeanischen Platte
erhöhte den Fluss in der Asthenosphäre der darüber liegenden europäischen Platte, was sich in
mit der Zeit abnehmenden
87
Sr/86Sr (0.70577 to 0.70373) und zunehmenden
143
Nd/144Nd
(0.51234 to 0.51264) Verhältnissen in manchen Segmenten widerspiegelt. Großräumige
Scherzonen und damit verbundene ‚strike-slip‘ sedimentäre Becken treten in den Panagyurishte
und Timok Segmenten gleichzeitig mit dem Magmatismus auf. Das weist darauf hin, dass in
diesen zentralen Segmenten milde Transtension vorherrschte, während das tiefmarine comagmatische Becken im Ost Srednogorie Segment durch starke othogonale Extension entstand.
Porphyrische und epithermale Kupfer Lagerstätten bildeten sich ausschliesslich in den
zentralen Segmenten, die nur milder Transtension während zeitgleicher Scherung ausgesetzt
vi
Zusammenfassung
waren. Dies begünstigte hoch-Druck Fraktionierung in der Unterkruste, die zur Ausbildung von
‚Adakit-ähnlichen‘ Signaturen führte, und Akkumulation von Volatilen und Metallen. Späteres
Umbiegen und damit zusammenhängende Extension verbargen die relativ einfache Geometrie
des magmatischen Gürtels, ermöglichten aber die Erhaltung der oberflächennahen Lagerstätten.
Das zweite Kapitel untersucht einen genetischen Zusammenhang zwischen dem Kreide
und dem miozänen kalk-alkalinen Magmatismus im Apuseni Gebirge. Der Kreide
Magmatismus im Apuseni Gebirge ist silizischer als der subduktionsbezogene Magmatismus
in den anderen Segmenten, und rhyolitische Ignimbrite treten nur in diesem Segment auf. Die
Mantelschmelzen waren vermutlich einer polybaren Entwicklung in der mittleren bis oberen
Kruste ausgesetzt. Niedrige Sr/Y Verhältnisse, abnehmende Sr Gehalte und Dy/Yb
Verhältnisse
deuten
auf
Fraktionierung
einer
Plagioklas-
und
Amphibol-haltigen
Mineralvergesellschaftung, die möglicherweise in der mittleren Kruste stattfand (~20 km).
Hohe
87
Sr/86Sr80Ma Verhältnisse (bis zu 0.716278) und niedrige
143
Nd/144Nd80Ma Verhältnisse
(0.152187) können durch Beimengung von bis zu 60% partieller Schmelzen kontinentaler
Kruste erklärt werden. Andesitische bis dazitische Schmelzen stiegen dann in seichtere
Krustenniveaus (<8 km) auf, wo sie sich zu rhyolitischen Schmelzen weiter entwickelten. Der
explosive Vulkanismus könnte einen schnellen Verlust von volatilen Phasen verursacht haben,
wodurch die Bildung porphyrischer Lagerstätten verhindert wurde. Im Miozän schmolz dann
der durch Subduktion veränderte Mantel erneut auf, was zur Bildung von kalk-alkalinem
Magmatismus und damit verbundenen epithermalen Au-Te und porphyrischen Cu-Au
Lagerstätten führte. Zwei Gruppen miozäner Magmen mit unterschiedlichen
87
Sr/86Sr
Verhältnissen und Spurenelementsignaturen wurden wahrscheinlich durch magmatische
Differentiation in unterschiedlichen Niveaus in der Kruste gebildet. Eine Gruppe mit hohen
87
Sr/86Sr Verhältnissen (0.706529-0.707596) assimilierte lokale Kruste und erlebte fraktionierte
Kristallisation vermutlich in der mittleren bis oberen Kruste. Eine Gruppe mit tiefen 87Sr/86Sr
Verhältnissen (0.703819-0.705431) zeigt starke Anreicherung in Sr, Ba und La und hat ‚Adakitähnliche‘ Spurenelementsignaturen (hohes Sr/Y und La/Yb). Diese Signaturen erwarben die
Magmen möglicherweise durch Aufnahme von niedrig gradigen partiellen Schmelzen
mafischer Kumulate, die sich während des Kreide Magmatismus in der Kruste gebildet hatten.
Beide Gruppen miozäner Magmen sind ungewöhnlich Gold-reich, was durch das erneute
Aufschmelzen von Au+(Cu, Te)-reichen Sulfiden erklärt werden könnte, die nach
Schmelzextraktion in der späten Kreide im Mantel zurückgelassen worden waren.
Das dritte Kapitel zielt darauf ab die Zeitabfolge und das tektonische Milieu von
jurassischen Ophioliten und kalk-alkalinen Magmen im Apuseni Gebirge zu erörtern. Die Süd
vii
Zusammenfassung
Apuseni-Ophiolite sind das nördlichste Vorkommen von Ophioliten, die heute die kontinentale
Dacia Einheit in Rumänien, Serbien, Mazedonien und Griechenland überlagern. Die Ophiolite
zeigen Affinität zu mittelozeanischen Rücken Basalten, aber leichte Anreicherungen von Th
und U und eine Verarmung an Nb könnte auf ihre Entstehung in einem randlichen oder ‚backarc‘ Becken hindeuten. Vier Gabbros aus den Ophioliten haben U-Pb Alter zwischen 158.9 und
155.9 Ma (später Jura). Kalk-alkaline Granitoide (158.6 – 152.9 Ma, später Jura), die die
Ophiolite intrudieren, haben Subduktions-bezogene Spurenelementsignaturen (hohe LILE,
tiefe HFSE). Die tiefen Sr und hohen Nd Isotopenverhältnisse schließen ihre Bildung aufgrund
von Obduktions-bedingtem Schmelzen von Metasedimenten des Kontinentalrandes oder
aufgrund von Kollision aus. Die Sr und Nd Isotope überlappen mit den denen der Ophiolite und
deuten darauf hin, dass die kalk-alkalinen Magmen in einem Inselbogen gebildet wurden. Nur
sehr geringe Mengen an subduzierten Sedimenten gelangten in diesem Milieu in die
Mantelquelle. Da kaum krustales Material die Magmenquelle kontaminierte, muss die
Inselbogenserie bereits die Ophiolite intrudiert haben, bevor sie gemeinsam mit diesen auf den
Kontinentalrand obduziert wurde. Darum ermöglicht der jüngste Inselbogengranitoid mit einem
Alter von ~153 Ma (spätes Kimmeridgium), das maximale Alter der Obduktion der Süd
Apuseni-Ophiolite abzuschätzen.
viii
Contents
Abstract ...................................................................................................................................... i
Zusammenfassung .................................................................................................................... v
1. Introduction ......................................................................................................................... 1
2. Tectonic, magmatic and metallogenic evolution of the Late Cretaceous arc in the
Carpathian-Balkan orogen ................................................................................................. 5
2.1 Abstract ............................................................................................................................. 5
2.2 Introduction....................................................................................................................... 6
2.3 Regional Geology ............................................................................................................. 8
2.3.1 Tectonic Units of the Carpathian-Balkan Orogen ..................................................... 9
2.3.2 Prearc Nappe Assemblage and Postarc Tectonic Modifications ............................. 10
2.3.3 The Late Cretaceous Magmatic Arc ........................................................................ 12
2.4 Results............................................................................................................................. 13
2.4.1 Sample Selection and Compilation of Data ............................................................. 13
2.4.2 Geochemical Results................................................................................................ 14
2.4.3 Age Constraints........................................................................................................ 19
2.5 Discussion ....................................................................................................................... 21
2.5.1 Tectonic Significance of Magma Geochemistry and Magmatic Ages .................... 22
2.5.2 Tectonic Significance of Comagmatic Sedimentary Basins and Shear Zones ........ 24
2.5.3 Ore Deposits: Regional Stress Regime and Preservation ........................................ 26
2.5.4 Reconstruction and Tectonic Model for the ABTS Belt.......................................... 28
2.6 Summary and Conclusions ............................................................................................. 34
3. The link between Late Cretaceous and Miocene magmatism in the Apuseni
Mountains, Romania ......................................................................................................... 37
3.1 Abstract ........................................................................................................................... 37
3.2 Introduction..................................................................................................................... 38
3.3 Geological Setting of the Apuseni Mountains ................................................................ 40
3.4 Results............................................................................................................................. 43
3.4.1 Field Relations and Description of the Late Cretaceous Samples ........................... 44
3.4.2 Petrography .............................................................................................................. 45
3.4.3 Major and Trace Element Characteristics ................................................................ 47
3.4.4 Radiogenic Isotope Characteristics .......................................................................... 50
3.5 Discussion ....................................................................................................................... 51
3.5.1 Late Cretaceous Subduction: Magmatic Preparation of the Lithosphere ................ 51
3.5.2 Origin and Evolution of the Miocene Magmatism .................................................. 56
3.5.3 High Pressure Fractional Crystallization versus Cumulate Melting ........................ 57
3.5.4 Implications for the Miocene Mineralization .......................................................... 59
3.6 Summary and Conceptual Model ................................................................................... 60
4. Tectonic significance of new U-Pb ages of Jurassic Ophiolites and associated
Granitoids in the South Apuseni Mountains, Romania ................................................. 65
4.1 Abstract ........................................................................................................................... 65
4.2 Introduction..................................................................................................................... 66
4.3 Geological Setting of the Apuseni Mountains ................................................................ 68
4.4 Methods .......................................................................................................................... 70
4.5 Results............................................................................................................................. 71
4.5.1 Sampling, Field Observations and Petrography ...................................................... 71
4.5.2 Major and Trace Element Characteristics ................................................................ 73
4.5.3 Sr and Nd Isotopes ................................................................................................... 75
4.5.4 In situ U-Pb LA-ICP-MS Zircon Dating ................................................................. 76
4.6 Discussion ....................................................................................................................... 77
4.6.1 Tectonic Significance of the Geochemical Data...................................................... 78
4.6.2 Tectonic Significance of the U-Pb Zircon Ages ...................................................... 80
4.6.3 Geodynamic Model.................................................................................................. 82
4.7 Conclusions..................................................................................................................... 84
5. General Conclusions and Outlook .................................................................................... 87
References ............................................................................................................................... 91
Acknowledgements ............................................................................................................... 102
Appendices ............................................................................................................................ 103
Appendix 1 Methods ........................................................................................................ 104
Appendix 2 Sample number, region, coordinates, type, lithology, mineralogy .............. 108
Appendix 3 Major and trace element composition of the studied samples ..................... 112
Appendix 4 Sr and Nd isotope data ................................................................................. 121
Appendix 5 Calculated mean 206Pb/238U ages .................................................................. 124
Appendix 6 Map and sampling locations for the Banat region and Apuseni Mountains 126
Appendix 7 Tectonic map of the ABTS belt summarizing the crystallization ages ........ 128
Appendix 8 Concordia and weighted mean 206Pb/238U age plots for LA-ICP-MS and
TIMS dating ................................................................................................. 129
Appendix 9 U-Pb dates of inherited zircons .................................................................... 139
Appendix 10 Hf isotope plots ............................................................................................ 140
Curriculum Vitae ................................................................................................................. 144
1.Introduction
1.Introduction
Continental magmatic arcs form along active subduction zones where a continental
margin overrides a subducting oceanic plate. Prime examples are the still active arcs along the
Andean margin, the extinct arcs in the North American Cordillera and mostly extinct magmatic
arcs along the Eurasian continental margin. Continental magmatic arcs are often segmented
along a subduction zone due to different styles of tectonic deformation, differences in preexisting geology, convergence rate and direction or heterogeneities within the subducting plate
[e.g. Hildreth and Moorbath, 1988; Kay et al., 1999]. More precisely, along-arc differences
amongst segments have been attributed to subducting ridges [e.g. von Huene and Ranero,
2009], slab tear [Wortel and Spakman, 2000; Rosenbaum et al., 2008], periods of flat
subduction of young oceanic lithosphere [Kay and Coira, 2009; Ramos and Folguera, 2009],
or interaction with different basement crustal domains [Wörner et al., 1992; Mamani et al.,
2008]. Aditionally, across-arc variations may occur due to the shallowing or steepening of the
subducting slab [Haschke et al., 2002; Trumbull et al., 2006]. The different processes occurring
in segmented arcs may show in the geochemical signatures of arc magmas, which are
accordingly modified by mineral fractionation and crustal assimilation processes in the mature
continental crust [de Paolo, 1981; Hildreth and Moorbath, 1988; Kay et al., 1999; Annen et al.,
2006]. Subduction related magmatic arcs are commonly associated with magmatichydrothermal porphyry-style Cu±Au±Mo and epithermal Au±Ag±Cu deposits [Sawkins, 1972;
Sillitoe, 1972; 2010]. These deposits are usually restricted to segments that experienced
compressional stress and do not extend along the entire length of magmatic arcs [Camus, 2002;
Richards, 2003; Rohrlach and Loucks, 2005; Sillitoe and Perelló, 2005; Sillitoe, 2010].
The Apuseni-Banat-Timok-Srednogorie (ABTS) magmatic arc formed at the European
continental margin during subduction of Neotethys ocean in the Late Cretaceous, and extends
over 1000 km from Romania, through Serbia and Bulgaria, to the Black Sea [e.g. Berza et al.,
1998; Popov et al., 2002]. The ABTS magmatic arc is divided into five segments, (1) the
Apuseni, (2) the Banat, (3) the Timok, (4) the Panagyurishte, and (5) the Eastern Srednogorie
segment, that show distinct magmatic and mineralization trends. The arc was intensely
deformed after its emplacement on a lithospheric scale [e.g. Fügenschuh and Schmid, 2005],
which makes the reconstruction and interpretation of arc magmatism and the associated
geotectonic setting more difficult. The timing and evolution of the arc magmatism are well
1
1.Introduction
studied in the central and eastern segments [von Quadt et al., 2005; Peytcheva et al., 2008;
Georgiev et al., 2009; Kouzmanov et al., 2009; Peytcheva et al., 2009; Georgiev et al., 2012;
Kolb et al., 2013], but information on the northernmost Banat and Apuseni segments is still
scarce. New U-Pb ages and geochemical whole rock data for the Banat and Apuseni segments
complete the data set collected during years of research within the ‘fluids and mineral resources’
group at ETH Zurich. Using the complete dataset and tectonic constraints, we aim to improve
reconstructions of the Late Cretaceous configuration and refine the tectono-magmatic history
of the ABTS magmatic arc to identify causes for the differences amongst the arc segments.
Calc-alkaline magmatism akin to arc magmatism and associated Au-rich epithermal and
porphyry-style deposits can also form in tectonic settings that are not related to a
contemporaneously active subduction zone [e.g. Richards, 2009; Richards, 2011]. In such postsubduction settings, the magmas are sourced from the lithospheric mantle or lower crust, which
have been metasomatised during a preceding subduction-process [Haschke and Ben-Avraham,
2005; Richards, 2009; Shafiei et al., 2009; Hou et al., 2015]. Post-subduction lithospheric
thickening, lithospheric extension or mantle lithosphere delamination [e.g. Richards, 2009]
may trigger re-melting of the source region. Tens to hundreds of millions of years may elapse
between subduction-related metasomatism and later re-melting of the lithospheric source
[Pettke et al., 2010]. This setting is found in the Apuseni Mountains, where post-subduction
Miocene calc-alkaline magmatism is superimposed on the barren Late Cretaceous arc
magmatism. The Miocene post-subduction magmas are associated with unusually Au- and Terich epithermal Au-Ag-Te and porphyry Cu-Au deposits [Udubasa et al., 1992; Alderton and
Fallick, 2000; Kouzmanov et al., 2005a; Kouzmanov et al., 2005b]. The formation of the
Miocene magmatism has been related to extension-induced asthenospheric upwelling and remelting of the pre-metasomatised lithospheric mantle [Seghedi et al., 1998; Rosu et al., 2004;
Neubauer et al., 2005; Harris et al., 2013]. The Late Cretaceous arc magmatism and the
Miocene post-subduction magmas are spatially overlapping, which indicates a common source
region, which was presumably metasomatised during subduction of Neotethys ocean in the Late
Cretaceous [Harris et al., 2013]. The generation of the barren Late Cretaceous and the
mineralized Miocene magmatism is probably linked through processes occurring in the mantle
to lower crust. To test this hypothesis, new geochemical major, trace element and isotopic data
for the Late Cretaceous arc magmatism are combined with previously published data for the
Miocene post-subduction magmatism [Harris et al., 2013].
Besides the Late Cretaceous arc magmatism and the Miocene post-subduction
magmatism, a third phase of calc-alkaline magmatism occurs in the Apuseni Mountains. This
2
1.Introduction
Jurassic calc-alkaline magmatism is associated with the Eastern Vardar ophiolitic unit, which
was presumably obducted onto the Dacia Mega-Unit in Late Jurassic to Early Cretaceous times
[Schmid et al., 2008; Kounov and Schmid, 2013]. The age of the ophiolitic and calc-alkaline
series is poorly constrained, and apart from K-Ar ages, which are readily disturbed, only scarce
Re-Os molybdenite ages [Zimmerman et al., 2008] are available for the calc-alkaline granitoids.
New U-Pb zircon ages might not only better constrain the formation age of the ophiolites and
granitoids, but also the timing of the obduction of the ophiolites. The calc-alkaline granitoids
intrude the ophiolitic series and may have formed in an island-arc setting [Bortolotti et al.,
2002; Nicolae and Saccani, 2003; Bortolotti et al., 2004]. However, calc-alkaline granitoids
might also form during or after collision, or due to obduction-induced melting of continentderived sedimentary material [Barbarin, 1999; Cox et al., 1999; Searle and Cox, 1999].
Recently, Šarić et al. [2009] identified different types of calc-alkaline granitoids in other parts
of the Eastern Vardar ophiolites based on Sr-Nd isotopic data, which are not yet available for
the granitoids in the Apuseni Mountains. We provide new U-Pb zircon ages and Sr-Nd isotope
data for the calc-alkaline granitoids and the ophiolites and refine published tectonic models
[Bortolotti et al., 2002; Ionescu et al., 2009; Kounov and Schmid, 2013; Reiser, 2015].
Organization of the thesis:
The first chapter of this thesis is a broader introduction to the formation of continental
magmatic arcs and associated porphyry-style and epithermal deposits. Moreover, it introduces
the notion that calc-alkaline magmatism can also form in tectonic settings of post-subduction
by the reactivation of a metasomatised source.
The second chapter examines the large-scale tectono-magmatic evolution of the Late
Cretaceous ABTS magmatic arc by combining whole rock geochemical data and U-Pb zircon
ages with tectonic constraints. This chapter was published in ‘Tectonics’.
The third chapter explores a genetic link between Late Cretaceous and Miocene calcalkaline magmatism, which occur in close spatial relationship in the Apuseni Mountains. The
focus of this chapter is resolving the magmatic evolution of the Late Cretaceous and Miocene
magmatism based on petrography and whole rock geochemical data. This chapter is intended
for publication in a petrology-related journal.
The fourth chapter aims at refining the timing and tectonic setting of Jurassic ophiolites
and calc-alkaline series in the Apuseni Mountains using U-Pb zircon ages and geochemical
whole rock data, and will be submitted to a journal with a focus on regional geology.
The fifth chapter is an overall conclusions section that also suggests approaches for
future research.
3
1.Introduction
4
Chapter 2
2. Tectonic, magmatic, and metallogenic
evolution of the Late Cretaceous arc in the
Carpathian-Balkan orogen
D. Gallhofer, A. von Quadt, I. Peytcheva, S.M. Schmid, C.A. Heinrich
Tectonics, 34(9), 1813-1836, doi: 10.1002/2015TC003834
2.1
Abstract
The Apuseni-Banat-Timok-Srednogorie Late Cretaceous magmatic arc in the
Carpathian-Balkan orogen formed on the European margin during closure of the Neotethys
Ocean. It was subsequently deformed into a complex orocline by continental collisions. The
Cu-Au mineralized arc consists of geologically distinct segments: the Apuseni, Banat, Timok,
Panagyurishte, and Eastern Srednogorie segments. New U-Pb zircon ages and geochemical
whole rock data for the Banat and Apuseni segments are combined with previously published
data to reconstruct the original arc geometry and better constrain its tectonic evolution. Trace
element and isotopic signatures of the arc magmas indicate a subduction-enriched source in all
segments and variable contamination by continental crust. The magmatic arc was active for 25
Myr (~92–67 Ma). Across-arc age trends of progressively younger ages toward the inferred
paleo-trench indicate gradual steepening of the subducting slab away from the upper plate
European margin. This leads to asthenospheric corner flow in the overriding plate, which is
recorded by decreasing 87Sr/86Sr (0.70577 to 0.70373) and increasing 143Nd/144Nd (0.51234 to
0.51264) ratios over time in some segments. The close spatial relationship between arc
magmatism, large-scale shear zones, and related strike-slip sedimentary basins in the Timok
and Pangyurishte segments indicates mild transtension in these central segments of the restored
arc. In contrast, the Eastern Srednogorie segment underwent strong orthogonal intraarc
extension. Segmental distribution of tectonic stress may account for the concentration of rich
porphyry Cu deposits in the transtensional segments, where lower crustal magma storage and
fractionation favored the evolution of volatile-rich magmas.
5
Chapter 2
2.2
Introduction
Magmatic arcs form above active subduction zones at convergent plate boundaries,
where a continental or oceanic plate margin overrides a subducting oceanic plate. Along a
subduction zone, continental and oceanic arcs generally form distinct segments [e.g., Mahlburg
Kay et al., 1982; Hildreth and Moorbath, 1988]. In continental-margin arcs, the style of tectonic
deformation may differ among segments along the arc and may also vary perpendicular to the
arc in response to differences in preexisting geology, convergence rate and direction, or
heterogeneities within the subducting plate. For example, along-arc differences among
segments have been attributed to subducting ridges [Cross and Pilger, 1982; von Huene and
Ranero, 2009], slab tear [Wortel and Spakman, 2000; Rosenbaum et al., 2008], or flat-slab
subduction of young oceanic lithosphere [Haschke et al., 2002; Kay and Coira, 2009; Ramos
and Folguera, 2009]. Across-arc variations in style and composition of magmatism may be
related to steepening or shallowing of the subducting slab [e.g., Trumbull et al., 2006], which
can be related to changing rates or angles of plate convergence. The composition of subductionrelated mantle magmas may vary as a result of heterogeneous source enrichment, and especially
in continental arcs, will be modified by mineral fractionation and crustal assimilation processes,
which occur primarily in the lower crust and on further ascent through the mature continental
crust [de Paolo, 1981; Hildreth and Moorbath, 1988; Annen et al., 2006]. Igneous geology and
geochemistry can be used to identify magma sources and evolution processes and thus serve as
evidence to interpret the large-scale plate-tectonic setting of complex arcs.
Subduction-related magmatic arcs are frequently endowed with magmatichydrothermal porphyry Cu ± Au ± Mo and epithermal Au ± Ag ± Cu deposits, which can
themselves be taken as tectonic indicators [Sawkins, 1972; Sillitoe, 1972; Groves and Bierlein,
2007]. These deposits usually occur in discrete belts and do not extend along the entire length
of magmatic arcs. Barren and mineralized segments are thought to be due to large-scale
variations in tectonic stress of the lithosphere, and well-endowed segments empirically correlate
with flat-slab subduction, subduction of oceanic ridges, or subduction reversals [Solomon,
1990; Cooke et al., 2005; Rohrlach and Loucks, 2005; Rosenbaum et al., 2005]. Major porphyry
deposits develop preferentially in arc segments that were subjected to a compressional stress
state during ore deposit formation [Sillitoe, 1997; Camus, 2002; Richards, 2003; Rohrlach and
Loucks, 2005; Sillitoe and Perelló, 2005; Sillitoe, 2010]. Horizontal compression can trap
magmas in a lower crustal magma chamber, where high-pressure magmatic differentiation and
cyclic replenishment lead to enrichment in volatiles and metal content. Compression also
influences the development of upper crustal magma chambers, thus preventing volcanic
6
Chapter 2
eruption and unfocused loss of volatiles but favoring focused fluid release through intensely
veined porphyry stocks [Rohrlach and Loucks, 2005; Richards, 2011; Loucks, 2014]. High
magmatic water contents favor the crystallization of hornblende and suppress plagioclase
crystallization in the lower crust [e.g., Burnham, 1979; Lang and Titley, 1998; Richards et al.,
2001; Richards, 2003; Rohrlach and Loucks, 2005; Chiaradia, 2009]. Magmas which have
evolved through these processes therefore have distinct geochemical signatures, generally
referred to as “adakite-like”, which are characterized by high Sr/Y ratios, low Y concentrations,
high light/heavy rare Earth element ratios (LREE/HREE), and weak or absent Eu anomalies
[Kay et al., 1999; Richards et al., 2001; Rohrlach and Loucks, 2005; Richards and Kerrich,
2007; Richards, 2011].
The Eurasian continental margin includes one of the world´s longest magmatic arc
systems [Jankovic, 1997; Perelló et al., 2008; Richards et al., 2012; Richards, 2015], second
only to the circum-Pacific region [Garwin et al., 2005; Sillitoe and Perelló, 2005]. Unlike the
circum-Pacific, which is dominated by long-lasting subduction of oceanic plates below
continents, the magmatic arcs of Eurasia are embedded in the Alpine-Himalayan intracontinental orogenic system (Figure 2.1). Arc magmatism was driven by subduction of the
Neotethys Ocean in Mesozoic to Tertiary times but terminated at the time of collision and was
subsequently heavily overprinted by major collision-related deformations [Dewey et al., 1973;
Schmid et al., 2008]. This collisional overprinting makes the reconstruction and interpretation
of arc magmatism and the associated geotectonic setting more difficult [Sosson et al., 2010;
Bouilhol et al., 2013]. The Late Cretaceous Apuseni-Banat-Timok-Srednogorie (ABTS) belt in
southeastern Europe is the westernmost arc in the Alpine-Himalayan orogenic system related
to the subduction of Neotethys [e.g., Berza et al., 1998; Popov et al., 2002]. This magmatic arc
extends over 1000 km length from the Apuseni Mountains of Romania, through Serbia and
Bulgaria to the Black Sea (Figure 2.1), finding a continuation not discussed in this contribution
all the way to Iran. It was deformed after emplacement on a lithospheric scale [e.g., Neubauer,
2002]. Five segments that show distinct magmatic and mineralization trends can be
distinguished along this arc (Figure 2.2). The timing and evolution of the magmatism and its
associated ore deposits are well studied in the central and eastern segments [von Quadt et al.,
2005; Kamenov et al., 2007; Peytcheva et al., 2008; Georgiev et al., 2009b; Kouzmanov et al.,
2009; Peytcheva et al., 2009; Georgiev et al., 2012; Kolb et al., 2013]. However, information
on the northwestern Banat and Apuseni segments is still scarce [Zimmerman et al., 2008]. This
hampers the establishment of a larger-scale model of the arc that also takes account of the
regional tectonic and geophysical constraints.
7
Chapter 2
Here we present new U-Pb ages, whole rock major and trace element analyses, and Sr
and Nd isotopic data for the northern Banat and Apuseni segments and combine these new
findings with those from previous studies. Comparing the geochemical characteristics along the
arc reveals similarities and differences between arc segments and identifies common magmatic
processes that were active in all segments. We then use our extensive geochronological data
set, to test and improve reconstructions of the Late Cretaceous paleotectonic evolution of the
belt [e.g., Neubauer, 2002; Fügenschuh and Schmid, 2005]. Furthermore, we combine
geochronological data with tectonic constraints derived from comagmatic sedimentary basins
and fault systems to refine the large-scale tectonic history of the ABTS belt.
Figure 2.1: Tectonic sketch map of western Eurasia (modified from Morelli and Barrier [2004]. Major
Late Tertiary to active thrust belts, active subduction zones, and recent arc volcanoes are shown in black;
the sutures of the Neotethys (green) and location of Mesozoic to Oligocene arc magmas (red) are
highlighted: ABTS Apuseni-Banat-Timok-Srednogorie belt, Alborz magmatic arc, Carpathian
magmatic arcs, Eastern Pontide magmatic arc, Kerman belt, Lesser Caucasus magmatic arc, SanandajSirjan magmatic arc, Urumieh-Dokhtar magmatic arc, Yüksekova-Baskil magmatic arc.
2.3
Regional Geology
The Late Cretaceous magmatic arc and associated metallogenic belt crop out in the
Balkan, Northern Rhodopes, South Carpathian, and Apuseni mountain ranges, which generally
rise to less than 2000 m above sea level and are partly obscured by Miocene to recent sediments
of the Pannonian Basin. The subduction-related igneous rocks intruded previously assembled
tectonic units, and they discordantly cross older nappe boundaries (Figure 2.2) (modified from
Schmid et al. [2008]). The arc was intensely deformed after its emplacement and bent around
the Moesian platform to create the present L shape of the arc [e.g., Ratschbacher et al., 1993;
8
Chapter 2
Fügenschuh and Schmid, 2005; Ustaszewski et al., 2008; van Hinsbergen et al., 2008], but this
deformation was not pervasive and was confined to brittle fault structures and ductile shear
zones at the current level of exposure. The complex interplay of compressional and extensional
tectonics which partly predated, and partly overprinted the Late Cretaceous magmatic arc, gave
rise to the distinct segments of ABTS belt: Apuseni, Banat, Timok, Panagyurishte, and Eastern
Srednogorie segments, which we define here based on geographic region and major crustal fault
zones (Figure 2.2).
2.3.1 Tectonic Units of the Carpathian-Balkan Orogen
The Carpathian-Balkan orogen formed due to subduction of oceans and collision of
continental blocks with the European continental margin, which was driven by the overall
convergence between the African and European plates [e.g., Boccaletti et al., 1974; Herz and
Savu, 1974; Csontos and Vörös, 2004; Schmid et al., 2008; Matenco et al., 2010; Schmid et al.,
2011]. The Moesian platform represents the undeformed European foreland and was
amalgamated with other units derived from the European margin (Tisza and Dacia Mega-Units)
in mid-Cretaceous times [Săndulescu, 1984; 1994; Schmid et al., 2008]. The nonophiolitic parts
of the Dinarides became detached from the Adriatic microplate that separated from the African
plate during the Mesozoic within the present-day South Mediterranean realm [Handy et al.,
2010; Marton et al., 2010]. All the continental units that host the ABTS belt are located north
and east of the Europe-Adria suture (Sava zone; see Figure 2.2) and were originally derived
from the European plate. This relationship holds for the Dacia Mega-Unit encompassing the
major units of the South Carpathians which continue into the Balkan orogen in Bulgaria
[Csontos and Vörös, 2004; Schmid et al., 2008] and for the continental Tisza Mega-Unit, poorly
exposed in isolated inselbergs within the Pannonian Basin and in the Apuseni Mountains of
Romania [Csontos and Vörös, 2004; Haas and Pero, 2004; Kounov and Schmid, 2013].
Although the origin of the Strandzha, Circum-Rhodope, and the Rhodope Units is somewhat
ambiguous, they certainly have been part of the European continental margin since at least
Cretaceous times [Okay et al., 2001; Schmid et al., 2008; Burg, 2011]. The Rhodope Unit is
separated from the Dacia Mega-Unit along the right-lateral Maritsa fault system that comprises
several NW-SE trending shear zones, among others the Iskar-Yavoritsa shear zone, which
became active in the Late Cretaceous [Georgiev et al., 2009a; Naydenov et al., 2013].
Two domains of oceanic lithosphere, Neotethys and Alpine Tethys, opened in Triassic
and Jurassic times, respectively. Their remnants denote distinct paleogeographic realms (Figure
2.1). Opening of the Alpine Tethys was kinematically linked to the opening of the central
9
Chapter 2
Atlantic. The main branch of Alpine Tethys forms an oceanic suture in the Alps and Western
Carpathians that links with the Neotethys suture of the Sava zone across the mid-Hungarian
shear zone. A second branch led to the suturing of the Ceahlau-Severin Ocean of the East and
South Carpathians that can only be followed until the Serbian-Bulgarian border (Figures 2.1
and 2.2) [Kräutner and Krstić, 2002; Schmid et al., 2008; Matenco et al., 2010]. The Alpine
Tethys was too narrow to give rise to subduction-related arc magmatism during its closure, and
it was proposed that intrusions in the Alps were related to slab break-off [cf. Davies and von
Blanckenburg, 1995]. Miocene and younger calc-alkaline magmas in the Carpathians are
related to postcollisional extension or lithospheric delamination [de Boorder et al., 1998;
Seghedi et al., 1998; Harangi and Lenkey, 2007; Fillerup et al., 2010; Seghedi and Downes,
2011]. The Neotethys suture is located between units of the European continental margin (Tisza
and Dacia Mega-Units) and the Adriatic margin (Dinarides) [Schmid et al., 2008]. Before final
collision between Europe and Adria, two different parts of the Neotethys were obducted during
the latest Jurassic. The Eastern Vardar ophiolitic sheets were obducted onto the Tisza and Dacia
Mega-Units, while the Western Vardar sheet was emplaced onto the Adria-derived Dinarides
[Csontos and Vörös, 2004; Schmid et al., 2008; Kounov and Schmid, 2013]. Final closure of
the main branch of Neotethys along the Sava Neotethys suture occurred at the end of the
Cretaceous [Pamic, 2002; Karamata, 2006; Ustaszewski et al., 2010], after a long period of
north dipping subduction of the remnants of the Neotethys Ocean, attached to the Adria margin
in the south, underneath the Europe-derived Tisza and Dacia Mega-Units. According to most
authors, it is this originally north dipping subduction of Neotethys that gave rise to the
magmatism in the ABTS belt analyzed in this study. After collision at the end of the Cretaceous,
intracontinental convergence continued in Cenozoic times, which led to the severe oroclinal
bending of the ABTS belt, as seen in Figure 2.2 [Fügenschuh and Schmid, 2005].
2.3.2 Prearc Nappe Assemblage and Postarc Tectonic Modifications
Because the Late Cretaceous ABTS belt cuts across the boundaries between different
tectonic units (Figure 2.2), it is obvious that this magmatic arc of the Carpathian-Balkan orogen
was preceded by several earlier and distinct compressional phases. These compressional phases
partly involved closure of small oceans leading to collision and nappe stacking within the
European continental margin but not to arc magmatism. The north facing Strandzha and
Circum-Rhodope Units were intensely deformed and regionally metamorphosed during a latest
Jurassic to earliest Cretaceous orogeny [Okay et al., 2001; Bonev and Stampfli, 2011]. At about
the same time, parts of the nappe stack of the Rhodope unit formed [Ricou et al., 1998; Bonev
10
Chapter 2
et al., 2006; Burg, 2011], associated with a first event of eclogitization within what is referred
to as the Rhodope Suture Zone or Nestos suture [Krenn et al., 2010; Turpaud and Reischmann,
2010].
Figure 2.2: Geological map of the Carpathian-Balkan orogen, modified from Schmid et al. [2008; 2011],
showing major tectonic units and the occurrences of Late Cretaceous igneous rocks and sedimentary
basins grouped into five segments of ABTS belt. These are, from NW to SE, the Apuseni, Banat, Timok,
Panagyurishte, and Eastern Srednogorie segments. The red bars are the reference lines approximating
the present-day orientation of the arc front in each of the five segments, based on geochronological data
derived in this and previous studies. MF = Maritsa fault system and TF = Timok fault are major
transverse structures used to separate the segments.
A second compressional phase in the Early Cretaceous (~130–100 Ma, “Austrian”
phase) led to subduction of the narrow Ceahlau-Severin Ocean (Alpine Tethys; Figure 2.2), to
collision associated with, in present-day coordinates, northeast facing nappe stacking within the
Dacia Mega-Unit [e.g., Schmid et al., 1998; Iancu et al., 2005] and south directed
synmetamorphic thrusting in the Rhodopes [Burg, 2011]. This event is recorded by 40Ar/39Ar
cooling ages in the Getic-Supragetic, Srednogorie, and Biharia Units [Dallmeyer et al., 1996;
Dallmeyer et al., 1999; Velichkova et al., 2004; Kounov et al., 2010]. A subsequent Cretaceous
11
Chapter 2
tectonic phase affected the Tisza Mega-Unit in the Turonian (~94 to 90 Ma) forming the
present-day nappe stack of the Apuseni Mountains, followed by extension that formed Gosautype postorogenic basins (Figure 2.2) from the Late Turonian onward [Schmid et al., 2008;
Kounov and Schmid, 2013], synchronous with the magmatic activity in the ABTS belt.
Thrusting during the latest Cretaceous was associated with the thrusting of the internal units of
the Dacia nappe stack eastward over the Danubian nappes and with only minor deformation in
the Apuseni Mountains and most parts of the Dacia Mega-Unit (Late CampanianMaastrichtian) [Săndulescu, 1984; Iancu et al., 2005; Schmid et al., 2008; Kounov and Schmid,
2013], which only mildly affected the ABTS belt. However, most of the oroclinal bending of
the Dacia and Tisza Mega-Units, together with the originally straight ABTS belt they host,
occurred in Cenozoic times [Fügenschuh and Schmid, 2005; Ustaszewski et al., 2008] and in
the framework of the invasion of the Tisza and Dacia Mega-Units into the Carpathian
embayment [Balla, 1987; Csontos and Vörös, 2004].
2.3.3 The Late Cretaceous Magmatic Arc
Calc-alkaline magmatism was initiated after the Austrian phase when the Lower
Cretaceous Europe-verging nappe stack preserved within Tisza and Dacia became an upper
plate continental unit above the north dipping subduction zone of the Neotethys lower plate.
The magmatic arc can be divided into five segments (Figure 2.2), based on structurally observed
boundaries and published geological mapping.
The northernmost, presently NNE-SSW trending (1) Apuseni segment extends from
volcanics outcropping within the Tisza Mega-Unit to small plutons emplaced in the Eastern
Vardar ophiolitic unit. The (2) Banat segment is located south of the Apuseni segments and the
Eastern Vardar ophiolitic unit. The mostly plutonic rocks of the Banat segment crop out
exclusively within the Dacia Mega-Unit of the Banat area in Romania. The Danube is the
natural boundary between the Banat segment and the (3) Timok segment. The adjacent (3)
Timok segment differs from the Banat segment in being preserved at a higher erosion level, as
indicated by abundant volcanic rocks and volcaniclastic basins, richly mineralized porphyries,
and the abundant occurrence of adakite-like rocks [Ciobanu et al., 2002; Kolb et al., 2013]. The
Timok and (4) Panagyurishte segments are divided from each other by the Cenozoic rightlateral Timok Fault that links with the Late Cretaceous right-lateral Maritsa fault system (Figure
2.2). All igneous products to the east of the Timok Fault are attributed to the Panagyurishte
segment. The Panagyurishte segment comprises West Bulgarian occurrences and stretches from
the Elatsite deposit in the north to the Rila granitic pluton in the Rhodopes in the south [von
12
Chapter 2
Quadt and Peytcheva, 2005; Peytcheva et al., 2007]. The separation of the Panagyurishte
segment from the (5) Eastern Srednogorie segment is defined by a gap in magmatic activity that
coincides with a reduction of crustal thickness toward the Eastern Srednogorie segment [Yosifov
and Pchelarov, 1977], probably induced during rifting of the Black Sea [Görür, 1988] and
persisting today. (5) The Eastern Srednogorie segment is the only segment that hosts potassiumrich primitive magmas [e.g., Georgiev et al., 2009b] extruded in a marine intraarc rift basin.
The Eastern Srednogorie segment is terminated by the Western Black Sea Fault, which shifted
the Istanbul zone, formerly connected to the Prebalkan, to the south during the opening of the
Black Sea [Okay et al., 1994]. The magmatic arc continues into the Pontides, Lesser Caucasus,
and into northern Iran outside the area of this study [e.g., Jankovic, 1997; Rice et al., 2009;
Sosson et al., 2010].
The central Timok and Panagyurishte segments host economically significant porphyrytype and epithermal Cu-Au deposits [Ciobanu et al., 2002; Heinrich and Neubauer, 2002],
whereas skarn-type (calc-silicate replacement) and polymetallic vein deposits prevail in the
three adjacent segments, the Apuseni, Banat and Eastern Srednogorie segments [Vlad, 1997;
Berza et al., 1998; Popov et al., 2002]. Broadly coeval volcano-sedimentary basins and
sedimentary basins (Gosau-type basins in Romanian literature) occur in all segments along the
magmatic arc [Georgiev et al., 2001; Willingshofer et al., 2001; Popov et al., 2002; Kräutner
and Krstić, 2003; Schuller et al., 2009].
