The Tectono-metamorphic Evolution

of the Tauern Window Eclogites,

Eastern Alps

Andrew Smye

Department of Earth Sciences

University of Cambridge

A thesis submitted for the degree of

Doctor of Philosophy

September 4, 2011

This dissertation is the result of my own work and includes nothing which is the

outcome of work done in collaboration except where specifically indicated in the text.

Andrew J. Smye, September 4, 2011

Acknowledgements

It would not have been possible to write this thesis without the financial support of a

NERC studentship (NE/007647/1), a NERC, NIGL-steering-committee grant (NE/1068/1108)

and a Magdalene College travel grant.

I am indebted to Tim and Mike for their foresight in proposing such an interesting PhD

title and their patient supervision throughout the duration of the thesis. Collectively,

they helped me foster an independence of thought and kindled my curiosity in orogenic

processes. In particular, Tim has graciously taught me the wonders and intricacies of

metamorphic petrology and how it can be used as a tool to investigate tectonic prob-

lems. Mike has encouraged me to think critically and to question the fundamental

tenets of a scientific problem; in particular Mike’s knowledge of thermal modeling and

geochronology was hugely appreciated. Both Tim and Mike are friends, in addition to

supervisors, for which i am very thankful.

The thesis has benefited from fruitful collaboration with Clare Warren (Open Univer-

sity) and Mark Caddick (ETHZ), both of whom have provided wise advice and stim-

ulating discussion concerning the Eastern Alps. Specifically, I am thankful to Clare

who introduced me to the world of 40Ar/39Ar geochronology, and to Mark, who guided

my thoughts concerning thermal modeling and helped me with numerous programming

bugs. Furthermore, Clare o↵ered constructive comments on the 40Ar/39Ar chapter.

I am hugely grateful to Randy Parrish, Dan Condon and Matt Horstwood at NIGL

for directing U–Th–Pb geochronology. Without Randy’s insight into dealing with non-

conventional petrochronometers such as allanite, and Dan’s TIMS expertise, the project

would not have succeeded. Vanessa Paschley and John Cottle are also thanked for

countless hours of LA–ICPMS supervision.

Marian Holness and Alan Smith are thanked for their roles as academic friends. Keith

Gray and Ian Millar (Kingston) provided high-quality thin sections on which the project

is based. Probe work was made possible by Chiara Petrone and Andy Buckley; Martin

Walker kindly tutored me in use of the Department’s SEM. My o�ce mates Niko and

Ben, in addition to a host of other Department members, are thanked for providing

much fun and interesting discussion. Ben Veitch is thanked for help with MATLAB.

Thanks, also, to the sta↵ at Magdalene College who provided a superb home from

which to study.

Summers spent in the Tauern were made so much more enjoyable by the company of

the following field assistants, who, in many cases, not only served to help with the prac-

ticalities of fieldwork, but contributed scientifically: Owen Weller, Lucy Greenwood,

Dave Wilson, Mike Kember and Alan Chetwynd. Frau Dorer and her family provided

much needed hospitality after many a hostile day in the Venedigergruppe.

Finally, special mention to Katie who graciously endured the brunt of my PhD expe-

rience and who continues to challenge me to see work in life’s perspective. Thank you

all.

Abstract

The Tauern Window exposes the largest continuous tract of eclogite-facies Penninic

crust in the Eastern Alps – the Eclogite Zone. Despite being amongst Earth’s best

studied orogens, little is understood about the rates of Alpine metamorphism within

the Eastern Alps. Recent Rb–Sr geochronology suggests East-Alpine eclogitization

occurred within the early Oligocene, whereas thermal modeling studies require HP -

metamorphic ages > 60 Ma, allowing a period of ⇠30 Ma for isotherm re-equilibration

following thrust-sheet emplacement. Given that the Tauern Window is a classic local-

ity for understanding rates of conductive thermal relaxation in tectonically thickened

crust, this paradox raises profound questions over the length scales of the mechanisms

responsible for heat transfer within orogenic crust.