2.4
Results
2.4.1 Sample Selection and Compilation of Data
We collected some 100 samples of Late Cretaceous igneous rocks from the
northernmost Banat and Apuseni segments. The Late Cretaceous plutons, subvolcanic stocks,
dikes, and volcanics are generally undeformed. The majority of samples are fresh and lack
alteration, and weathered surfaces were removed prior to sample processing. Loss on ignition
(LOI) values for the Banat and Apuseni samples range from 0.45 to 6.57 wt %, and 15 out of
the 87 analyzed samples have a LOI higher than 3 wt %, which indicates moderate alteration.
Higher LOI values might affect mobile elements (e.g., Sr, Pb, K, and Ba) but the variations of
these elements are probably primary magmatic, because they do not correlate systematically
with LOI. Whole rock samples were analyzed for major and trace elements and for Sr and Nd
isotope ratios (see Text S1 in the supporting information for details). Additionally, we compiled
geochemical whole rock data from previous studies of the Timok segment [Kolb et al., 2013],
the Panagyurishte segment [Kamenov et al., 2003; Stoykov et al., 2004; Chambefort et al., 2007;
13
Chapter 2
Kamenov et al., 2007; Peytcheva et al., 2008; Kouzmanov et al., 2009; Peytcheva et al., 2009]
and the Eastern Srednogorie segment [Georgiev et al., 2009b]. Major and trace element data
are reported in Table S1, and Sr and Nd isotopic data are reported in Table S2.
Samples showing clear temporal relations in the field, deduced from crosscutting
relations or wall-rock contacts, were preferentially chosen for U-Pb age dating (see Text S1 for
details). Single zircon crystals of 54 Late Cretaceous igneous rocks from the Banat and Apuseni
segments were dated by the laser ablation–inductively coupled plasma–mass spectrometry (LAICP-MS) and/or thermal ionization mass spectrometry (TIMS) method. For both methods,
crystallization ages were calculated from the dates of single zircon crystals and given as
206
Pb/238U ages. Our new data are complemented by U-Pb ages for the Timok segment [Kolb,
2011; Kolb et al., 2013], the Panagyurishte segment [von Quadt et al., 2002; Kamenov et al.,
2003; Stoykov et al., 2004; von Quadt et al., 2005; von Quadt and Peytcheva, 2005; Chambefort
et al., 2007; Peytcheva et al., 2007; Peytcheva et al., 2008; Kouzmanov et al., 2009; Peytcheva
et al., 2009; Atanasova-Vladimirova et al., 2010; Nedkova et al., 2012; Bidzhova et al., 2013]
and the Eastern Srednogorie segment [Georgiev et al., 2012]. Calculated crystallization ages
for all segments are reported in Table S3. A tectonic map of the ABTS belt summarizing the
crystallization ages of the magmatic samples is provided in Appendix 7. LA-ICP-MS single
zircon dates for the Banat and Apuseni segments are reported in Table S4, TIMS data are given
in Table S5. Concordia and 206Pb/238U age weighted average plots are given in Appendix 8. We
report 2 standard deviations (2σ) of overlapping and concordant TIMS or LA-ICP-MS ages in
a population of analyses of each sample, as a conservative measure of age uncertainty, rather
than standard errors of the mean that become unrealistically small in the case of numerous point
analyses [von Quadt et al., 2011].
2.4.2 Geochemical Results
2.4.2.1 Major Elements and Classification
We classify volcanic, shallow intrusive (porphyritic to hypabyssal) and plutonic rocks
from the ABTS belt according to their whole rock geochemistry. In the A = Na2O + K2O, F =
FeO, M = MgO (AFM) plot (Figure 2.3a) [Kuno, 1968; Irvine and Baragar, 1971] nearly all
Late Cretaceous samples from the ABTS belt fall into the field of calc-alkaline rocks, with the
exception of primitive magmas from the Eastern Srednogorie segment, which fall into the
tholeiite field on account of their exceptionally high Fe + Mg. For easier comparison, all
volcanic and intrusive samples were plotted in the same total alkalis versus silica (TAS)
diagram, subdivided and labelled for volcanic rock names only (Figure 2.3b) [Le Maitre et al.,
14
Chapter 2
1989]. Additional information about the igneous rock series can be inferred from the K2O
versus SiO2 diagram (Figure 2.3c) [Rickwood, 1989]. The samples show a wide compositional
range from basalt to rhyolite with a predominance of andesite to dacite magmas. The majority
of Apuseni samples have SiO2 contents over 60 wt %, fall into the dacite and rhyolite fields,
and belong to the high-K calc-alkaline series. The Banat samples are mostly intermediate to
acid rocks and plot in the calc-alkaline and high-K calc-alkaline fields. Timok segment samples
are predominantly intermediate in silica but comprise calc-alkaline to shoshonite series [Kolb
et al., 2013]. The Panagyurishte samples show a wide compositional variation and fall into the
fields of calc-alkaline and high-K calc-alkaline series. Basic to intermediate magmas
predominate in the Eastern Srednogorie segment. Samples from the central part of this segment
are particularly enriched in alkalis (Na2O + K2O) and hence belong to the high-K calc-alkaline
and shoshonite series [Georgiev et al., 2009b].
2.4.2.2 Trace Element Characteristics
The individual arc segments have different concentrations of trace elements which may
be explained by mineral fractionation (see section 2.5.1 for discussion) (Figure 2.4). Light rare
Earth elements (LREE, e.g., La and Ce) are moderately enriched in the ABTS belt samples and
pronounced Eu anomalies are largely absent (Eu/Eu* 0.8–1.2) for the majority of samples with
less than 65 wt % SiO2 (Figure 2.4a). In normal mid-ocean ridge basalt (N-MORB)-normalized
[Sun and McDonough, 1989] trace element plots (Figure 2.4b), samples from all five segments
show enrichment in large-ion lithophile elements (LILEs), such as Ba, K, Sr, and Pb; somewhat
less enrichment in U and Th; and depletion of Nb, Ta, and other high field strength elements
(HFSEs, e.g., Zr and Hf).
The plots of La/Yb versus Yb and Sr/Y versus Y (Figures 2.4c and 2.4d) aim at
distinguishing normal-arc magmas from adakite-like signatures defined by La/Yb and Sr/Y
ratios in excess of 20 and Yb and Y contents below 1.9 and 18 ppm, respectively [Defant and
Drummond, 1993; Richards and Kerrich, 2007]. Apart from some Timok, Banat, and
Panagyurishte samples, most samples plot in the normal-arc field in the La/Yb plot (Figure
2.4c). The Sr/Y plot, however, shows more variation among the arc segments. The majority of
samples from Panagyurishte and Banat segments fall into both fields, close to the field limits;
some Panagyurishte samples with low Y contents fall below the adakite-like Sr/Y. Eastern
Srednogorie and Timok samples partly fall into the normal-arc field but also display adakitelike affinities with higher Sr/Y ratios. Apuseni samples mainly plot in the normal-arc field. Rare
Earth element ratios, (La/Sm)N and (Dy/Yb)N, plotted versus SiO2 can give insights into mineral
fractionation (Figures 2.4e and 2.4f). The (La/Sm)N ratios of ABTS belt samples generally
15
Chapter 2
increase with increasing SiO2; the Apuseni samples show a decrease at SiO2 contents higher
than 70 wt %. (Dy/Yb)N ratios decrease with increasing SiO2 in all arc segments.
Figure 2.3: Chemical classification of Late
Cretaceous igneous rocks from different
segments of the ABTS belt (see inset of
Figure 2.3b). The open symbols denote the
volcanic and shallow intrusive rocks, and
the filled symbols indicate the plutonic
rocks. The larger symbols indicate the
samples related to ore deposits. (a) AFM
diagram showing the boundary between
calc-alkaline and tholeiitic series after
Kuno [1968] (solid line) and Irvine and
Baragar [1971] (stippled line). (b) Total
alkalis versus silica (TAS) diagram [Le
Maitre et al., 1989]. TB: trachybasalt,
BTA: basaltic trachyandesite, TA:
trachyandesite. (c) K2O versus SiO2
classification diagram for subalkalic rocks,
boundary bands after Rickwood [1989].
Analyses in Figures 2.3b and 2.3c were
recalculated to 100% on an H2O-free basis.
16
Chapter 2
Figure 2.4: Trace element characteristics of Late Cretaceous igneous rocks from ABTS belt. (a)
Chondrite-normalized rare Earth elements. All samples from Banat segment are plotted (red lines); for
the other arc segments the overall range of data is represented by lines for the maximum and minimum
values. (b) N-MORB-normalized trace element patterns. Normalizing values for Figures 2.4a and 2.4b
from Sun and McDonough [1989]. Trace element ratios: (c) La/Yb versus Yb. (d) Sr/Y versus Y (ppm)
and the fields for adakite-like and normal-arc magma compositions (from Defant and Drummond
[1993]) and qualitative differentiation paths of various minerals (from Richards and Kerrich [2007]).
(e) (La/Sm)N and (f) (Dy/Yb)N versus SiO2 (wt%) as a measure of fractionation, with mineral
fractionation paths from Davidson et al. [2007].
17
Chapter 2
2.4.2.3 Isotope Geochemistry
Age-corrected Sri isotope ratios and εNdi values show a significant variation among the
arc segments. Nevertheless, individual segments partly overlap. The data lie between the midocean ridge basalt (MORB)-type mantle source field (Timok) and that of Variscan granitoids
(Apuseni) from the prearc European basement [Duchesne et al., 2008]. The Timok samples are
the least radiogenic (highest εNdi: +4.4 to +5.4 and lowest 87Sr/86Sri: 0.70339 to 0.70375) [Kolb
et al., 2013]; the more radiogenic Timok samples overlap with the isotopically primitive Banat
samples. Banat samples extend to highly radiogenic isotopic compositions (low εNdi: -1.8 to 3.1 and high 87Sr/86Sri: 0.70598 to 0.70707) and partly overlap with the basement granitoids.
Panagyurishte samples are also part of this array and overlap with or are slightly more
radiogenic than the Banat samples. However, a few Panagyurishte samples are shifted off this
trend toward more primitive Sri isotope ratios at low εNdi values (-1.7 and lower). Agecorrected Sri isotope ratios (not plotted in Figure 2.5) for the Eastern Srednogorie segment range
from 0.70392 to 0.70590 and overlap with the other segments.
Figure 2.5: Initial Sr and Nd isotope ratios for samples from the Late Cretaceous ABTS belt. Samples
from Banat, Timok, and Eastern Srednogorie segments were backcalculated to 80 Ma based on analyzed
bulk Rb and Sm concentrations; Panagyurishte samples were corrected for 85 Ma and 90 Ma (Table
S2). (a) εNdi versus 87Sr/86Sri. Field for MORB from Stracke et al. [2005], composition of depleted
MORB mantle (DMM) and enriched DMM from Workman and Hart [2005], field for Variscan
granitoids (corrected to 80 Ma) from Duchesne et al. [2008] and Peytcheva et al. [2008], and
composition of Serbian flysch sediments from Prelevic et al. [2008]. Two-component isotopic mixing
for mantle source enrichment between isotopically unmodified, least radiogenic Timok sample (MM38)
[Kolb et al., 2013] and Serbian flysch (06FL03) [Prelevic et al., 2008]. Mixing line for crustal
contamination shows percent addition of representative basement gneiss (R4626) [Duchesne et al.,
2008] to the presumably isotopically unmodified, least radiogenic Timok sample (MM38) [Kolb et al.,
2013]; see text for discussion.
18
Chapter 2
2.4.3 Age Constraints
2.4.3.1 Timing of Magmatic Activity
All arc segments were simultaneously active over the time period of 81 to 75 Ma
(Campanian), but onset and termination of magmatic activity occurred at different times in the
individual segments (Figure 2.6 and Table S3). Magmatic activity along the ABTS belt started
as early as ~92 Ma (Turonian) in the northern Panagyurishte segment [von Quadt et al., 2002;
Stoykov et al., 2004; Chambefort et al., 2007; Kamenov et al., 2007]; it set in slightly later,
between 90 and 87 Ma, in the Timok and Eastern Srednogorie segments [Kolb, 2011; Georgiev
et al., 2012]. In the Banat and Apuseni segments, magmatism apparently did not start before
~84 and ~81 Ma, respectively. The youngest igneous products occur in the Banat and Timok
segments (74 to 71 Ma). However, still younger magmatism is found in the Rhodope Unit,
which forms the southern parts of the Panagyurishte and Eastern Srednogorie segments (67-71
Ma) [von Quadt and Peytcheva, 2005; Marchev et al., 2006; Peytcheva et al., 2007]. Our
compilation includes only the Late Cretaceous magmatism in the Rhodope Unit, although
magmatism continued beyond the Cretaceous-Palaeogene boundary, used as a somewhat
arbitrary limit of our compilation in this area.
2.4.3.2 Across-Arc Age Progression and Isotope Variation
Some arc segments show distinct across-arc age variations, whereby geochemical
distribution patterns vary in correlation with the age of magma emplacement. However, because
trace elements might be affected by differentiation, we focus on isotopic variations only. We
use the red reference lines approximating the orientation of the arc front in each segment (Figure
2.2) to sort the ages and isotopic compositions within the arc segments. We start with samples
close to these lines and move perpendicular to the reference line toward northwest in the
Apuseni segment, to the west in the Banat and Timok segments, to the southwest in the
Panagyurishte segment, and to the south in the Eastern Srednogorie segment (Figure 2.6, from
left to right).
The Timok and Pangyurishte segments show the most pronounced progression of
magmatic activity (Figure 2.6), which is also clearly visible in the age-distribution map included
in Appendix 1 of this paper. The oldest ages occur closest to the interpreted arc magmatic front;
in each segment, the youngest magmatic ages are found furthest away from the arc front [von
Quadt et al., 2005; Kolb et al., 2013]. In the Eastern Srednogorie segment, ages show no clear
overall trend, but the oldest intrusions tend to occur in the southern periphery (Strandzha Unit)
of the region exposing overall younger, submarine extrusive rocks [Georgiev et al., 2012].
19
Chapter 2
Figure 2.6: (a) Compilation of 206Pb/238U ages (in Ma) for ABTS belt and Rhodopes (see Table S3 for
sources). The five segments of ABTS belt are arranged along the arc from northwest (Apuseni) to
southeast (Eastern Srednogorie). Within the individual segments, data points are sorted in terms of
increasing distance in km from the inferred arc front (red reference lines in Figure 2.2), which
corresponds to a sorting from paleo-north to paleo-south after restoration of the L-shaped bend in the
magmatic arc (see text for explanation). The grey bands indicate the more restricted intervals of
magmatic-hydrothermal Cu ± Au mineralization in the Timok and Panagyurishte segments. The green
frames indicate the sedimentation intervals in comagmatic basins. (b) Corresponding across-arc trends
of 87Sr/86Sri ratios. This compilation also includes 87Sr/86Sri ratios of samples for which no U-Pb ages
are available. Two Apuseni samples (DG094 and DG112) fall outside the plotted Sr isotope ratio range,
further emphasizing the unusually wide range in this segment. (c) Corresponding across-arc trends of
143
Nd/144Ndi ratios.
20
Chapter 2
Magmatism in the Rhodope Unit (67–71 Ma) [von Quadt and Peytcheva, 2005;
Marchev et al., 2006; Peytcheva et al., 2007] is interpreted as the continuation of the magmatic
activity in the Panagyurishte and Eastern Srednogorie segments. Within the Banat segment
younging of magmatism toward the west from 83.9 to 70.2 Ma is only observed in the area that
is currently adjacent to Serbia. By contrast, two samples from the northern part located close to
the inferred arc front yield younger ages (75–76 Ma), and the large Bocşa pluton and close-by
smaller intrusions that occur furthest away from the inferred arc front yield distinctly older ages
(76 to 80 Ma). In the Apuseni segment, where the youngest ages (75.5 to 78 Ma) are found
closest to the inferred arc front, the trend in age progression appears to be reversed.
The across-arc age trends are accompanied by distinct geochemical changes, which are
most clearly traced by isotopic signatures. The Panagyurishte magmatism evolves toward lower
87
Sr/86Sri ratios and higher εNdi south-wards, and the most depleted isotopic compositions are
found in the youngest samples [von Quadt et al., 2005]. A similar evolution to more primitive
Sr and Nd isotopic ratios away from the arc front is found in the Banat segment. Magmatism in
the Timok segment, however, evolves towards more radiogenic Sr and less radiogenic Nd
isotope ratios away from the arc front, as well as in time, but is briefly interrupted by a shift to
lower 87Sr/86Sri ratios and higher εNdi at ~83 Ma [Kolb et al., 2013]. The 87Sr/86Sri ratios in the
Eastern Srednogorie segment are relatively low and vary little in the central rifted basin,
whereas higher ratios are found in the southern part (Strandzha Unit), which includes older
intrusives in this segment [Georgiev et al., 2009b]. No clear isotopic trend is observed in the
Apuseni samples. Some volcanic rocks have particularly high
87
Sr/86Sri ratios, whereas the
youngest rocks coincide with relatively low 87Sr/86Sri ratios.
2.5
Discussion
This discussion aims at supporting and refining the idea that the geochemical
characteristics of the different arc segments are best explained by a single north dipping
subduction zone of the Neotethys Ocean (also referred to as the “Vardar” Ocean) that was active
during the Late Cretaceous and was closed by collision of the Adriatic microcontinent with the
European plate [e.g., von Quadt et al., 2005; Georgiev et al., 2012; Kolb et al., 2013].
Subduction induced magmatism within this northerly adjacent continental upper plate but also
created distinct magmatic and metallogenetic characteristics for each segment of the magmatic
arc. First, we will discuss the tectonic significance of our compilation of geochemical and
geochronological data. Next, we will refine the tectonic reconstruction of this arc (Figure 2.7),
21
Chapter 2
prior to its large-scale bending into the present-day L shape by deformations occurring during
and after continental collision.
2.5.1 Tectonic Significance of Magma Geochemistry and Magmatic Ages
Subduction-related mantle melts that ascend through mature continental crust undergo
fractionation and crustal contamination, which modifies their original composition [de Paolo,
1981; Hildreth and Moorbath, 1988; Thirlwall et al., 1996]. Therefore, tracing the primary
source(s) of magmas in an evolved arc is difficult. Magmatism in all segments of the ABTS
belt shows a clear “arc signature”, i.e., enrichment in LILE (Ba, K, Sr, and Pb) and depletion in
HFSE (Nb, Ta, Zr, and Hf; Figurse 2.4a and 2.4b), which is also prominent in the less evolved
basaltic andesites and indicates a subduction-related origin of the magmas.
The observed isotopic compositions in the ABTS arc magmas indicate interaction with
continental crust (Figure 2.5). Theoretically, some crustal contamination might have already
taken place in the mantle source via addition of subducted sediment [Elliott, 2003]. In a
continental arc, however, evolved isotopic compositions are more likely acquired during
interaction and assimilation of the local continental crust [Hildreth and Moorbath, 1988;
Wörner et al., 1992]. The local continental crust [Duchesne et al., 2008] and subducted
sediments [Prelevic et al., 2008] are isotopical rather similar in this region, which impede any
further distinction, but most of the contamination might have occurred during storage and
melting, assimilation, storage, and homogenization processes (MASH) [Hildreth and
Moorbath, 1988] in the continental crust. The ABTS belt magmas form an array between
primitive isotopic compositions similar to MORB (high εNdi and low
87
Sr/86Sri ratios), as
observed for the Timok segment, and more evolved isotopic compositions (low εNdi and high
87
Sr/86Sri ratios), which partly overlap with the field of local Variscan granitoids as observed
for the Apuseni segment. This array can be approximated by admixing of the least radiogenic
sample, which presumably represents the isotopically unmodified mantle in the case of the
Timok segment [Kolb et al., 2013], with varying amounts of basement granitoid. Depending on
the choice of crustal assimilant, 10 to several tens of percent of crustal melt were added to the
mantle-derived parental melt. Assimilation of local crustal basement is also indicated by the
presence of inherited zircons in igneous rocks from most arc segments, recording Jurassic,
Carboniferous, Ordovician, and older crystallization ages [Georgiev et al., 2012; Kolb et al.,
2013].
The ABTS belt magmas have normal-arc as well as adakite-like geochemical signatures
(Figure 2.4). Adakite-like signatures (La/Yb and Sr/Y > 20, Yb < 1.9 ppm, and Y < 18 ppm),
22
Chapter 2
[Defant and Drummond, 1993] are frequently associated with economic porphyry Cu-Au and
epithermal Cu-Au-Mo deposits [e.g., Rohrlach and Loucks, 2005; Richards and Kerrich, 2007]
but probably do not indicate slab melting as suggested in the original interpretation of adakite
by Defant and Drummond [1993]. Instead, adakite-like geochemical signatures indicate
plagioclase-absent fractionation of amphibole at high pressure (> 0.8–1 GPa and 25–30 km)
[e.g., Alonso-Perez et al., 2009] and thus reflect intermediate storage of increasingly hydrous
magmas in the lower crust [Rohrlach and Loucks, 2005; Richards, 2011]. Yttrium as well as
middle rare Earth elements (e.g., Sm and Dy) are preferentially incorporated into amphiboles
[Davidson et al., 2007; Richards and Kerrich, 2007], and the observed increasing (La/Sm)N and
decreasing (Dy/Yb)N ratios in magmas from all arc segments of the ABTS belt are most likely
to indicate widespread hornblende fractionation (Figures 2.4e and 2.4f). Kolb et al. [2013]
proposed that adakite-like signatures next to normal-arc signatures in the Timok magmas were
caused by high-pressure amphibole fractionation in the lower crust, followed by variable
proportions of upper crustal plagioclase fractionation (incorporating Sr) and assimilation of
local basement. This interpretation might also be applicable to adakite-like rocks from the other
segments of ABTS belt. We therefore conclude that all ABTS belt magmas tapped a mantle
wedge source that was enriched by subduction. During ascent, the mantle magmas were
modified to variable degrees by fractionation in the lower and/or the upper continental crust
and varying degrees of assimilation of the local upper crust of dominantly Variscan (Late
Paleozoic) age.
The ABTS belt shows pronounced across-arc age and isotopic variations during a period
of 25 Myr (~92 to 67 Ma, Turonian to Maastrichtian). Although different segments record
different periods of initiation and conclusion of arc magmatism, no major breaks in magmatic
activity are observed within any of the arc segments, and all arc segments were magmatically
active within a common time window of ~83 to ~75 Ma (Figure 2.6). Age progression in
magmatic arcs away from the arc front toward the inferred subduction trench, i.e., narrowing
of the arc-trench gap, is commonly interpreted in terms of progressive steepening of the
subducting slab. Such steepening can be associated with slab rollback; i.e., trench migration
and slab hinge retreat away from the upper plate due to increasing slab pull forces as the age of
the subducted oceanic slab increases, inducing back-arc extension in the upper plate [Heuret
and Lallemand, 2005]. Trends of younging magmatic ages from paleo-north to paleo-south
(that is, toward the paleo-trench), which is pronounced in the case of the Banat, Timok, and
Panagyurishte segments, indicate narrowing of the arc-trench gap by up to 100 km (Figure 2.2),
consistent with steepening of the subducting Neotethys slab and intraarc extension in the upper
23
Chapter 2
plate. The southward migration is particularly pronounced in the Panagyurishte segment and
even more so if the plutons within the northernmost Rhodopes are included. A systematic age
progression is missing in the Eastern Srednogorie segment, where magmatism in a central deep
marine basin (Yambol-Burgas Basin) [Georgiev et al., 2001] yields ages between 81.2 and 78.0
Ma, while older intrusions occur in the southerly adjacent region [Georgiev et al., 2012].
Magmatism in the Apuseni segment neither shows any systematic trend (or even a reversed
trend?) in the age pattern. At the same time, the age range of magmatism in the Apuseni segment
(75.5–80.8), as well as that of the East Srednogorie segment, overlaps with magmatism in the
central segments. This indicates that magmatism was relatively stationary within the upper plate
during the process of subduction at both ends of ABTS belt.
Slab rollback or slab steepening also enhances corner flow within the subcontinental
mantle wedge, associated with asthenospheric upwelling and partial melting [Gvirtzman and
Nur, 1999]. Therefore, an increasing mantle input to the later and more southerly magmas over
time, only traced by isotopic compositions in the case of the Banat and Panagyurishte segments,
is in good agreement with the proposed slab steepening (Figure 2.6b). Many of the Apuseni
magmas, however, show a distinctly higher degree of crustal contamination. In contrast to the
other segments, the Timok magmas originally derived from a mantle-type source with
negligible crustal contamination and show increasing crustal contamination in the younger
magmas only [Kolb et al., 2013]. Locally deviating, strongly crustal-influenced isotope ratios
in the Timok and Eastern Srednogorie segments have been explained by a higher degree of
crustal assimilation, perhaps due to thicker crust or a longer residence time in the crust relative
to magmas from elsewhere in the ABTS belt [Georgiev et al., 2009b; Kolb et al., 2013].
2.5.2 Tectonic Significance of Comagmatic Sedimentary Basins and Shear Zones
Arc magmatism in the ABTS belt is intimately associated with Late Cretaceous
sedimentary basins (Figure 2.2), which provide additional constraints on the state of stress in
different segments of the magmatic arc. Some of them contain volcanic and volcaniclastic
materials, while others lack such material. It is important to note that not all of these Late
Cretaceous sedimentary basins are associated with ABTS belt magmatism. Late Cretaceous
postorogenic basins unrelated to ABTS magmatism are also widespread in the Tisza and Dacia
Mega-Units and are often referred to as Gosau-type basins [Willingshofer et al., 1999; Schuller
et al., 2009]. They formed during collapse of overthickened and gravitationally unstable
continental crust and seal former nappe contacts formed during the Early Cretaceous
(“Austrian”) orogeny. Sedimentation in these older basins typically starts in Albian24
Chapter 2
Cenomanian times [Kounov and Schmid, 2013]. Only the age range of comagmatic basins (94–
72 Ma) whose formation is related to ABTS magmatism rather than orogenic collapse is
indicated in Figure 2.6.
In the northern parts of the Eastern Srednogorie segment, layers of andesitic pyroclastics
and reworked tephra deposits already appear in the Turonian (94–90 Ma) sedimentary sequence
[Nachev and Dimitrova, 1995], but the most pronounced magmatic activity of submarine
extrusive arc magmas coincides with the most intense phase of intraarc extension and crustal
thinning in the Campanian 81–78 Ma [Georgiev et al., 2012]. Magmatism and sedimentation
abruptly stopped at around 72 Ma ago. The volcano-sedimentary basins east of Sofia cover the
entire time span between Late Turonian and the end of Campanian (92–72 Ma) [Popov et al.,
2012]. Sedimentation of volcaniclastic material in the Timok segment west of Sofia starts in
the Turonian and ends before the deposition of Late Campanian to Maastrichtian clastic and
reefal sediments [Banješević, 2010]. These biostratigraphic constraints agree with the ~90 to
79 Ma age interval for igneous rocks associated with the volcaniclastic basin in the eastern part
of the Timok segment [Kolb, 2011]. In contrast to the Eastern Srednogorie Basin, the basins in
the Panagyurishte and Timok segments are marine pull-apart basins, which partly formed
during dextral shearing along the Iskar-Yavoritsa shear zone [Georgiev et al., 2009a; Naydenov
et al., 2013]. This dextral shear zone is syntectonic with felsic and mafic plutons emplaced
during the 86–75 Ma time interval. Hence, pluton emplacement partly overlaps with the age
span of the volcano-sedimentary parts of the Panagyurishte and Timok Basins. This suggests
opening of the Panagyurishte and Timok Basins in a scenario of crustal-scale dextral strike-slip
motion, interpreted in terms of dextral transpression by previous authors [Georgiev et al.,
2009a; Naydenov et al., 2013; Georgiev et al., 2014]. However, a transtensional rather than
transpressional setting is indicated for two reasons: (1) the Iskar-Yavoritsa shear zone is
contemporaneous with the opening of comagmatic transtensional basins [Naydenov et al., 2013]
and (2) the Iskar-Yavoritsa shear zone is confined to the area of deposition of the Panagyurishte
and Timok Basins and has no further continuation to the east. The age of volcaniclastic
sediments deposited in sedimentary basins in the Banat region is not well constrained [Barzoi
and Seclaman, 2010], but magmatic dikes (76–75 Ma) are associated with SantonianCampanian sediments (~84–72 Ma). In the Apuseni segment, moderate intraarc extension
probably caused deepening within the Gosau-type orogenic collapse basins in the Campanian
[Kounov and Schmid, 2013].
In summary, the onset of comagmatic sedimentation in the Late Cretaceous volcanosedimentary basins (Figure 2.6) starts at ~94 Ma in the east (Eastern Srednogorie) and
25
Chapter 2
systematically becomes younger to the west and north (82 Ma in the Apuseni segment) and
coincides with the onset of magmatic activity along the ABTS belt. This extension is therefore
not related to orogenic collapse but rather to the subduction of the Neotethys Ocean triggering
comagmatic basin formation along the entire continental margin of the European upper plate.
Massive orthogonal intraarc extension is indicated for the Eastern Srednogorie segment, while
the tectonic setting in the Panagyurishte and the Timok segments was transtensive and
associated with relatively moderate extension [Georgiev et al., 2009a]. Moderate extension in
the peripheral Apuseni and Banat segments was of shorter duration.
2.5.3 Ore Deposits: Regional Stress Regime and Preservation
The regional stress regime of the crust is considered to be a critical factor in generating
different styles of ore deposits. Giant porphyry Cu-Au deposits in the circum-Pacific arcs
preferentially formed in arc segments that underwent contractional pulses [e.g., Sillitoe and
Perelló, 2005]. Geodynamically induced horizontal compression inhibits propagation of
subvertical dikes and keeps buoyant magmas trapped in sheet-like subhorizontal chambers
[Tosdal and Richards, 2001; Richards, 2003; Rohrlach and Loucks, 2005]. Prolonged storage
of magmas in lower crustal magma chambers is crucial for allowing enrichment in water, other
volatiles, and possibly ore metals [e.g., Richards, 2003; Rohrlach and Loucks, 2005; Richards
and Kerrich, 2007]. Stress relaxation eventually facilitates rapid ascent of fertile hydrous
magma into the upper crust, where volatiles exsolve from the magma to form ore deposits.
Magma ascent is focused in localized ascent paths such as strike-slip faults [Sillitoe and Perelló,
2005]. However, a change to strong extension is not favorable for porphyry deposit formation,
as it would result in volcanic eruption rather than magma storage in the upper crust [Tosdal and
Richards, 2001; Richards, 2003].
In the Late Cretaceous ABTS arc, indications for comagmatic compression have been
reported for the Panagyurishte segment [Naydenov et al., 2013], but crustal-scale dextral
transtension probably prevailed during the activity of the Maritsa fault system and the opening
of pull-apart comagmatic basins in the Panagyurishte and Timok segments. Significant
porphyry-type ore deposits occur only in these two segments, in association with magmas
exhibiting adakite-like trace element characteristics [e.g., von Quadt et al., 2005; Kolb et al.,
2013]. A near-neutral stress state of the crust with mild transtension might also explain why
igneous rocks with adakite-like signatures, derived by lower crustal high-pressure amphibole
fractionation, and normal-arc signatures, obtained by upper crustal assimilation and fractional
crystallization, occur in spatially overlapping areas in the Timok segment [Kolb et al., 2013].
26
Chapter 2
The peripheral Eastern Srednogorie segment, by contrast, was under extreme extension and did
not form significant porphyry-style deposits.
The NNW-SSE alignment of ore deposits in the Panagyurishte segment (between
Elatsite and Elshitsa; see Figure 2.2) [e.g. Popov et al., 2002; Moritz et al., 2004] is conspicuous
and was probably controlled by a deep crustal fault. The fact that the southeastern end of this
linear array abuts the dextral Iskar-Yavoritsa shear zone (which itself contains synkinematic
intrusions of Late Cretaceous age) [Georgiev et al., 2009a] suggests that the linear
Panagyurishte array may follow a tensional fault oriented parallel to the σ1/σ2 plane. First-rank
shear zones parallel to the WNW-ESE trending Iskar-Yavoritsa shear zone and NNW-SSE
tensional faults paralleling σ1 probably focused magma ascent and fluid flow to the sites of ore
deposit formation, as ore deposits in the Panagyurishte segment are primarily found along these
secondary oblique and cross-arc faults [Drew, 2005; Georgiev et al., 2014]. Relationships
between fault zones and ore deposit formation are less evident in the Timok segment but may
be obscured by more intense later tectonic overprint and poor exposure.
Partial preservation of a comparatively old porphyry Cu-Au district like the ABTS belt
depends critically on limited postemplacement uplift and erosion [Groves et al., 2005; Kesler
and Wilkinson, 2006]. This points to postemplacement processes that prevented or counteracted
substantial crustal thickening in the continental host units, which would have favored complete
erosion of the volcano-sedimentary successions, as it currently happens in the Andes where
preserved porphyry deposits are much younger [Sillitoe and Perelló, 2005]. Preservation of
shallow volcanics in the Apuseni, Banat, and Timok segments is probably due to extensional
and transtensional postemplacement processes associated with Eocene to Miocene bending
around the Moesian platform and extrusion of the Carpathian-Balkan orogen into the still open
Carpathian embayment [e.g., Schmid et al., 1998]. Extensional tectonics followed accretion in
the Aegean region from Paleogene times onwards [e.g., Burchfiel et al., 2008] and resulted in
core complex formation and tectonic denudation in the Rhodopes [Burg, 2011; Kaiser
Rohrmeier et al., 2013] (Figure 2.2). Extension may have also caused removal of shallow ore
deposits and volcanics, and exposure of deeper crustal levels in the southern parts of the
Panagyurishte and Eastern Srednogorie segments.
27
Chapter 2
Figure 2.7: Restoration of the configuration at the onset of arc magmatism around 90 Ma ago, partly
based on retrodeformations of Ustaszewski et al. [2008] and Fügenschuh and Schmid [2005] but
modified to restore the magmatic arc to a gently curved line, consistent with paleomagnetic data. TF =
Timok fault, MF = Maritsa fault system. The straight red lines are the reference lines for the individual
arc segments, with the same orientation relative to outcropping magmatic bodies as shown in Figure 2.2.
2.5.4 Reconstruction and Tectonic Model for the ABTS Belt
All reconstructions of the Carpathian-Balkan orogen are speculative to some extent due
to the intense later tectonic overprint in this region [Neugebauer et al., 2001; Csontos and
Vörös, 2004; Stampfli and Borel, 2004; Ustaszewski et al., 2008], but considering the tectonic
scenario that generated the extensive arc magmatism described in this paper provides important
additional constraints. Figure 2.7 shows a modified paleotectonic map of the Carpathian-Balkan
orogen for the Turonian (~90 Ma), indicating the Late Cretaceous position of continental blocks
hosting the arc-related magmatic rocks along a gently curved magmatic arc. The reconstruction
is still consistent with the block rotations proposed by Ustaszewski et al. [2008], but extended
backward in time to indicate the location of a continuous active plate boundary at the northern
28
Chapter 2
margin of Neotethys, such that the observed subduction magmatism can be at least qualitatively
explained.
We used the Miocene restoration by Ustaszewski et al. [2008] as a starting point for the
reconstruction and integrated a previous retrodeformation to the Late Cretaceous situation
proposed by Fügenschuh and Schmid [2005, their Figure 9]. Our restoration is still tentative,
because exact amounts of shortening or extension and changes in the geometric configuration
of the tectonic units cannot be quantified. For this new reconstruction, arc segments were
rotated and translated individually, but outlines of magmatic bodies and their distribution within
the segments, were left unchanged, to facilitate comparison with the present-day configuration.
This is reasonable because contacts between intrusions and wall rocks are generally
undeformed, indicating that displacements were concentrated along larger block faults even
though these segment boundaries could rarely be defined in the field. Significant deformation
within synkinematic igneous intrusions related to ABTS magmatism is directly observed only
along the Iskar-Yavoritsa shear zone [Georgiev et al., 2009a; Naydenov et al., 2013], which
was therefore used as one of the major segment boundaries in our retrodeformation.
The European foreland including Moesia is fixed in its present-day position.
Paleomagnetic data provide major constraints regarding rotation when restoring the Late
Cretaceous situation of the Tisza Mega-Unit of the European continental margin.
Counterclockwise back rotation of the Apuseni Mountains from their respective Miocene
position is required to account for the presently observed total of ~90° clockwise rotation
[Pǎtraşcu et al., 1990; Pǎtraşcu et al., 1992; Panaiotu, 1998; Marton et al., 2007; van
Hinsbergen et al., 2008]. A smaller amount of total rotation was adopted for the Banat and
Timok segments for geometric reasons. Only areas to the west of an inferred link between the
Maritsa fault system and the future Timok-Cerna-Jiu fault system were rotated around a rotation
pole fixed to the European margin, consistent with previous restorations [e.g., Schmid et al.,
1998]. The Getic-Supragetic-Srednogorie Units (Dacia) and the arc segments were shifted to
the south by 50 km to account for Maastrichtian and Cenozoic thrusting of the Srednogorie Unit
with respect to Moesia [Doglioni et al., 1996; Banks, 1997; Stuart et al., 2011]. Our
reconstruction also retrodeforms N-S extension in the Rhodopes that occurred in mid-Eocene
to Miocene times [e.g., Brun and Sokoutis, 2007; Burg, 2011; Kaiser Rohrmeier et al., 2013].
Measured in a N-S direction across the center of the Rhodopian core complex we restored 125
km of extension [Brun and Sokoutis, 2007; van Hinsbergen and Schmid, 2012]. This extension
continuously decreases westward (see rotation model of Brun and Sokoutis [2007]), which
results in an increase in the gap between trench and arc toward west (Figure 2.7). For easier
29
Chapter 2
comparison with Figure 2.2 we left the outlines of the Rhodope and Strandzha Units unchanged,
being aware that substantial portions of the Rhodopes were still covered by Circum-Rhodope
and Strandzha Units in the north, east, and south and by Danubian, Getic, and SerboMacedonian Units in the west during Late Cretaceous times. The intra-Turonian nappe stacking
in the Tisza Mega-Unit [Kounov and Schmid, 2013] was retro-deformed by taking back 150 km
of shortening. In order to achieve a better fit for the magmatic arc, the tectonic units of the
Apuseni Mountains had to be shifted further to the west, compared to the restoration by
Fügenschuh and Schmid [2005].
The width of the Neotethys remnant basin, i.e., the ocean that closed to form the SavaIzmir-Ankara suture between Adria and Europe in the Late Cretaceous [Schmid et al., 2008;
Schmid et al., 2011], is not well constrained and differs widely among published reconstructions
[e.g., Neugebauer et al., 2001; Csontos and Vörös, 2004; Stampfli and Borel, 2004]. Our
interpretation of the ABTS belt as a magmatic arc above a subduction zone places an additional
constraint on the minimum width of Neotethys, because a mature subduction slab must be
established before arc magmatism can start. Transporting hydrated oceanic lithosphere to a
typical depth required to initiate partial melting of the overriding mantle (100-120 km), and
choosing the arc-trench gap of 250 km resulting from our reconstruction at 90 Ma ago, a
minimum slab length of 270 km must have been subducted to trigger the onset of magmatism.
This corresponds to a rather flat subduction zone with a shallow angle of 22–26°. There is
considerable uncertainty on these estimates, and a shorter arc-trench gap or a higher depth of
melting would result in increases or decreases in the subducted slab length, respectively, and a
steeper initial subduction angle at the onset of magmatism. Given the Africa-Europe
convergence rate of 15 km/Ma before 90 Ma, calculated for the N-S direction [Rosenbaum et
al., 2002], subduction of 270 km of oceanic lithosphere would have been initiated some 18 Ma
earlier, i.e., in the Albian, which corresponds to the end of the Early Cretaceous (“Austrian”)
orogeny in which the prearc nappe system along the European margin was established. An
additional width of 300 km of oceanic lithosphere, remaining at 90 Ma as depicted in Figure
2.7, is necessary to sustain subduction-related magmatism until ~70 Ma, assuming an
unchanged plate motion speed of 15 km/Ma [Rosenbaum et al., 2002] until closure of the
Neotethys.
The southward migration of magmatic activity within the central segments of the arc is
best interpreted as resulting from a gradual increase in the subduction angle associated with a
reduction of the width of the arc-trench gap. For the particularly well-documented
Panagyurishte segment, 100 km of across-arc migration of magmatism from 92 to 75 Ma (67
30
Chapter 2
Ma including the Rhodopes) corresponds to a steepening of the subduction angle from an initial
22–26° to 34–39° and a contraction of the arc-trench gap from an initial 250 km to 150 km.
Slab rollback and southward migration of both trench and arc relative to fixed Europe is a rather
unlikely alternative to explaining across-arc migration of magmatism [cf. von Quadt et al.,
2005], given the evidence for continued Africa-Europe convergence. Consequently, we
interpret the comagmatic sedimentary basins as intraarc rift basins rather than back-arc rift
basins.
The magmatic arc of the ABTS belt has natural terminations on its two ends. It
terminates west of the Apuseni segment because there is no along-strike continuation of the
Neotethys (Sava) Ocean, due to a change in subduction polarity between Alps and Dinarides
along a transform fault that approximately coincides with the present-day mid-Hungarian shear
zone (Figure 2.7) [Schmid et al., 2008; Schmid et al., 2011]. In the east the magmatic arc
terminates at the West Black Sea Fault, a transform fault delimiting the oceanic Black Sea backarc basin to the west [Okay et al., 1994].
The evolution of the ABTS belt in Late Cretaceous times can be subdivided into the
following three stages illustrated in Figure 2.8 for three representative segments of the ABTS
belt:
Active continental margin at ~110 Ma. As discussed above north dipping subduction of
the Neotethys Ocean along the Sava trench must have started some time before the onset of arc
magmatic activity, most likely during the Albian. Sediment accumulation in subduction-related
basins started between 100 and 90 Ma along the full length of the European continental margin
(Figures 2.8a–2.8c). The formation of strike-slip and pull-apart basins in the Panagyurishte
segment indicates that the dextral Maritsa fault system has already been active at that time
(Figure 2.8b). Toward the end of this stage the mantle source was geochemically enriched by
subduction fluids and/or melts to generate the characteristic subduction-like signature of arc
magmas. A lower crustal magma chamber, where the first magmas were further enriched in
volatiles and metal content, might have already existed below the Panagyurishte and Timok
segments.
Initiation of magmatic activity, steepening of the subduction zone, and ore deposit
formation (~92 to 75 Ma). The earliest upper crustal magmatism is recorded by intrusive rocks
from the northern Panagyurishte segment and indirectly by comagmatic sediments preserved in
the Eastern Srednogorie segment (~92 Ma; Figures 2.8e and 2.8f). The onset of magmatic
activity systematically became younger toward the west (~89 to 82 Ma; Figure 2.6) in the other
31
Chapter 2
segments of ABTS belt (Figures 2.6 and 2.8d–2.8f). The ascent of magmas to the upper crust
might have been facilitated by the steepening of the subduction zone and was partly focused by
pull-apart structures, e.g., along the Panagyurishte lineament associated with strike-slip faulting
along the Maritsa fault system (Figure 2.8e). At the same time, magmatic activity shifted
continuously to the south in all the arc segments except for the Apuseni segment, as is evidenced
by progressively younger magmatic ages toward south. This age shift away from the continent
toward the paleo-trench is the most compelling evidence for continuous steepening of the
subduction zone, probably because of the increasing magnitude of slab pull forces.
Additionally, the trend to less radiogenic, more mantle-like Sr and Nd isotope ratios in most
segments and the deepening of the volcano-sedimentary basins support steepening of the
subducting Neotethys slab. Economic porphyry Cu and epithermal Cu ± Au deposits coincide
with early stages of magmatism in the Panagyurishte and Timok segments.
End of active subduction and arc magmatism by continental collision (~72 to 67 Ma).
Arc magmatism within or near to the intraarc basins ceased at ~72 Ma in all the segments, but
younger plutons occur further south within the Rhodopes and Strandzha Units south of the
Panagyurishte and Eastern Srednogorie segments (69–67 Ma; Figures 2.8g and 2.8h). These
latest plutons probably reflect the termination of active subduction of the Sava branch of
Neotethys Ocean and likely mark the collision between Adria and Europe at the end of
Maastrichtian (~66 Ma) [e.g., Schmid et al., 2008]. Younger plutons intruded the Rhodopes
only after a significant gap of some 10 Ma (55-56 Ma) [Soldatos et al., 2008; Jahn-Awe et al.,
2010; Marchev et al., 2013]. They probably intruded in a postcollisional setting after subduction
in the Sava trench had stalled and the subduction zone had shifted to a new trench further south.
Still younger Eocene to Oligocene (~42 to 26 Ma) magmatism in the Rhodope and Dacia Units
either formed due to postcollisional slab break-off or mantle delamination [Schefer et al., 2010;
Marchev et al., 2013] or was driven by a subduction zone now located further south in the
Aegean region [Lehmann et al., 2013]. Active subduction migrating southward matches a longlived environment of slab rollback on the larger scale, progressing since the Early Cretaceous
to present-day Crete, which is supported by an ~1600 km long tomographically imaged slab
beneath the Aegean region [Bijwaard et al., 1998; van Hinsbergen et al., 2005].
Figure 2.8: Schematic tectonic model for three representative segments of the ABTS belt. Late
Cretaceous tectonic history of (a, d, and g) the Banat segment, (b, e, and h) the Panagyurishte segment,
and (c, f, and i) the Eastern Srednogorie segment. The colors of tectonic units are the same as in Figure
2.7. Blue drops = mantle source enrichment and melting, grey lines = magma ascent paths; black =
active magma chambers and volcanoes; grey = extinct magma chambers and volcanoes; light green =
sedimentary basins; light pink = lower crustal magma chamber. The small insets in lower left corners
show a magnification of the crustal magmatic activity.
32
Chapter 2
33
Chapter 2
2.6
Summary and Conclusions
In this study we have attempted to resolve the tectonic history of the Late Cretaceous
magmatic arc embedded in the Carpathian-Balkan orogen, by comparing magma-chemical
signatures and age trends in distinct segments along and across the arc. Based on geochemical
characteristics, the Apuseni-Banat-Timok-Srednogorie (ABTS) belt can be interpreted as a
typical subduction-related magmatic arc that formed on a continental margin. The arc was
active for 25 Ma, and across-arc younging of the magmatic products toward the paleo-trench
provides clear evidence for gradual steepening of the subducting Neotethys slab. This north to
south age progression is accompanied by distinct isotopic trends in the respective arc segments,
generally indicating an increasing contribution of mantle melts, which probably results from
increasing asthenospheric corner flow. The contemporaneous formation of sedimentary and
volcano-sedimentary basins is likely due to the same tectonic processes, i.e., subduction and
slab steepening leading to intraarc extension. Economic deposits preferentially formed in the
central arc segments, because these were subjected to only mild transtension during
contemporaneous shearing, favoring high-pressure amphibole fractionation and accumulation
of magmatic volatiles. Collision with the Adriatic plate terminated active subduction in the Sava
trench and arc magmatism. Postemplacement bending of the entire arc and associated
extensional tectonics partly concealed the rather simple and typical geometry of this continental
magmatic arc but favored the preservation of near-surface ore deposits and shallow volcanosedimentary basins in this relatively old metallogenic belt.
Appendix
The distribution patterns of magmatic crystallization ages can yield important additional
constraints for plate tectonic reconstructions. All available magmatic ages were plotted in the
present-day tectonic map (Appendix 7) to identify any systematic variations of the magmatic
ages along and across the magmatic arc. Pronounced across-arc variations are observed in the
Panagyurishte, Timok, and Banat segments, where they change from north to south younging
trends (Panagyurishte) to east to west younging trends (Timok and Banat). The rotation, which
has been inferred from paleomagnetic data [e.g., Marton et al., 2007; van Hinsbergen et al.,
2008], is therefore also indicated by the magmatic ages. Calculated crystallization ages for all
segments are reported in Table S3. For the Banat and Apuseni segments, 2 standard deviations
of overlapping and concordant TIMS or LA-ICP-MS ages in a population of analyses of each
sample are reported, as a conservative measure of age uncertainty, rather than standard errors
of the mean that become unrealistically small in the case of numerous point analyses.
34
Chapter 2
Acknowledgements
This study was supported by the Swiss National Science Foundation grants 200020146681 and 20021-146651 and SNF scopes projects JRP 7BUPJ062396 and IZ73ZO_128089
and incorporates results by Melanie Kolb, Svetoslav Georgiev, and Majka Kaiser-Rohrmaier as
early pioneers of this multi-PhD project. Ioan Seghedi provided essential help during joint field
work in Romania, and we are most grateful for his regional geological insight. We thank Ramon
Aubert, Markus Wälle, Marcel Guillong, Lydia Zehnder, and Muhammed Usman for their
support in the laboratories. All original geochemical data used in this study, including a
compilation of results from M. Kolb and S. Georgiev, are provided in digital form as supporting
information. Douwe van Hinsbergen is thanked for sharing his ideas concerning plate tectonic
reconstructions, influencing parts of our Figure 2.7.We thank Jeremy Richards and Iain Neill
for their constructive reviews, which considerably improved this manuscript.
35
Chapter 2
36
Chapter 3
3. The link between
Late Cretaceous and Miocene magmatism
in the Apuseni Mountains, Romania
3.1
Abstract
Subduction-related arc magmatism, and magmatism generated through re-melting of the
subduction-modified mantle-source after several millions of years can have similar
geochemical signatures and metal endowment. A particular tectonic setting of Late Cretaceous
barren arc magmatism superimposed by later post-subduction Miocene magmatism associated
with Au-Te epithermal and Cu-Au porphyry deposits is found in the Apuseni Mountains of
Romania. Major, trace element and Sr-Nd isotope data of the Late Cretaceous and Miocene
calc-alkaline magmatism are compared in this study.
The Late Cretaceous arc magmatism in the Apuseni Mountains is more silicic than
coeval magmas in the other segments of this arc and rhyolitic ignimbrites exclusively occur in
this arc segment. Precursor basaltic melts were presumably extracted from the metasomatized
asthenospheric mantle and underwent a polybaric evolution in the mid to upper crust. Low Sr/Y
ratios, decreasing Sr contents and Dy/Yb ratios indicate that the mantle-derived melts
fractionated a plagioclase- and amphibole-bearing assemblage probably at mid crustal levels
(~20 km). The andesitic to dacitic melts then presumably ascended to shallow crustal levels and
evolved to high silica rhyolitic melts. Fractional crystallization, however, did not occur in a
closed system, because the partly high
143
87
Sr/86Sr80Ma ratios (up to 0.7163) and low
Nd/144Nd80Ma ratios (as low as 0.51218) require addition of partial melts of continental crust.
Rough estimates based on the Neodymium crustal index (NCI) indicate addition of a maximum
of 60% of crustal melt to the magmatic system. Explosive volcanism led to a rapid loss of
volatiles and this may have prevented the formation of porphyry-style ore deposits.
Extension in the Miocene triggered re-melting of the subduction-modified lithosphere
and led to calc-alkaline magmatism akin to subduction-related magmatism. The Miocene
37
Chapter 3
magmatism can be divided into a low and a high 87Sr/86Sr group that show distinct trace element
ratios and presumably evolved via distinct pathways in the continental crust. The older high
87
Sr/86Sr (0.7065-0.7076) group assimilated local crust and possibly fractionated at mid to upper
crustal levels. The low 87Sr/86Sr (0.7038-0.7054) group has ‘adakite-like’ high Sr/Y and La/Yb
ratios and shows extreme enrichments in Sr, Ba and La. We speculate that these signatures were
acquired by addition of small-degree partial melts of hydrous mafic cumulates formed earlier
during the Late Cretaceous arc magmatism to primary magmas produced by melting of
previously metasomatized mantle. Since both groups of Miocene magmas are associated with
ore deposits, we assume that the metals were sourced in the metasomatized lithospheric mantle.
Re-melting of Au+(Te,Cu)-rich sulfides left in the lithospheric mantle by the Late Cretaceous
arc magmas might give rise to the unusually Au-rich Miocene magmas.
3.2
Introduction
Calc-alkaline magmatic arcs generally form above active subduction zones and are
frequently associated with porphyry Cu-Au and epithermal Au-Ag deposits [Sillitoe, 1972;
Richards, 2003; Sillitoe, 2010]. During subduction, slab-derived fluids enrich the overlying
asthenospheric mantle wegde with volatiles and incompatible trace elements [Gill, 1981;
Tatsumi et al., 1986; Hawkesworth et al., 1993; Elliott et al., 1997]. Addition of the volatiles in
conjunction with decompression attending corner flow leads to partial melting of the hydrous
metasomatized asthenosphere. The basaltic magmas subsequently undergo fractional
crystallization, possibly associated with assimilation of lower, middle or upper crust [de Paolo,
1981; Hildreth and Moorbath, 1988; Annen et al., 2006]. At upper crustal levels, the magmas
eventually form magmatic arcs, and in favourable tectonic settings also porphyry Cu-Au and
epithermal Au-Ag deposits [e.g. Sillitoe, 1972; Richards, 2003]. However, geochemically
similar magmas and ore deposits can also form in tectonic settings that are not related to a
contemporaneously active subduction zone in areas where subduction has long ceased [e.g.
Richards, 2009; Richards, 2011a]. In such post-subduction settings, the magmas are sourced
from the lithospheric mantle or lower crust, which have presumably been pre-enriched during
a previous subduction process [Haschke and Ben-Avraham, 2005; Richards, 2009; Shafiei et
al., 2009; Hou et al., 2015]. Post-subduction lithospheric thickening, mantle lithosphere
delamination or lithospheric extension [e.g. Kay and Mahlburg Kay, 1993; Richards, 2009] can
trigger re-melting of the source region and result in magmatism similar to arc magmatism and
ore deposit formation. Millions to billions of years can elapse between enrichment by
subduction and re-melting of the lithospheric source [e.g. Pettke et al., 2010; Hou et al., 2015].
38
Chapter 3
A particular setting of a subduction-related magmatic arc superimposed by clearly postsubduction magmas, with associated magmatic-hydrothermal Cu-Au deposits, is found in the
Apuseni Mountains (Romania) (Figures 3.1, 3.2). Late Cretaceous subduction-related arc
magmatism formed during north-dipping subduction of the Neotethys ocean beneath the
European continental margin [Gallhofer et al., 2015]. Although other segments of the Late
Cretaceous Apuseni-Banat-Timok-Srednogorie (ABTS) arc host large porphyry-style and
epithermal Cu-Au deposits, only minor Zn-Pb-Fe Skarns are associated with the Late
Cretaceous Apuseni magmatism [Jankovic, 1997; Berza et al., 1998; Ciobanu et al., 2002]. The
superimposed younger phase of Miocene calc-alkaline magmatism shows typical subductionlike trace element signatures [Rosu et al., 2004; Seghedi et al., 2004; Harris et al., 2013], and
is associated with rich epithermal Au-Ag-Te and porphyry Cu-Au deposits in the “Golden
Quadrangle” [Udubasa et al., 1992; Alderton and Fallick, 2000; Kouzmanov et al., 2003, 2005a,
2007]. The Miocene magmas are unlikely to have formed during active subduction, as they
presumably occur too far away from the nearest contemporaneous subduction zone of Alpine
Tethys located in the Carpathian embayment (Figure 3.1) [e.g. Mason et al., 1998; Wortel and
Spakman, 2000]. Instead, their formation has been related to extensional melting. Paleogene to
Neogene rotation presumably lead to extension, upwelling of asthenospheric mantle and remelting of the pre-metasomatized source region producing the Miocene magmas [Seghedi et
al., 1998; Rosu et al., 2004; Neubauer et al., 2005; Harris et al., 2013].
The Late Cretaceous arc magmatism and the Miocene post-subduction magmas are
spatially overlapping, which makes a common source region in the mantle likely. Here, we test
the hypothesis that this mantle source was metasomatized during subduction and closure of the
Neotethys ocean in the Late Cretaceous, about 50 Ma prior to Miocene extensional re-melting
[e.g. Harris et al., 2013]. Accordingly, generation of the barren Late Cretaceous and the orehosting Miocene magmatism is linked through processes occurring in the mantle to lower crust.
By combining bulk rock geochemical and isotope data for the Late Cretaceous arc-related
[Gallhofer et al., 2015] with the Miocene post-subduction magmatism [Harris et al., 2013], we
first establish the tectonic setting, and propose a polybaric magmatic evolution for the Late
Cretaceous Apuseni magmas. We then reassess trace element characteristics of the Miocene
magmatism, and propose a conceptual model that links the Late Cretaceous and Miocene
magmatism, and can account for the observed geochemical trends and metal endowment.
39
Chapter 3
Figure 3.1: Geological map of the Alpine-Carpathian-Dinaride orogen modified from Schmid et al.
[2008]. Also shown are the Late Cretaceous Apuseni-Banat-Timok-Srednogorie (ABTS) magmatic arc
and Neogene magmatism in the Circum-Carpathian region. The red lines denote the suture zones of the
Alpine Tethys and Neotethys oceans.
3.3
Geological Setting of the Apuseni Mountains
The Apuseni Mountains have an isolated position within the Alpine-Carpathian-
Dinaride orogen, amidst the Pannonian basin to the north and west, the Transylvanian basin to
the east and the South Carpathians to the south (Figures 3.1, 3.2). Continuous convergence
between the European and African plates shaped this region and lead to closure of ocean basins
and collision of smaller continental units with Europe [Herz and Savu, 1974; Csontos and
Vörös, 2004; Schmid et al., 2008]. The Apuseni Mountains are located at the boundary between
the continental Tisza and Dacia tectonic Mega-Units (Figures 3.1, 3.2) [e.g. Csontos and Vörös,
2004; Schmid et al., 2008]. Additionally, the obducted Eastern Vardar ophiolitic unit, a former
branch of the Neotethys ocean, crops out in the southern Apuseni Mountains [Schmid et al.,
2008]. The Adria-derived ALCAPA unit borders the Tisza Mega Unit to the north (Figure 3.1).
A suture zone of the Alpine Tethys (the Ceahlau-Severin ocean) divides the Dacia Mega-Unit
from the Moesian platform, which is a part of the stable European continent, to the south
[Schmid et al., 2008].
The continental Dacia Mega Unit is Europe-derived, whereas the Tisza Mega Unit
shows mixed European and Adriatic sedimentary affinities [e.g. Csontos and Vörös, 2004; Haas
and Pero, 2004; Iancu et al., 2005]. Both Mega-Units comprise Variscan-metamorphosed
40
Chapter 3
basement with Neoproterozoic crustal components and Late Paleozoic to Mesozoic cover
sediments [Pana et al., 2002; Balintoni et al., 2009; Balintoni et al., 2010]. The Biharia nappes
(Figure 3.2) of the Apuseni Mountains have previously been regarded as part of the Tisza Mega
Unit [e.g. Csontos and Vörös, 2004], but have recently been assigned to the Dacia Mega Unit
[Schmid et al., 2008]. The correlation with the Dacia Mega Unit is largely based on the
observation that the Eastern Vardar ophiolitic unit elsewhere overlies nappes belonging to this
tectonic unit [Schmid et al., 2008].
The South Apuseni ophiolites (Figure 3.2), which are part of the Eastern Vardar
ophiolitic unit, crop out in the Apuseni Mountains and consist of Jurassic ophiolites intruded
by a Late Jurassic island arc [Bortolotti et al., 2002; Nicolae and Saccani, 2003; Bortolotti et
al., 2004]. The sequence was presumably obducted onto the Dacia Mega-Unit in the Late
Jurassic, which has been inferred from shallow water limestones overlying both the ophiolitic
unit and Dacia [e.g. Schmid et al., 2008; Kounov and Schmid, 2013]. The present-day nappe
stack of the Apuseni Mountains formed during the Turonian orogeny and NW-ward
backthrusting of the Biharia system and the overlying Eastern Vardar ophiolitic unit over the
Bihor and Codru nappes of the Tisza Mega Unit [Schmid et al., 2008; Kounov and Schmid,
2013]. Uppermost Turonian Gosau-type sediments seal the Turonian nappe stack [Schuller,
2004; Schuller et al., 2009]. A latest Cretaceous compressional phase (Late Campanian to
Maastrichtian) controlled enhanced subsidence of the Gosau-type basins and doming and
exhumation in the Bihor nappe system [Schuller et al., 2009; Merten et al., 2011; Kounov and
Schmid, 2013].
In the Late Cretaceous, the Apuseni Mountains were located on the upper plate during
NE-dipping subduction of the remnant Neotethys ocean along the Sava trench (Figure 3.1),
which triggered calc-alkaline magmatism in the Apuseni Mountains [e.g. Berza et al., 1998;
Ciobanu et al., 2002]. This magmatism is part of a continental magmatic arc and associated
metallogenic belt that extends from the Apuseni Mountains, through the Banat region and
Timok zone in the South Carpathians to the Srednogorie zone in the Balkanides (Figure 3.1)
[Heinrich and Neubauer, 2002; von Quadt et al., 2005; Georgiev et al., 2012; Kolb et al., 2013].
Shallow intrusive and plutonic rocks, especially smaller stocks and dykes, occur in all tectonic
units of the Apuseni Mountains (Figure 3.2) and are frequently associated with Gosau-type
sedimentary basins [e.g. Stefan et al., 1992; Schuller, 2004]. Additionally, a more than 200 km2
wide volcano-plutonic complex of mainly silicic volcanism is exposed in the Vladeasa massif
in the Tisza Mega Unit [Istrate, 1978; Stefan, 1980], and volcaniclastic deposits overlie the
South Apuseni ophiolites [e.g. Constantina et al., 2009]. Recent U-Pb zircon ages show that
41
Chapter 3
subduction-related arc magmas in the Apuseni Mountains were emplaced between 81 and 75
Ma [Gallhofer et al., 2015]. Subduction in the Sava trench ceased after collision of the TiszaDacia unit with the Dinarides at the end of the Maastrichtian (~66 Ma) [Karamata, 2006;
Schmid et al., 2008].
Figure 3.2: Geological map of the Apuseni Mountains, redrawn from official Romanian geological maps
(scale 1:200000) and modified after Balintoni [1994] and Kounov and Schmid [2013].
In the Paleogene and Neogene, subduction of the Alpine Tethys ocean occurred in the
outer Carpathian arc along the margin of the Dacia and ALCAPA Units (Figure 3.1). Miocene
magmatic rocks forming a linear array roughly parallel to this subduction zone did presumably
not form during active subduction of the Alpine Tethys, but are rather related to post-collisional
extension, slab break-off or lithospheric delamination [de Boorder et al., 1998; Seghedi et al.,
1998; Harangi and Lenkey, 2007; Fillerup et al., 2010; Seghedi and Downes, 2011]. The
42
Chapter 3
Miocene magmatism in the Apuseni Mountains, however, occurs far beyond this magmatic
front, and is therefore likely not associated with subduction of the Alpine Tethys ocean.
During Paleogene to Neogene times, the Tisza-Dacia unit translated NE-wards, rotated
clockwise around the protruding corner of the Moesian platform and invaded the still open
Carpathian embayment [Ratschbacher et al., 1993; Linzer et al., 1998; Schmid et al., 1998;
Fügenschuh and Schmid, 2005]. The Apuseni Mountains underwent substantial clockwise
rotation up to 90° after Cretaceous times [Pǎtraşcu et al., 1990; Pǎtraşcu et al., 1992; Pǎtraşcu
et al., 1994; Márton et al., 2007]. However, the exact timing of most of this rotation is
controversial. Paleomagnetic data indicate that most of the rotation (more than 68°) occurred
in the Middle Miocene (15 to 10 Ma) [Panaiotu, 1998; van Hinsbergen et al., 2008]. Geological
models, in contrast, predict smaller amounts of rotation (approximately 40°) during the
Miocene and propose that much of the rotation pre-dates the Miocene [Fügenschuh and Schmid,
2005; Ustaszewski et al., 2008].
Miocene rotation in the Apuseni Mountains resulted in extension and formation of NWSE trending Miocene grabens [e.g. Balintoni, 1994; Neubauer et al., 2005]. Associated elevated
heat flow presumably triggered re-melting of lithospheric mantle and generation of the Miocene
post-subduction magmatism [e.g. Seghedi et al., 1998; Rosu et al., 2004; Seghedi et al., 2004].
Miocene igneous rocks crop out in the southern Apuseni Mountains, primarily in the Dacia
Mega-Unit and the overlying South Apuseni ophiolites (Figure 3.2). The available K-Ar ages
indicate Miocene magmatic activity from ~15 to 7 Ma [Pécskay et al., 1995; Rosu et al., 2004].
Zircon U-Pb emplacement ages are only available for Rosia Montana, Rosia Poieni and the
Barza magmatic center and range from 13.6 to 9.2 Ma [Kouzmanov et al., 2005, 2006].
Subvolcanic rocks with porphyritic textures prevail and their geochemical signature changes
from normal calc-alkaline to ‘adakite-like’ over time [Rosu et al., 2004]. Moreover, the
influence of crustal assimilation processes diminishes gradually with increasing extension
[Rosu et al., 2004; Harris et al., 2013]. Harris et al. [2013] established a model for ore deposit
formation. They correlated initial extension with the occurrence of the largest Au deposits (e.g.
Rosia Montana) in the Apuseni Mountains, progressing mantle input then led to the formation
of Te-rich Au-Ag epithermal deposits (e.g. Sacarimb), and continued extension and partial
melting of the mantle resulted in Cu-Au porphyry deposits [Harris et al., 2013].
3.4
Results
We combined our previously published whole rock geochemical data for the Apuseni
and Banat segments of the Late Cretaceous magmatic arc [Gallhofer et al., 2015] with available
43
Chapter 3
data for the Miocene extension-related magmatism of the Apuseni region [Harris et al., 2013].
We use the data of the Late Cretaceous magmatism in the Banat segment mainly to cover the
full geochemical range, and to expand the data of the Late Cretaceous Apuseni segment to less
evolved compositions. In most of the diagrams, the Late Cretaceous Banat and Apuseni samples
form a partly overlapping compositional trend, but the Banat data are not discussed in detail.
3.4.1 Field Relations and Description of the Late Cretaceous Samples
We sampled occurrences of Late Cretaceous igneous rocks throughout the Apuseni
Mountains to provide a representative geochemical and geochronological dataset [Gallhofer et
al., 2015]. Here, we briefly describe which types of igneous rocks occur in the major tectonic
host units, and emphasize differences in their mode of occurrence. Besides small sub-volcanic
stocks and dykes intruding the South Apuseni ophiolites, this tectonic unit is overlain by a Late
Cretaceous pyroclastic breccia in the Mures valley [Constantina et al., 2009], and remnants of
a Late Cretaceous tuff near Zlatna (Figure 3.2). Both volcanic units have been mapped as
Miocene on official Romanian maps (scales 1:50000 and 1:100000), but our U-Pb ages, 79 Ma
for an andesite in the pyroclastic breccia and 80 Ma for the tuff, confirm a Late Cretaceous age.
The tuff and a syeno-dioritic dyke intruding the ophiolites occur in vicinity of Miocene igneous
rocks. The Dacia and Tisza Mega-Units host shallow intrusions up to 10s of square kilometres
in size. The continental Dacia and Tisza Mega-Units as well as the South Apuseni ophiolites
are overlain by Gosau-type sedimentary basins, which are pierced by numerous shallow
intrusives and subvolcanic dykes ranging from dioritic to granitic compositions. Shallow
granodiorites piercing Gosau-type sediments at the eastern margin of the Apuseni Mountains
yielded mean 206Pb/238U ages from 77.4 to 75.5 Ma.
In the Vladeasa massif in the northern Apuseni Mountains (Figure 3.2), shallow marine
Gosau-type sediments [Schuller et al., 2009] underlie a more than 300 km2 wide volcanoplutonic complex [Istrate, 1978; Stefan, 1980]. Istrate [1978] and Stefan [1980] describe an
early cycle of minor andesitic and prevailing silicic volcanism, which was followed by a later
cycle of intrusives. The silicic volcanic products include rhyolites with eutaxitic texture (welded
ignimbrites), massive rhyolites and dacites [Istrate, 1978; Stefan, 1980; Stefan et al., 1992].
The intrusives are mainly shallow granites to granodiorites, and diorites with porphyritic to
holocrystalline textures occur as dykes and stocks [Stefan et al., 1992]. We have dated volcanic
and intrusive rocks from the Vladeasa massif to determine age differences of the igneous rocks
previously established from field relationships and hydrothermal effects at the contacts [Istrate,
1978; Stefan, 1980]. The oldest intrusives, which were both sampled in the northern Vladeasa
44
Chapter 3
massif, have mean 206Pb/238U ages of 80.7 and 80.8 Ma, and two minimally younger intrusives
have mean
206
yielded mean
Pb/238U ages of 80.2 and 80.3 Ma. Two volcanics, an ignimbrite and a dacite,
206
Pb/238U ages of 79.8 and 80.3 Ma, respectively. The single zircon 206Pb/238U
ages of the volcanic rocks obtained by laser ablation inductively-coupled-plasma mass
spectrometry (LA-ICP-MS) show a considerable spread. Eruption ages are unfortunately not
available for the ignimbrites. The mean
206
Pb/238U ages of the silicic volcanics and shallow
intrusives in the Vladeasa massif overlap within error. However, even if volcanics and shallow
intrusives probably did not occur contemporaneously, magmatic activity in the Vladeasa massif
seems to be restricted to a relatively narrow time range.
3.4.2 Petrography
In this section, we briefly describe the mineralogy and textures found in the Late
Cretaceous igneous rocks, the Miocene igneous rocks have been characterized elsewhere in
more detail [e.g. Rosu et al., 2004; Seghedi et al., 2007]. The Miocene igneous rocks are mainly
subvolcanic andesites, with abundant plagioclase, amphibole and occasionally clinopyroxene,
orthopyroxene and biotite phenocrysts in a fine-grained plagioclase-dominated matrix [Rosu et
al., 2004]. Some amphibole phenocrysts are opacitised, and in Deva they show evidence for
two episodes of resorption [Seghedi et al., 2007; Harris et al., 2013]. Garnet has been reported
as a rare accessory phase [Rosu et al., 2004; Seghedi et al., 2007].
The Late Cretaceous igneous rocks were grouped in volcanic, shallow intrusive and
plutonic rocks. The volcanic rocks are often porphyritic with an aphanitic groundmass.
Occasionally, oriented plagioclase laths in the groundmass indicate flow banding, vesicles are
rarely observed. Compared to the volcanic rocks, the porphyritic shallow intrusive rocks have
a slightly coarser groundmass and sometimes contain more phenocrysts. The plutonic rocks
show phaneritic textures with equigranular to inequigranular grain sizes. The least evolved
sampled Late Cretaceous igneous rock is a diorite, and consists of plagioclase, idiomorphic
clinopyroxene and poikilitic biotite (Figure 3.3a). Andesites contain phenocrysts of plagioclase,
clinopyroxene or amphibole, and biotite in a glassy or plagioclase-dominated groundmass.
Rocks containing clinopyroxene and amphibole are rare and usually show signs of
disequilibrium, i.e. clinopyroxene that is replaced by amphibole or opaque rims around
amphibole crystals (Figures 3.3b, c). Clinopyroxenes rarely occur in the more evolved rocks,
but sometimes orthopyroxenes are observed in amphibole-bearing rocks. The more evolved
granodiorites, granites and dacites consist of plagioclase, biotite and amphibole occurring as
phenocrysts, interstitial phases or poikilitic (Figure 3.3d), alkali-feldspar, and quartz.
45
Chapter 3
Figure 3.3: Textures of the Late Cretaceous igneous rocks from the Apuseni Mountains. (a) Diorite
(DG106) with an inequigranular texture of plagioclase, poikilitic biotite and clinopyroxene, and small
secondary interstitial biotite. Photomicrograph was taken under plane polarized light. (b) Porphyritic
andesite (DG078) with phenocrysts of plagioclase, clinopyroxene, and amphibole. Amphibole crystals
show thick opaque rims and internal corrosion. Crossed polarizers. (c) Syeno-diorite (DG113) consisting
of plagioclase, alkali-feldspar, biotite and interstitial amphibole. Green amphibole replaces
clinopyroxene. Plane polarized light. (d) Mafic phenocrysts in the matrix are mostly altered to chlorite,
but amphibole is still preserved within plagioclase phenocryst (DG121). Plane polarized light. (e)
Welded ignimbrite (DG093) with elongate fiamme, which are partly altered to epidote, and quartz
fragments. Plane polarized light. (f) Unwelded tuff (DG115) with abundant glass shards, embayed quartz
and a lithoclast. Plane polarized light. Abbreviations: amph: amphibole, ap: apatite, bt: biotite, chl:
chlorite, cpx: clinopyroxene, ep: epidote k-fsp: alkali-feldspar, mt: magnetite, plag: plagioclase, qtz:
quartz.
46
Chapter 3
Porphyritic rhyolites and welded ignimbrites showing eutaxitic textures are amongst the
most evolved samples. Besides fiamme, which are partly altered to epidote and quartz, quartz
and alkali-feldspar fragments, and quartz-rich lithic fragments occur in a glassy matrix (Figure
3.3e). An unwelded tuff from close to Zlatna (Figure 3.2) consists of abundant glass-shards,
embayed quartz crystals and rare lithic fragments (Figure 3.3f). Magnetite is a ubiquitous
accessory phase in most of the Late Cretaceous igneous rocks. Additionally, apatite, titanite and
opaque phases occur. Observed alteration features include beginning sericitization of
plagioclase and alkali-feldspar phenocrysts, partial replacement of biotite and other mafic
minerals by chlorite and rarely epidote, and opacitization of mafic phenocrysts. More detailed
descriptions of Late Cretaceous igneous rocks from the Vladeasa massif are provided in Stefan
et al. [1992], Istrate [1978] and Stefan [1980].
3.4.3 Major and Trace Element Characteristics
The major element characteristics show some differences between the Late Cretaceous and
Miocene calc-alkaline magmas (Figure 3.4). The Late Cretaceous Apuseni samples cover a wider range
of SiO2 content (~53-79 wt%) compared to the more intermediate Miocene Apuseni samples (~49-64
wt%). The majority of the Late Cretaceous samples contain more than 63 wt% SiO2 and have a dacitic
to rhyolitic composition. Nearly all Late Cretaceous samples plot in the calc-alkaline field according to
Miyashiro [1974], but the Miocene samples cross the division line between the calc-alkaline and
tholeiitic trend (Figure 3.4). Late Cretaceous Apuseni samples are predominantly high-K calc-alkaline
[Rickwood, 1989], whereas the Miocene igneous rocks partly have lower K2O contents and mainly plot
in the calc-alkaline series field. In the Late Cretaceous samples, TiO2, MgO, Fe2O3 and CaO (not shown)
contents continuously decrease with increasing SiO2 content. Al2O3 and P2O5 contents are relatively
constant up to 59 wt% SiO2 and then decrease. The Late Cretaceous Banat samples lie on a trend with
the Apuseni samples. Miocene samples show a negative correlation of Fe2O3 and TiO2 with SiO2,
correlations with SiO2 content are less obvious in Al2O3, MgO, P2O5 and CaO (not shown) contents.
Some enclaves from the Miocene igneous rocks have lower MgO and TiO2 contents than would be
expected at low SiO2 contents.
Trace elements follow distinct trends in the Late Cretaceous and Miocene magmas,
which may be explained by mineral fractionation (Figures 3.5, 3.6, 3.7). Incompatible Rb shows
a weak positive correlation with SiO2 in Late Cretaceous and Miocene igneous rocks. Barium
and La slightly increase, and Sr decreases with increasing SiO2 in Late Cretaceous samples, but
these elements scatter considerably in the Miocene samples. Miocene Apuseni samples with
SiO2 ranging from 56 to 61 wt% show the highest concentrations of large ion lithophile
elements (LILE, e.g. Ba, Sr, Pb) and light rare earth elements (LREE, e.g. La), whereas more
evolved samples have lower contents (Figure 3.5).
47
Chapter 3
Figure 3.4: (a)-(f) Variation diagrams for major element oxides vs. SiO2 (wt%). Division line in plot (a)
is from Miyashiro [1974]. Dividing bands for the different subalkalic series in plot (b) are from
Rickwood [1989]. Abbreviations: p plutonic rocks, si shallow intrusive rocks, v volcanic and subvolcanic
rocks, d dyke rocks.
Zirconium scatters in both igneous suites, but seems to reach a maximum content at ~65
wt% SiO2 in the Late Cretaceous samples. Yttrium decreases up to ~65 wt% SiO2, and then
abruptly rises in the Late Cretaceous samples. The LILE Rb and K are, except for one Miocene
sample (RM-04-CH-32), more enriched in the Late Cretaceous Apuseni rocks, whereas the
LILE Ba, Sr and Pb are highly enriched (up to 2500 ppm) in Miocene samples from Deva,
Sacarimb, Rosia Poieni and Baia de Aries magmatic centers (low
87
Sr/86Sr group) (Figures
3.6a,b). High field strength elements (HFSE, e.g. Nb, Ta, Hf, Zr) and Ti are depleted in both
rock suites and form troughs in N-MORB normalized trace element plots (Figure 3.6b). Light
rare earth elements (LREE) are enriched over heavy rare earth elements (HREE) in the Late
Cretaceous and Miocene samples (Figure 3.6a).
The Miocene low
87
Sr/86Sr samples have lower HREE contents and particularly high
La/Yb ratios (up to ~50, Figures 3.5, 3.6a). The La/Sm ratio representing a light to middle rare
earth element ratio (LREE/MREE) slightly increases with increasing SiO2 content in Late
Cretaceous Apuseni samples, then decreases at ~70 wt% SiO2, and scatters in the Miocene
Apuseni samples (Figure 3.5h). The Dy/Yb ratio (MREE/HREE) slightly decreases with
increasing SiO2 content in both igneous suites (Figure 3.5i). The Miocene Apuseni magmas,
with the exception of a few high
87
Sr/86Sr Rosia Montana and Barza samples, exhibit only
weakly positive and negative Eu anomalies. The high silica (> 70 wt% SiO2) Late Cretaceous
Apuseni samples, especially from the Vladeasa massif and other occurrences in the northern
48
Chapter 3
Apuseni Mountains, show pronounced negative Eu anomalies (Eu/Eu*< 0.8) (Figures 3.6a,
3.7b).
Figure 3.5: (a)-(f) Variation diagrams for selected compatible and incompatible trace elements vs. SiO2
(wt%) as a measure of differentiation. (g)-(i) Rare earth element ratios vs. SiO2 (wt%). Fractionation
trends are from Davidson et al. [2007]. Abbreviations: p plutonic rocks, si shallow intrusive rocks, v
volcanic and subvolcanic rocks, d dyke rocks.
Figure 3.6: Trace element characteristics of Late Cretaceous and Miocene igneous rocks. (a) Chondrite
normalized rare earth element (REE) patterns. (b) N-MORB normalized trace element patterns.
Normalization values for (a) and (b) from Sun and McDonough [1989]. The irregular patterns in the
heavy REE composition of some Late Cretaceous Apuseni and Banat samples is due to low
concentrations close to the detection limit.
49
Chapter 3
Miocene samples, particularly from the Deva, Sacarimb, Rosia Poieni and Baia de Aries
magmatic centers (low 87Sr/86Sr group), yield high ‘adakite-like’ Sr/Y ratios (Figure 3.7a) [Rosu
et al., 2004; Harris et al., 2013]. The highest Sr/Y ratios coincide with none to slightly positive
Eu anomalies (Eu/Eu*=0.9 - 1.2) (Figure 3.7b). The Late Cretaceous Apuseni rocks mainly
have low Sr/Y and high Y contents of ‘normal-arc’ magmas, except for a few samples with
mildly ‘adakite-like’ signatures (Sr/Y ≤ 40) from the southern Apuseni Mountains (Figure 3.7).
The Late Cretaceous Banat samples, by comparison, more frequently show adakite-like
signatures.
Figure 3.7: ‘Adakite-like’ characteristics in Miocene post-subduction magmas and Late Cretaceous arc
magmas. (a) Sr/Y vs. Y (ppm). Fields for ‘adakite-like’ and normal arc magmas are from Richards and
Kerrich [2007]. (b) Sr/Y vs. Europium anomaly (Eu/Eu*). Abbreviations: p plutonic rocks, si shallow
intrusive rocks, v volcanic and subvolcanic rocks, d dyke rocks.