Field-based structural measurements are consistent with the interpretation that the

Eclogite Zone was tectonically juxtasposed between the European basement complex

– the Venediger nappe, and a dismembered ophiolite sequence – the Glockner nappe,

during exhumation. Conventional thermobarometry and pseudosection calculations on

a garnet–chloritoid–kyanite-bearing metasediment collected from the Eclogite Zone,

show that the Eclogite Zone was subducted to the verge of UHP conditions: ⇠26 kbar

and ⇠540�C. Furthermore, such assemblages are shown to be common throughout

(U)HP terranes and representative of a state of enhanced detachment for slab-top ma-

terial. During exhumation the Eclogite Zone was subjected to partial recrystallisation

in the blueschist facies before its upper-structural layers were penetratively a↵ected by

regional Barrovian metamorphism. Coexisting jadeitic-pyroxene and glaucophane pre-

served as rare symplectite cores show that the Glockner nappe experienced a similar P–

T path from blueschist facies metamorphism between ⇠9–13 kbar and and 350–450�C

to Barrovian conditions of ⇠5–8 kbar at 450–550�C. The Venediger nappe lacks demon-

strative evidence for subduction-related HP metamorphism. Phase-diagram modeling

of sub-millimetric garnet, biotite, muscovite and epidote, which grew at the expense

of plutonic plagioclase in the crystalline basement complex, shows that the internal

Venediger nappe experienced high-grade Barrovian metamoprhism, close to solidus con-

ditions: 10–13 kbar and 550–620�C. Collectively, these data are interpreted as showing

that the Penninic-nappe stack was assembled under blueschist-facies conditions prior

to behaving as a coherent unit during concomitant decompression and heating to Bar-

rovian conditions.

LA–MC–ICPMS and ID–TIMS U–Th–Pb ages of metamorphic allanite from the Eclog-

ite Zone, which when coupled with rare earth element analysis and thermobarometric

modeling, demonstrate that the European continental margin was subducted to be-

tween 8–13 kbar (30-45 km) by 34.2±3.6 Ma. These data define: (i.) an upper limit on

the timing of eclogite facies metamorphism at 26.2±1.8 kbar (70–80 km) and 553±12�C,

(ii.) plate velocity (1–6 cm.a�1) exhumation of the Eclogite Zone from mantle to mid-

crustal depths, and (iii.) a maximum duration of 10 Ma (28–38 Ma) for juxtaposition

of Alpine upper-plate and European basment units and subsequent conductive heating

thought to have driven regional Barrovian (re)crystallisation at ca. 30 Ma.

U–Th–Pb geochronologic data are augmented by 40Ar/39Ar single-grain fusion data

from individual muscovite and paragonite collected from representative lithologies of

each of the Pennine nappes. Apparent ages range from 23 Ma to 90 Ma – generally

older than the U–Th–Pb allanite age. Textural and chemical evidence show that the

majority of micas grew or equilibrated close to the peak of HP metamorphism. Nu-

merical modeling of open-system Ar di↵usion in white mica subjected to nappe-specific

T–t paths suggest that 40Ar/39Ar ages constraining the timing of cooling during ex-

humation should lie between 26 and 28 Ma. The discrepancy between measured- and

model-40Ar/39Ar ages confirms the spatially-variably presence of excess-Ar. Modeling

of open and closed system mica–fluid 40Ar exchange shows that intra-sample 40Ar/39Ar

age di↵erences are best reproduced by a finite volume grain boundary network with

centimeter-scale connectivity. Mafic rocks show larger intra-sample age ranges because

they do not de-volatise during burial.

One-dimensional modeling shows that the thermal evolution of the Tauern Window

cannot be explained by conductive heating of overthickened crust on lithosphere with

normal continental thermal gradients. Either the thermal evolution took place with a

significant thermal contribution from the mantle, for which there is little supporting

evidence, or heating occurred during thrusting with heat provided by an overlying

thrust sheet. It is possible that syn–thrusting heating by rapid emplacement of a hot

overthrust sheet, overlooked in previous thermal models, could plausibly have played

an influential role in the tectono–metamorphic evolution of the Eastern Alps and in

overthrust terrains in general.