3.4.4 Radiogenic Isotope Characteristics
Here, we compare the isotopic characteristics of Late Cretaceous arc magmas and
Miocene post-subduction magmas. The Late Cretaceous samples show a spread in Sr and Nd
isotopic compositions (Figure 3.8a). Four samples from the southern Apuseni Mountains have
the least radiogenic age-corrected (80 Ma)
87
Sr/86Sr ratios (0.70424-0.70530) and slightly
positive εNd values (0.9 to 3.1), but none of these samples overlaps with mid ocean ridge basalt
(MORB)-type mantle. All other Late Cretaceous Apuseni samples, particularly the SiO2-rich
samples from the Vladeasa massif, exhibit 87Sr/86Sr80Ma ratios in excess of 0.705301 and εNd
lower than -1.5, and partly overlap Variscan granitoids from the Dacia basement in the South
Carpathians [Duchesne et al., 2008]. An increase of 87Sr/86Sr ratios and decrease of εNd values
with increasing SiO2 contents is also observed in the Late Cretaceous Apuseni samples. The
Late Cretaceous Banat samples mainly occupy an intermediate position between the high and
low Sr isotope ratios of the Apuseni samples. The
87
Sr/86Sr ratios of the Miocene Apuseni
samples were age-corrected for 11 Ma, but the initial ratio does not differ significantly from the
50
Chapter 3
measured ratio in such young samples. The Miocene post-subduction magmas can be divided
in two groups based on their
87
Sr/86Sr ratios (Figure 3.8b). The Rosia Montana and Barza
samples show high 87Sr/86Sr ratios (0.70659-0.70777) and correspondingly negative εNd (0 to
-3.6), whereas the remaining samples yielded considerably lower
87
Sr/86Sr ratios (0.70383-
0.70548) and positive εNd values (0.4-3.3) that approach MORB-type mantle. The Sr and Nd
isotope ratios of the Rosia Montana and Barza (high 87Sr/86Sr group) samples overlap Variscan
granitoids from the Dacia basement.
Figure 3.8: Sr and Nd isotope ratios for Late Cretaceous arc magmas and Miocene post-subduction
magmas. (a) 143Nd/144Nd vs. 87Sr/86Sr at 80 Ma. (b) 143Nd/144Nd vs. 87Sr/86Sr at 11 Ma. Two Late
Cretaceous samples have higher 87Sr/86Sr11Ma ratios than the range plotted. Field for MORB-type mantle
(mid ocean ridge basalt) from Stracke et al. [2005]; field for Variscan granitoids (age-corrected) from
Banat region from Duchesne et al. [2008]; basanites from Banat region [Downes et al., 1995; Tschegg
et al., 2010]. Abbreviations: p plutonic rocks, si shallow intrusive rocks, v volcanic and subvolcanic
rocks, d dyke rocks.
3.5
Discussion
Although the Late Cretaceous and Miocene Apuseni igneous suites were formed in
different tectonic settings, they show almost complete overlap in their geochemical and isotopic
characteristics. However, the two suites of magmatic rocks also differ in some key trace element
ratios (e.g. Sr/Y, La/Yb) and in their metal contents. In the following, we discuss the
geochemical characteristics in the light of possible sources and geochemical processes that
could potentially create these rocks in the same geographic area, but at different times. We
develop a conceptual model that explicitly links the Miocene magmatism to the geochemical
preparation of the lithosphere by Late Cretaceous tectonics, and explains the trace element
trends and metal endowment.
3.5.1 Late Cretaceous Subduction: Magmatic Preparation of the Lithosphere
In the Late Cretaceous, the Apuseni Mountains were a part of the European continental
margin, which was in an upper plate position during NE-dipping subduction of the Neotethys
51
Chapter 3
ocean. This subduction triggered magmatic activity along the entire continental margin and
gave rise to the more than 1000 km long ABTS (Apuseni-Banat-Timok-Srednogorie) magmatic
arc [e.g. Gallhofer et al., 2015]. The large amounts of dacites and rhyolites cropping out in the
Apuseni Mountains are unique within the ABTS belt. Some of the silicic volcanics are welded
ignimbrites with fiamme and according to earlier geological mapping [Istrate, 1978; Stefan,
1980] the rhyolites and dacites cover an area of approximately 300 km2 in the Vladeasa massif.
Additionally, rhyolites also occur in isolated positions north of the Vladeasa massif and close
to Zlatna (Figure 3.2). Only vague estimates exist for the thickness of the volcanic sequence
and the extent of erosion in the area is unknown, which impedes the calculation of magma
volumes for the Vladeasa massif. The areal extent is, however, smaller than that of large silicic
volcanic fields in the Altiplano-Puna of the Andes or the Cordillera of North America that cover
1000s of km2 [e.g. de Silva et al., 2006; Bachmann et al., 2007; Lipman, 2007; Kay et al., 2010].
The occurrence of silicic explosive volcanics in the Apuseni Mountains might be related to only
minor erosion in this segment of the ABTS belt [Gallhofer et al., 2015].
The Late Cretaceous Apuseni magmas range from basaltic andesite to rhyolite, but
silicic granodiorites, dacites and rhyolites clearly prevail. The mafic minerals observed in the
Late Cretaceous rocks change from pyroxene and biotite in the least evolved samples to
amphibole and biotite in the silicic varieties. The occurrence of abundant hydrous minerals
indicates that the magmas from which they crystallized were likely cold-wet-oxidized magmas,
which are typically found in subduction zone settings [e.g. Bachmann and Bergantz, 2008;
Deering and Bachmann, 2010]. The magmas have low FeOt/MgO ratios and follow the calcalkaline trend (Figure 3.4a), probably due to early crystallization of magnetite [Miyashiro,
1974], which is a common phase in the Late Cretaceous Apuseni samples. The high FeOt/MgO
ratios in the rhyolites might be explained by silica saturation [Sisson et al., 2005]. The least
evolved samples (< 65 wt% SiO2) show enrichments in LILE, LREE and depletions in HFSE,
which point to a subduction-enriched mantle source of the parental magmas. Even the least
evolved Sr and Nd isotope ratios of basaltic andesite magmas (Figure 3.8a) do not overlap
MORB-type mantle, which might reflect subduction-enrichment of the mantle source. Based
on the above observations, we conclude that the parental melts of the Late Cretaceous Apuseni
magmas are mantle-derived hydrous basalts generated in a subduction setting.
During ascent through the continental crust, the basaltic mantle-derived melts
underwent crystal fractionation associated with melting and assimilation of wall rocks [e.g. de
Paolo, 1981; Hildreth and Moorbath, 1988; Annen et al., 2006], which is supported by the
observed major and trace element trends, and increasingly radiogenic isotope ratios. The major
52
Chapter 3
and trace element evolution, and the observed mineral assemblages of the Late Cretaceous
samples provide valuable information on the potential fractionating phases. Although Sr is
slightly enriched in the least evolved Apuseni samples, the Sr content is generally lower than
in the Late Cretaceous samples from the adjacent Banat arc segment (Figures 3.5, 3.6).
Strontium is readily partitioned into plagioclase, and the decreasing Sr contents might be
interpreted in terms of plagioclase fractionation. Additionally, the decreasing Sr/Y ratios and
increasing La/Sm ratios with increasing differentiation (Figure 3.5h) indicate that plagioclase
was likely stable over a wide interval of fractional crystallization. The marked decrease in Al2O3
contents at 59 wt% SiO2 probably marks a higher proportion of plagioclase compared to mafic
phases in the crystallizing mineral assemblage. Decreasing Dy/Yb (Figure 3.5i) and Sm/Yb (not
shown) with increasing differentiation point to fractionation of amphibole, and indicate that
garnet did not crystallize [Davidson et al., 2007]. The sudden increase of Y at ~65 wt% SiO2
probably indicates the decreasing role of amphibole, and its replacement by another mafic
phase, possibly biotite, in the fractionation sequence. Given the generally low Sr/Y ratios, we
infer that plagioclase crystallized concomitantly with amphibole, and probably other mafic
phases, over the entire interval of fractional crystallization.
Since the deep root of the Late Cretaceous magmatic arc in the Apuseni Mountains is
not exposed and mafic cumulate xenoliths have not been observed, we can only speculate about
the type and amount of the cumulates formed during fractional crystallization. Deep arc roots
in the lower to middle crust typically consist of ultramafic to mafic cumulates or restites, which
are followed upward by mafic to felsic cumulates in the middle to shallow portions of the crust
[e.g. Jagoutz, 2010; Dessimoz et al., 2012; Ducea et al., 2015]. Vander Auwera et al. [2015]
have shown that the differentiation trend of the Late Cretaceous Apuseni magmas can be
predicted by fractional crystallization of gabbronoritic to dioritic cumulates in the upper crust.
Based on trace element signatures, they inferred that plagioclase was not suppressed in the
fractionation sequence, and assumed that gabbronoritic cumulates might have also formed in
the lower crust due to a low water content (<2.5 wt% H2O) of the mantle-derived melts.
However, early onset of plagioclase crystallization is not only influenced by the water content
of a melt, but might also be influenced by the depth of crystallization. Fractional crystallization
experiments of hydrous basaltic melts (≥3 wt% H2O) produce plagioclase early before
amphibole in the crystallization sequence only at mid-crustal or shallower pressure levels (~
0.7 GPa, 20 km) [e.g. Sisson et al., 2005; Blatter et al., 2013; Nandedkar et al., 2014]. The
experimental cumulates of Nandedkar et al. [2014] evolve from dunites, websterites and
plagioclase-bearing pyroxenites to amphibole-gabbroic cumulates.
53
Chapter 3
For the Late Cretaceous Apuseni magmas, we assume a polybaric evolution. As mafic
cumulates are generally not common in the upper crust, we assume that fractional crystallization
of amphibole-gabbroic cumulates at ~20 km depth might have driven the mantle-derived melts
to andesitic to dacitic compositions. The melts were then extracted and ascended to shallower
crustal levels (< 250 MPa, < 8 km) to form the high silica rhyolites [e.g. Gualda and Ghiorso,
2013]. The upper crustal cumulate was probably dominated by plagioclase, which is indicated
by the negative Eu-anomalies (Figures 3.6a, 3.7b) and very low Sr contents of the rhyolites,
apatite, which might explain the low P contents, and quartz. Such an evolution would be
consistent with models proposed by Deering and Bachmann [2010] and Lee and Bachmann
[2014]. However, fractional crystallization did not occur in a closed system, because the
evolved Sr and Nd isotope ratios observed in the Late Cretaceous Apuseni magmas require a
contribution of old crustal components.
To assess the amount of the crustal contribution, the isotopic composition of the mantle
and the potential assimilant need to be identified. We are aware that isotopic compositions
inherited from a subduction-enriched mantle source are only reliably detectable in primitive
basaltic rocks that have not undergone subsequent crustal assimilation [e.g. Davidson, 1996].
In a continental arc, interaction of the mantle-derived magmas with the local basement will
largely conceal subduction-enrichment and the isotopic ratios will likely reflect the local
basement [e.g. Wörner et al., 1992; Mamani et al., 2008]. Nevertheless, as the least radiogenic
Sr and Nd isotope ratios observed in the Late Cretaceous magmas fall within the mantle array,
they might represent the subduction-enriched mantle beneath the Apuseni Mountains. The
basement rocks in the Apuseni Mountains are Paleozoic gneisses and amphibolites, which are
intruded by Variscan granitoids and overlain by Permian to Lower Cretaceous sediments [Pana
et al., 2002; Balintoni et al., 2009; Balintoni et al., 2010]. The gneisses and granitoids have
present-day εNd values as low as -15 (average εNd0Ma = -7) [Pana et al., 2002], but their Sr
isotopic ratios have not been determined. However, Sr isotope ratios are available for Variscan
granitoids in the basement of the Dacia Mega-Unit in the Banat region, which have similar ages
and Nd isotope compositions as the ones occurring in the Apuseni Mountains [Duchesne et al.,
2008].
We conducted simple mixing calculations to estimate the amount of crustal contribution
(Figure 3.9). A sample with low Sr and high Nd isotope ratios (DG077) was taken as the mantle
endmember. The isotopic composition of the assimilant was chosen to overlap the range of
granitoid compositions in the Banat region (87Sr/86Sr=0.718,
143
Nd/144Nd=0.5121) [Duchesne
et al., 2008]. With these assumptions, the Late Cretaceous Apuseni samples, apart from the
54
Chapter 3
high-silica rhyolites, can be generated by addition of 25-40 % of the crustal assimilant. Vander
Auwera et al. [2015] have modeled assimilation and fractional crystallization of the Late
Cretaceous Apuseni magmas using the isotopic composition of a near-primary magma from
Mount Shasta as a starting composition (87Sr/86Sr=0.703789,
143
Nd/144Nd=0.512880) and a
granitoid from the Mecsek Mountains of the Tisza Mega Unit as the crustal assimilant
(87Sr/86Sr=0.7134, 143Nd/144Nd=0.5121). Although they could reproduce most of the observed
isotopic compositions of the Apuseni magmas, their model cannot account for the highly
evolved isotopic compositions of the high silica rhyolites from our dataset that require a
considerably more radiogenic assimilant. To better quantify the crustal addition to the rhyolites,
we
additionally
calculated
the
Neodymium
component)/(εNdcrustal component-εNdmantle component)]
crustal
index
[NCI=(εNdrock-εNdmantle
[de Paolo et al., 1992] using the same mantle
endmember (εNd= 3.0), but the crustal assimilant with the lowest Nd isotopic ratios (εNd80Ma=
-14.5) from Pana et al. [2002]. The calculated NCI values are 0.4 to 0.6 for the rhyolites, i.e.
40 to 60 % of crust was added in their generation. Such high degrees of assimilation can only
occur in relatively thick crust. Our results are in line with studies from continental arcs that
indicate that crustal melts contribute around 50% on average to the magmatic system [de Silva
et al., 2006; Ducea and Barton, 2007; Ducea et al., 2015].
Figure 3.9: Estimate of the crustal addition to the Late Cretaceous magmatism. The mixing line was
calculated for the isotopically most depleted Late Cretaceous sample (DG077), which might be
representative for the subduction-enriched Apuseni mantle, and an assimilant (87Sr/86Sr= 0.718; Sr= 200
ppm; 143Nd/144Nd= 0.5121; Nd= 35 ppm) overlapping the range of granitoids in the Dacia basement
[Duchesne et al., 2008]. Tick marks denote 10% intervals of crustal addition. p plutonic rocks, si shallow
intrusive rocks, v volcanic and subvolcanic rocks.
55
Chapter 3
3.5.2 Origin and Evolution of the Miocene Magmatism
The depleted Sr and Nd isotope ratios of the Miocene Apuseni magmas indicate that
they originally derived from a mantle source [e.g. Rosu et al., 2004; Harris et al., 2013].
Additionally, the Miocene Apuseni magmas show trace element patterns (enrichment in LILE,
LREE, and depletion in HFSE) (Figure 3.6), which are commonly characteristic for subductionrelated magmatism, and suggest that the mantle source was modified by subduction-enrichment
[Rosu et al., 2004; Seghedi et al., 2004; Harris et al., 2013]. Because the only known active
subduction zone at that time, the Alpine Tethys, was located far from the Apuseni Mountains
(Figure 3.1), the subduction-related trace element signature must have been stored for 10s of
millions of years and imparted to the Miocene post-subduction magmas by later re-melting of
the source [Rosu et al., 2004; Harris et al., 2013]. Subduction-derived trace elements and
volatiles are readily dispersed in the convecting asthenospheric mantle, which might therefore
be excluded as a potential source, but these elements might be stored in hydrous minerals in the
sub-continental lithospheric mantle [e.g. Richards, 2009; Pettke et al., 2010; Richards, 2011a].
Based on their isotopic composition, the Miocene magmas can be divided into two
groups, one showing high 87Sr/86Sr ratios (Barza, Rosia Montana), and the other low 87Sr/86Sr
ratios (Deva, Rosia Poieni and others) (Figure 3.8b) [Rosu et al., 2004; Harris et al., 2013].
This has been interpreted in terms of different sources for the two groups [Seghedi et al., 2007],
but we assume that both groups derived from a common metasomatized mantle source, and
were subsequently modified by additional processes. The high 87Sr/86Sr ratio group might be
explained by addition of crustal partial melts to the mantle-derived parental melt [e.g. Harris et
al., 2013]. Additionally, these samples have the lowest Sr (200-500 ppm), Ba (200-500 ppm)
and La (10-20 ppm) contents amongst the Miocene magmas (Figure 3.6), which might be
attributed to fractional crystallization of plagioclase during differentiation of these magmas.
The low
87
Sr/86Sr group shows moderate to extreme enrichments in LILE and LREE
elements (Sr= 500->2000 ppm; Ba= 500->2000 ppm; La= 25-95 ppm) and highly variable Sr/Y
ratios (30-300) (Figures 3.5, 3.6, 3.7). The observed variations would require a spatially rather
heterogeneous mantle source and variable degrees of partial melting [Rosu et al., 2004].
However, a heterogeneous mantle source can probably not account for the very high LILE and
LREE contents of the Miocene magmas. We speculate that subduction-induced metasomatism
might result in a similar distribution of LILE and LREE in the convecting asthenospheric and
the overlying sub-continental lithospheric mantle. The average continental arc basalt contains
425 ppm Sr [Kelemen et al., 2014], and Sr contents in lithospheric mantle-derived basaltic melts
will probably not exceed the maximum Sr contents observed in arc basalts (~700-1000 ppm).
56
Chapter 3
In the following section, we therefore discuss different processes, which may give rise to the
high contents of LILE and LREE observed in the low 87Sr/86Sr group of the Miocene magmas.
3.5.3 High Pressure Fractional Crystallization versus Cumulate Melting
Apart from moderate to high Sr/Y ratios, the low 87Sr/86Sr Miocene magmas have high
La/Yb ratios, constant to moderately decreasing Dy/Yb ratios and none to a slightly positive
Europium anomaly (Eu/Eu*= 1±0.1) (Figures 3.4, 3.7). High Sr/Y (> 20-40 ppm) and La/Yb
(> 20 ppm) ratios are commonly characteristic for ‘adakite-like’ rocks [Defant and Drummond,
1990, 1993]. Originally, ‘adakite-like’ signatures have been interpreted in terms of melting of
a subducted oceanic slab [Kay, 1978; Defant and Drummond, 1990]. Defant and Drummond
[1990] argued that the absence of plagioclase and residual garnet in the metamorphosed slab
drives the generated melts to low Y contents and high Sr/Y ratios, as Y and heavy REEs are
preferentially incorporated in garnet and Sr is released into the melt. More recently, however,
other processes have been recognized that can create the same geochemical characteristics, and
might as well occur in post-subduction settings. These include partial melting of thickened
mafic lower crust in the presence of residual garnet [e.g. Kay et al., 1999; Kay et al., 2005],
partial melting of normal lower crustal lithologies [e.g. Qian and Hermann, 2013], and high
pressure fractional crystallization of hydrous mantle-derived arc magmas [e.g. Castillo et al.,
1999; Macpherson et al., 2006; Richards and Kerrich, 2007; Chiaradia, 2009; Richards,
2011b].
Fractional crystallization yields high Sr/Y ratios when fractionation of plagioclase is
suppressed, which is achieved, if (1) the magma crystallizes at high pressures (>0.8 GPa, ~25
km), or if (2) the magma is rich in H2O (>3 wt%). At higher pressures and water contents,
amphibole and garnet will appear earlier in the crystallization sequence than plagioclase, which
increases the Sr content and decreases the Y content of the melt [Burnham, 1979; Müntener et
al., 2001; Richards and Kerrich, 2007; Alonso-Perez et al., 2009]. High pressure fractional
crystallization of abundant amphibole and/or garnet in subduction-related arcs is also consistent
with observations from exposed lower crustal sections of island and continental arcs, which
consist of cumulitic garnetites, hornblendites and amphibole-gabbros [Greene et al., 2006;
Jagoutz, 2010; Dessimoz et al., 2012]. In the case of the post-subduction Miocene magmas,
however, plagioclase-absent fractional crystallization was probably not the crucial process for
generating the high Sr/Y ratios. There is no clear correlation between Sr/Y ratios and SiO2
contents, which would be expected during progressive differentiation, and some of the high
Sr/Y magmas have rather high Mg-numbers (up to Mg#= 0.65).
57
Chapter 3
Another possibility to increase the Sr/Y ratios, LILE and LREE contents of the mantlederived melts is addition of partial melts from the lower crust, underplated basalts or lower
crustal cumulates [Haschke et al., 2002; Haschke and Ben-Avraham, 2005; Richards and
Kerrich, 2007; Richards, 2009; Shafiei et al., 2009; Mamani et al., 2010; Hou et al., 2015]. Due
to the low 87Sr/86Sr ratios of the Miocene samples, the contaminant probably had mantle-like
Sr isotope ratios. Lower crustal cumulates would fulfill this criterion and probably crystallized
from mantle-derived melts during the earlier stage of Late Cretaceous arc magmatism.
Although the nature of these cumulates is fairly speculative, ultramafic and hydrous mafic
lithologies, comparable to those in exposed lower crustal arc sections [Greene et al., 2006;
Jagoutz, 2010; Dessimoz et al., 2012], might be expected. The presence of accessory garnet in
a few Miocene rocks [Rosu et al., 2004] indicates that melting might have taken place in the
garnet stability field. Moreover, ambient noise tomography predicted an anomalously thickened
crust beneath the Apuseni Mountains (30-40 km) [Ren et al., 2013]. Experiments of partial
melting of amphibolite, representative of metamorphosed hydrous gabbro, produce residual
assemblages that can account for ‘adakite-like’ trace element signatures [Wolf and Wyllie, 1994;
Rapp, 1995; Rapp and Watson, 1995]. Partial melting might produce garnet-, clinopyroxene
and amphibole-bearing residues with negligible amounts of plagioclase, which results in high
Sr/Y ratios of the partial melts, at pressures not only higher than 1.0-1.5 GPa (> 30-45 km)
[Wolf and Wyllie, 1994; Rapp and Watson, 1995], but also at 1-1.25 GPa (30-40 km) [Qian and
Hermann, 2013].
cpx
amph
gt
KdLa
0.0154
0.12
0.022
KdYb
2
1.7
24
KdSr
0.17
0.28
0.015
KdY
2.6
2.46
2.9
DbulkLa
DbulkYb
DbulkSr
DbulkY
cpx:gt:amph
50:30:20
0.04
8.54
0.15
2.66
amph:gt
90:10
0.11
3.93
0.25
2.50
Table 3.1: Partition coefficients used for batch melting. All Kds are from the GERM database. Amph
amphibole, cpx clinopyroxene, gt garnet.
To test, whether addition of a cumulate melt could account for the elevated LILE and
LREE contents of the Miocene magmas, we calculated the La/Yb and Sr/Y ratios of such melts
using simple batch melting [Rollinson, 1993] (Figure 3.10). We used the composition of a
cumulitic amphibole gabbro from a lower crustal arc section exposed in the Cascades (Sr=695.9
ppm, Y=18.29 ppm, La=9.72 ppm, Yb=1.41 ppm) [Dessimoz et al., 2012] as a substitute for
58
Chapter 3
the unknown cumulates in the Apuseni Mountains. We chose residual assemblages containing
variable proportions of amphibole, garnet and clinopyroxene. Partition coefficients (Kd) for
dacitic to rhyolitic liquids were taken from the GERM database (http://earthref.org/KDD/).
Partial melting leaving a clinopyroxene-dominated, garnet and amphibole-bearing residue
(50:30:20) yields higher Sr/Y and La/Yb ratios than an amphibolitic residue containing 90%
amphibole and 10% garnet. Both residues require a rather high degree of partial melting (>
40%) to directly produce the high Sr/Y Miocene magmas. Such melting degrees would probably
not be achievable by pure dehydration melting without addition of external fluid [e.g. Qian and
Hermann, 2013]. Therefore, the high Sr/Y Miocene magmas were likely not produced by high
degree partial melting of lower crustal cumulates. Instead, the observed trace element signatures
might be explained by addition of 20-30% of only small degree cumulate melts (< 5% partial
melting) to the mantle-derived melt. One sample plotting at Sr/Y=309 (GM-04-CH-42E) falls
outside the general trend. This sample is a quench-textured enclave rich in plagioclase [Harris
et al., 2013], which might explain the unusually high Sr/Y ratio.
Figure 3.10: La/Yb versus Sr/Y for low 87Sr/86Sr Miocene igneous rocks. Lines represent batch melting
curves of hypothetical lower crustal cumulates, leaving residues containing amphibole (amph), garnet
(gt) and clinopyroxene (cpx). Tick marks indicate 10% melting steps.
3.5.4 Implications for the Miocene Mineralization
Late Cretaceous and Miocene igneous rocks are spatially overlapping in the Apuseni
Mountains (Figure 3.2), but while the Late Cretaceous magmatism is barren, Miocene magmas
are associated with rich epithermal Au-Ag-Te and Cu-Au porphyry-style deposits [Udubasa et
al., 1992; Alderton and Fallick, 2000; Wallier et al., 2006; Harris et al., 2013]. The spatial
relation indicates that the Late Cretaceous subduction-enrichment and arc magmatic activity
were probably a prerequisite for generating the Au- and Cu-rich Miocene magmas. The link
between the arc magmatism and the post-subduction magmatism might ultimately be the
59
Chapter 3
metasomatized mantle source [Harris et al., 2013]. A similar genetic link between barren
subduction-related and later post-subduction mineralized magmatism has also been inferred in
other parts of the Tethyan continental margin [e.g. Shafiei et al., 2009; Hou et al., 2015].
Melts derived from the metasomatized mantle above subduction zones are commonly
hydrous, oxidized, and enriched in S and metals [de Hoog et al., 2001; Richards, 2003; Jugo,
2009; Kelley and Cottrell, 2009; Jenner et al., 2010; Zimmer et al., 2010; Richards, 2011a;
Richards, 2014]. Although the Late Cretaceous igneous rocks are not associated with porphyrystyle deposits, we assume that their parental melts might have had the potential to form such
deposits. One explanation for the lack of Late Cretaceous porphyry deposits in the Apuseni
Mountains might be that the porphyry deposits are still buried at depth. It would also be possible
that the considerable addition of crustal melts, which is evidenced by the evolved
87
Sr/86Sr
ratios, strongly diluted the metal contents of the mantle-derived melts. The third and probably
most appropriate explanation might be the occurrence of explosive volcanism, resulting from
exsolution and rapid loss of volatiles (SO2, H2O) from the magma, which were then no longer
available for porphyry-style ore formation [Cloos, 2001; Richards, 2003].
Since the high and low
87
Sr/86Sr Miocene magmas are associated with Au-rich ore
deposits, and both groups evolved differently in the crust, the source of the metals was probably
the lithospheric mantle. This is compatible with a model proposed by Richards [2009; 2011a]
that invokes a second stage of melting of subduction-modified lithosphere for generating
particularly Au-rich porphyry deposits. Richards [2009] argues that small amounts of sulfide
phases can be left in the asthenosphere or lithosphere after a first stage of melting, despite the
high oxidation states of metasomatized mantle melts. This will only mildly deplete normal arcmagmas in Cu content, but it will virtually strip the magmas of highly siderophile elements like
Au [Richards, 2009]. Second stage melting of the subduction-modified lithospheric mantle
under sulfur-undersaturated conditions would then give rise to Au-rich magmas that can form
Au-Cu porphyry and epithermal Au deposits [Richards, 2009]. The change from Au-rich to AuTe-rich and increasingly Cu-rich ore deposits has been explained by continuous melting of the
source region during progressive rotation-induced extension, upwelling of the asthenosphere
and melting of the lithospheric mantle [Harris et al., 2013].
3.6
Summary and Conceptual Model
Here, we summarize the findings and suggest a conceptual model (Figure 3.11) that
links the Late Cretaceous and Miocene magmatism through lithospheric processes, and explains
the observed geochemical trends and metal endowment.
60
Chapter 3
Figure 3.11: Conceptual model for magma generation in the Apuseni Mountains (a) in the Late
Cretaceous and (b) in the Miocene. The schematic cross sections have a different scale for crust and
mantle. (a) Subduction of the Neotethys ocean triggered mantle metasomatism and partial melting of
the asthenosphere. During ascent of the mantle melts, Au+(Cu,Te)-rich sulfides were supposedly left in
the lithospheric mantle. Ultramafic to hydrous mafic cumulates might have formed in the lower crust.
The magmas presumably pooled in the middle crust, where they assimilated local crust and fractionated
a plagioclase-bearing assemblage. The thereby generated andesitic to dacitic melts ascended to shallow
crustal levels and partly evolved to rhyolites. Eruption of ignimbrites probably led to a loss of volatiles,
which prevented the formation of ore deposits. (b) Rotation of the continental units in the Miocene
induced extension, which presumably led to upwelling of the asthenosphere [e.g. Rosu et al., 2004;
Seghedi et al., 2004; Neubauer et al., 2005]. The upwelling asthenosphere possibly triggered re-melting
of the metasomatized lithospheric mantle [Rosu et al., 2004; Harris et al., 2013]. Re-melting of the
Au+(Cu, Te)-bearing sulfides [e.g. Richards, 2009] might explain the occurrence of unusually Au-rich
epithermal and porphyry-style deposits associated with the Miocene magmatism. The high 87Sr/86Sr
magmas possibly formed via assimilation and fractional crystallization at mid-crustal levels. The low
87
Sr/86Sr magmas probably interacted with small degree partial melts of previously formed hydrous
mafic cumulates, which gave rise to the high Sr, Ba and La contents and high ‘adakite-like’ Sr/Y ratios.
61
Chapter 3
Active subduction of the Neotethys ocean in the Late Cretaceous triggered arc
magmatism in the Apuseni Mountains and in adjacent arc segments at the European continental
margin [von Quadt et al., 2005; Georgiev et al., 2012; Kolb et al., 2013; Gallhofer et al., 2015].
The arc magmatism in the Apuseni Mountains is more silicic than the magmatism in the other
arc segments. Besides intermediate to silicic plutonic and shallow intrusive rocks, dacites and
rhyolitic ignimbrites occur only in this arc segment. Basaltic precursor melts were supposedly
generated by melting of the metasomatized mantle wedge (Figure 3.11a), and ascended into the
overlying continental crust, where they presumably underwent a polybaric evolution. The crust
in the Apuseni Mountains might have been relatively thick due to a preceding compressional
phase [e.g. Kounov and Schmid, 2013], but magmatic processes at lower crustal levels are not
recorded by the trace element signatures of the magmatism. However, this does not necessarily
mean that fractionation at lower crustal levels did not occur. Trace element signatures might
have been overprinted by the onset of plagioclase fractionation or by assimilation of wall rocks.
Hence, ultramafic and hydrous mafic cumulates might have also formed at lower crustal levels
in the Apuseni Mountains. The ascending mantle-melts probably stalled at mid crustal levels
(~20 km), and underwent fractional crystallization in the presence of plagioclase, coupled with
extensive assimilation of wallrocks consisting of gneisses or metapelites. Assimilation of
crustal wall-rocks was important in the formation of the Late Cretaceous magmas and estimates
based on isotopic compositions indicate addition of up to 60% of crustal melt. Batches of the
evolved andesitic to dacitic melts then ascended further into the upper crust, where they pooled
in a shallow magma chamber (< 6-8 km). The shallow crustal crystal mush might have been
periodically reheated due to incremental emplacement of the melt batches, which favored the
generation and expulsion of high silica rhyolitic melts [e.g. Bachmann et al., 2007]. We
speculate that the eruption of ignimbrites led to a loss of volatiles, which prevented the
formation of ore deposits.
In the Miocene, rotation of the Tisza and Dacia Mega-units, which host the Apuseni
Mountains (Figure 3.2), around the fixed Moesian platform led to extension and formation of
NW-SE striking grabens in the southern Apuseni Mountains [Rosu et al., 2004; Seghedi et al.,
2004; Neubauer et al., 2005]. The extensional processes presumably triggered asthenospheric
upwelling and remelting of the subduction-modified lithospheric mantle [Rosu et al., 2004;
Harris et al., 2013]. In our model, the high and low 87Sr/86Sr groups of Miocene rocks were
presumably both sourced in the lithospheric mantle, but evolved via distinct pathways in the
overlying continental crust (Figure 3.11b). The first batches of mantle melt assimilated crustal
material, which gave rise to high 87Sr/86Sr ratios, and presumably fractionated a plagioclase62
Chapter 3
bearing assemblage in mid to upper crustal levels. Successive mantle melts passed the crust
without considerable crustal contamination probably due to proceeding extension [Rosu et al.,
2004; Harris et al., 2013], and yielded the low
87
Sr/86Sr Miocene rocks. Instead, the mantle
melts might have assimilated small degree partial melts of originally mantle-derived hydrous
mafic cumulates (amphibole-gabbro or hornblendites) in the lower crust, which would explain
the rather extreme concentrations of Sr, Ba and La, and their ‘adakite-like’ high Sr/Y ratios.
Melting of the cumulates was probably assisted by pre-heating by the early pulses of Miocene
mantle melts ascending to shallower crustal levels. The high 87Sr/86Sr magmas are associated
with Au-rich deposits, which might be explained by remelting of Au-bearing sulfides left in the
lithospheric mantle after a first phase of mantle melting in the Late Cretaceous [e.g. Richards,
2009]. The low
87
Sr/86Sr rocks are progressively enriched in Te and Cu owing to continued
melting of the lithospheric mantle [Harris et al., 2013].
Acknowledgements
Acknowledgements: This study was supported by the Swiss National Science
Foundation grants 200020-146681 and 20021-146651 and SNF scopes projects JRP
7BUPJ062396 and IZ73ZO_128089. Ioan Seghedi provided essential help during joint field
work in Romania and we are most grateful for his regional knowledge of the Miocene
magmatism. This study incorporates results from Caroline Harris, a previous PhD student
working on the Miocene magmatism. We thank Ramon Aubert, Markus Wälle, Marcel
Guillong, Lydia Zehnder and Muhammed Usman for support in the laboratories.
63
Chapter 3
64
Chapter 4
4. Tectonic significance of new U-Pb ages of
Jurassic Ophiolites and associated
Granitoids in the South Apuseni Mountains,
Romania
4.1
Abstract
Ophiolites in the South Apuseni Mountains belong to the Eastern Vardar ophiolitic unit
and represent remnants of the Neotethys ocean. The Jurassic ophiolitic unit contains tholeiitic
and calc-alkaline magmas and presently overlies the continental Dacia Mega-Unit. New U-Pb
zircon ages, and Sr and Nd isotope ratios for the two distinct magmatic series give new insights
into their tectono-magmatic history. The tholeiitic ophiolites show dominantly MORB-type
affinities, but occasionally are slightly enriched in Th and U, and depleted in Nb, which
indicates that they probably formed in a marginal or back-arc basin. Four gabbros from the
ophiolites yielded U-Pb ages between 158.9 and 155.9 Ma (Late Jurassic). The tholeiitic
ophiolites are intruded by granitoids and volcanic equivalents of the calc-alkaline series, which
show trace element signatures characteristic for subduction-enrichment (high LILE, low
HFSE). The low 87Sr/86Sr ratios (0.703836-0.704550) and high
143
Nd/144Nd ratios (0.512599-
0.512616) of the calc-alkaline series overlap with the ophiolites (0.703863-0.704303 and
0.512496-0.512673), and exclude its generation due to obduction-induced melting of
metasedimentary material deposited on the continental margin, or in a collisional or postcollisional setting. Instead, the isotope systematics of the calc-alkaline series are consistent with
an origin due to intra-oceanic subduction in an island arc setting. The island arc granitoids
intruded the ophiolites between 158.6 and 152.9 Ma (Late Jurassic). The island arc granitoids
do not show substantial crustal input and thus they must have already been emplaced in the
ophiolites before the entire sequence was obducted onto the continental margin. Hence, the age
65
Chapter 4
of the youngest granitoid yields an estimate for the maximum age of obduction of the South
Apuseni nappes (~153 Ma, Late Kimmeridgian).
4.2
Introduction
The Neotethys ocean is essential for unraveling the Mesozoic tectonic evolution of the
Carpathian-Dinaride-Balkan orogen. It formerly divided the European continental margin from
the Adriatic microplate that was attached to the African plate [e.g. Schmid et al., 2008].
Convergence between the European and the African plate, which presently is still going on, led
to closure of the Neotethys ocean during several stages in the Mesozoic. Remnants of the
Neotethys ocean are therefore preserved in different tectonic positions, which previously led
workers to distinguish several ocean basins in the Carpathian-Dinaride-Balkan orogen [e.g.
Csontos and Vörös, 2004; Robertson et al., 2009]. Here, we adopt the “one-ocean” concept that
was first formulated by Bernoulli and Laubscher [1972] and more recently illustrated in Schmid
et al. [2008].
According to this concept, the three different types of ophiolitic units occurring between
the Carpathian-Balkan and Dinaride orogens are sourced in the same long-lived Neotethys
ocean. The Sava zone [sensu Schmid et al., 2008] is the only real suture zone and is related to
the final closure of the Neotethys ocean in this region at the end of the Cretaceous [Karamata,
2006; Ustaszewski et al., 2010] (Figure 4.1). The large sheets of ophiolites preserved to the east
and west of this suture zone (the Eastern and Western Vardar ophiolitic units sensu Schmid,
Figure 4.1) lie on top of continental units. Although still disputed, many authors agree now that
the Eastern and Western Vardar ophiolitic units were obducted onto the continental units during
Late Jurassic to Early Cretaceous times [Pamić et al., 2002; Karamata, 2006; Schmid et al.,
2008; Hoeck et al., 2009; Ionescu et al., 2009; Kounov and Schmid, 2013].
The obducted Eastern Vardar ophiolitic unit forms only a narrow strip parallel to the
Sava suture zone in Serbia, Macedonia (FYROM) and Greece, but crops out in a larger area in
the South Apuseni nappes of the Apuseni Mountains of Romania. The South Apuseni nappes
overlie the continental Dacia Mega-Unit [Schmid et al., 2008] (Figure 4.1) and consist of
Middle Jurassic ophiolites sensu strictu and Late Jurassic granitoids and their volcanic
equivalents presumably formed in an island arc setting [Savu et al., 1981; Bortolotti et al., 2002;
Nicolae and Saccani, 2003; Bortolotti et al., 2004]. The ages of the magmatic sequences and
the timing of obduction have been inferred mainly from stratigraphic evidence, and apart from
K-Ar ages, which can easily be disturbed, only scarce Re-Os ages [Zimmerman et al., 2008] are
available for molybdenite veins occurring in the granitoids. Hence, more reliable
66
Chapter 4
geochronological constraints are necessary to refine the tectonic model for the emplacement of
the granitoids and the timing of obduction of the South Apuseni nappes in the Late Jurassic to
Early Cretaceous.
The distinction between the ophiolitic and island arc series in the South Apuseni nappes
is based on detailed petrographic observations and geochemistry [Nicolae and Saccani, 2003;
Bortolotti et al., 2004]. However, radiogenic isotope ratios, which would give additional
insights into the source of the two magmatic series, are not available for the South Apuseni
nappes. Radiogenic isotope ratios have been used to detect crustal assimilation in granitoids
from other parts of the Eastern Vardar ophiolitic unit, and led to a fundamentally different
interpretation of the tectonic setting of some of the alleged island arc granitoids [Saric et al.,
2009]. Hence, Sr and Nd isotope analyses of the granitoids in the Apuseni Mountains are
essential to confirm or exclude their origin in an island arc setting.
Here, we present new U-Pb zircon ages for the Apuseni ophiolites and associated
granitoids. We obtained whole rock major and trace element data, and Sr and Nd isotope ratios
of the same samples and combined our geochemical data with a previously published dataset
[Bortolotti et al., 2004]. We discuss the tectonic significance of the geochemical data and the
new U-Pb ages, and refine the existing geodynamic models [Bortolotti et al., 2002; Ionescu et
al., 2009; Kounov and Schmid, 2013; Reiser, 2015] for the obduction of the South Apuseni
nappes.
Figure 4.1: Geological map of the Alpine-Carpathian-Dinaride orogen [modified from Schmid et al.,
2008]. The red lines denote the suture zones of the Alpine Tethys and Neotethys oceans.
67
Chapter 4
4.3
Geological Setting of the Apuseni Mountains
The Apuseni Mountains are offset from the arcuate Carpathian orogen and surrounded
by the Pannonian basin to the west and north, the Transylvanian basin to the east, and the South
Carpathians to the south (Figure 4.1). The Apuseni Mountains are situated at the contact
between the continental Tisza and Dacia Mega-Units, and in their southern part the obducted
Eastern Vardar ophiolitic unit overlies the Dacia Mega-Unit [e.g. Csontos and Vörös, 2004;
Schmid et al., 2008]. All tectonic units are unconformably covered by Late Cretaceous posttectonic sediments (Gosau-type sediments) [Schuller et al., 2009], and intruded by Late
Cretaceous and Miocene calc-alkaline igneous suites (Figure 4.2).
Figure 4.2: Geological map of the Apuseni Mountains, redrawn from official Romanian geological maps
(scale 1:200000) and modified after Balintoni [1994] and Kounov and Schmid [2013].