Contents

1 Introduction 11.1 Project Aims . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11.2 The Alps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2

1.2.1 Lithotectonic units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41.3 The Problem . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51.4 Fieldwork . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61.5 Thesis structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7

2 Geology of the Tauern Window 82.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82.2 Geologic Setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 82.3 Field Area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112.4 Litho-tectonic units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11

2.4.1 Venediger nappe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112.4.1.1 Lithologies of the Venediger nappe . . . . . . . . . . . . . . . . . . . . . . . . 112.4.1.2 Structures of the Venediger nappe . . . . . . . . . . . . . . . . . . . . . . . . 14

2.4.2 Eclogite Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 162.4.2.1 Lithologies of the Eclogite Zone . . . . . . . . . . . . . . . . . . . . . . . . . 162.4.2.2 Structures of the Eclogite Zone . . . . . . . . . . . . . . . . . . . . . . . . . . 19

2.4.3 Rote-Wande nappe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 232.4.4 Glockner nappe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23

2.4.4.1 Lithologies of the Glockner nappe . . . . . . . . . . . . . . . . . . . . . . . . 232.4.4.2 Structures of the Glockner nappe . . . . . . . . . . . . . . . . . . . . . . . . 24

2.4.5 Matrei Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 272.4.6 Altkristallin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28

2.5 Fabric correlation across units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30

3 Conditions of Metamorphism 333.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33

3.1.1 Application of equilibrium thermodynamics . . . . . . . . . . . . . . . . . . . . . . . . 333.1.2 Pseudosection modeling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 353.1.3 Average P–T . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36

3.2 Eclogite Zone P–T evolution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 373.2.1 Previous work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38

vii

3.2.2 Sample petrography . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 393.2.3 Mineral chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 403.2.4 Metamorphic modeling of TH-680 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42

3.2.4.1 Activity–composition models . . . . . . . . . . . . . . . . . . . . . . . . . . . 433.2.4.2 TH–680 P–T pseudosection . . . . . . . . . . . . . . . . . . . . . . . . . . . . 443.2.4.3 Average P–T calculations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 453.2.4.4 Influence of Garnet a–X models . . . . . . . . . . . . . . . . . . . . . . . . . 473.2.4.5 P–T evolution of TH–680 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48

3.3 Garnet–chloritoid–kyanite assemblages . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 503.3.1 Metamorphic modeling of garnet + chloritoid + kyanite assemblages . . . . . . . . . . 52

3.3.1.1 The KFMASH subsystem . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 533.3.1.2 The NCKFMASH subsystem . . . . . . . . . . . . . . . . . . . . . . . . . . . 543.3.1.3 The NCKFMASHO subsystem . . . . . . . . . . . . . . . . . . . . . . . . . . 543.3.1.4 The MnNCKFMASHO system . . . . . . . . . . . . . . . . . . . . . . . . . . 54

3.3.2 Bulk compositional controls . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 563.3.3 Comparison of peak P–T conditions . . . . . . . . . . . . . . . . . . . . . . . . . . . . 583.3.4 Tectonic implications . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60

3.4 Glockner nappe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 623.4.1 Previous work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 623.4.2 Field relations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 633.4.3 Petrology of breakdown textures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 643.4.4 Mineral chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 653.4.5 Metamorphic modeling of gl+jd assemblages . . . . . . . . . . . . . . . . . . . . . . . 66

3.4.5.1 a–X models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 683.4.5.2 Estimating a reactive bulk composition . . . . . . . . . . . . . . . . . . . . . 693.4.5.3 Breakdown reactions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 713.4.5.4 Calibrating the clinopyroxene model . . . . . . . . . . . . . . . . . . . . . . . 753.4.5.5 Peak conditions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 763.4.5.6 Average P–T . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 77