The continental Mega-Units comprise poly-phase (Variscan and Alpine) metamorphic
basement and sedimentary cover nappes, which partly show European faunal affinity [Csontos
68
Chapter 4
and Vörös, 2004; Haas and Pero, 2004; Iancu et al., 2005]. The lowermost Bihor and Codru
nappe systems are ascribed to the Tisza Mega-Unit [e.g. Csontos and Vörös, 2004] and consist
of partly Neoproterozoic basement intruded by Paleozoic granites, and Late Paleozoic to
Mesozoic cover sediments [Pana et al., 2002; Balintoni et al., 2009; Balintoni et al., 2010].
The uppermost Biharia nappe system has previously been regarded as integral part of the Tisza
Mega-Unit [e.g. Csontos and Vörös, 2004], but is now assigned to the Dacia Mega-Unit
[Schmid et al., 2008]. This is due to the observation that the Eastern Vardar ophiolitic unit
commonly overlies nappes of the Dacia Mega-Unit [Schmid et al., 2008]. The Biharia nappe
system is composed of polymetamorphic Variscan basement [Dallmeyer et al., 1999] and postVariscan to Mesozoic metasediments [Csontos and Vörös, 2004]. Parts of the Biharia nappe
were buried and metamorphosed during the Early Cretaceous [e.g. Dallmeyer et al., 1999].
The Eastern Vardar ophiolitic unit overlies the Biharia nappe system of the Apuseni
Mountains [e.g. Schmid et al., 2008]. The part of the Eastern Vardar ophiolitic unit cropping
out in the Apuseni Mountains is also known as South Apuseni nappes or ophiolites [Bortolotti
et al., 2004; Kounov and Schmid, 2013]. In this contribution we use the term South Apuseni
ophiolitic unit for this composite unit that is subdivided into a presumably Middle Jurassic
tholeiitic series (the ophiolites sensu strictu) and Late Jurassic calc-alkaline granitoids and their
volcanic equivalents, which probably originated in an island arc setting [Savu et al., 1981;
Bortolotti et al., 2002; Nicolae and Saccani, 2003; Bortolotti et al., 2004]. In the Mureş valley,
the South Apuseni ophiolitic unit is strongly dismembered into several tectonic slices, and
according to Bortolotti et al. [2002] the MORB-type ophiolite sequence is ~2000 m thick. From
bottom to top, the ophiolites comprise an intrusive section of layered and isotropic gabbros and
rare ultramafic cumulates, which are followed upwards by a sheeted dyke complex, and an
extensive volcanic sequence. The volcanic sequence includes massive and pillow-lava basalts
[Bortolotti et al., 2002]. The age of the ophiolites is poorly constrained by a wide spread of KAr whole rock ages (~168-139 Ma) [Nicolae et al., 1992] and Callovian to Oxfordian (~166157 Ma) radiolarites deposited on top of the ophiolites [Lupu et al., 1995]. The ophiolites are
intruded by calc-alkaline granitoid bodies (mainly granites and granodiorites, subordinately
diorites) and overlain by massive lava flows ranging in composition from basalts to rhyolites
[Nicolae and Saccani, 2003; Bortolotti et al., 2004]. Two Re-Os ages (~159 and 160 Ma)
[Zimmerman et al., 2008] have been obtained from molybdenite associated with granitoids in
the Mureş valley. According to Romanian geological maps (1:50000), the volcanic cover is
supposedly older than the granitoids. Late Jurassic (Kimmeridgian-Tithonian, ~154-145 Ma) to
69
Chapter 4
Early Cretaceous shallow water limestones [Bortolotti et al., 2002; Şerban et al., 2004] and
Early Cretaceous flysch overlie the ophiolites and granitoids in the Mureş valley.
The Alpine tectonic evolution started in the Late Jurassic with the tectonic emplacement
of the South Apuseni nappes onto the continental Biharia nappe system [e.g. Schmid et al.,
2008; Kounov and Schmid, 2013]. Late Jurassic to Early Cretaceous shallow water carbonates
unconformably overlie the South Apuseni nappes as well as the continental basement units in
the Trascau mountains (east of the study area) [Bucur and Săsăran, 2005], which implies that
obduction must have occurred before deposition of these carbonates probably in the Late
Oxfordian [Kounov and Schmid, 2013]. Late Jurassic obduction was probably triggered by
closure of the Eastern Vardar branch of Neotethys, which was rooted between the Tisza and
Dacia continental units [e.g. Kounov and Schmid, 2013]. Alternatively, Late Cretaceous
narrowing of the Neotethys embayment due to opening of the Alpine Tethys might have led to
simultaneous obduction of the Eastern and Western Vardar ophiolitic units [Reiser, 2015].
Subsequent Early Cretaceous (‘Austrian’) internal nappe stacking and sinistral strike-slip
movements might have facilitated the transport of the obducted South Apuseni nappes over a
seemingly large distance [Reiser, 2015].
In a later, intra-Turonian phase that formed the present-day nappe stack in the Apuseni
Mountains, the Dacia Mega-Unit and the ophiolitic unit were thrust over the Tisza Mega-Unit
[Haas and Pero, 2004; Schmid et al., 2008; Kounov and Schmid, 2013]. The entire Apuseni
nappe stack underwent substantial clockwise rotation (~90°) in the Cenozoic [Pǎtraşcu et al.,
1990; Pǎtraşcu et al., 1994; Márton et al., 2007], when it moved around the Moesian platform
and invaded the still open Carpathian embayment [e.g. Fügenschuh and Schmid, 2005;
Ustaszewski et al., 2008].
4.4
Methods
Fused glass beads of rock powder mixed with Lithium-Tetraborate (1:5) were analysed
for major element oxides by x-ray fluorescence (XRF) using an Axios PANalytical WD-XRF
spectrometer at ETH Zürich. Trace elements and rare earth elements were determined on freshly
broken surfaces of the same glass beads by laser-ablation inductively-coupled-plasma mass
spectrometry (LA-ICP-MS). For more details on the analytical procedure see Appendix 1 of
this thesis. Strontium and Neodymium isotopes were analysed on 50 to 70 mg whole rock
powder digested in HF and HNO3. Sr and Nd were subsequently separated by ion-exchange
chromatography in columns with SrSpec , TRUSpec and LnSpec Eichrom resins [Pin et al.,
1994]. Strontium was loaded onto outgassed Re single filaments with HNO3 and Ta emitter,
70
Chapter 4
whereas the Nd fraction was loaded onto double filaments with 2N HCl. Both were analysed
with a Thermo Scientific TritonPlus mass spectrometer at ETH Zürich. Repeated measurements
of NBS 987 and JNd-i yielded a 87Sr/86Sr mean ratio of 0.710234 ± 0.000004 and a 143Nd/144Nd
mean ratio of 0.512100 ± 0.000003 (2σm), respectively.
Following standard zircon separation, the zircons were pre-treated by chemical abrasion
to remove those domains of zircon grains that have lost Pb [Mattinson, 2005]. The abraded
zircons were mounted in epoxy resin and polished to expose the grain centre.
Cathodoluminescence (CL) pictures were taken to resolve inherited cores and freshly grown
rims of single zircon grains using a FEI Quanta 200 FEG at the Scope M facility at ETH Zürich.
In-situ laser-ablation inductively-coupled-plasma mass spectrometry (LA-ICP-MS) was
performed with an Element-XR SF-ICP-MS (Thermo Fisher, Bremen, Germany) coupled to an
193 nm Excimer laser (Resonetics Resolution S155-LR). The Excimer laser was operated at 5
Hz with a spot size of 30 µm [detailed setup in von Quadt et al., 2014]. Blocks of several 100
analyses were run including a minimum of 15 analyses of the primary standard zircon GJ-1.
Secondary standard zircons Plesovice, 91500 and Temora were analysed for data quality
control. No common lead correction was performed on the data acquired with this system owing
to low count rates of 204Pb. The obtained raw data was imported into Iolite [Paton et al., 2010;
Paton et al., 2011] and reduced and corrected for downhole fractionation using the VizualAge
data reduction scheme [Petrus and Kamber, 2012]. The software Isoplot 3.75 [Ludwig, 2012]
was used to prepare Concordia diagrams and calculate weighted average ages. We report
weighted average
206
Pb/238U ages and calculate the uncertainty as 2 standard deviation of
concordant and overlapping LA-ICP-MS ages in a population of analyses of each samples, as
a conservative measure of age uncertainty.
4.5
Results
Our geochemical dataset of the Jurassic South Apuseni ophiolites and granitoids with
their volcanic equivalents is supplemented with geochemical data previously published by
Bortolotti et al. [2004] in order to show differences between the Jurassic ophiolites and
granitoids. Additionally, we present Sr and Nd isotope ratios for the ophiolites and the
granitoids. Mean 206Pb/238U ages were obtained by in-situ LA-ICP-MS dating of zircons from
the ophiolites and granitoids.
4.5.1 Sampling, Field Observations and Petrography
For this study, the western occurrences of the South Apuseni ophiolites and granitoids
in the Mureş valley were sampled (Figures 4.2, 4.9). The South Apuseni ophiolites comprise an
71
Chapter 4
intrusive section, a sheeted dyke complex and pillow lavas [Bortolotti et al., 2002; Bortolotti et
al., 2004]. The intrusive ophiolites consist of layered and isotrope gabbros (samples DG064,
DG068, DG082), and occasionally microgabbros occur. Apart from one basalt with columnar
jointing, the sheeted dykes, massive and pillow-lava basalts were not sampled due to the smaller
chance of finding zircons in mafic volcanics. The gabbros have a coarse crystalline texture and
consist of twinned plagioclase and interstitial to hypidiomorphic clinopyroxene, which is
partially replaced by chlorite and opaque phases (Figure 4.3a). Two dioritic samples consist of
plagioclase, chlorite and amphibole, which are probably both secondary phases replacing
primary pyroxenes, and accessory apatite and opaques. There are three major granitoid bodies
intruding the ophiolites, the Săvârşin (sample DG062, DG065), Cerbia (DG069, DG072) and
Căzăneşti bodies (Figure 4.9). They mainly consist of granites and granodiorites, and the
marginal zones are sometimes made up of quartz-diorites. Additionally, felsic volcanics occur
west of the Săvârşin intrusive body on top of the ophiolites, and we sampled a dyke-like
porphyritic rhyolite (DG067). An andesite (DG081) was sampled close to Birtin. The granites
to granodiorites have (in)equigranular textures and consist of plagioclase, which partly occurs
as large phenocrysts, and alkalifeldspar and quartz, which frequently show granophyric
intergrowths (Figure 4.3b). Biotite and occasionally green amphibole are the dominant mafic
phases and are partially replaced by secondary chlorite. Apatite and opaques are common
accessory phases.
Figure 4.3: Textures of ophiolites and associated granitoids. (a) Coarse crystalline gabbro (DG064)
consisting of plagioclase and clinopyroxene, which is partially replaced by chlorite. Photomicrograph
was taken under plane polarized light. (b) Granite (DG072) showing granophyric intergrowth of kfeldspar and quartz. Photomicrograph was taken under crossed polarizers. bt: biotite, chl: chlorite, cpx:
clinopyroxene, k-fsp: alkali-feldspar, plag: plagioclase, qtz: quartz.
72
Chapter 4
4.5.2 Major and Trace Element Characteristics
The major and trace element whole rock characteristics generally show differences
between the ophiolites and the granitoids. On Miyashiro’s [1974] discrimination plot of
FeOt/MgO versus SiO2 the granitoids and their volcanic equivalents plot in the calc-alkaline
field, whereas most ophiolites plot in the tholeiitic field due to their higher FeOt contents
(Figure 4.4).
Figure 4.4: FeOt/MgO versus SiO2 classification diagram after Miyashiro [1974]. Jurassic ophiolites
and associated granitoids are plotted for comparison [including data from Bortolotti et al., 2004].
In major element variation plots (Figure 4.5), the granitoid samples show a decrease of
MgO, Al2O3 and CaO with increasing SiO2. Na2O and P2O5 initially increase and then decrease
with increasing SiO2, whereas Al2O3 is constant until 60 % SiO2 and decreases at higher SiO2
contents. The ophiolites generally have lower SiO2 contents, lower Al2O3 and K2O contents,
and higher CaO and MgO contents than the calc-alkaline granitoids and their volcanic
equivalents. Na2O and P2O5 contents increase with increasing SiO2 to partly higher values than
those observed in the calc-alkaline series. The calc-alkaline series and the tholeiitic ophiolites
are clearly distinct in the TiO2 and FeOt versus Mg-number plots (Figure 4.6). The ophiolites
trend towards high TiO2 and FeOt contents with decreasing Mg-number, whereas TiO2 and
FeOt contents decrease with decreasing Mg-number in the calc-alkaline series.
Chondrite normalized rare earth element (REE) plots (Figure 4.7a) show an elevation
of light REE (LREE), relatively flat heavy REE and none to slightly negative Eu anomalies
(Eu/Eu*=0.6 to 1.0) in the calc-alkaline granitoids. By contrast, the normalized ophiolite
patterns are flat and LREE are slightly depleted or have the same abundance as HREE (Figure
4.7b).
73
Chapter 4
Figure 4.5: Variation diagrams for major element oxides versus SiO2. Jurassic ophiolites and associated
granitoids are plotted [including data from Bortolotti et al., 2004].
In N-MORB (normal mid ocean ridge basalt) normalized trace element plots, the calcalkaline series shows an enrichment in large ion lithophile elements (LILE, e.g. Ba, K, Pb) and
depletions in high field strength elements (HFSE, e.g. Nb, Ta, Ti) (Figure 4.7c). The ophiolites
generally have lower trace element contents. Their LILE contents are generally low, but Th and
U are slightly enriched, and they lack a strong depletion in P (Figure 4.7d).
74
Chapter 4
Figure 4.6: (a) TiO2 versus Mg# and (b) FeOt versus Mg# show distinct trends of the Jurassic ophiolites
and associated granitoids [including data from Bortolotti et al., 2004].
Figure 4.7: Trace element characteristics of the Jurassic ophiolites and associated granitoids. (a) and (b)
Chondrite normalized rare earth element (REE) patterns. (c) and (d) N-MORB normalized trace element
patterns. Gray bands indicate the range of data from Bortolotti et al. [2004]. Normalization values are
from Sun and McDonough [1989].
4.5.3 Sr and Nd Isotopes
The initial Sr and Nd isotopes of the Jurassic ophiolites and granitoids from the South
Apuseni nappes partly overlap (Figure 4.8). The granitoids have a range of age corrected
87
Sr/86Sr ratios from 0.703836 to 0.704550, and the
75
143
Nd/144Nd ratios vary only little from
Chapter 4
0.512599-0.512616. The gabbros from the ophiolitic series have
0.704303) similar to the granitoids, but their
143
87
Sr/86Sri ratios (0.703863-
Nd/144Ndi ratios have a wider spread from
0.512496-0.512673 (Figure 4.8). The ophiolites as well as the granitoids have less radiogenic
Nd isotope ratios than present-day mid ocean ridge basalt (MORB)-type mantle [Stracke et al.,
2005]. The gabbros have initial
143
Nd/144Nd ratios similar to the Jurassic Demir Kapija
ophiolites (Macedonia, FYROM) [Božović et al., 2013]. The granitoids from the South Apuseni
nappes overlap with the isotopic composition of andesitic and adakite-like volcanic rocks from
FYROM [Božović et al., 2013]. The Fanos granite (northern Greece) has slightly higher
87
Sr/86Sri and lower
143
Nd/144Ndi ratios than the South Apuseni granitoids. Granitoids from
Serbia and FYROM, in contrast, have distinctly higher 87Sr/86Sr ratios and lower 143Nd/144Nd
ratios [Saric et al., 2009].
Figure 4.8: Initial Sr and Nd isotope ratios for the Jurassic ophiolites and associated granitoids from the
South Apuseni nappes. The isotope ratios were age corrected for 155 Ma using the whole rock Sr, Rb,
Sm and Nd concentrations. Field for present-day MORB from Stracke et al. [2005]; fields of Jurassic
ophiolites and andesitic and ‘adakite-like’ volcanic rocks from FYROM (former Yugoslavian Republic
of Macedonia) from Božović et al. [2013]; high 87Sr/86Sr granitoids from FYROM, fields for the Fanos
granite (northern Greece) and low and high 87Sr/86Sr granitoids from Serbia are from Saric et al. [2009].
4.5.4 In situ U-Pb LA-ICP-MS Zircon Dating
In order to further improve the succession of ophiolites and granitoids inferred from
field relations, we dated 4 ophiolites and 5 granitoids. Two ophiolites, a gabbro (DG064) and a
76
Chapter 4
diorite (DG063), sampled close to the Săvârşin granitoid body (Figure 4.9), yielded mean
206
Pb/238U ages of 157.7±5.0 Ma (2 stdev) and 155.9±5.4 Ma, respectively. An ophiolitic gabbro
(DG068) from the Cerbia area has a mean
206
Pb/238U age of 158.9±5.2 Ma. Another gabbro
(DG082) from the ophiolitic series that occurs in association with the Căzăneşti granitoid
(Figure 4.9) yielded a mean
206
Pb/238U age of 158.2±4.4 Ma. The ophiolitic suite has an age
range from 158.9 to 155.9 Ma. The granitoids and their volcanic equivalents are slightly
younger than the ophiolites, except for the sample from the Săvârşin body (DG062) that yielded
a mean 206Pb/238U age of 158.6±2.9 Ma. Two granitoids from the Cerbia body (DG069, DG072)
have mean 206Pb/238U ages of 154.8±2.9 and 152.9±4.0 Ma, respectively. An andesite sampled
east of Căzăneşti (DG081, Figure 4.9) is 154.5±4.5 Ma old. A rhyolite sample (DG067) from
the volcanic cover overlying the ophiolites yielded a mean 206Pb/238U age of 155.0±3.0 Ma. The
granitoids have an age range from 158.6 to 152.9 Ma.
Figure 4.9: Geological map of the South Apuseni nappes in the Mureş valley, redrawn from official
Romanian geological maps (scale 1:200000: sheet 17 Brad; scale 1:50000: sheets 72c Săvârşin, 72d
Roşia Nouă, 73a Hălmagiu, 88b Lăpugi-Coştei and 89a Gurasada). Purple stars and numbers indicate
sample location and mean 206Pb/238U ages of granitoids, green stars and numbers indicate sample
location and mean 206Pb/238U ages of gabbros from the ophiolitic series.
4.6
Discussion
In this section, the tectonic implications of the geochemical and geochronological data
will be discussed. In accordance with previous studies, we will show that the ophiolites and
granitoids are genetically unrelated, and we will resolve the tectonic setting of the granitoids
77
Chapter 4
and their volcanic equivalents. The new age constraints provide valuable information about the
timing of events, and will be used to refine existing geodynamic models [e.g. Bortolotti et al.,
2002; Ionescu et al., 2009] for the obduction of the Eastern Vardar ophiolitic unit.
4.6.1 Tectonic Significance of the Geochemical Data
The major element trends observed in the ophiolites and granitoids (Figures 4.4, 4.5,
4.6), especially the differences in the TiO2 and FeOt evolution with increasing differentiation
[Bortolotti et al., 2004], clearly indicate that they were not derived by fractionation from the
same parental magma. The ophiolites follow a tholeiitic trend with high FeOt/MgO ratios
(Figure 4.4), whereas most granitoids and their volcanic equivalents are depleted in FeOt, which
is generally characteristic for calc-alkaline suites that experience early crystallization and
removal of Fe-Ti oxides [Miyashiro, 1974]. The ophiolitic suite is characterized by relatively
flat REE patterns (Figure 4.7b) similar to modern mid ocean ridge basalts, but occasionally has
slightly elevated LILE (Th and U) and a Nb depletion, which are more typical for suprasubduction zone ophiolites [e.g. Dilek and Furnes, 2011]. In contrast to the ophiolitic suite, the
calc-alkaline granitoids are conspicuously enriched in LILE (e.g. Ba, K, Pb) (Figure 4.7c). They
also show depletions in HFSE (Nb, Ta) and pronounced negative Ti-anomalies. These particular
trace element patterns are commonly taken as indicators of a subduction-enriched mantle source
[e.g. Hawkesworth et al., 1997; Woodhead et al., 2001; Elliott, 2003]. In summary, although
the major and trace element characteristics of the ophiolites point to a mid ocean ridge setting
[Bortolotti et al., 2002], they also show some features of supra-subduction zone ophiolites and
probably formed in a marginal to back arc basin. The geochemical characteristics of the
granitoids indicate their subduction-related origin [Bortolotti et al., 2002; Nicolae and Saccani,
2003].
The geochemical characteristics are the most compelling evidence for an island-arc
setting of the calc-alkaline granitoids [Nicolae and Saccani, 2003; Bortolotti et al., 2004].
Moreover, Nicolae and Saccani [2003] have shown that the observed mineral compositions and
whole rock geochemical trends can be explained by closed-system fractional crystallization of
a mantle-derived melt. Nevertheless, a possible crustal contribution might have been missed,
because the major and trace element composition of continental crust is similar to that of
subduction-related rocks and crustal contamination can only be reliably detected by radiogenic
isotopes, which have so far not been available for Apuseni granitoids. Calc-alkaline granitoids
associated with obducted ophiolites can form due to a variety of tectonic processes, and might
be subduction-related, form during obduction or subsequent collision [e.g. Barbarin, 1999].
78
Chapter 4
Recently, Saric et al. [2009] have studied Jurassic granitoid intrusions from the southern parts
of the Eastern Vardar ophiolitic units in Serbia, Macedonia and Greece, and have identified
distinct tectonic settings of granitoid formation based on the isotopic compositions.
In contrast to the high 87Sr/86Sr granitoids in FYROM and northern Greece (Figure 4.8),
which intrude continental basement and the overlying ophiolites, the granitoids and their
volcanic equivalents in the Mureş valley are exclusively associated with ophiolites [Bortolotti
et al., 2002]. Granitoids intrude and crosscut all levels of the ophiolitic sequence, and basaltic
to rhyolitic lava flows are deposited on top of the ophiolites [Bortolotti et al., 2002]. The
Jurassic granitoids are, however, never associated with the underlying continental Biharia
nappe system and they do not contain any xenoliths of continental basement, which excludes
their generation in a post-collisional setting after ophiolite emplacement. Moreover, if collision
between the Biharia nappe system (Dacia) and the Tisza Mega-Unit occurred at all [Reiser,
2015], it presumably occurred later in the Early Cretaceous [Schmid et al., 2008; Kounov and
Schmid, 2013], well after intrusion of the Jurassic granitoids. This still leaves two alternative
possibilities, (1) the granitoids intruded the ophiolitic sequence before emplacement onto the
continental Biharia nappe system in a subduction-related setting, or (2) the granitoids formed
due to obduction-induced melting of the continental Biharia nappe system or continent-derived
sedimentary material deposited in the trench.
Granitoids formed due to process (2) generally have high
143
87
Sr/86Sr ratios and low
Nd/144Nd ratios characteristic for continental crust or sediments, and have been identified in
the Oman ophiolites [Cox et al., 1999; Searle and Cox, 1999] and in the Serbian part of the
Eastern Vardar ophiolitic unit [Saric et al., 2009]. The initial Sr and Nd isotope ratios of the
granitoids from the Mureş valley, however, do not indicate any considerable crustal
contribution (Figure 4.8). The initial
143
87
Sr/86Sr ratios (0.703836 to 0.704550) and initial
Nd/144Nd ratios (0.512599 to 0.512616) overlap with the gabbros from the ophiolitic series,
and might thus reflect the subduction-enriched mantle source of the parental melt. Moreover,
the Sr and Nd isotope systematics of the granitoids from the Mureş valley resemble Jurassic
subduction-related rocks associated with the Demir Kapija ophiolites in Macedonia (FYROM)
(Figure 4.8) [Božović et al., 2013]. Hence, the Sr and Nd isotope systematics are consistent with
scenario (1), in which the calc-alkaline granitoids and volcanics in the Mureş valley formed in
a subduction-related island arc setting [Bortolotti et al., 2002; Nicolae and Saccani, 2003]. This
implies that the granitoids intruded the ophiolite sequence before their joint emplacement onto
the continental Biharia nappe system.
79
Chapter 4
4.6.2 Tectonic Significance of the U-Pb Zircon Ages
Mean
206
Pb/238U ages obtained for gabbros from the ophiolitic sequence of the South
Apuseni nappes range from 159 to 156 Ma (Oxfordian-Early Kimmeridgian). The calc-alkaline
granitoids (the Săvârşin and Cerbia bodies) yielded mean 206Pb/238U ages between 159 and 153
Ma (Oxfordian-Late Kimmeridgian), and a calc-alkaline rhyolite and an andesite yielded ages
of 156 Ma and 154 Ma (Kimmeridgian), respectively. The new U-Pb zircon ages for the Cerbia
and Săvârşin granitoids are partly younger than the Re-Os molybdenite ages (159.1 and 159.8
Ma) reported by Zimmerman et al. [2008]. The ages of the ophiolites and the granitoids overlap
within error, which indicates that ophiolite formation and intrusion of the subduction-related
granitoids is restricted to a narrow time span. Nevertheless, the granitoids tend to have slightly
younger mean ages (Figure 4.10).
Radiogenic age constraints are also available for ophiolites and granitoids in the Greek
and Macedonian (FYROM) parts of the Eastern Vardar ophiolitic unit (Figure 4.10). Samples
from the ophiolitic series in Greece yielded zircon U-Pb ages of 166.6±1.8 Ma, 169.2±1.4 Ma,
160.0±1.2 and 165.3±2.2 Ma for the Guevgueli, Thessaloniki and Sithonia ophiolites,
respectively [Zachariadis, 2007]. Additionally, Bonev et al. [2015] reported U-Pb ages for a
layered gabbro (158.4±1.9 Ma) and a rhyolite (148.9±1.0 Ma) of the Sithonia ophiolite.
Gabbros from the Greek Evros and Samothraki ophiolites yielded zircon U-Pb ages of 169±2
Ma and 160±5 Ma, respectively [Koglin et al., 2007]. A U-Pb age of 166.4±1.2 Ma has been
reported for a gabbro from the Demir Kapija ophiolite in Macedonia (FYROM) [Božović et al.,
2013]. Anders et al. [2005] have reported U-Pb ages of 158±1 Ma for the calc-alkaline Fanos
granite and 164±2 Ma for the Mikro Dassos rhyolite, which intrude the Guevgueli ophiolites.
The Gerakini diorite (Greece) yielded a U-Pb age of 172.8±1.2 Ma [Bonev et al., 2015]. The
Monopigadon granite intruding the Thessaloniki ophiolite (Greece) yielded a U-Pb age of
159±1 Ma [Meinhold et al., 2009] and a diorite from the nearby Chortiatis unit is 159.1±4.2 Ma
old [Zachariadis, 2007]. Ar-Ar dating of feldspar of a subduction-related rock intruding the
Demir Kapija ophiolite yielded an age of 164±0.5 Ma [Božović et al., 2013]. In summary, the
ophiolites and granitoids from the southern parts of the Eastern Vardar ophiolitic unit tend to
be older than their counterparts in the Apuseni Mountains. The mid to Late Jurassic radiogenic
ages of the ophiolites indicate that their formation did not occur simultaneously in the back-arc
or marginal basins of the Eastern Vardar branch of Neotethys. Although the granitoids formed
at different tectonic stages during subduction, collision or obduction, they mostly intruded the
ophiolites in the Late Jurassic.
80
Chapter 4
Figure 4.10: Mean 206Pb/238U ages of the ophiolites (blue error bars) and associated granitoids and their
volcanic equivalents (red error bars) of the Eastern Vardar ophiolitic unit. The data is sorted
geographically from north (Apuseni Mountains) to south (Greece). 206Pb/238U ages for the South Apuseni
ophiolites (SAp oph.) and the Săvârşin, Cerbia and Căzăneşti granitoids and one rhyolite are from this
study. U-Pb ages of the Sithonia, Metamorphosis, Guevgueli and Thessaloniki ophiolites (Greece) are
from Zachariadis [2007] and Bonev et al. [2015]. U-Pb ages of the Samothraki and Evros ophiolites
(Greece) are from Koglin et al. [2007]. U-Pb age of the Demir Kapija ophiolite (FYROM) and Ar-Ar
Fsp age of a calc-alkaline rock associated with the Demir Kapija ophiolite are from [Božović et al.,
2013]. Re-Os ages for the Cerbia and Săvârşin granitoids are from Zimmerman et al. [2008]. U-Pb ages
of the Fanos granite and the Mikro Dassos rhyolite (Greece) are from Anders et al. [2005], of the
Monopigadon granite (Greece) is from Meinhold et al. [2009], of one Chortiatis diorite (Greece) is from
Zachariadis [2007], and of the Gerakini diorite is from Bonev et al. [2015]. Grey bar indicates the
stratigraphic age of shallow water limestones deposited on top of the South Apuseni nappes [Şerban et
al., 2004; Bucur and Săsăran, 2005]. The maximum age of obduction (~153 Ma) of the South Apuseni
nappes was inferred from the youngest calc-alkaline intrusive (DG072). Note that the errors reported
for literature data are standard-errors-of-the-mean, whereas we report 2 standard deviation (2σ) of
overlapping and concordant TIMS or LA-ICP-MS ages in a population of analyses of each sample. This
leads to seemingly larger errors of samples from this study.
The timing of obduction of the Eastern Vardar ophiolitic unit is poorly constrained,
because a metamorphic sole that could probably be dated has so far not been detected.
According to Kounov and Schmid [2013], the obduction must have occurred before the
deposition of Late Jurassic to Early Cretaceous shallow water carbonates [Bucur and Săsăran,
2005], which unconformably overlie the South Apuseni nappes as well as the continental
basement units. The new U-Pb ages of the calc-alkaline series put an additional constraint on
the timing of obduction. As discussed earlier, the calc-alkaline granitoids are clearly
81
Chapter 4
subduction-related and have low initial 87Sr/86Sr and high initial
143
Nd/144Nd ratios. The non-
crustal Sr and Nd isotope ratios exclude their generation via obduction-induced melting of
continental sedimentary material or in a post-collisional setting. This means that the granitoids
in the Mureş valley must have already been in the ophiolites at the time of obduction. Obduction
can therefore only have started after emplacement of the last calc-alkaline granitoid. Hence, the
age of the youngest granitoid, i.e. 152.9±4.0 Ma (Late Kimmeridgian), is the maximum age of
obduction of the South Apuseni nappes. The maximum age of obduction approximately fits the
deposition age of the shallow water limestones in the Mureş valley (Kimmeridgian-Tithonian)
[Şerban et al., 2004].
4.6.3 Geodynamic Model
Here, we compare and discuss several existing tectonic models for the Jurassic evolution
of the South Apuseni nappes [Bortolotti et al., 2002; Ionescu et al., 2009; Kounov and Schmid,
2013; Reiser, 2015] and include our findings to improve the models. Based on the new
geochemical data and U-Pb ages, we distinguish three distinct events in the Jurassic, (1)
ophiolite formation in the Oxfordian to Early Kimmeridgian, (2) intra-oceanic subduction and
generation of a calc-alkaline island arc mainly in the Kimmeridgian, and (3) obduction starting
in the Latest Kimmeridgian to Tithonian (Figures 4.10, 4.11).
The geochemical data of the South Apuseni ophiolitic series occasionally show a subtle
subduction influence, which indicates that the ophiolites probably did not form in a typical mid
ocean ridge setting [Saccani et al., 2001; Bortolotti et al., 2002], but in a marginal or remnant
back arc basin of the Neotethys ocean [Bortolotti et al., 2004; Ionescu et al., 2009; Kounov and
Schmid, 2013]. This marginal basin might have been rooted between the continental Tisza and
Dacia Mega-Units (Figure 4.11a) [Săndulescu, 1984; Schmid et al., 2008; Ionescu et al., 2009;
Kounov and Schmid, 2013]. However, the back-arc ophiolites might have also been generated
within the main Neotethys ocean, similar to the setting observed further south in the Demir
Kapija and Guevgueli ophiolites (Figures 4.10, 4.11) [Zachariadis, 2007; Božović et al., 2013].
Back-arc spreading apparently started earlier in the southern parts of the Eastern Vardar
ophiolitic unit, at ~169 Ma, and was restricted to the Middle Jurassic [Koglin et al., 2007;
Zachariadis, 2007; Božović et al., 2013], whereas the oceanic lithosphere preserved in the South
Apuseni ophiolites is Late Jurassic. The reason for this north-ward (in present-day coordinates)
propagation of ocean floor formation is unclear.
82
Chapter 4
Figure 4.11: (a) Paleogeographic sketch map, and (b), (c), (d) schematic cross-sections (not to scale) for
the Late Jurassic. (a) The sketch map shows the distribution of continental and oceanic units in
southeastern Europe in the Late Jurassic (Oxfordian, ~160 Ma). Modified from Schmid et al. [2008] and
Reiser [2015]. (b), (c), (d) The cross-sections depict the evolution of a marginal or back-arc basin, onset
of intra-oceanic subduction and island arc formation, and obduction of the ophiolitic and island arc series
onto the continental Dacia Mega-Unit, similar to the evolution in the southern parts of the Eastern Vardar
ophiolitic unit [Zachariadis, 2007; Božović et al., 2013]. In contrast to the sketch presented in Kounov
and Schmid [2013, their Figure 12], the South Apuseni nappes are not rooted between the Tisza and
Dacia Mega-Unit [Reiser, 2015] in these cross-sections, which renders the Tisza Mega-Unit “invisible”
in these profiles. Modified from Kounov and Schmid [2013], Schmid et al. [2008] and Božović et al.
[2013].
Shortly after formation of the ophiolitic series, the granitoids formed in an intra-oceanic
subduction zone superimposed on the ophiolitic series (Figure 4.11) [Bortolotti et al., 2002;
Ionescu et al., 2009]. The granitoids exclusively intrude the ophiolitic series and probably
represent an early island arc sequence. The granitoids show trace element signatures
characteristic of subduction-related igneous rocks, and were presumably formed by closedsystem fractional crystallization of a parental mantle melt [Nicolae and Saccani, 2003], which
is also consistent with the observed Nd and Sr isotope ratios. The dip direction of the intraoceanic subduction zone is somewhat controversial. Several authors assume that the subducting
oceanic lithosphere was dipping towards NE to E (Jurassic coordinates) beneath oceanic
lithosphere and the adjacent continental Dacia Mega-Unit [e.g. Bortolotti et al., 2002; Božović
et al., 2013]. However, with regard to the later obduction, southwestward subduction of oceanic
lithosphere beneath oceanic lithosphere seems to be more appropriate to explain the island arc
granitoids (Figure 4.11) [Schmid et al., 2008; Ionescu et al., 2009; Kounov and Schmid, 2013].
83
Chapter 4
After island arc magmatic activity, the ophiolites and island arc granitoids were
presumably emplaced onto the continental Biharia nappe system of the Dacia Mega-Unit
(Figure 4.11) [Kounov and Schmid, 2013]. The calc-alkaline granitoids and volcanics in the
Mureş valley are clearly subduction-related and none of the granitoids formed due to obductioninduced melting or in a post-collisional setting. Therefore, the obduction onto the continental
margin can only have started after intrusion of the youngest island arc granitoid, and the inferred
maximum age of obduction of the South Apuseni nappes is ~153 Ma (Latest Kimmeridgian).
This age fits fairly well with the age of obduction inferred from the deposition of shallow water
carbonates (Late Kimmeridgian-Tithonian, ~154-145 Ma) [Şerban et al., 2004] on top of the
ophiolitic unit in the Mureş valley and further east also on continental basement.
The island arc series in the Mureş valley is slightly younger than subduction-related and
post-collisional igneous rocks in the southern parts of the Eastern Vardar ophiolitic unit (Figure
4.10) [Anders et al., 2005; Božović et al., 2013]. The observed north-ward (present-day
coordinates) propagation of the obduction, which can be inferred from the progressively
younger ages of the calc-alkaline suites, would be consistent with a model recently proposed
by Reiser [2015]. He suggested that the Eastern Vardar ophiolitic unit was obliquely obducted
onto the entire length of the Dacia Mega-Unit due to the narrowing of the Meliata bay during
concomitant opening of the Alpine Tethys ocean further north. The model by Reiser [2015] can
explain emplacement of the Eastern Vardar ophiolites by a single process operating in the entire
Neotethys realm. In contrast to most of the previous models [Săndulescu, 1984; Ionescu et al.,
2009; Kounov and Schmid, 2013], Reiser’s model does not require a marginal basin rooted
between the Dacia and Tisza Mega-Units. He explains further transport of the South Apuseni
nappes over the Dacia Mega-Unit by Early Cretaceous internal nappe stacking in the Dacia
Mega-Unit assisted by strike-slip movements [Reiser, 2015]. The schematic cross-sections
(Figures 4.11b, c, d) incorporates this hypothesis and places the South Apuseni nappes in a
back-arc position close to the Dacia Mega-Unit, similar to a model proposed by Božović et al.
[2013] for the Demir-Kapija ophiolites (FYROM).
4.7
Conclusions
This contribution aims at clarifying the Late Jurassic tectonic evolution of the ophiolites
and associated granitoids and their volcanic equivalents exposed in the South Apuseni nappes
of Romania, which are considered to be the northern continuation of the Eastern Vardar
ophiolitic unit [e.g. Schmid et al., 2008]. The geochemical data coupled with the Sr and Nd
isotopes indicate that the calc-alkaline granitoids derived from a mantle source that was
84
Chapter 4
enriched in subduction components (e.g. LILE). Because their isotopic composition overlaps
with that of the ophiolites, we infer that the calc-alkaline series was formed in an island arc
setting with none to only minor contribution of subducted sediment to the mantle source. The
low Sr and high Nd isotope ratios of the granitoids furthermore suggest that they probably did
not form due to obduction-induced melting of metasediments deposited on the continental
margin, or in a collisional to post-collisional setting. Based on occasionally observed slightly
elevated LILE (Th and U) and Nb depletions in the ophiolites, we interpret the ophiolites not
as pristine MORB-type, but place them in a marginal to back-arc setting. The new U-Pb ages
confirm that the ophiolites are slightly older than the island-arc granitoids. For the abovediscussed geochemical reasons, the island-arc granitoids must have intruded the ophiolites
before their joint emplacement on top of the continental margin. Therefore, the age of the
youngest granitoid places a constraint on the timing of obduction. The maximum age of
obduction of the South Apuseni nappes is ~153 Ma (Late Kimmeridgian). We assume that the
tectonic evolution of ocean spreading, island arc formation and obduction in the South Apuseni
ophiolitic unit was similar to that in the southern parts of the Eastern Vardar ophiolitic unit.
Hence, we suggest that the South Apuseni ophiolites formed in a marginal to back-arc basin
close to the continental Dacia Mega-Unit.
Acknowledgement
This study was supported by the Swiss National Science Foundation grants 200020146681 and 20021-146651 and SNF scopes projects JRP 7BUPJ062396 and IZ73ZO_128089.
Ioan Seghedi provided essential help during joint field work in Romania. We are most grateful
for Stefan M. Schmid’s regional geological insight. We thank Ramon Aubert, Markus Wälle,
Marcel Guillong, Lydia Zehnder and Muhammed Usman for support in the laboratories.
85
Chapter 4
86
5.General Conclusions and Outlook
5. General Conclusions and Outlook
This thesis presents major and trace element data, radiogenic isotope (Sr-Nd) data and
U-Pb zircon ages for the Late Cretaceous magmatism in the Banat region and Apuseni
Mountains in Romania. Combining the new data with data previously collected in a series of
studies within the ‘fluids and mineral resources’ group at ETH allowed us to resolve the
tectono-magmatic history of the ABTS magmatic arc on a regional scale and use these data for
an overall analysis of the subduction magmatism in the realm of the Neotethys in southeastern
Europe.
The reconstruction of the Late Cretaceous situation presented in chapter 2 shows that
the magmatic arc had a rather simple geometry typical of a segmented continental magmatic
arc. Subduction-related magmatism was approximately contemporaneous in all arc segments,
but did not start at exactly the same time in each segment. It occurred over 25 Ma, from 92.2 to
66.8 Ma. Across-arc younging of the magmatic products towards the paleo-trench provides
clear evidence for the gradual steepening of the subducting Neotethys slab. This north to south
age progression is accompanied by distinct isotopic trends in some segments, which point to an
increasing contribution of mantle melts, and probably resulted from asthenospheric corner flow.
The segmented nature of the arc may be explained by different states of tectonic stress in the
arc segments, which has been inferred from concomitant shear zones and sedimentary basins.
The centrally located Panagyurishte and Timok segments were subjected to only mild
transtension during contemporaneous shearing, which favoured the formation of long-lived
lower crustal magma chambers and high-pressure amphibole fractionation, which ultimately
led to an enrichment in volatiles and metals. Hence, porphyry-style and epithermal ore deposits
preferentially occur in these central arc segments. By contrast, the Eastern Srednogorie segment
underwent strong orthogonal extension, which prevented the formation of such deposits. Postemplacement deformation of the entire arc and associated extensional tectonics favoured the
preservation of ore deposits in this relatively old metallogenic belt.
The third chapter infers a genetic link between the Late Cretaceous arc magmatism and
the Miocene post-subduction magmatism in the Apuseni Mountains. Re-melting of subductionmodified mantle and cumulates which were left in the lower crust by the Late Cretaceous arc
magmatism gave rise to the distinct geochemical characteristics of the Miocene magmas. The
Late Cretaceous arc magmatism presumably underwent a polybaric evolution from mid to upper
87
5.General Conclusions and Outlook
crustal levels. Lower crustal processes might have been overprinted by subsequent fractionation
of plagioclase or by assimilation of wall rocks. At mid crustal levels mantle-derived melts
fractionated a plagioclase- and amphibole-bearing assemblage. Here, assimilation of a
maximum of 60% partial melts of crustal rocks, which is clearly indicated by Sr and Nd isotope
ratios, occurred. Evolved andesitic to dacitic melts then ascended to the upper crust, where high
silica rhyolitic melts presumably formed in a mush zone. Eruptive volcanism might have
prevented the formation of ore deposits in the Late Cretaceous. The Miocene post-subduction
magmatism presumably formed due to extension-induced re-melting of the lithospheric mantle,
which was responsible for the unusually Au-rich nature of the Miocene magmas. Two groups
of Miocene magmas, characterized by their different Sr isotope ratios, evolved via distinct paths
during their ascent in the continental crust. (1) A high 87Sr/86Sr group assimilated local crustal
wall rocks and fractionated a plagioclase-bearing assemblage probably at mid crustal levels. (2)
A low
87
Sr/86Sr group possibly assimilated small-degree melts of hydrous mafic cumulates
which were left in the crust by the Late Cretaceous arc magmatism. This gave rise to extreme
enrichments in LILE and LREE and ‘adakite-like’ trace element signatures.
The fourth chapter aims to refine the tectonic setting and evolution of the Jurassic calcalkaline granitoids and ophiolites, which crop out in the southern Apuseni Mountains. The
radiogenic isotope composition (Sr-Nd) of the calc-alkaline series overlaps that of the ophiolitic
series, and does not show a significant crustal input. This suggests that the granitoids cannot
have formed due obduction-induced melting of sediments deposited on the continental margin
or in a collisional to post-collisional setting. The isotope ratios support intra-oceanic subduction
and a previously proposed island arc setting. The new U-Pb zircon ages for the ophiolites (158.9
- 155.9 Ma) and calc-alkaline series (158.6 - 152.9 Ma) show that both were formed in the Late
Jurassic. However, the calc-alkaline series has slightly younger ages than the ophiolites. The
lack of crustal contamination of the calc-alkaline granitoids indicates that they must have
already been emplaced in the ophiolites before the entire sequence was obducted onto the
continental margin. Therefore, the age of the youngest granitoid places a constraint on the
timing of obduction. The maximum age of obduction of the sequence is ~153 Ma (Late
Kimmeridgian).
Recommendations for future research:

Detailed investigation of the sedimentary basins associated with the Late Cretaceous
magmatic arc would yield insights into their formation. Notably, basin-bounding structures
88
5.General Conclusions and Outlook
in relation to volcanic and intrusive magma emplacement could resolve the origin of basin
opening and give insight into the stress state of the crust, which is important for subvolcanic
ore formation. Additionally, detailed U-Pb age dating of primary in-situ volcanics and
plutonic rocks interlayered with the sediments are recommended to better constrain the
timing of basin formation.

A detailed mapping and sampling campaign with a focus on the volcanic products in the
Vladeasa massif would be strongly recommended to better define the timing and volume
of the volcanic activity. It is still unclear, where the caldera is located and whether one or
more eruptions from more than one caldera occurred.
40
Ar-39Ar dates for the volcanic
products might be a useful tool to address this question and determine eruption ages.

I would recommend determining the geochemistry of mineral phases and least squares
modeling combined with a trace element FC model and AFC calculations for isotope ratios
to better constrain differentiation by fractional crystallization of the Late Cretaceous
Apuseni magmas.

Further isotope work (whole rock Pb isotopes, Hf and oxygen isotopes on zircons, and
possibly oxygen isotopes on pyroxenes) should be performed to strengthen the genetic link
between Late Cretaceous and Miocene Apuseni magmas proposed in this thesis.
89
5.General Conclusions and Outlook
90
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Acknowledgements
Acknowledgements
I am grateful to many people who supported me during my PhD. First of all, I would
like to thank Albrecht von Quadt for his supervision throughout the years. He introduced me to
U-Pb zircon dating, isotope geochemistry and mass spectrometry and assisted me with the
measurements. I also profited from many fruitful discussions about the Bulgarian and Serbian
geology and his thorough knowledge of the ABTS belt. Albrecht knows nearly every single
result of age dating ever done in that region and is a ‘living geochronology encyclopedia’.
I would also like to thank Christoph Heinrich, who gave me the opportunity to start this
project in the ‘fluids and mineral deposits group’. I am grateful for the numerous scientific
discussions we had and for his guidance while writing the chapters. He taught me how to more
clearly structure and express my ideas.
Irena Peytcheva introduced me to the secrets of the clean lab. Without her assistance,
the practical part of this thesis would not have been possible. I also appreciate her support during
field work, her patience and constant geniality.
Ioan Seghedi provided essential help during joint field work in Romania and I am
grateful that he took the time for assisting me.
I would also like to thank Stefan M. Schmid for joining us in the Apuseni Mountains
and sharing his immense knowledge and unique understanding about the geology in the AlpineCarpathian-Dinaide realm. I am grateful for his thorough review of my second chapter and the
long and fruitful geological discussions.
I would like to thank Ramon Aubert for introducing me to the art of zircon separation
from scratch, which includes a variety of different lab methods. I greatly appreciate Muhammed
Usman’s help with sample separation. I am grateful for the assistance of Markus Wälle and
Marcel Guillong with LA-ICP-MS work. Lydia Zehnder’s assistance with XRF analyses is very
much appreciated. I would also like to thank Britt Meyer for her help with all kinds of things.
I am grateful to all present and former ‘fluid and mineral deposits group’ members and
visiting guests for the good working atmosphere and the stimulating discussions over lunch.
Last, but not least, I would like to thank my parents and family for encouraging and
supporting me during my education.
102
Appendices
Appendices
Appendix 1
Methods.
Appendix 2
Sample number, region, coordinates, type, lithology, mineralogy and texture.
Appendix 3
Major (wt%) and trace (ppm) element composition of the studied samples.
Appendix 4
Sr and Nd isotope data.
Appendix 5
Calculated mean 206Pb/238U ages.
Appendix 6
Map and sampling locations for the Banat region and Apuseni Mountains.
Appendix 7
Tectonic map of the ABTS belt summarizing the crystallization ages.
Appendix 8
Concordia and weighted mean 206Pb/238U age plots for LA-ICP-MS and TIMS
dating.
Appendix 9
U-Pb dates of inherited zircons.
Appendix 10 Hf isotope plots.
Electronic Appendices:
Single zircon dates obtained by LA-ICP-MS and TIMS dating.
Hf isotope data.
CL pictures of zircons.
103
Appendix 1
Appendix 1: Methods.
Samples were crushed by high voltage fragmentation (SelFrag) using a 2 mm sieve and
a fraction of ca. 60 g was representatively divided with a sample splitter and subsequently
pulverised in a tungsten-carbide mill. Fused glass beads of rock powder mixed with LithiumTetraborate (mixing ratios 1:5 to 1:8 for siliceous rocks) were analysed for major element
oxides by x-ray fluorescence (XRF) using an Axios PANalytical WD-XRF spectrometer at
ETH Zürich. Loss on ignition (LOI) was determined at 1050°C. Trace elements and rare earth
elements were determined on freshly broken surfaces of the same glass beads by laser ablationinductively coupled plasma-mass spectrometry (LA-ICP-MS). The LA-ICP-MS system
comprises an Excimer 193 nm laser system operated with a 60 µm laser beam diameter and
10Hz repetition rate connected to a quadrupole ICP-MS system (Elan6100). NIST 610 glass
reference material was used as an external calibration standard and major element oxide
concentrations (CaO or Al2O3), previously determined by XRF, were used as internal standards.
Analytical reproducibility is better than 5 to 10% for concentrations above 1 ppm and better
than 15% for lower concentrations [Günther et al., 2001]. The SILLS software package was
applied for data reduction [Guillong et al., 2008]. For every sample, three spot analyses on the
glass bead were averaged to obtain the reported trace element concentration. Major and trace
element data are reported in Supporting Information Table S1.
Strontium and Neodymium isotopes were analysed on 50 to 70 mg whole rock powder
digested in HF and HNO3. Sr and Nd were subsequently separated by ion-exchange
chromatography in columns with SrSpec , TRUSpec and LnSpec Eichrom resins [Pin et al.,
1994]. Strontium was loaded onto outgassed Re single filaments with HNO3 and Ta emitter,
whereas the Nd fraction was loaded onto double filaments with 2N HCl. Both were analysed
with a Thermo Scientific TritonPlus mass spectrometer operated in static mode at ETH Zürich.
Repeated measurements of NBS 987 and JNd-i yielded a
87
Sr/86Sr mean ratio of
0.710234±0.000004 and a 143Nd/144Nd mean ratio of 0.512100±0.000003, respectively. Sr and
Nd isotopic ratios are reported in Supporting Information Table S2.
Zircons for U-Pb dating were liberated from whole rocks by high voltage fragmentation
(SelFrag), concentrated by density separation using methylene-iodide and hand-picked under a
binocular. All zircons for both in-situ Laser ablation inductively-coupled-plasma mass
spectrometry (LA-ICP-MS) dating and high precision isotope dilution-Thermal Ionisation Mass
Spectrometry (ID-TIMS) dating were pre-treated by chemical abrasion to remove those
domains of zircon grains that have lost Pb [Mattinson, 2005]. Zircons were annealed at 860°C
for 48 hours, transferred to 3 ml screw-top PFA Savillex vials and leached in HF in a PARR
104
Appendix 1
digestion vessels at 180° for 12 to 15 hours. The leachate was pipetted out and the zircons were
washed in 6.2 N HCl on a hotplate at 85°C for approximately 24 hours and rinsed with ultrapure
H2O and double-distilled acetone.
For in-situ LA-ICP-MS dating, the abraded zircons were mounted in epoxy resin and
polished to expose the grain centre. Cathodoluminescence (CL) and back-scattered electron
(BSE) pictures were acquired with a FEI Quanta 200 FEG at EMEZ/Scope M ETH Zürich to
resolve inherited cores and freshly grown rims. Data was acquired using two distinct LA-ICPMS systems at ETH Zürich, an Elan 6100 ICP-MS (PerkinElmer, Norwalk, CT, USA) coupled
to an 193 nm ArF-Excimer laser ablation system and an Element-XR SF-ICP-MS (Thermo
Fisher, Bremen, Germany) coupled to an 193 nm Excimer laser (Resonetics Resolution S155LR). The ArF-Excimer laser was operated at 10 Hz with a spot size of 40 µm, the Excimer laser
was operated at 5 Hz with a spot size of 30 µm [detailed setup in von Quadt et al., 2014]. On
the first system, measurements were performed in blocks of 20-24 analyses, bracketed before
and after by three analyses of the primary standard zircon GJ-1 [602.9± 2.3 Ma, Jackson et al.,
2004]. Plesovice [Slama et al., 2008] was used as secondary standard for data quality control.
The data was not corrected for common lead due to low 204Pb count rates. Data reduction and
fractionation correction was performed using Glitter software [van Achterbergh et al., 2001].
On the second system, longer blocks of up to several 100 analyses were run, including a
minimum of 15 analyses of the primary standard zircon GJ-1. Secondary standard zircons
Plesovice, 91500 and Temora were analysed for data quality control. No common lead
correction was performed on the data acquired with this system owing to low count rates of
204
Pb. Raw data obtained by this system were imported into Iolite [Paton et al., 2010; Paton et
al., 2011] and reduced and corrected for downhole fractionation using the VizualAge [Petrus
and Kamber, 2012] data reduction scheme. Single zircon dates are reported in Supporting
Information Table S4. The software Isoplot 3.75 [Ludwig, 2012] was used to prepare Concordia
diagrams and calculate weighted average ages (Appendix 8). It has been shown that weighted
average ages of chemically abraded zircon crystals overlap with ages obtained by ID-TIMS
method and can therefore be interpreted as geologically accurate ages [von Quadt et al., 2014].
We report weighted average
206
Pb/238U ages and calculate the uncertainty as 2 standard
deviation, as a conservative measure of age uncertainty. Clearly discordant and apparently older
and younger zircons were rejected from the weighted average calculation. We interpret slightly
older zircons as antecrysts and assume that chemical abrasion was not entirely effective in
younger zircons.
105
Appendix 1
For ID-TIMS dating single zircon grains were selected, weighed and loaded into Teflon
vessels or microcapsules. Zircons were spiked with the 202-205Pb/233-235U spike of the Earthtime
(ET) Working Group (http://www.earth-time.org) and dissolved in concentrated HF and 7N
HNO3 for 5-6 days (Teflon vessels) or 3 days (microcapsules). After dissolution, zircons were
dried down and re-dissolved in 3N HCl. Pb and U were separated by anion exchange
chromatography and loaded on outgassed re-filaments for analysis with a Thermo Scientific
TritonPlus TIMS equipped with a digital ion counting system of a MasCom multiplier. Both Pb
and U (as UO2) isotope ratios were measured sequentially with the electron multiplier. The
mass fractionation of Pb and U was corrected through the double ET
202-205
Pb/233-235U spike.
The composition of the total procedural Pb blank of the Teflon vessels was
206
Pb/204Pb
18.6±0.71 (1σ%), 207Pb/204Pb 15.62±1.03 and 208Pb/204Pb 38.08±0.98. The microcapsules had
a total procedural Pb blank of
208
206
Pb/204Pb 18.3±0.15 (1σabs),
207
Pb/204Pb 15.51±0.15 and
Pb/204Pb 37.75±2.41. The model Th/U[zircon] was calculated from radiogenic 208Pb/206Pb ratio
assuming concordance. Measured ratios were reduced using REDUX. The 206Pb/238U ratio was
corrected for initial
230
Th disequilibrium using Th/U[magma] of the whole rock. Single zircon
TIMS data is reported in Supporting Information Table S5. Concordia ages were calculated
using Isoplot 3.75 [Ludwig, 2012] from zircon dates overlapping within error. In cases of zircon
dates that did not overlap within error and showed a larger scatter, we interpreted the youngest
concordant zircon date as crystallisation age.
Guillong, M., D. L. Meier, M. M. Allan, C. A. Heinrich, and B. W. D. Yardley (2008), Appendix A6: SILLS: A Matlab-based
program for the reduction of Laser Ablation ICP-MS data of homogeneous materials and inclusions, in Mineralogical
Association of Canada Short Course, edited by P. Sylvester, pp. 328-333, Mineralogical Association of Canada,
Vancouver, B. C.
Günther, D., A. von Quadt, R. Wirz, H. Cousin, and V. 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, Mikrochimica Acta, 136(3-4), 101-107, doi: 10.1007/s006040170038.
Jackson, S. E., N. J. Pearson, W. L. Griffin, and E. A. Belousova (2004), The application of laser ablation-inductively coupled
plasma-mass spectrometry to in situ U-Pb zircon geochronology, Chem. Geol., 211(1-2), 47-69, doi:
10.1016/j.chemgeo.2004.06.017.
Ludwig, K. J. (2012), User's Manual for Isoplot 3.75 A Geochronological Toolkit for Microsoft Excel, in Berkeley
Geochronology Center Special Publication, edited, p. 75, Berkeley Geochronolgy Center, Berkeley CA.
Mattinson, J. M. (2005), Zircon U–Pb chemical abrasion (“CA-TIMS”) method: Combined annealing and multi-step partial
dissolution analysis for improved precision and accuracy of zircon ages, Chem. Geol., 220(1–2), 47-66, doi:
10.1016/j.chemgeo.2005.03.011.
Paton, C., J. Hellstrom, B. Paul, J. Woodhead, and J. Hergt (2011), Iolite: Freeware for the visualisation and processing of mass
spectrometric data, Journal of Analytical Atomic Spectrometry, 26(12), 2508-2518, doi: 10.1039/C1JA10172B.
Paton, C., J. D. Woodhead, J. C. Hellstrom, J. M. Hergt, A. Greig, and R. Maas (2010), Improved laser ablation U-Pb zircon
geochronology through robust downhole fractionation correction, Geochemistry, Geophysics, Geosystems, 11(3),
Q0AA06, doi: 10.1029/2009GC002618.
Petrus, J. A., and B. S. Kamber (2012), VizualAge: A Novel Approach to Laser Ablation ICP-MS U-Pb Geochronology Data
Reduction, Geostandards and Geoanalytical Research, 36(3), 247-270, doi: 10.1111/j.1751-908X.2012.00158.x.
106
Appendix 1
Pin, C., D. Briot, C. Bassin, and F. Poitrasson (1994), Concomitant separation of strontium and samarium-neodymium for
isotopic analysis in silicate samples, based on specific extraction chromatography, Analytica Chimica Acta, 298(2),
209-217, doi: 10.1016/0003-2670(94)00274-6.
Slama, J., et al. (2008), Plesovice zircon - A new natural reference material for U-Pb and Hf isotopic microanalysis, Chem.
Geol., 249(1-2), 1-35, doi: 10.1016/j.chemgeo.2007.11.005.
van Achterbergh, E., C. Ryan, S. E. Jackson, and W. L. Griffin (2001), Appendix 3: data reduction software for LA-ICP-MS,
in Mineralogical Association of Canada Short Course, edited by P. Sylvester, pp. 239-243, Mineralogical Association
of Canada, Vancouver, B. C.
von Quadt, A., D. Gallhofer, M. Guillong, I. Peytcheva, M. Waelle, and S. Sakata (2014), U-Pb dating of CA/non-CA treated
zircons obtained by LA-ICP-MS and CA-TIMS techniques: impact for their geological interpretation, Journal of
Analytical Atomic Spectrometry, 29(9), 1618-1629.
107
Appendix 2: Sample number, region, coordinates, type, lithology, mineralogy and texture.
Region1
Longitude
Latitude
Type2
Lithology
Mineralogy3
Texture4
DG001a
DG001b
DG002
DG003
DG004
DG005
DG006
DG007
DG008
DG009
DG010
DG011
DG012
DG013
DG014
DG015
DG016
DG017
DG018
DG019
DG020
DG021
DG022
DG023
DG024
DG025
DG026
DG027
DG028
DG029
DG030
DG031
DG032
DG033
DG034
DG035
EBR
EBR
EBR
EBR
EBR
EBR
EBR
CBR
CBR
CBR
CBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
44.90342
44.90242
45.04177
45.04523
45.14360
45.14223
45.14223
44.98033
44.70908
44.71462
44.72030
45.39073
45.38213
45.38485
45.39048
45.37973
45.38038
45.37973
45.37973
45.38158
45.38230
45.38230
45.43167
45.50685
45.49330
45.49330
45.42353
45.42353
45.42353
45.42353
45.34135
45.34123
45.28318
45.22977
45.23403
45.25503
22.36840
22.36882
22.25618
22.24733
22.26335
22.26425
22.26425
22.09930
21.91625
21.91753
21.92815
21.77433
21.75392
21.75597
21.68745
21.69708
21.69667
21.69708
21.69708
21.70378
21.70587
21.70587
21.73883
21.78150
21.78112
21.78112
21.88387
21.88387
21.88387
21.88387
21.77267
21.77238
21.74907
21.60528
21.61143
21.58283
bsm
bsm
d
si
bsm
d
si
si
si
si
si
i
bsm
i
i
d
i
d
i
i
i
i
i
d
i
bsm
i
d
i
d
si
si
i
d
i
i
por. granite
por. granite
andesite
granodiorite
gneiss
lamprophyre dyke
por. granodiorite
por. granodiorite
por. granodiorite
por. granodiorite
por. diorite
granodiorite
gneiss
granodiorite
qtz monzonite
por. diorite
monzogranite
aplite dyke
monzonite
monzogranite
qtz monzonite
monzogranite
granodiorite
dacite
microgabbro
gabbroic dyke
granodiorite
lamprophyre dyke
diorite
gabbro
por. syeno-diorite
por. gabbro
granodiorite
diorite
gabbro
granodiorite
plag, ksp, mafics repl. by chl, GM plag-dom.; acc. ap, op
plag, ksp, mafics repl. by chl, GM plag-dom.; sec. carb; styloliths
zoned plag, green amph in fine GM (plag,chl, qtz); acc. ap, op; sericit., chlorit.
plag (sericit.); mafics repl. by chl, op, ser, carb
plag, qtz (inequigranular); bt, amph in layers, beginning chlorit.
plag, mafics repl. by chl
zoned plag, bt (partly repl. by chl) in fine GM (plag-dom.) acc. ap, op
plag, green amph, bt, qtz in GM (plag, bt, qtz), acc. ap, op
plag (ser.), qtz, mafics repl. by chl+op in medium GM (pl, qtz), op
coarse plag (sericit), green amph, bt (partly repl.by chl) in medium GM, op
plag, bt (partly chlorit.) in medium GM (pl-dom.), acc. ap, op
coarse plag, ksp, qtz, green amph, bt
xenomorph qtz and plag, layers of bt (partly chlorit.)
coarse plag, ksp (ser.), qtz, bt (chlorit.), other mafics (repl. by chl+op)
plag, qtz, interstitial bt (partly repl. by chl.), cpx, acc. op, ap
qtz, plag (sericit.), mafics repl. by chl+op in fine GM
coarse plag, ksp, qtz, green amph, bt (partly chlorit.), acc. tit, op, ap
medium qtz, ksp, plag, bt (repl. by op.), granophyric qtz+ksp
plag, qtz, bt, green amph, acc. op
plag, qtz, ksp, bt (partly chlorit.)>green amph, ksp (perthitic)
plag, qtz, bt, green amph, cpx, acc. op, ap
plag, qtz, ksp, bt (partly chlorit.), green amph, cpx, acc. op, ap
altered, plag (sericit.), mafics repl. by op
altered, sericitised, mafics repl. by ser+chl with op. rims
plag, cpx, GM (lath shaped plag, bt), op
por.
por.
por.
equ.
fol.
por.
por.
por.
por.
por.
por.
equ.
fol.
equ.
equ.
por.
equ.
equ.
equ.
equ.
equ.
equ.
equ.
por.
equ.
idiomorphic cpx in GM (lath shaped plag), mafics repl. by sec. chl, carb
poikilitic plag, qtz, green amph, op in medium GM (plag, qtz, amph)
fine bt, plag (laths), cpx, op, mafics repl. by chl, carb veinlets
plag, green amph, rare bt in fine GM (plag, amph), acc. op
green amph (poikilitic, along rims sec. chl), cpx, zoned plag in fine matrix
bt, green amph, zoned plag, ksp, acc. ap, op
poikilitic plag, bt, poikilitic cpx, acc. ap, op, tit
plag (beginning sericit.) amph, bt (partly repl. by chl), acc. op
plag, ksp, qtz, bt (partly replaced by chl), cpx
equ.
equ.
poik.
equ.
por.
por.
equ.
equ.
equ.
equ.
108
Appendix 2
108
Sample#
Region1
Longitude
Latitude
Type2
Lithology
DG036
DG037
DG038
DG039
DG040
DG041
DG042
DG043
DG044
DG045
DG046
DG047
DG048
DG049
DG050
DG051
DG052
DG053
DG054
DG055
DG056
DG057
DG058
DG059
DG060
DG061
DG062
DG063
DG064
DG065
DG066
DG067
DG068
DG069
DG070
DG071
WBR
SWBR
SWBR
SWBR
SWBR
SWBR
CBR
PR
PR
PR
PR
PR
PR
PR
PR
PR
PR
PR
PR
PR
PR
PR
SA
SA
PR
PR
SA
SA
SA
SA
SA
SA
SA
SA
SA
SA
45.25443
45.04588
45.06473
45.05822
44.86438
44.72872
44.92145
45.53118
45.60933
45.60427
45.61250
45.52905
45.52905
45.58887
45.68897
45.69118
45.74475
45.87953
45.82560
45.88673
45.88543
45.87947
45.95298
45.95088
45.76602
45.88885
46.04486
46.04486
46.02864
46.02688
46.02688
46.06277
46.03670
46.04880
46.04880
46.05705
21.58330
21.73188
21.72152
21.72255
21.73062
21.70613
21.92040
22.39663
22.33803
22.32917
22.35858
22.32430
22.32430
22.14705
22.07718
22.15722
22.18503
22.63757
22.52752
22.54935
22.54815
22.55077
22.35808
22.71883
22.76350
22.89778
22.27600
22.27600
22.26728
22.26510
22.26510
22.24113
22.44645
22.45410
22.45410
22.49377
i
i
i
si
i
si
d
si
i
si
i
si
d
i
si
i
d
bas
bas
v
v
v
i
v
bas
nv
IAG
oph
oph
IAG
oph
IAG
oph
IAG
IAG
IAG
diorite
gabbro
granodiorite
por. granodiorite
granodiorite
por. granodiorite
andesite
por. diorite
granodiorite
por. granite
granodiorite
por. granodiorite
basaltic andesite
granodiorite
por. diorite
diorite
granodiorite
basanite
basanite
andesite
andesite
dacite
microdiorite
andesite
basanite
andesite
granite
diorite
gabbro
granite
basalt
por. granite
gabbro
granodiorite
diorite
granite
Mineralogy3
Texture4
lath-shaped plag, bt (partly replaced by chl), cpx, amph (interstitial)
plag, qtz, bt, green amph, acc. op, ap
coarse plag (sericite rims), bt, green amph (partly poikilitic)
coarse plag, qtz, bt, green amph (partly chlorit.), fine interstitial plag, acc. op
plag, bt, op in fine grained GM (plag, bt, op), acc. ap, chlorit., sericit.
zoned plag (ser rims), bt, green amph, op in fine GM, weak chlorit.
plag, amph, cpx (mafics partly repl. by chl+op), ap, GM (pl, chl, ser)
plag, qtz, green amph in medium GM, sec. chl, ser, ep
plag (sericit), embayed qtz, mafics repl. by chl+op+ser, fine GM
plag, qtz, bt, green amph, acc. op
qtz, plag (sericit.) in fine GM (pl laths), sec. chl, ser, carb
plag, qtz, cumulophyric cpx, vesicles (chl, qtz), acc. op
qtz, zoned plag, Kfsp, bt, green amph, acc. op, tit
altered, plag, qtz, de-vitrified glass, vesicles (chl+qtz), sec. ser+carb
plag, bt, cpx, sec. chl, acc. ap, op
plag, qtz, mafics chlorit., sec. ser
cpx, ol in fine dark matrix
cpx, ol in fine dark matrix (plag laths)
plag, cpx, de-vitrified shards
cpx, amph, plag in fine matrix (plag laths)
plag, cpx, op in very fine GM
plag, cumulophyric cpx+bt in medium GM (pl)
cpx, op replace mafics, fine GM (plag laths)
cpx, ol in fine GM
plag, green amph, bt, op
plag, kfsp, qtz, bt+green amph (repl. by chl), acc. op, ap
plag, chl, amph, op
plag, cpx, chl, op
equ.
equ.
por.
por.
por.
por.
por.
por.
por.
equ.
por.
por.
equ.
por.
equ.
por.
por.
por.
por.
por.
por.
equ.
por.
por.
por.
equ.
equ.
equ.
plag, cpx repl. by chl+op
plag, kfsp+qtz granophyre, bt (chl.), op, ap
equ.
equ.
109
Appendix 2
109
Sample#
Region1
Longitude
Latitude
Type2
Lithology
Mineralogy3
Texture4
DG072
DG073
DG074
DG075
DG076
DG077
DG078
DG079
DG080
DG081
DG082
DG083
DG084
DG085
DG086
DG087
DG088
DG089
DG090
DG091
DG092
DG093
DG094
DG095
DG096
DG097
DG098
DG099
DG100
DG101
DG102
DG103
DG104
DG105
DG106
DG107
DG108
DG109
SA
SA
SA
SA
SA
SA
SA
SA
SA
SA
SA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
46.05543
45.92358
45.92358
45.92358
45.92713
45.92713
45.96798
45.96918
46.03785
46.15100
46.15680
46.27125
46.27390
46.29065
46.30185
46.26832
46.27103
46.26378
46.58778
46.68133
46.69060
46.70152
46.71088
46.71413
46.71367
46.75410
46.81893
46.74418
46.64710
46.58858
46.57568
46.47505
46.88542
46.84370
46.84370
46.82887
46.81280
46.87327
22.46230
22.69925
22.69925
22.69925
22.74107
22.74107
22.69712
22.69568
22.79517
22.62997
22.47152
22.62265
22.62520
22.62408
22.62165
22.66240
22.66495
22.66232
22.56492
22.54875
22.58228
22.63315
22.62062
22.59040
22.58512
22.55622
22.58397
23.37145
23.19155
23.45845
23.45433
23.46813
22.87755
22.85953
22.85953
22.87615
22.89360
22.87838
IAG
sed
v
v
i
v
v
v
si
IAG
oph
v
i
i
i
i
bsm
i
i
v
v
v
v
sed
sed
v
v
i
si
i
si
si
i
v
i
v
si
si
granite
sediment
volcanic clast
tuffite
granodiorite
andesite
andesite
andesite
trachyandesite
andesite
gabbro
sediment
granite
granodiorite-granite
diorite
microdiorite
basement
granite
granodiorite
rhyolite
dacite
rhyolite
rhyolite
sediment
sediment
rhyolite
andesite
granodiorite
por. granodiorite
diorite-granodiorite
por. granodiorite
por. granodiorite
granodiorite
dacite
diorite
dacite
granodiorite
por. granite
qtz+kfsp (granophyre), plag, bt, op, ap
breccia, qtz-rich clasts in shale
equ.
cpx, mafics partly repl. by carb, chl, in fine GM (plag laths)
plag, cpx, amph (opaque rims) in fine matrix of felty plag
plag, bt (inclusions of cpx), amph?, cpx in very fine GM, sec. chl
plag (sericit.), qtz, mafics repl. by op., chl., ser.
plag, mafics repl. by sec. chl., calc.
plag, amph, chl, qtz, ap, op
fine grained, tuffitic
por.
por.
por.
por.
por.
equ.
zoned plag, bt (partly repl. by chl), ksp, qtz, green amph, acc. op
plag, bt, opx, green amph, chlorit., acc. ap, op
altered, fine grained, plag, op
equ.
equ.
qtz, plag, ksp, bt, mica in medium matrix
plag, bt, green amph, qtz, ksp, op, sec. chl, carb, ep
equ.
equ.
qtz, plag, ksp (ser.,carb.), mafics repl. by chl+op, ser. fiamme in glassy GM
fiamme filled by ep+qtz, embayed qtz, ksp, plag, qtz-rich clasts
flattened pumice clasts, embayed qtz, ksp, de-vitrified shards, rare op, sec. chl, ep
volcaniclastic sediment
quartzite and red limestone clasts in fine grey matrix (contact Gosau-Ignimbrite)
de-vitrified shards, plag, ksp (sec. ser), embayed qtz, acc.op
plag, qtz, ksp, green-brown amph (op. rims), bt, vesicles (chl), sec. carb
por.
eut.
eut.
qtz, plag, green amph, bt repl. by chl, acc. op,ap
plag, qtz, ksp, bt, green amph in fine GM (plag, qtz, bt), acc. op, ap
plag, qtz, bt (poikilitic) in fine matrix (plag dom.)
equ.
por.
por.
plag fragments, opx, green amph in glassy GM
medium plag, poikilitic cpx, bt, interstitial chl, sec. ser
plag fragments, ksp, qtz, green amph, bt, sec. chl+ep in fissures, acc. ap, op
plag, embayed qtz, bt replag. by chl+op, acc. ap,op
plag, ksp, qtz, bt+amph: replag. by chl, qtz, acc. op,ap
por.
equ.
por.
por.
por.
110
eut.
por.
Appendix 2
110
Sample#
Sample#
Region1
Longitude
Latitude
Type2
Lithology
Mineralogy3
Texture4
DG110
DG111
DG112
DG113
DG114
DG115
DG116
DG117
DG118
DG119
DG120
DG121
DG122
NA
NA
NA
NA
NA
SA
SA
SA
SA
SA
SA
SA
SA
46.91063
47.03518
47.20660
46.45937
46.34180
46.09067
46.15900
46.14723
46.12723
46.13393
46.17658
46.22072
46.30295
22.86430
22.63967
23.11995
22.72925
23.00620
23.25670
22.70505
22.69970
22.69532
22.69842
23.42310
23.46440
23.51783
si
v
v
i
v
v
si
si
i
nv
si
si
si
por. granodiorite
rhyolite
rhyolite
syeno-diorite
trachydacite
tuff
por. syeno-diorite
por. granodiorite
diorite
andesite
por. diorite
por. granodiorite
por. diorite
plag, embayed qtz, bt, green amph (rims of chl+op), acc. ap, zirc, op
por.
plag, ksp+qtz micrographic, bt, green amph, remnant cpx, op, sec. chl
equ.
glass shards, embayed qtz, plag+ksp fragments, qtz-rich lithoclasts
nw
green amph, plag, ksp, qtz, bt in fine matrix, sec. chl, ser, acc. op, ap
plag (ser.), bt+amph replag. by chl, amph preserved in plag, acc. op
por.
por.
por.
1
111
Appendix 2
111
CBR Central Banat region, EBR East Banat region, NA North Apuseni Mountains, PR: Poiana Rusca, SA South Apuseni Mountains, SWBR South-west Banat
region, WBR West Banat region,
2
v volcanic, i intrusive, si shallow intrusive/subvolcanic, bas basanite, nv Neogene volcanic, sed sedimentary rock, bsm basement, IAG island arc granitoid, oph
ophiolite
3
amph amphibole, ap apatite, bt biotite, calc calcite, carb carbonate, chl chlorite/chloritised, cpx clinopyroxene, ksp alkali-feldspar, op opaque phase, opx
orthopyroxene, plag plagioclase, qtz quartz, ser sericite/sericitised, acc accessory, GM ground mass, repl replaced, sec secondary,
4
equ equigranular, eut eutaxitic, fol foliated, nw non welded, poik poikilitic, por porphyritic,
Appendix 3: Major (wt%) and trace (ppm) element composition of the studied samples.
112
Ni
Ga
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
DG002
d
Banat
EBR
Cretaceous
62.78
0.51
16.15
5.19
0.12
2.13
4.73
3.53
2.74
0.19
1.81
99.88
12.1
117.6
19.5
10.5
14.4
67.5
595.9
15.3
112.1
6.6
2.1
608.5
21.3
44.0
5.0
19.2
3.9
1.0
2.7
0.5
2.9
0.6
2.0
0.3
1.9
0.3
3.4
0.6
9.6
7.4
2.8
DG003
d
Banat
EBR
Cretaceous
59.76
0.56
16.88
5.40
0.11
2.04
5.19
3.85
2.16
0.20
3.63
99.79
11.7
127.6
15.0
9.1
14.2
62.5
672.0
16.8
116.6
6.2
5.3
564.3
20.7
41.0
4.8
19.4
4.4
1.1
3.1
0.5
3.2
0.6
2.1
0.3
1.8
0.3
3.1
0.5
6.4
5.9
1.9
DG006
i
Banat
EBR
Cretaceous
63.04
0.49
17.04
5.13
0.06
1.55
5.03
3.60
2.24
0.19
1.42
99.79
7.6
89.3
10.5
7.9
15.8
61.6
675.4
16.1
126.0
7.3
2.3
621.4
22.1
46.4
5.2
20.7
4.3
1.1
2.7
0.6
3.1
0.6
1.8
0.2
1.3
0.4
3.5
0.6
6.2
5.9
2.1
DG007
si
Banat
CBR
Cretaceous
63.73
0.40
17.04
4.09
0.10
1.55
4.41
4.24
2.01
0.15
2.17
99.90
8.3
95.4
15.3
7.3
14.3
49.6
653.9
13.1
88.6
6.4
0.9
530.3
17.9
33.7
3.7
14.5
3.0
0.9
1.8
0.3
2.6
0.6
1.3
0.2
1.5
0.3
2.8
0.6
6.4
5.2
1.6
DG008
si
Banat
CBR
Cretaceous
65.81
0.32
17.82
3.22
0.01
0.95
2.75
4.39
2.36
0.12
2.50
100.27
6.9
69.2
15.0
12.7
11.6
67.4
762.9
10.3
94.8
4.6
1.5
704.3
18.6
35.0
3.8
14.2
3.5
0.8
1.4
0.3
2.2
0.3
1.2
0.1
1.1
0.3
2.9
0.2
2.9
4.3
1.3
DG009
si
Banat
CBR
Cretaceous
64.29
0.34
18.30
3.61
0.08
1.27
3.97
4.26
2.45
0.15
1.51
100.21
6.2
75.2
12.8
5.7
14.8
64.3
806.5
11.5
101.7
5.4
4.2
655.6
16.3
32.1
3.3
14.1
2.7
0.8
2.0
0.3
1.9
0.5
1.5
0.2
0.9
0.2
3.1
0.4
3.7
4.4
1.3
DG010
si
Banat
CBR
Cretaceous
60.42
0.53
17.42
5.90
0.13
2.31
6.12
3.55
2.31
0.22
1.03
99.96
12.7
140.0
22.1
11.8
15.2
62.0
774.1
15.8
95.1
4.8
0.9
516.4
16.9
33.7
4.1
15.6
3.5
1.0
2.9
0.4
2.6
0.6
1.8
0.2
1.6
0.2
2.7
0.4
3.4
4.9
1.5
112
DG011
i
Banat
WBR
Cretaceous
64.62
0.56
15.76
4.70
0.08
2.59
3.85
3.60
2.96
0.17
1.11
100.00
11.3
98.6
45.1
13.7
DG013
i
Banat
WBR
Cretaceous
63.74
0.58
15.69
4.68
0.07
2.25
3.96
3.64
3.25
0.18
2.06
100.10
10.6
92.7
41.4
12.0
DG014
i
Banat
WBR
Cretaceous
59.97
0.73
17.01
5.98
0.11
2.38
5.02
3.81
4.19
0.27
0.56
100.03
13.0
130.4
7.5
15.3
DG015
d
Banat
WBR
Cretaceous
55.84
0.86
17.85
8.91
0.06
3.15
2.56
3.49
1.87
0.23
4.79
99.59
15.1
187.7
108.1
20.7
DG016
i
Banat
WBR
Cretaceous
63.66
0.56
16.25
4.54
0.09
1.84
4.03
3.73
4.29
0.20
0.63
99.81
9.4
89.3
29.0
12.2
DG017
d
Banat
WBR
Cretaceous
76.40
0.13
13.04
0.67
0.01
0.00
0.69
3.65
4.37
0.03
0.74
99.72
14.1
109.8
477.9
15.8
129.1
11.2
3.0
528.4
29.5
54.6
5.9
21.8
4.7
1.0
3.3
0.5
3.0
0.6
1.6
0.3
1.5
0.2
3.6
1.0
10.5
14.3
3.1
9.0
14.6
112.9
470.3
15.4
135.4
11.9
2.4
556.4
28.2
56.2
6.2
22.8
4.2
0.9
3.5
0.5
2.6
0.6
1.6
0.3
1.5
0.2
3.9
1.2
11.9
17.1
2.0
8.8
16.6
169.0
657.1
25.5
103.7
12.8
8.2
447.8
34.3
74.8
9.1
34.8
7.5
1.4
5.9
0.9
5.1
0.9
2.8
0.4
2.6
0.5
2.8
1.1
21.2
16.9
5.9
37.2
19.5
93.2
498.3
21.0
111.0
7.2
4.1
1057.1
24.2
51.5
6.0
25.2
5.1
1.1
4.4
0.6
3.9
0.7
1.9
0.3
1.8
0.2
3.2
0.5
11.1
6.6
4.0
14.4
139.9
568.1
32.5
228.5
15.3
4.3
521.5
31.7
75.2
9.4
35.9
7.2
1.3
6.1
0.8
5.1
1.2
3.2
0.5
3.5
0.4
6.4
1.7
17.5
19.9
6.1
10.9
124.5
171.3
8.2
73.8
8.4
1.4
602.4
27.1
44.5
4.7
16.6
3.3
0.7
8.2
5.8
0.4
1.6
0.3
1.8
3.4
1.2
9.3
19.7
5.0
Appendix 3
Sample
type
segment
locality
Age
SiO2
TiO2
Al2O3
Fe2O3t
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Sum
Sc
V
Cr
Co
DG018
i
Banat
WBR
Cretaceous
52.45
0.97
16.09
9.61
0.23
6.36
5.87
3.12
2.58
0.20
2.38
99.85
23.9
182.2
349.9
29.1
Ni
Ga
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
63.1
17.6
142.5
579.8
19.3
83.2
8.4
5.4
469.0
25.3
54.1
6.2
26.9
5.6
1.3
4.7
0.6
3.8
0.7
2.1
0.3
1.7
0.2
2.5
0.5
12.0
4.6
2.1
DG019
i
Banat
WBR
Cretaceous
70.02
0.28
15.86
1.70
0.03
0.60
2.51
3.76
4.37
0.08
0.67
99.86
3.3
30.1
7.3
DG020
i
Banat
WBR
Cretaceous
64.22
0.49
16.68
4.05
0.08
1.74
3.94
4.10
3.80
0.17
0.75
100.04
9.7
77.8
26.9
10.3
DG021
i
Banat
WBR
Cretaceous
64.60
0.54
16.53
4.32
0.08
1.64
3.73
3.84
4.02
0.19
0.70
100.18
8.9
80.6
32.8
9.7
DG024
i
Banat
WBR
Cretaceous
51.31
1.00
16.59
8.77
0.20
6.42
7.49
2.94
3.21
0.52
1.68
100.13
20.3
227.1
180.6
27.9
DG026
i
Banat
WBR
Cretaceous
67.61
0.36
15.55
2.97
0.06
1.55
3.53
3.75
3.00
0.12
1.23
99.72
7.4
65.8
51.5
8.9
DG027
d
Banat
WBR
Cretaceous
46.38
0.84
13.86
8.75
0.16
12.39
8.59
1.65
0.88
0.20
6.57
100.26
26.2
171.2
640.3
42.7
DG028
i
Banat
WBR
Cretaceous
57.26
0.88
15.58
7.23
0.17
5.05
6.20
3.72
1.91
0.21
1.91
100.11
22.5
165.1
127.1
21.5
DG029
d
Banat
WBR
Cretaceous
48.85
0.94
14.34
8.79
0.19
9.92
9.34
1.80
0.94
0.26
4.71
100.08
29.7
188.5
422.9
33.8
DG030
si
Banat
WBR
Cretaceous
61.23
0.69
16.35
5.04
0.08
2.89
5.21
4.11
3.10
0.26
0.95
99.92
11.6
122.6
13.1
11.9
DG031
si
Banat
WBR
Cretaceous
49.87
1.28
16.62
8.03
0.21
6.56
11.50
3.28
0.86
0.25
1.74
100.20
29.8
269.4
68.6
24.0
DG032
i
Banat
WBR
Cretaceous
63.03
0.60
16.12
5.17
0.09
2.63
4.62
3.67
2.82
0.20
0.78
99.72
12.6
112.2
49.8
12.1
DG033
d
Banat
WBR
Cretaceous
56.83
0.92
17.04
7.58
0.13
3.34
6.35
3.69
3.48
0.35
0.45
100.17
16.1
189.1
44.6
19.4
14.6
153.9
528.4
15.8
127.6
8.3
4.2
534.0
35.7
63.5
6.9
25.4
4.6
1.1
4.0
0.5
2.3
0.5
1.6
0.4
1.8
0.3
4.3
0.8
16.7
20.5
2.8
15.3
131.8
504.9
12.4
145.0
10.0
6.2
503.7
25.2
51.9
5.4
20.8
3.9
0.9
2.8
0.4
2.4
0.5
1.5
0.2
1.3
0.2
4.1
0.8
14.8
11.1
2.6
12.0
16.2
151.8
563.0
16.9
233.6
11.3
5.3
578.8
36.2
70.1
7.3
28.1
4.7
1.3
4.1
0.5
3.4
0.7
2.2
0.3
2.3
0.3
6.0
0.7
15.7
18.7
3.7
63.5
17.1
104.9
1006.4
21.4
168.8
7.6
6.4
839.0
38.4
85.9
10.3
42.8
7.7
2.2
6.8
0.7
4.5
0.8
2.5
0.3
2.0
0.2
3.8
0.5
6.6
12.5
3.7
13.3
14.2
99.4
607.6
9.5
82.6
6.6
1.8
598.5
25.9
43.7
4.2
15.2
2.0
1.1
1.9
0.3
2.2
0.3
1.5
0.1
1.9
0.1
3.1
0.4
12.4
9.5
1.8
251.1
14.1
40.4
240.4
16.1
78.8
10.4
0.5
198.9
13.0
29.3
3.6
15.3
3.4
1.1
3.6
0.5
3.0
0.6
1.5
0.2
1.6
0.3
1.9
0.6
2.3
2.4
0.8
34.3
15.8
73.9
607.4
24.6
111.7
16.6
1.4
393.8
39.1
84.8
9.2
36.7
6.9
2.0
5.6
0.8
4.2
0.8
2.5
0.3
2.5
0.4
3.6
1.2
10.2
10.3
4.7
143.0
13.2
53.4
358.1
16.9
100.5
10.8
2.3
237.6
21.8
44.4
4.9
20.3
4.0
1.1
3.8
0.5
3.2
0.6
1.5
0.2
1.3
0.3
2.5
0.6
2.6
4.5
1.2
10.8
15.0
83.4
600.6
17.5
144.7
11.6
2.7
735.1
34.7
69.4
7.7
28.5
5.9
1.3
4.1
0.6
3.4
0.6
1.8
0.2
1.5
0.3
3.5
0.9
19.9
15.8
3.3
40.6
17.6
27.3
548.6
22.1
91.1
7.8
2.3
246.8
23.9
52.9
6.8
29.8
6.3
1.7
4.7
0.8
4.6
0.9
2.4
0.3
1.9
0.3
2.4
0.5
25.1
4.7
1.2
10.2
15.1
106.9
559.7
16.1
132.9
11.1
6.7
634.5
30.4
59.1
6.2
24.2
3.5
1.1
3.7
0.5
2.8
0.6
1.8
0.2
2.1
0.3
3.8
1.0
12.4
14.7
3.8
15.9
17.4
120.1
869.0
23.1
193.2
10.2
5.0
535.9
34.8
77.8
9.3
36.8
7.0
1.8
6.6
0.8
4.3
0.9
2.5
0.3
2.2
0.3
5.6
0.7
20.1
13.7
5.1
113
Appendix 3
113
Sample
type
segment
locality
Age
SiO2
TiO2
Al2O3
Fe2O3t
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Sum
Sc
V
Cr
Co
DG034
i
Banat
WBR
Cretaceous
51.13
1.12
17.52
9.96
0.15
4.57
8.21
3.58
1.96
0.45
1.31
99.95
21.1
255.3
45.1
25.7
DG035
i
Banat
WBR
Cretaceous
63.57
0.59
16.22
4.78
0.09
1.91
4.28
3.82
3.55
0.20
0.89
99.89
9.1
95.9
46.8
11.0
DG036
i
Banat
WBR
Cretaceous
55.72
0.83
17.55
7.22
0.15
3.68
6.16
4.00
2.85
0.26
1.73
100.16
18.3
205.3
22.2
18.2
DG037
i
Banat
SWBR
Cretaceous
49.77
0.83
18.02
9.02
0.16
6.40
11.20
3.17
0.50
0.27
0.83
100.16
23.9
255.9
39.2
28.3
DG038
i
Banat
SWBR
Cretaceous
63.57
0.53
16.56
4.65
0.07
2.46
4.99
3.90
2.48
0.16
0.90
100.27
10.2
81.9
16.8
15.5
DG039
si
Banat
SWBR
Cretaceous
55.25
0.96
17.90
7.25
0.13
4.18
7.00
4.00
1.93
0.30
0.89
99.78
17.2
161.5
55.3
20.6
DG040
i
Banat
SWBR
Cretaceous
65.51
0.50
16.26
3.94
0.08
2.03
4.32
3.98
2.59
0.13
0.70
100.06
8.2
73.2
20.0
11.9
DG041
i
Banat
SWBR
Cretaceous
64.02
0.55
16.21
4.87
0.06
2.67
4.51
3.63
1.66
0.18
1.88
100.25
11.4
95.3
25.7
13.4
DG042
d
Banat
CBR
Cretaceous
62.10
0.45
17.85
5.29
0.11
2.45
5.50
3.34
1.95
0.12
1.25
100.41
12.1
130.9
8.1
11.9
DG043
si
Banat
PR
Cretaceous
61.91
0.65
15.80
5.11
0.09
2.89
5.37
3.71
2.66
0.17
1.50
99.86
11.7
109.0
57.8
13.4
DG044
i
Banat
PR
Cretaceous
64.42
0.61
16.30
4.52
0.11
2.23
3.88
3.69
2.68
0.14
1.16
99.75
7.3
90.6
28.5
9.9
DG045
si
Banat
PR
Cretaceous
69.25
0.33
14.96
3.15
0.08
1.61
1.62
3.26
3.87
0.13
1.66
99.90
3.7
52.3
23.1
7.5
DG046
i
Banat
PR
Cretaceous
64.67
0.58
15.91
4.67
0.07
2.33
3.91
3.70
3.07
0.14
0.86
99.90
6.1
89.4
27.2
11.5
Ni
Ga
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
20.2
18.6
58.4
1062.6
24.2
92.6
5.4
2.2
609.4
30.4
66.1
7.9
35.6
7.4
2.0
6.7
0.7
4.8
0.9
2.8
0.3
2.1
0.3
2.6
0.3
11.3
6.6
1.7
15.7
16.1
129.1
566.9
19.0
215.6
11.0
4.6
561.1
34.4
68.5
7.6
28.7
5.6
1.3
4.3
0.5
3.8
0.6
2.1
0.3
2.3
0.3
5.5
0.8
15.7
16.3
5.3
17.1
123.0
652.2
22.0
154.7
11.3
5.6
562.0
29.0
63.8
7.5
29.2
6.4
1.5
5.0
0.7
3.6
0.8
2.0
0.3
2.0
0.3
3.6
0.5
18.8
8.3
3.9
27.7
17.7
13.2
1475.5
13.3
24.8
2.4
1.3
264.8
15.4
34.9
4.5
19.2
4.4
1.4
3.6
0.5
2.9
0.4
1.3
0.1
1.2
0.2
0.9
0.2
7.3
1.3
0.4
16.3
13.7
91.4
577.9
14.8
120.1
8.6
5.3
529.9
25.5
48.3
4.9
18.5
3.5
0.8
3.1
0.4
2.5
0.5
1.3
0.2
1.4
0.2
3.2
0.8
10.4
9.4
3.8
24.2
17.0
59.4
788.1
19.5
147.5
22.1
3.5
566.2
28.4
57.3
6.5
25.8
4.7
1.3
4.6
0.6
3.6
0.8
2.0
0.2
1.9
0.3
3.6
1.2
9.1
5.5
2.0
8.3
14.7
81.7
483.6
10.4
100.1
10.6
2.3
508.3
21.5
41.5
4.3
15.6
2.6
0.8
2.4
0.3
1.9
0.4
0.6
0.2
1.4
0.2
2.5
0.9
11.7
9.1
2.4
9.5
14.4
46.0
449.6
14.0
109.9
11.1
1.4
375.8
21.2
42.1
4.5
16.9
3.2
0.9
3.0
0.5
2.5
0.5
1.2
0.2
1.4
0.3
3.0
0.8
5.3
8.7
2.1
4.3
14.1
55.3
471.7
15.9
78.2
5.8
4.1
647.0
15.4
31.2
3.4
13.5
2.8
0.8
2.5
0.4
2.9
0.6
1.6
0.2
1.8
0.3
2.2
0.5
8.7
5.9
2.1
18.5
15.1
78.4
526.9
14.3
123.9
8.3
2.1
612.4
25.6
47.7
5.1
20.1
3.6
1.1
3.3
0.4
2.7
0.5
1.7
0.2
1.3
0.2
3.7
0.6
13.6
10.2
2.2
14.6
65.9
579.5
12.5
116.5
11.0
1.9
612.5
38.6
67.9
6.8
24.9
3.7
1.1
4.0
0.5
2.5
0.4
1.5
0.1
0.9
0.2
3.0
0.9
182.8
10.8
2.1
11.4
117.9
262.2
7.9
115.7
8.8
2.6
518.3
23.2
44.3
4.0
14.7
2.5
0.6
1.5
0.3
1.8
0.3
1.0
14.6
98.2
513.1
13.2
111.5
12.1
2.9
640.9
28.5
52.6
5.7
22.5
4.1
1.0
3.7
0.5
2.8
0.3
1.9
0.1
1.2
0.2
3.3
0.9
11.0
11.3
2.6
114
0.2
3.7
0.7
11.9
10.4
2.5
Appendix 3
114
Sample
type
segment
locality
Age
SiO2
TiO2
Al2O3
Fe2O3t
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Sum
Sc
V
Cr
Co
DG047
si
Banat
PR
Cretaceous
66.88
0.40
14.75
2.95
0.05
1.31
2.85
4.45
2.56
0.11
3.54
99.84
3.1
57.2
28.2
7.0
DG048
d
Banat
PR
Cretaceous
53.25
0.98
15.92
8.13
0.14
4.88
8.18
2.35
2.57
0.38
3.22
99.99
14.7
182.9
110.5
22.4
DG049
i
Banat
PR
Cretaceous
64.84
0.53
16.05
4.31
0.07
2.24
4.21
3.59
3.09
0.14
0.95
100.02
5.3
88.9
37.5
11.5
DG051
i
Banat
PR
Cretaceous
55.54
0.84
16.72
8.68
0.17
4.25
7.67
3.24
1.95
0.22
0.54
99.81
18.6
205.9
32.7
20.3
DG052
d
Banat
PR
Cretaceous
63.49
0.39
14.76
3.79
0.07
1.74
4.10
3.33
2.17
0.13
6.08
100.04
6.7
81.7
26.5
10.2
DG055
v
Banat
PR
Cretaceous
56.85
0.89
18.32
5.26
0.15
3.27
8.44
3.23
2.22
0.25
0.84
99.72
20.1
176.8
65.5
13.4
DG056
v
Banat
PR
Cretaceous
61.41
0.59
15.61
5.13
0.07
2.64
4.91
2.85
3.58
0.23
2.82
99.84
9.9
113.6
62.1
14.4
DG057
v
Banat
PR
Cretaceous
63.89
0.74
18.33
3.72
0.06
0.39
5.34
3.65
2.98
0.27
0.77
100.15
10.4
126.2
23.5
8.5
DG058
i
Apuseni
SA
Cretaceous
56.53
0.76
16.05
7.39
0.13
4.86
7.46
3.34
2.59
0.28
0.55
99.96
19.2
174.5
133.3
21.3
DG059
v
Apuseni
SA
Cretaceous
57.79
0.75
17.18
6.31
0.12
4.36
7.00
3.90
1.82
0.29
0.50
100.03
14.8
124.5
120.9
16.8
DG076
i
Apuseni
SA
Cretaceous
62.85
0.65
16.70
3.16
0.04
0.62
5.41
3.87
3.86
0.27
2.81
100.24
9.2
69.6
64.3
22.6
DG077
v
Apuseni
SA
Cretaceous
59.79
0.66
16.29
5.48
0.08
4.18
6.38
3.75
2.01
0.22
1.15
99.98
14.3
117.4
146.5
16.7
DG078
v
Apuseni
SA
Cretaceous
60.55
0.67
17.18
5.18
0.10
2.50
5.41
3.68
2.16
0.19
2.12
99.75
12.8
94.6
12.4
13.3
Ni
Ga
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
12.9
113.3
325.8
12.5
121.4
11.0
6.1
418.1
26.4
48.1
4.8
18.3
2.4
0.9
3.3
0.5
2.2
0.4
1.4
0.1
0.7
0.2
3.4
0.8
7.2
9.4
2.3
24.4
15.8
67.5
761.5
23.7
208.4
10.4
3.0
654.4
37.1
80.1
9.5
39.2
6.5
1.6
6.5
0.8
4.4
0.8
2.4
0.2
2.0
0.3
5.0
0.5
19.0
12.7
3.4
12.9
88.2
528.3
13.8
104.4
10.6
1.3
583.3
31.0
57.9
5.9
23.4
3.7
0.9
2.9
0.5
2.5
0.4
1.5
0.1
0.9
0.2
2.5
0.7
13.0
11.6
1.5
15.9
52.6
635.2
23.8
96.2
4.8
2.2
407.0
18.3
39.5
4.8
21.4
4.0
1.3
4.3
0.7
4.2
0.8
2.3
0.2
2.7
0.4
2.5
0.2
12.3
5.1
1.3
10.9
12.8
66.3
318.0
10.9
101.1
5.4
4.6
70.0
25.0
49.2
5.2
19.5
3.3
0.9
2.8
0.3
2.1
0.5
1.2
0.1
1.3
0.2
2.9
0.5
12.3
10.6
2.1
18.1
16.4
55.0
737.8
19.6
121.9
13.3
2.2
702.6
24.8
51.0
6.0
25.1
5.4
1.5
4.5
0.6
4.3
0.7
2.3
0.3
1.8
0.3
3.4
0.9
12.9
6.8
2.5
22.5
15.6
169.1
527.2
16.4
147.5
9.5
5.2
673.5
31.0
61.2
6.9
27.7
5.1
1.2
3.9
0.5
3.5
0.6
1.7
0.2
1.6
0.2
3.7
0.7
19.6
9.7
3.6
16.1
87.0
470.1
30.0
169.3
8.7
2.8
698.0
32.0
58.0
7.2
30.3
5.9
1.5
6.2
0.8
4.7
1.1
3.1
0.4
3.0
0.4
4.2
0.5
13.5
8.5
2.7
27.8
15.5
90.3
583.6
22.3
136.9
11.6
4.9
528.7
31.0
65.2
7.4
31.6
6.2
1.7
5.5
0.7
4.5
0.8
2.3
0.3
2.1
0.4
3.5
0.9
14.7
11.5
3.2
45.3
14.2
45.0
580.4
19.5
150.0
12.7
1.0
437.3
29.6
59.3
6.9
26.7
5.7
1.4
4.7
0.6
3.2
0.7
1.8
0.3
2.0
0.3
3.8
0.9
9.9
8.1
2.5
68.6
16.0
146.2
602.3
15.7
247.7
12.3
3.9
597.0
39.6
77.6
8.4
32.6
5.7
1.4
4.8
0.6
2.8
0.5
1.6
0.2
1.6
0.3
6.1
0.9
20.8
22.1
5.5
49.8
13.6
54.7
518.7
16.6
131.8
12.1
1.2
452.2
27.9
54.4
6.0
23.9
4.5
1.2
3.7
0.6
2.9
0.5
1.7
0.3
1.6
0.3
3.3
0.8
10.8
8.4
2.5
15.8
181.6
453.4
18.8
155.7
8.2
3.9
495.3
22.7
46.6
5.2
20.5
4.3
1.0
3.7
0.5
3.5
0.7
2.1
0.3
1.9
0.3
4.0
0.6
15.5
8.4
2.3
115
Appendix 3
115
Sample
type
segment
locality
Age
SiO2
TiO2
Al2O3
Fe2O3t
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Sum
Sc
V
Cr
Co
DG079
v
Apuseni
SA
Cretaceous
59.33
0.62
16.92
6.54
0.09
2.75
5.35
3.12
1.93
0.23
2.79
99.67
14.1
107.3
DG080
si
Apuseni
SA
Cretaceous
60.85
0.64
17.65
5.20
0.09
3.31
1.04
2.46
4.22
0.13
4.31
99.91
14.3
141.6
DG084
i
Apuseni
NA
Cretaceous
68.74
0.40
14.57
3.26
0.05
1.26
2.55
3.18
4.05
0.11
1.41
99.58
8.4
62.1
DG085
i
Apuseni
NA
Cretaceous
67.95
0.45
14.66
3.58
0.04
1.60
2.67
3.24
3.82
0.12
1.82
99.96
9.8
71.8
DG086
i
Apuseni
NA
Cretaceous
57.41
0.85
16.57
7.06
0.13
4.09
7.09
3.16
2.67
0.25
0.70
99.97
24.7
187.2
DG090
i
Apuseni
NA
Cretaceous
67.54
0.53
15.46
3.69
0.08
1.67
2.96
4.11
2.95
0.12
1.04
100.14
9.7
67.4
DG091
i
Apuseni
NA
Cretaceous
73.30
0.35
13.75
2.31
0.05
0.41
1.27
3.72
3.70
0.05
1.21
100.11
7.1
17.1
DG092
v
Apuseni
NA
Cretaceous
65.46
0.51
14.36
3.79
0.07
1.24
4.16
0.94
3.57
0.13
5.64
99.86
10.5
46.2
DG093
v
Apuseni
NA
Cretaceous
74.46
0.19
13.09
2.38
0.05
0.27
1.34
3.11
4.31
0.03
0.95
100.18
11.0
14.1
DG094
v
Apuseni
NA
Cretaceous
75.53
0.15
13.33
1.48
0.04
0.30
0.61
3.12
4.32
0.02
1.04
99.93
15.9
11.1
DG097
v
Apuseni
NA
Cretaceous
76.13
0.09
12.65
1.27
0.04
0.07
0.78
3.62
4.51
0.01
0.81
99.99
11.1
1.8
DG099
i
Apuseni
NA
Cretaceous
63.83
0.62
16.69
3.71
0.09
1.58
5.14
3.41
2.53
0.14
2.54
100.28
10.9
96.8
DG100
si
Apuseni
NA
Cretaceous
67.13
0.43
14.90
3.47
0.07
1.57
3.14
3.33
3.26
0.12
2.78
100.20
7.7
63.9
Cr
Co
39.8
14.7
35.6
16.2
27.2
13.3
11.9
30.4
24.5
23.6
8.3
3.6
20.3
6.6
4.0
5.9
4.5
51.1
8.7
9.6
14.0
141.4
47.2
40.4
126.6
13.4
4.0
728.9
33.2
68.1
8.1
31.1
6.7
1.1
7.7
1.1
7.8
1.4
4.6
0.4
4.6
0.6
4.9
1.2
21.8
15.8
4.1
35.8
15.0
85.6
381.5
17.7
160.9
9.7
3.0
585.2
25.0
45.0
4.9
18.8
3.7
1.0
3.3
0.5
3.2
0.6
1.7
0.3
1.9
0.3
4.2
0.8
12.5
10.5
3.2
15.5
136.7
250.7
19.4
148.0
9.9
6.1
509.3
30.4
54.0
5.8
20.3
3.5
0.6
3.9
0.5
3.2
0.7
1.6
0.4
2.2
0.3
4.1
1.0
18.1
15.6
3.3
Ni
Ga
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
15.4
129.8
550.9
19.8
195.2
8.7
7.8
641.5
30.0
56.6
6.3
24.6
4.7
1.2
4.3
0.6
3.2
0.7
1.9
0.3
1.6
0.4
4.8
0.6
19.3
11.9
3.1
14.6
100.0
94.7
17.6
104.3
1.7
0.5
483.3
9.9
21.9
2.7
11.1
2.9
0.9
2.7
0.4
2.7
0.6
2.1
0.3
2.1
0.3
2.8
0.2
189.7
5.5
2.2
14.7
170.4
285.0
18.0
160.5
9.1
8.1
714.5
32.0
60.2
6.6
23.4
3.7
0.9
4.1
0.5
2.9
0.5
2.2
0.3
1.8
0.4
5.0
0.9
21.0
15.5
2.6
14.6
152.4
266.1
21.8
178.2
8.5
8.1
617.8
36.0
65.0
7.1
25.0
5.1
1.2
4.7
0.7
3.1
0.8
2.4
0.4
2.5
0.5
4.6
0.7
16.3
19.5
4.2
15.6
82.1
490.8
22.3
125.8
7.4
4.7
578.7
24.6
51.8
6.1
24.0
5.7
1.3
5.0
0.7
4.2
0.8
2.1
0.3
2.1
0.3
3.3
0.5
22.6
7.4
1.7
16.2
115.3
229.3
23.4
178.4
11.7
5.0
561.8
26.8
53.5
5.6
20.9
4.5
0.9
3.6
0.6
4.2
0.8
2.5
0.4
2.5
0.4
4.5
0.8
15.0
12.2
3.6
14.7
119.0
112.6
25.0
178.0
14.8
3.1
529.9
32.1
65.0
7.5
27.6
5.2
1.0
5.1
0.8
4.5
0.8
2.6
0.6
2.5
0.4
4.8
1.2
12.0
11.7
2.8
116
14.4
154.8
139.1
35.1
256.5
12.5
8.5
532.1
37.9
72.3
8.4
30.9
6.6
1.5
6.2
1.0
5.9
1.2
3.8
0.5
3.3
0.5
6.5
0.9
24.7
12.7
3.2
14.7
154.4
87.2
39.3
224.1
16.1
3.1
872.0
51.6
98.9
11.9
44.1
9.5
1.2
9.9
1.1
7.7
1.7
4.3
0.8
4.1
0.6
6.4
1.1
20.9
15.3
4.4
16.0
203.9
39.2
55.5
145.8
19.9
11.0
167.9
18.6
44.1
5.9
24.1
7.7
0.5
8.0
1.4
9.1
1.8
6.2
0.8
5.1
0.8
5.5
1.4
28.1
18.7
6.6
Appendix 3
116
Sample
type
segment
locality
Age
SiO2
TiO2
Al2O3
Fe2O3t
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Sum
Sc
V
DG101
i
Apuseni
NA
Cretaceous
62.91
0.59
16.04
4.43
0.09
2.78
4.68
3.37
2.82
0.17
1.97
99.85
11.7
90.8
DG102
i
Apuseni
NA
Cretaceous
66.86
0.53
15.64
3.83
0.05
1.74
3.77
3.06
3.25
0.14
1.34
100.21
8.7
73.5
DG103
si
Apuseni
NA
Cretaceous
67.99
0.44
15.66
3.19
0.05
0.80
3.83
3.03
2.94
0.14
2.09
100.17
7.1
47.4
DG104
i
Apuseni
NA
Cretaceous
66.74
0.56
15.76
4.01
0.06
1.59
3.78
3.99
2.69
0.12
0.65
99.95
9.9
65.9
DG105
v
Apuseni
NA
Cretaceous
64.67
0.51
15.87
4.54
0.09
1.24
3.66
4.22
2.87
0.13
2.16
99.95
12.9
40.8
DG106
i
Apuseni
NA
Cretaceous
53.11
0.70
18.45
7.69
0.14
4.67
9.20
3.06
1.78
0.17
1.17
100.13
27.5
210.0
DG107
v
Apuseni
NA
Cretaceous
66.49
0.49
16.11
4.29
0.08
1.29
3.03
4.27
2.88
0.12
1.08
100.13
12.3
46.8
DG108
i
Apuseni
NA
Cretaceous
67.79
0.44
15.14
3.41
0.06
1.07
3.06
3.68
3.34
0.11
2.18
100.29
6.9
51.9
DG109
si
Apuseni
NA
Cretaceous
65.17
0.62
15.74
4.28
0.09
1.74
3.82
3.73
3.49
0.16
1.06
99.90
12.4
72.7
DG110
si
Apuseni
NA
Cretaceous
67.96
0.48
15.44
3.50
0.07
1.43
3.12
4.17
3.06
0.11
0.87
100.20
9.2
56.3
DG111
v
Apuseni
NA
Cretaceous
79.15
0.07
11.29
0.86
0.08
0.08
0.31
2.68
4.10
0.07
1.09
99.80
3.2
1.6
DG112
v
Apuseni
NA
Cretaceous
77.86
0.14
11.96
1.65
0.03
0.12
0.18
3.48
3.26
0.03
1.43
100.14
8.7
8.1
DG113
i
Apuseni
NA
Cretaceous
60.05
0.58
16.07
5.78
0.11
3.52
4.52
3.72
3.19
0.23
2.16
99.93
17.9
128.4
Cr
Co
41.7
10.9
8.9
7.8
29.1
10.3
24.8
11.6
25.3
25.2
9.6
8.1
28.4
9.7
14.8
8.0
3.4
4.0
83.3
14.3
Ni
Ga
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
13.0
14.7
90.3
418.2
17.7
133.4
10.6
3.1
617.9
26.5
49.7
5.6
21.0
4.0
1.0
3.4
0.5
3.3
0.6
1.8
0.3
1.8
0.3
3.9
0.8
17.9
10.9
3.8
15.6
116.7
318.7
16.8
124.4
11.8
2.6
525.0
27.5
50.1
5.3
21.7
4.9
1.1
2.9
0.6
3.2
0.6
1.9
0.1
2.6
0.4
4.1
1.1
9.3
15.3
5.2
13.3
97.4
257.8
20.5
123.0
9.7
4.0
525.2
34.9
63.0
6.9
25.7
5.4
1.2
4.6
0.6
3.9
0.8
2.1
15.5
93.0
254.4
23.1
165.7
11.7
3.0
540.1
27.9
55.0
5.9
22.7
4.9
1.1
4.5
0.7
4.1
0.7
2.6
0.4
2.2
0.4
4.2
0.9
36.3
11.5
3.8
21.4
17.1
101.1
259.3
34.5
261.1
12.8
4.9
589.4
33.1
66.3
7.4
30.3
6.5
1.4
6.3
1.0
6.2
1.2
4.0
0.5
3.3
0.5
6.5
0.9
17.5
11.4
3.5
15.2
48.3
403.9
16.4
87.2
4.1
2.3
651.3
19.1
37.5
4.5
18.9
4.1
1.1
3.4
0.6
3.0
0.6
1.6
0.3
1.6
0.3
2.3
0.3
11.7
6.2
1.9
17.1
97.1
251.7
35.9
260.4
13.5
4.6
619.3
36.0
67.9
7.7
31.4
6.2
1.8
5.5
1.1
5.8
1.3
3.6
0.6
3.8
0.8
7.4
0.7
17.7
12.6
2.5
14.9
122.3
223.5
28.0
171.7
11.5
7.3
613.9
31.4
55.8
6.5
24.0
5.1
0.7
4.9
0.7
5.3
0.9
2.8
0.3
3.3
0.5
4.5
0.6
15.6
13.9
4.4
15.6
109.3
321.5
30.2
219.7
11.0
4.6
664.1
31.6
60.3
7.0
27.3
6.2
1.5
5.8
0.9
5.5
1.1
3.3
0.4
3.2
0.4
6.1
0.8
13.1
7.9
2.4
14.3
110.8
230.0
23.8
169.3
10.7
3.7
558.0
27.7
53.5
5.8
20.4
4.3
0.9
3.9
0.6
3.9
0.8
2.3
0.3
2.4
0.4
4.4
0.9
16.1
12.2
4.4
10.0
118.7
49.1
16.2
42.6
24.4
2.2
826.9
6.9
12.6
1.6
5.0
1.2
0.4
1.7
0.4
2.3
0.5
1.7
0.2
2.1
0.3
2.0
1.6
19.8
7.3
4.2
10.4
91.7
84.1
29.3
185.9
11.6
1.0
1188.1
37.2
70.6
8.2
31.0
6.3
1.1
5.6
0.9
5.2
1.3
3.3
0.3
4.0
0.7
5.7
1.1
13.6
13.0
3.4
19.3
14.7
95.0
597.3
18.5
121.9
6.1
2.7
612.5
23.9
47.6
5.5
21.9
4.3
1.1
3.9
0.6
3.5
0.6
1.8
0.3
2.0
0.3
3.4
0.4
10.9
8.4
2.3
2.6
0.2
3.6
1.0
15.0
13.7
3.5
117
Appendix 3
117
Sample
type
segment
locality
Age
SiO2
TiO2
Al2O3
Fe2O3t
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Sum
Sc
V
DG114
v
Apuseni
NA
Cretaceous
68.82
0.25
17.09
2.57
0.08
0.66
0.28
8.30
0.64
0.11
1.33
100.13
4.1
22.3
DG115
v
Apuseni
SA
Cretaceous
75.22
0.09
11.23
0.22
0.00
0.47
2.68
1.11
2.69
0.01
5.98
99.70
9.5
Cr
Co
4.4
12.5
25.8
357.2
9.5
112.5
5.1
0.9
125.1
10.9
23.5
2.1
8.2
1.6
0.6
1.9
0.2
1.6
0.2
1.3
Ni
Ga
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
0.8
0.3
3.2
0.2
15.3
5.6
2.1
DG116
si
Apuseni
SA
Cretaceous
57.13
0.75
16.99
6.64
0.11
3.55
3.79
6.34
2.45
0.20
2.12
100.07
15.5
136.1
DG117
si
Apuseni
SA
Cretaceous
63.97
0.37
15.27
3.83
0.07
2.29
4.61
3.24
2.20
0.10
4.09
100.05
9.6
74.7
DG118
i
Apuseni
SA
Cretaceous
60.41
0.53
17.01
5.09
0.13
2.00
6.59
3.14
1.24
0.18
3.52
99.84
11.5
100.4
DG120
si
Apuseni
SA
Cretaceous
65.69
0.52
15.65
4.30
0.08
2.00
3.84
3.15
3.97
0.14
0.74
100.07
11.9
93.8
DG121
i
Apuseni
SA
Cretaceous
63.86
0.47
15.50
4.19
0.07
2.72
4.01
3.33
2.95
0.16
3.11
100.37
12.7
78.2
DG122
si
Apuseni
SA
Cretaceous
57.92
0.52
15.80
5.90
0.11
3.78
5.41
2.67
3.47
0.25
3.95
99.78
15.9
123.1
DG062
IAG
Apuseni
SA
Jurassic
70.35
0.41
14.82
3.17
0.08
0.92
2.55
3.81
2.70
0.13
1.28
100.21
10.6
55.6
DG063
oph
Apuseni
SA
Jurassic
56.36
0.66
16.89
6.30
0.15
5.36
7.75
3.71
0.68
0.19
1.81
99.87
22.2
134.4
DG064
oph
Apuseni
SA
Jurassic
48.37
0.57
18.93
6.64
0.16
7.44
14.20
1.67
0.20
0.04
1.99
100.21
36.3
189.7
DG065
IAG
Apuseni
SA
Jurassic
74.25
0.20
13.49
1.48
0.04
0.26
0.96
3.78
4.85
0.04
0.49
99.84
2.6
13.2
DG066
oph
Apuseni
SA
Jurassic
50.14
1.16
15.19
10.87
0.21
7.60
10.78
3.06
0.35
0.10
0.88
100.33
42.2
275.9
1.6
31.9
17.3
42.7
13.3
12.5
10.8
96.0
12.0
69.2
14.6
9.0
192.1
21.0
240.2
22.7
7.8
124.5
41.4
13.4
87.2
907.4
22.7
153.8
12.8
2.2
1516.9
33.1
61.9
7.5
31.2
7.0
1.2
5.3
0.9
4.9
1.0
3.3
0.3
2.8
0.4
6.3
1.1
33.4
15.1
2.3
14.0
69.1
248.5
19.4
124.1
10.4
0.2
846.7
21.9
44.0
5.1
20.2
4.1
1.1
3.9
0.6
3.5
0.7
2.1
0.3
2.0
0.4
3.2
0.6
10.0
5.6
1.7
13.9
70.1
386.2
12.9
91.7
5.3
6.8
591.6
13.9
26.4
2.9
13.1
2.9
0.7
2.6
0.3
2.4
0.5
1.7
15.9
32.0
413.4
21.4
124.4
5.7
1.0
393.4
13.3
29.5
3.5
15.2
3.4
1.0
3.8
0.5
3.8
0.7
2.9
0.4
1.9
0.4
4.0
0.5
6.0
3.9
1.5
14.4
148.4
371.0
21.2
138.5
8.6
9.0
672.1
30.7
55.8
6.7
24.4
4.5
1.2
3.9
0.7
4.1
0.8
1.5
0.3
2.5
0.4
3.6
0.9
15.9
11.5
2.3
14.4
95.9
536.0
14.9
135.6
6.1
1.6
606.6
23.6
45.4
5.0
19.4
3.3
0.9
3.2
0.5
3.0
0.4
1.5
0.3
0.9
0.2
3.7
0.5
17.1
9.7
3.0
8.8
15.0
110.9
515.3
18.2
137.1
6.4
2.3
1029.0
26.4
52.0
6.0
23.6
4.6
1.3
3.5
0.5
3.2
0.7
1.8
0.3
1.7
0.3
3.5
0.4
12.4
9.8
2.7
12.8
70.6
320.5
23.1
174.1
4.5
0.3
552.1
25.3
50.4
5.8
21.5
4.8
1.2
3.8
0.6
3.7
0.8
3.3
0.6
3.7
0.5
5.0
0.3
1.5
8.1
2.3
57.6
11.6
11.8
363.1
16.2
68.6
2.1
0.4
75.0
12.0
25.8
3.1
14.0
3.3
1.1
3.3
0.5
2.9
0.6
2.0
0.3
1.4
0.2
1.7
0.2
1.5
3.6
0.7
54.5
11.6
6.3
370.3
11.5
36.8
1.0
0.2
23.5
4.5
11.4
1.7
6.9
2.2
0.7
2.3
0.3
1.9
0.3
1.4
0.2
1.0
0.1
1.1
0.1
1.6
1.0
0.5
2.0
0.3
3.3
0.5
28.0
5.2
2.8
118
12.6
154.4
176.9
17.7
148.8
17.3
0.9
581.8
49.7
77.4
7.1
23.5
4.2
0.7
3.0
0.5
2.8
0.7
1.2
3.2
0.5
4.6
1.3
5.6
42.1
9.4
22.5
13.3
8.4
145.7
26.3
63.0
1.1
0.1
41.1
3.6
8.9
1.5
8.0
3.0
1.0
3.7
0.7
4.2
0.9
3.2
0.4
2.8
0.4
2.2
0.1
2.0
0.3
0.1
Appendix 3
118
Sample
type
segment
locality
Age
SiO2
TiO2
Al2O3
Fe2O3t
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Sum
Sc
V
Sample
type
segment
locality
Age
SiO2
TiO2
Al2O3
Fe2O3t
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Sum
Sc
V
Cr
Co
6.9
DG068
oph
Apuseni
SA
Jurassic
51.32
0.35
16.30
6.58
0.16
8.60
12.38
2.80
0.16
0.02
1.42
100.08
38.7
144.5
142.0
37.0
12.5
131.1
167.7
15.7
175.2
16.3
0.4
958.8
60.6
95.9
8.5
27.4
3.7
0.7
3.2
0.5
2.9
0.6
1.9
0.4
2.2
0.4
4.7
1.1
6.9
35.8
4.6
48.5
9.7
5.8
214.3
10.5
16.9
0.2
0.7
22.1
0.8
2.4
0.4
2.6
1.0
0.4
1.4
0.2
1.9
0.4
1.4
0.2
1.3
0.2
0.6
0.1
1.6
0.1
0.1
DG072
IAG
Apuseni
SA
Jurassic
76.33
0.14
12.71
0.78
0.01
0.03
0.56
3.24
4.93
0.03
0.68
99.45
9.1
5.6
11.6
119.2
143.8
13.7
93.6
12.8
0.9
1086.5
35.5
58.8
6.4
18.8
3.2
0.6
2.8
0.5
2.6
0.6
1.7
0.4
2.2
0.3
3.8
0.9
7.8
27.1
4.4
DG081
IAG
Apuseni
SA
Jurassic
59.26
0.60
17.11
3.42
0.05
1.54
4.88
6.53
1.32
0.28
4.93
99.91
9.3
86.8
18.6
8.5
DG082
oph
Apuseni
SA
Jurassic
54.85
2.95
15.61
2.94
0.11
4.81
9.63
5.47
0.56
0.04
3.02
100.01
20.4
221.9
141.3
8.8
11.0
20.4
573.3
22.6
192.1
7.1
0.5
329.1
43.7
75.4
8.9
31.8
5.8
1.7
4.9
0.7
4.2
0.8
2.1
0.3
2.2
0.3
4.6
0.4
11.4
13.0
3.5
15.5
14.4
7.4
258.2
78.0
76.6
10.0
2.4
60.3
24.3
82.2
12.1
56.9
14.1
2.5
15.1
2.4
15.1
3.1
9.2
1.2
8.0
1.1
2.3
0.5
0.4
4.5
0.9
DG119
nv
Apuseni
SA
Miocene
60.75
0.48
16.26
4.63
0.14
1.97
5.46
3.14
1.29
0.18
5.67
99.97
9.3
76.7
9.9
DG001a
bsm
Banat
EBR
Carboniferous
72.01
0.25
14.63
1.98
0.06
0.98
0.69
6.86
0.88
0.06
1.44
99.84
3.26
19.84
18.74
7.90
14.4
39.2
357.5
20.3
128.8
6.3
3.9
300.3
14.3
30.2
3.8
16.3
3.1
1.0
3.5
0.4
3.4
0.6
2.5
0.3
2.4
0.4
4.3 0.4
8.7
4.6
1.6
13.83
10.01
35.44
202.43
6.65
107.55
5.95
1.08
233.47
11.84
19.64
2.2
8.1
2.0
0.4
1.0
0.3
1.6
0.2
1.1
0.1
1.0
0.2
3.1
0.5
6.9
4.2
1.9
119
DG001b
bsm
Banat
EBR
Carb-Perm
70.98
0.26
15.36
1.70
0.02
0.91
0.57
6.05
2.60
0.07
1.23
99.77
3.04
24.51
24.13
7.95
12.08
98.79
264.57
7.04
108.43
6.36
1.59
739.57
11.93
19.69
2.2
7.7
2.4
0.5
1.2
0.3
1.5
0.2
0.8
0.7
0.2
3.0
0.5
5.2
4.4
2.7
DG004
bsm
Banat
EBR
Carboniferous
57.26
0.81
19.48
5.78
0.07
3.21
6.34
4.11
1.81
0.07
0.99
99.93
16.61
143.60
8.68
22.48
DG012
bsm
Banat
WBR
Carboniferous
64.15
0.84
16.78
5.21
0.07
1.78
2.55
3.30
3.70
0.16
1.37
99.91
15.64
75.96
56.20
13.00
DG025
bsm
Banat
WBR
Ordovician
77.90
0.48
10.65
3.61
0.04
1.24
0.63
3.14
0.96
0.14
1.45
100.23
8.06
50.12
37.39
8.77
DG053
bas
Banat
PR
Paleogene
43.00
2.15
14.79
10.30
0.18
9.26
10.89
4.74
1.54
0.99
2.07
99.92
18.65
196.38
176.42
36.93
DG054
bas
Banat
PR
Paleogene
44.21
2.03
14.86
10.01
0.17
9.62
10.68
4.39
1.34
0.82
1.60
99.74
19.23
197.52
226.62
38.35
24.07
17.36
59.07
441.49
20.67
87.46
6.02
1.40
533.71
29.17
62.49
7.0
27.8
6.0
1.4
4.9
0.7
4.0
0.7
2.2
0.3
1.8
0.2
2.5
0.5
5.6
10.5
1.7
6.91
18.86
108.54
327.26
37.42
336.43
15.40
5.05
1932.95
66.30
136.38
15.2
58.8
10.5
1.9
9.3
1.1
7.1
1.5
4.6
0.6
3.9
0.6
9.2
1.0
15.2
26.6
1.6
9.01
31.91
112.48
16.12
130.73
6.46
1.31
200.24
20.24
39.97
4.8
18.9
4.2
1.0
3.6
0.7
3.2
0.6
2.4
0.4
1.5
0.4
3.0
0.4
6.4
5.7
0.7
135.18
16.74
42.20
1074.32
25.78
209.75
111.56
0.96
1163.26
64.89
114.71
11.9
46.5
8.7
2.4
7.5
1.0
5.5
1.0
2.7
0.3
2.2
0.3
4.2
5.6
3.2
10.0
2.6
143.03
16.23
59.59
1004.59
22.43
191.15
89.35
0.85
957.66
55.99
103.59
11.0
42.0
7.6
2.2
6.4
0.8
5.0
0.9
2.5
0.3
2.0
0.2
4.0
4.7
4.6
9.4
2.4
Appendix 3
119
Ni
Ga
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
DG067
IAG
Apuseni
SA
Jurassic
72.25
0.26
14.19
1.95
0.06
0.42
0.68
4.00
5.11
0.07
0.92
99.91
2.2
24.7
DG060
bas
Banat
PR
Paleogene
42.22
2.32
13.72
10.71
0.17
12.01
10.78
3.76
1.18
0.62
2.25
99.74
22.93
215.26
294.58
44.58
DG061
nv
Apuseni
SA
Miocene
59.39
0.46
17.47
4.55
0.10
2.01
6.87
3.67
2.28
0.22
2.65
99.67
13.94
151.77
13.12
9.16
Ni
Ga
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
219.86
14.72
27.46
754.64
19.72
155.82
81.27
0.74
790.40
39.67
73.01
7.7
32.5
6.0
2.1
5.6
0.8
4.2
0.7
2.0
0.3
1.2
0.2
3.7
4.5
1.7
6.2
1.7
15.33
46.06
1568.41
13.48
82.35
7.74
1.09
2128.71
62.28
101.11
9.7
32.7
4.5
1.2
3.2
0.4
2.3
0.4
1.5
0.2
1.3
0.2
2.3
0.4
27.4
16.5
4.3
DG073
sed
Apuseni
SA
Ordovician
66.67
0.59
6.23
3.38
0.19
1.01
10.25
0.55
1.72
0.07
9.30
99.97
Appendix 3
120
Sample
type
segment
locality
Age
SiO2
TiO2
Al2O3
Fe2O3t
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Sum
Sc
V
Cr
Co
120
Appendix 4: Sr and Nd isotope data.
locality
Age
DG002
DG003
DG006
DG007
DG008
DG009
DG010
DG011
DG013
DG014
DG015
DG016
DG017
DG018
DG019
DG020
DG021
DG024
DG026
DG027
DG028
DG029
DG030
DG031
DG032
DG033
DG034
DG035
DG036
DG037
DG038
DG039
DG040
DG041
DG042
DG043
DG044
EBR
EBR
EBR
CBR
CBR
CBR
CBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
WBR
SWBR
SWBR
SWBR
SWBR
SWBR
CBR
PR
PR
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Rb_ppm Sr_ppm
67.5
62.5
61.6
49.6
67.4
64.3
62.0
109.8
112.9
169.0
93.2
139.9
124.5
142.5
153.9
131.8
151.8
104.9
99.4
40.4
73.9
53.4
83.4
27.3
106.9
120.1
58.4
129.1
123.0
13.2
91.4
59.4
81.7
46.0
55.3
78.4
65.9
595.9
672.0
675.4
653.9
762.9
806.5
774.1
477.9
470.3
657.1
498.3
568.1
171.3
579.8
528.4
504.9
563.0
1006.4
607.6
240.4
607.4
358.1
600.6
548.6
559.7
869.0
1062.6
566.9
652.2
1475.5
577.9
788.1
483.6
449.6
471.7
526.9
579.5
87
Rb/86Sr
0.3197
0.2624
0.2575
0.2143
0.2494
0.2251
0.2261
0.6485
0.6776
0.7259
0.5279
0.6951
2.0514
0.6937
0.8222
0.7369
0.7608
0.2942
0.4617
0.4743
0.3434
0.4217
0.3919
0.1410
0.5389
0.3901
0.1552
0.6429
0.5325
0.0253
0.4464
0.3500
0.4768
0.2894
0.3309
0.4199
0.3212
87
Sr/86Sr
2σ
87
Sr/86Sr80Ma
0.705683
± 35
0.705320
0.705872
0.705227
0.706479
0.706036
0.704839
0.705673
0.706062
0.706037
0.705468
0.705619
0.707254
0.705361
0.705891
0.705684
0.705147
0.706127
0.705542
0.705223
0.705492
0.705041
0.705956
0.707234
0.705581
0.705041
0.704872
0.705567
0.705519
0.704459
0.705373
0.704641
0.705101
0.704942
0.706359
0.705446
0.705140
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
0.705579
0.704983
0.706196
0.705780
0.704582
0.704936
0.705292
0.705212
0.704868
0.704829
0.704922
0.704573
0.704956
0.704846
0.704282
0.705793
0.705017
0.704684
0.705102
0.704562
0.705511
0.707074
0.704969
0.704598
0.704696
0.704836
0.704914
0.704430
0.704866
0.704243
0.704559
0.704613
0.705983
0.704969
0.704775
4
19
15
21
26
21
11
15
11
9
12
4
11
3
7
19
10
11
22
15
17
18
12
16
9
10
5
5
11
41
13
2
19
4
6
Nd_ppm Sm_ppm
121
19.2
19.4
20.7
14.5
14.2
14.1
15.6
21.8
22.8
34.8
25.2
35.9
16.6
26.9
25.4
20.8
28.1
42.8
15.2
15.3
36.7
20.3
28.5
29.8
24.2
36.8
35.6
28.7
29.2
19.2
18.5
25.8
15.6
16.9
13.5
20.1
24.9
3.9
4.4
4.3
3.0
3.5
2.7
3.5
4.7
4.2
7.5
5.1
7.2
3.3
5.6
4.6
3.9
4.7
7.7
2.0
3.4
6.9
4.0
5.9
6.3
3.5
7.0
7.4
5.6
6.4
4.4
3.5
4.7
2.6
3.2
2.8
3.6
3.7
147
Sm/144Nd
0.1279
0.1415
0.1317
0.1311
0.1552
0.1204
0.1413
0.1358
0.1160
0.1357
0.1275
0.1263
0.1244
0.1311
0.1152
0.1191
0.1054
0.1135
0.0829
0.1400
0.1184
0.1241
0.1304
0.1332
0.0907
0.1205
0.1316
0.1234
0.1380
0.1443
0.1192
0.1147
0.1050
0.1193
0.1306
0.1134
0.0944
143
Nd/144Nd
2σ
0.512561
0.512533
0.512518
0.512493
0.512462
0.512471
0.512595
0.512623
0.512611
0.512629
0.512664
0.512626
0.512598
0.512629
0.512539
0.512594
0.512626
0.512640
0.512625
0.512736
0.512652
0.512718
0.512643
0.512512
0.512603
0.512619
0.512633
0.512618
0.512603
0.512693
0.512627
0.512623
0.512645
0.512714
0.512442
0.512602
0.512573
± 4
± 2
± 5
± 3
± 7
± 5
± 3
± 2
± 4
± 2
± 3
± 2
± 3
± 3
± 12
± 3
± 5
± 6
± 4
± 5
± 6
± 5
± 5
± 44
± 4
± 5
± 2
± 3
± 5
± 3
± 3
± 32
± 8
± 7
± 3
± 3
± 8
143
Nd/144Nd80Ma
εNd80Ma
εNd0Ma
0.512494
0.512459
0.512449
0.512424
0.512381
0.512408
0.512521
0.512552
0.512550
0.512558
0.512597
0.512560
0.512533
0.512560
0.512479
0.512532
0.512571
0.512581
0.512582
0.512663
0.512590
0.512653
0.512575
0.512442
0.512556
0.512556
0.512564
0.512553
0.512531
0.512617
0.512565
0.512563
0.512590
0.512652
0.512374
0.512543
0.512524
-0.8
-1.5
-1.7
-2.2
-3.0
-2.5
-0.3
0.3
0.3
0.4
1.2
0.5
0.0
0.5
-1.1
-0.1
0.7
0.9
0.9
2.5
1.1
2.3
0.8
-1.8
0.4
0.4
0.6
0.4
-0.1
1.6
0.6
0.5
1.1
2.3
-3.1
0.1
-0.2
-1.5
-2.0
-2.3
-2.8
-3.4
-3.3
-0.8
-0.3
-0.5
-0.2
0.5
-0.2
-0.8
-0.2
-1.9
-0.9
-0.2
0.0
-0.3
1.9
0.3
1.6
0.1
-2.5
-0.7
-0.4
-0.1
-0.4
-0.7
1.1
-0.2
-0.3
0.1
1.5
-3.8
-0.7
-1.3
Appendix 4
121
sample
locality
Age
DG045
DG046
DG047
DG048
DG049
DG050
DG051
DG052
DG055
DG056
DG057
DG058
DG059
DG076
DG077
DG078
DG079
DG080
DG084
DG085
DG086
DG090
DG091
DG092
DG093
DG094
DG097
DG099
DG100
DG101
DG102
DG103
DG109
DG110
DG112
DG113
DG115
DG116
PR
PR
PR
PR
PR
PR
PR
PR
PR
PR
PR
SA
SA
SA
SA
SA
SA
SA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
NA
SA
SA
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Cretaceous
Rb_ppm Sr_ppm
117.9
98.2
113.3
67.5
88.2
262.2
513.1
325.8
761.5
528.3
52.6
66.3
55.0
169.1
87.0
90.3
45.0
146.2
54.7
181.6
129.8
100.0
170.4
152.4
82.1
115.3
119.0
154.8
154.4
172.8
129.7
85.6
136.7
90.3
116.7
97.4
109.3
110.8
91.7
95.0
87.2
69.1
635.2
318.0
737.8
527.2
470.1
583.6
580.4
602.3
518.7
453.4
550.9
94.7
285.0
266.1
490.8
229.3
112.6
139.1
87.2
38.1
40.4
381.5
250.7
418.2
318.7
257.8
321.5
230.0
84.1
597.3
907.4
248.5
87
Rb/86Sr
1.2693
0.5404
0.9818
0.2501
0.4712
0.8068
0.2339
0.5886
0.2103
0.9055
0.5226
0.4366
0.2186
0.6853
0.2978
1.1305
0.6652
2.9801
1.6873
1.6167
0.4723
1.4185
2.9822
3.1406
4.9955
13.1400
9.3120
0.6331
1.5392
0.6092
1.0332
1.0662
0.9596
1.3585
3.0808
0.4491
0.2712
0.7849
87
Sr/86Sr
2σ
87
Sr/86Sr80Ma
0.705290
0.705790
0.705077
0.705642
0.705798
0.704936
0.706049
0.705228
0.706701
0.706657
0.705631
0.704484
0.706011
0.704660
0.706586
0.706439
0.712464
0.708562
0.708528
0.706355
0.707828
0.710317
0.711134
0.715551
0.731214
0.720063
0.707382
0.709127
0.707446
0.708133
0.708334
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
11
4
15
13
12
24
15
9
9
9
11
3
19
17
2
6
7
3
5
13
7
72
16
11
12
17
9
45
24
2
3
0.704676
0.704674
0.704793
0.705106
0.704881
0.704670
0.705380
0.704989
0.705672
0.706063
0.705135
0.704236
0.705232
0.704321
0.705301
0.705683
0.709077
0.706644
0.706690
0.705819
0.706216
0.706927
0.707564
0.709873
0.716278
0.709479
0.706662
0.707377
0.706754
0.706959
0.707122
0.707881
0.715006
0.706636
0.708879
0.706914
±
±
±
±
±
21
49
28
15
9
0.706337
0.711504
0.706126
0.708571
0.706022
Nd_ppm Sm_ppm
122
14.7
22.5
18.3
39.2
23.4
2.5
4.1
2.4
6.5
3.7
21.4
19.5
25.1
27.7
30.3
31.6
26.7
32.6
23.9
20.5
24.6
11.1
23.4
25.0
24.0
20.9
27.6
30.9
44.1
4.0
3.3
5.4
5.1
5.9
6.2
5.7
5.7
4.5
4.3
4.7
2.9
3.7
5.1
5.7
4.5
5.2
6.6
9.5
18.8
20.3
21.0
21.7
25.7
27.3
20.4
31.0
21.9
31.2
20.2
3.7
3.5
4.0
4.9
5.4
6.2
4.3
6.3
4.3
7.0
4.1
147
Sm/144Nd
0.1056
0.1141
0.0817
0.1044
0.0987
0.1274
0.1175
0.1058
0.1342
0.1150
0.1226
0.1234
0.1350
0.1095
0.1175
0.1127
0.1195
0.1643
0.1000
0.1273
0.1491
0.1351
0.1185
0.1336
0.1359
0.1899
0.1386
0.1248
0.1075
0.1182
2.8075
2.9861
0.1422
0.1321
0.1276
0.1245
0.1403
0.1268
143
Nd/144Nd
2σ
143
Nd/144Nd80Ma
εNd80Ma
εNd0Ma
0.512526
0.512579
0.512539
0.512587
0.512579
0.512689
0.512611
0.512667
0.512527
0.512509
0.512497
0.512646
0.512767
0.512642
0.512750
0.512518
0.512514
0.512578
0.512471
0.512449
0.512542
0.512503
0.512393
0.512436
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
16
17
25
20
9
22
8
6
4
4
6
3
2
6
8
5
8
69
13
79
13
9
5
17
0.512471
0.512519
0.512496
0.512532
0.512527
0.512622
0.512550
0.512612
0.512457
0.512449
0.512433
0.512581
0.512696
0.512585
0.512688
0.512459
0.512451
0.512492
0.512419
0.512382
0.512464
0.512432
0.512331
0.512366
-1.3
-0.3
-0.8
-0.1
-0.1
1.7
0.3
1.5
-1.5
-1.7
-2.0
0.9
3.1
1.0
3.0
-1.5
-1.6
-0.8
-2.3
-3.0
-1.4
-2.0
-4.0
-3.3
-2.2
-1.2
-1.9
-1.0
-1.2
1.0
-0.5
0.6
-2.2
-2.5
-2.8
0.2
2.5
0.1
2.2
-2.3
-2.4
-1.2
-3.3
-3.7
-1.9
-2.6
-4.8
-3.9
0.512286
0.512322
0.512508
0.512488
0.512445
0.512381
0.512348
0.512526
0.512498
0.512292
0.512535
0.512289
0.512528
± 8
± 9
± 16
± 5
± 5
± 6
± 12
± 9
± 9
± 33
± 7
± 8
± 9
0.512187
0.512249
0.512443
0.512432
0.512383
0.512011
0.511978
0.512452
0.512429
0.512225
0.512470
0.512216
0.512462
-6.8
-5.6
-1.8
-2.0
-3.0
-4.4
-5.0
-1.6
-2.1
-6.0
-1.3
-6.2
-1.4
-6.9
-6.2
-2.5
-2.9
-3.8
-5.0
-5.7
-2.2
-2.7
-6.7
-2.0
-6.8
-2.1
Appendix 4
122
sample
locality
Age
DG062
DG063
DG064
DG067
DG068
DG072
DG081
DG082
DG119
DG001a
DG001b
DG004
DG012
DG025
DG053
DG054
DG060
DG061
DG073
DG083
SA
SA
SA
SA
SA
SA
SA
SA
SA
EBR
EBR
EBR
WBR
WBR
PR
PR
PR
SA
SA
NA
Jurassic
Jurassic
Jurassic
Jurassic
Jurassic
Jurassic
Jurassic
Jurassic
Miocene
Carboniferous
Carb-Perm
Carboniferous
Carboniferous
Ordovician
Paleogene
Paleogene
Paleogene
Miocene
Ordovician
Ordovician
Rb_ppm Sr_ppm
70.6
11.8
6.3
131.1
5.8
119.2
20.4
7.4
39.2
35.4
98.8
59.1
108.5
31.9
42.2
59.6
27.5
46.1
54.1
53.1
320.5
363.1
370.3
167.7
214.3
143.8
573.3
258.2
357.5
202.4
264.6
441.5
327.3
112.5
1074.3
1004.6
754.6
1568.4
326.9
1000.8
87
Rb/86Sr
0.6200
0.0900
0.0500
2.2100
0.0800
2.3400
0.1000
0.0800
0.3100
0.4900
1.0500
0.3800
0.9400
0.8000
0.1100
0.1700
0.1000
0.0800
0.4700
0.1500
87
Sr/86Sr
2σ
0.705206
0.704091
0.703969
0.708983
0.704242
0.709569
0.704771
0.704481
0.705091
0.707875
0.708410
0.709143
0.712949
0.718762
0.703393
0.703507
0.703285
0.704523
6
9
16
16
8
11
40
18
11
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
0.708292 ±
87
Sr/86Sr155Ma
Nd_ppm Sm_ppm
0.703836
0.703889
0.703863
0.704121
0.704074
0.704414
0.704550
0.704303
42
18
11
5
3
9
6
8
123
21.5
14.0
6.9
27.4
2.6
18.8
31.8
56.9
16.4
8.1
7.7
27.8
58.8
18.9
46.5
42.0
32.5
32.7
15.4
27.5
4.8
3.3
2.2
3.7
1.0
3.2
5.8
14.1
3.1
2.0
2.4
6.0
10.5
4.2
8.7
7.6
6.0
4.5
2.8
6.0
147
Sm/144Nd
0.1395
0.1470
0.1972
0.0849
0.2417
0.1070
0.1143
0.1554
0.1206
0.1555
0.1963
0.1359
0.1125
0.1399
0.1183
0.1135
0.1161
0.0866
0.1127
0.1370
143
Nd/144Nd
0.512755
0.512645
0.512781
0.512685
0.512851
0.512724
0.512732
0.512831
0.512741
0.512564
0.512591
0.512341
0.512151
0.512001
0.512875
0.512859
0.512971
0.512649
0.512307
0.512229
2σ
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
4
16
7
4
18
6
4
6
9
3
4
5
4
6
11
4
14
19
5
4
143
Nd/144Nd155Ma
0.512614
0.512496
0.512581
0.512599
0.512606
0.512615
0.512616
0.512673
εNd155Ma εNd0Ma
3.4
1.1
2.8
3.1
3.3
3.5
3.5
4.6
2.3
0.1
2.8
0.9
4.2
1.7
1.8
3.8
2.0
-1.4
-0.9
-5.8
-9.5
-12.4
4.6
4.3
6.5
0.2
-6.5
-8.0
Appendix 4
123
sample
Appendix 5
Appendix 5: Calculated mean 206Pb/238U ages.
ID
type1
locality
Age [Ma]2
uncertainty [Ma]
error type
method3
DG005
d
DG006
age type4
MSWD
EBR
79.6
si
EBR
DG002
d
DG003
si
DG007
4.3
2 sd
LA-ICP-MS
WTD
1.6
26
20 c
80.6
5.2
2 sd
LA-ICP-MS
WTD
1.8
20
19 c
EBR
80.8
EBR
79.0
4.4
2 sd
LA-ICP-MS
WTD
1.9
38
33 c
4.8
2 sd
LA-ICP-MS
WTD
2.1
33
27 c
si
CBR
79.9
4.8
2 sd
LA-ICP-MS
WTD
1.3
39
34 c
DG042
d
CBR
79.5
4.5
2 sd
LA-ICP-MS
WTD
1.9
48
45 c
DG008
si
CBR
DG009
si
CBR
81.3
3.7
2 sd
LA-ICP-MS
WTD
1.4
39
29 c
83.98
0.28
2 sigma
ID-TIMS
3
2 t
DG010
si
CBR
79.5
5
2 sd
LA-ICP-MS
WTD
1.08
40
35 c
DG038
i
SWBR
71.2
4.1
2 sd
LA-ICP-MS
WTD
2
45
33 c
DG039
si
SWBR
73.9
3.2
2 sd
LA-ICP-MS
WTD
1.7
48
40 c
DG040
i
SWBR
71.3
4.3
2 sd
LA-ICP-MS
WTD
1.7
33
27 c
DG041
si
SWBR
70.2
4.5
2 sd
LA-ICP-MS
WTD
0.79
39
33 c
DG030
si
WBR
77.0
3.3
2 sd
LA-ICP-MS
WTD
2.7
24
17 t
DG032
i
WBR
75.1
3.2
2 sd
LA-ICP-MS
WTD
1.12
43
34 c
DG013
i
WBR
75.2
2.3
2 sd
LA-ICP-MS
WTD
3.9
47
38 t
DG033
d
WBR
76.6
3.7
2 sd
LA-ICP-MS
WTD
1.6
45
35 c
DG034
i
WBR
78.3
1.6
2 sd
LA-ICP-MS
WTD
2.3
49
42 c
DG035
i
WBR
78.0
2.8
2 sd
LA-ICP-MS
WTD
1.1
46
37 c
DG019
i
WBR
79.4
2.1
2 sd
LA-ICP-MS
WTD
1.9
47
37 c
DG020
i
WBR
78.3
2.1
2 sd
LA-ICP-MS
WTD
1.9
55
43 c
DG021
i
WBR
79.8
1.6
2 sd
LA-ICP-MS
WTD
1.6
49
42 c
CA
N5 N*6 quality7
DG014nonCA
i
WBR
77.8
2.5
2 sd
LA-ICP-MS
WTD
2.1
39
33 t
DG016
i
WBR
78.9
1.8
2 sd
LA-ICP-MS
WTD
1.5
48
38 c
DG022
i
WBR
77.5
2.4
2 sd
LA-ICP-MS
WTD
1.2
47
46 c
DG023
d
WBR
76.6
7.5
2 sd
LA-ICP-MS
WTD
2
14
14 t
DG024
i
WBR
76.9
2.9
2 sd
LA-ICP-MS
WTD
0.93
20
16 t
DG026
i
WBR
76.45
0.09
2 sigma
ID-TIMS
CA
5
4 c
DG027
d
WBR
76.49
0.09
2 sigma
ID-TIMS
CA
4
2 t
DG028
i
WBR
76.43
0.10
2 sigma
ID-TIMS
CA
6
2 c
DG029
d
WBR
77.00
0.16
2 sigma
ID-TIMS
CA
3
2 t
DG044
i
PR
76.06
0.16
2 sigma
ID-TIMS
CA
5
4 c
DG045
si
PR
75.06
0.18
2 sigma
ID-TIMS
CA
5
4 c
DG049
i
PR
77.15
0.10
2 sigma
ID-TIMS
CA
4
3 t
DG052
d
PR
78.0
3.1
2 sd
LA-ICP-MS
46
42 c
DG077
v
SA
77.85
0.07
2 sigma
ID-TIMS
4
1 t
DG079
v
SA
79.0
2.7
2 sd
LA-ICP-MS
WTD
1.4
25
17 t
DG115
v
SA
80.0
1.7
2 sd
LA-ICP-MS
WTD
3.2
46
29 t
DG116
si
SA
80.3
0.9
2 sd
LA-ICP-MS
WTD
1.5
48
33 c
DG085
i
NA
79.5
3.2
2 sd
LA-ICP-MS
WTD
2.8
31
21 t
DG113
i
NA
80.3
1.6
2 sd
LA-ICP-MS
WTD
2.8
32
24 t
DG090
i
NA
80.3
1.3
2 sd
LA-ICP-MS
WTD
2.8
53
36 c
DG091
v
NA
80.2
1.7
2 sd
LA-ICP-MS
WTD
2.8
39
31 c
DG092
v
NA
79.8
1.3
2 sd
LA-ICP-MS
WTD
2.2
49
36 c
124
WTD
1.9
yc
Appendix 5
ID
type1
locality
Age [Ma]2
uncertainty [Ma]
error type
method3
DG093
v
NA
80.3
DG098
v
SA
DG109
si
DG110
si
DG112
4.4
2 sd
LA-ICP-MS
WTD
5.4
42
31 t
80.8
1.4
2 sd
LA-ICP-MS
WTD
2.2
49
33 t
NA
80.3
1.2
2 sd
LA-ICP-MS
WTD
1.8
40
32 c
NA
80.7
0.8
2 sd
LA-ICP-MS
WTD
0.98
55
30 c
v
NA
80.8
1.5
2 sd
LA-ICP-MS
WTD
2.5
46
35 t
DG099
i
NA
79.9
1.2
2 sd
LA-ICP-MS
WTD
1.3
47
44 c
DG100
si
NA
77.5
1.2
2 sd
LA-ICP-MS
WTD
2.7
49
31 t
DG101
i
NA
77.4
1.2
2 sd
LA-ICP-MS
WTD
2.6
43
34 t
DG102
si
NA
76.0
2.3
2 sd
LA-ICP-MS
WTD
1.3
50
38 c
DG103
si
NA
75.5
1.9
2 sd
LA-ICP-MS
WTD
1.6
53
35 c
DG062
IAG
SA
158.6
2.9
2 sd
LA-ICP-MS
WTD
2.3
41
25 t
DG063
oph
SA
155.9
4.9
2 sd
LA-ICP-MS
WTD
1.3
47
39 c
DG064
oph
SA
157.7
5.0
2 sd
LA-ICP-MS
WTD
1.3
50
38 c
DG067
IAG
SA
155.0
3.0
2 sd
LA-ICP-MS
WTD
3.3
50
20 t
DG068
oph
SA
158.9
5.2
2 sd
LA-ICP-MS
WTD
1.5
24
19 c
DG069
IAG
SA
154.8
2.9
2 sd
LA-ICP-MS
WTD
3.5
46
21 t
DG072
IAG
SA
152.9
4
2 sd
LA-ICP-MS
WTD
1.08
34
25 c
DG081
IAG
SA
154.5
4.5
2 sd
LA-ICP-MS
WTD
1.2
45
35 c
DG082
oph
SA
158.2
4.4
2 sd
LA-ICP-MS
WTD
1.09
46
41 c
DG119
nv
SA
11.5
0.216
2 sd
LA-ICP-MS
WTD
1.1
40
36 c
DG025
bsm
WBR
449.6
57.48
2 sigma
TIMS
yc
4
1 t
DG073
sed
SA
442.9
3.6
LA-ICP-MS
tuffzirc
69
16 t
DG083
sed
NA
454.3
7
LA-ICP-MS
tuffzirc
58
7 t
DG004
bsm
EBR
317.1
13.2
LA-ICP-MS
WTD
42
41 c
2 sd
age type4
MSWD N*6 N5
quality7
CA (chemical abrasion)-treated zircons were preferred over non-CA-treated zircons.
1
v volcanic; i intrusive; si shallow intrusive; d dyke; nv neogene volcanic; IAG island arc granitoid; oph
ophiolite; bsm basement; sed sediment
2
calculated 206Pb/238U crystallization age
3
TIMS Thermal Ionization Mass Spectrometry; LA-ICP-MS Laser-ablation inductively-coupled plasmamass-spectrometry
4
CA Concordia Age; WTD weighted average; yc youngest concordant; tuffzirc age calculated with
zircon age extraction algorithm of ISOPLOT
5
Number of zircons analyzed
6
Number of zircons used for age calculation
7
c confident; t tentative
125
Appendix 6
Appendix 6: Map and sampling locations for the Banat region and Apuseni Mountains.
Appendix Figure A1 (a): Geological map of Banat region with sampling points.
126
Appendix 6
Appendix Figure A1 (b): Geological map of Apuseni Mountains with sampling points.
127
Appendix 7
Appendix 7: Tectonic map of ABTS belt summarizing the crystallization ages.
Appendix Figure A2: Tectonic map of ABTS belt with mean 206Pb/238U ages of Late Cretaceous igneous
rocks.
128
Appendix 8
Appendix 8: Concordia and weighted mean 206Pb/238U age plots for LA-ICP-MS and TIMS
dating.
Appendix Figure A3: Concordia and weighted mean
Banat region and Apuseni Mountains.
129
206
Pb/238U age plots for all dated samples from
Appendix 8
130
Appendix 8
131
Appendix 8
132
Appendix 8
133
Appendix 8
134
Appendix 8
135
Appendix 8
Plots for TIMS dating for Late Cretaceous igneous rocks (DG026, DG027, DG028, DG029, DG009,
DG077, DG044, DG049, DG045).
136
Appendix 8
137
Appendix 8
138
Appendix 9
Appendix 9: U-Pb dates of inherited zircons.
Appendix Figure A4 (a) and (b): Histograms of in-situ LA-ICP-MS single grain 206Pb/238U ages of
concordant inherited zircons for (a) the Banat segment, and (b) the Apuseni segment. Plotted are only
zircons younger than 1 Ga, because older zircons are mostly discordant. Bin width is 20 Ma. Two tuffites
(DG073, DG083) from the Apuseni segment with mainly inherited zorcons are not included in this plot
and shown separately (c, d). Peaks occurring at 80-100 Ma are not necessarily zircons inherited from
basement lithologies, but might indicate prolonged crystallisation in a crustal magma chamber.
Interestingly, although isotope data of the samples from the Apuseni Mountains indicate abundant
crustal contribution, the Apuseni samples contain less zircons inherited from crustal basement rocks.The
most prominent peaks occur at 200 Ma (Upper Triassic to Lower Jurassic), 300-360 Ma (Carboniferous;
Variscan orogeny), 420-480 Ma (Ordovician-Silurian; Caledonian orogeny), 500-540 Ma (Cambrian;
Cadomian orogeny) and 620-660 Ma (Neoproterozoic). Zircons older than 1 and 2 Ga are more
frequently observed in the Apuseni segment. Compared to the Timok (Kolb 2011) and Eastern
Srednogorie (Georgiev et al. 2012) segments, the Ordovician-Silurian and the Neoproterozoic peaks are
more prominent in the Banat and Apuseni segments, whereas Jurassic peaks are largely missing.
Appendix Figure A4 (c) and (d): In-situ LA-ICP-MS single grain zircon 206Pb/238U ages of tuffite
samples DG073 and DG083 (Apuseni Mountains). Shown are only zicons younger than 1 Ga. Two
prominent peaks at 400-450 Ma and 600-650Ma. DG083 from the northern Apuseni Mountains also
contains ~300 Ma old zircons.
139
Appendix 10
Appendix 10: Hf isotope plots.
Appendix Figure A4: Hf isotope diagram for single zircons from the Banat region and Apuseni
Mountains. All εHf values were age-corrected to 80 Ma. Uncertainty in the εHf is less than 1 epsilon
unit (2σ). The samples were grouped into regions for easier comparison. (a) εHf80Ma of single zircons
versus 206Pb/238U weighted average age of the host rock. (b) εHf80Ma of single zircons versus 206Pb/238U
single zircon age. Error bar in lower right corner shows average 1σ error of the single zircon ages.
140
Appendix 10
Appendix Figure A4: Hf isotope diagram for single zircons from the Banat region and Apuseni
Mountains. All εHf values were age-corrected to 80 Ma. Uncertainty in the εHf is less than 1 epsilon
unit (2σ). The samples were grouped into regions for easier comparison. (c) εHf80Ma of single zircons
versus 87Sr/86Sr80Ma of the host rock. (d) εHf80Ma of single zircons versus εNd80Ma of the host rock.
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Appendix Figure A4: Hf isotope diagram for single zircons from the Banat region and Apuseni
Mountains. All εHf values were age-corrected to 80 Ma. Uncertainty in the εHf is less than 1 epsilon
unit (2σ). The samples were grouped into regions for easier comparison. (a) εHf80Ma of single zircons
versus 206Pb/238U weighted average age of the host rock. (e) εHf80Ma of single zircons versus across-arc
distance (km) for samples from the Banat region.
(a) Samples older than 76 Ma from the Banat region have generally positive εHf values
indicating a contribution of mantle. The scatter of εHf values is considerable for the samples from the
Bocşa intrusion in the west Banat (e.g. DG026, DG028, DG029, DG023, DG014). The youngest igneous
rocks sampled in the southwestern Banat region (DG038, DG039, DG040, DG041) evolve to
increasingly more positive mantle-like εHf values. The two samples from the Apuseni Mountains have
negative εHf values pointing towards substantial crustal contamination.
(b) The bulk of the 82 to 74 Ma old single zircons has εHf values between ~0 and +4. Especially
the younger zircons from the southwestern Banat trend towards highly positive εHf values. Single
zircons from the Apuseni Mountains have negative εHf values.
(c) The εHf80Ma of single zircon grains show a negative correlation with 87Sr/86Sr80Ma of the host
rock, particularly visible in the samples from the central Banat region. Samples from the SW Banat seem
to show a positive correlation between εHf and 87Sr/86Sr. The scatter in εHf data is considerable between
87
Sr/86Sr ratios of 0.7045 and 0.7050. The lowest 87Sr/86Sr ratios (DG021, DG039) do not coincide with
the highest εHf values. Negative εHf values coincide with high 87Sr/86Sr ratios (~0.7070) in the two
samples from the Apuseni Mountains.
(d) The εHf80Ma of single zircon grains increase with increasing εNd80Ma of the host rock.
Samples with positive εNd values show a large scatter from highly positive to slightly negative εHf
values. The sample with the highest εHf values (DG041, SW Banat) also has a high εNd value. Samples
from the central and east Banat have positive εHf values indicating mantle-derivation, but their whole
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rock εNd values are negative pointing to crustal assimilation. The samples from the Apuseni Mountains
have negative εHf and εNd values characteristic for crustal input.
(e) Samples from the Banat region are plotted versus their distance from the inferred arc front
(for explanations see chapter 2, Figure 2.6). Closer to the arc front, the zircons have lower εHf80Ma values,
the values increase until ~50 km are reached. Here, a scatter sets in that is largely due to the Bocşa
massif and other intrusions in the west Banat. After ~65 km, the εHf80Ma values decrease again.
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Curriculum Vitae
Personal Information
Name: Daniela Gallhofer
E-mail: [email protected]
Date of birth: 12th January 1985
Nationality: Austria
Education
2011-2015
Doctoral studies at the Institute of Geochemistry and Petrology, ETH Zürich,
Switzerland
Thesis: Magmatic geochemistry and geochronology in relation to the
geodynamic and metallogenic evolution of the Banat Region and the Apuseni
Mountains of Romania
2007-2010
MSc in Applied Geosciences, Montanuniversität Leoben, Austria
Thesis: Lithological and geochemical characterization of the Breitenau
magnesite deposit (Paleozoic of Graz/Eastern Alps)
2003-2007
BSc in Applied Geosciences, Montanuniversität Leoben, Austria
Thesis 1: Geology in the area of St.Jakob/Breitenau (Styria, Eastern Alps)
Thesis 2: Exploration, mineralogy and petrology of Wanko quarry, Lower
Austria
2003
A-levels at BG/BRG Mürzzuschlag
Work experience
Internships:RHI: Breitenau mine
core logging and mapping in magnesite underground and
open pit mine
KMI: Waldenstein mine core logging and mapping in specularite underground
mine
OMV
Teaching assistant for bachelor’s and master’s classes
Skills
Languages: German (Native)
English (B2)
Italian and Spanish (basic knowledge)
Software: Microsoft Office
Adobe Illustrator
ioGAS
Corel Draw
SigmaPlot
excellent knowledge
excellent knowledge
good knowledge
good knowledge
good knowledge
144
EndNote
Iolite
Surpac
good knowledge
good knowledge
basic knowledge
Equipment and analyses: TIMS (thermal ionisation mass spectrometry)
SEM (scanning electron microscopy)
LA-ICP-MS (laser-ablation inductively-coupled-plasma mass
spectrometry)
ion exchange chemistry (U-Pb zircon, Sr-Nd-Pb whole rock, Lu-Hf
zircon)
XRF (x-ray fluorescence)
microscopy, mineral separation (SelFrag, heavy liquid separation)
Driving license: B
Publications
Gallhofer, D., A. v. Quadt, I. Peytcheva, S. M. Schmid, and C. A. Heinrich (2015), Tectonic,
magmatic, and metallogenic evolution of the Late Cretaceous arc in the Carpathian-Balkan
orogen, Tectonics, 34, doi:10.1002/2015TC003834.
von Quadt, A., D. Gallhofer, M. Guillong, I. Peytcheva, M. Waelle, and S. Sakata (2014), UPb dating of CA/non-CA treated zircons obtained by LA-ICP-MS and CA-TIMS techniques:
impact for their geological interpretation, Journal of Analytical Atomic Spectrometry, 29(9),
1618-1629. doi:10.1039/C4JA00102H
Letsch, D., W. Winkler, A. von Quadt, and D. Gallhofer (2015), The volcano-sedimentary
evolution of a post-Variscan intramontane basin in the Swiss Alps (Glarus Verrucano) as
revealed by zircon U–Pb age dating and Hf isotope geochemistry, Int. J. Earth Sci., 104(1),
123-145, doi: 10.1007/s00531-014-1055-0.
Gatzoubaros, M., A. von Quadt, D. Gallhofer, and R. Rey (2014), Magmatic evolution of
pre-ore volcanics and porphyry intrusives associated with the Altar Cu-porphyry prospect,
Argentina, Journal of South American Earth Sciences, 55, 58-82, doi:
10.1016/j.jsames.2014.06.005.
Lehmann, S., J. Barcikowski, A. von Quadt, D. Gallhofer, I. Peytcheva, C. A. Heinrich, and
T. Serafimovski (2013), Geochronology, geochemistry and isotope tracing of the Oligocene
magmatism of the Buchim–Damjan–Borov Dol ore district: Implications for timing, duration
and source of the magmatism, Lithos, 180–181(0), 216-233, doi:
10.1016/j.lithos.2013.09.002.
Referees
Prof. Dr. Christoph A. Heinrich
Fluids and Mineral Resources Group, Institute of Geochemistry and Petrology, ETH Zürich
E-mail: [email protected]
Dr. Albrecht von Quadt
Fluids and Mineral Resources Group, Institute of Geochemistry and Petrology, ETH Zürich
E-mail: [email protected]
145