3.4.6 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 793.4.7 Tectonic implications . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 81

3.5 Venediger nappe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 833.5.1 Previous work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 833.5.2 Sample petrography and mineral chemistry . . . . . . . . . . . . . . . . . . . . . . . . 84

3.5.2.1 Samples TH–519 and ASA–41a . . . . . . . . . . . . . . . . . . . . . . . . . . 843.5.2.2 Sample SP614 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 863.5.2.3 Sample ASA–38a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 863.5.2.4 Sample ASA–42a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87

3.5.3 Average P–T . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 873.5.3.1 Sample TH–519 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 873.5.3.2 Sample SP614 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 883.5.3.3 Sample ASA–38a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 883.5.3.4 Sample ASA–42a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88

3.5.4 Metamorphic modeling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 89

viii

3.5.5 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 903.5.6 Tectonic implications . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 93

3.6 P–T evolution of the Pennine nappes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 953.7 Chapter Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 96

4 U–Th–Pb Isotope Geochronology 994.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 994.2 Aims . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1004.3 Previous work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1004.4 Geochronology of high–P metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1024.5 U-Th-Pb geochronology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104

4.5.1 Principles and methodology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1044.5.2 Accessory minerals in Eclogite Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105

4.5.2.1 Techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1054.5.2.2 Mafic lithologies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1064.5.2.3 Metasedimentary lithologies . . . . . . . . . . . . . . . . . . . . . . . . . . . 107

4.6 Allanite and REE-rich epidote . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1094.7 Techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 111

4.7.1 Grain selection . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1114.7.2 LA–ICPMS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1124.7.3 ID–TIMS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 113

4.8 Results and discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1144.8.1 Allanite LA–ICPMS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1144.8.2 Allanite ID–TIMS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 119

4.9 Allanite petrogenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1224.9.1 Textural observations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1224.9.2 REE+Y modeling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1224.9.3 Metamorphic modeling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126

4.10 Metamorphic rates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1284.11 Tectonic Implications . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1294.12 Chapter Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 131

5 40Ar/39Ar Geochronology 1335.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1335.2 40Ar/39Ar geochronology in the Tauern Window . . . . . . . . . . . . . . . . . . . . . . . . . 1345.3 Methodology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 136

5.3.1 Sample description and mica chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . 1365.3.1.1 Venediger nappe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1375.3.1.2 Eclogite Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1375.3.1.3 Rote-Wand nappe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1395.3.1.4 Glockner nappe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 140

5.3.2 Step-heating versus single grain fusion data . . . . . . . . . . . . . . . . . . . . . . . . 1405.4 40Ar/39Ar analytical technique . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1425.5 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143

ix

5.6 Numerical Modelling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1445.6.1 Numerical techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1475.6.2 Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1475.6.3 Venediger nappe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1475.6.4 Eclogite Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1485.6.5 Rote-Wand nappe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1495.6.6 Glockner nappe . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1495.6.7 Uncertainty associated with di↵usion parameters . . . . . . . . . . . . . . . . . . . . . 1515.6.8 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 151

5.7 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1525.7.1 40ArE partitioning . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1535.7.2 Open system . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1555.7.3 Closed system . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1585.7.4 40ArE as a tracer for metamorphic porosity . . . . . . . . . . . . . . . . . . . . . . . . 1605.7.5 Tectono-metamorphic significance of 40Ar/39Ar ages . . . . . . . . . . . . . . . . . . . 1605.7.6 Suitability of 40Ar/39Ar geochronology under HP conditions . . . . . . . . . . . . . . 162

5.8 Chapter Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163

6 Thermal modeling 1646.1 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1646.2 Previous work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1656.3 Aim . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1666.4 The model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 168

6.4.1 Numerical techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1686.4.2 Model constraints . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 169

6.5 Solutions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1756.5.1 Overthrust sheet thickness . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1756.5.2 Crustal heat production and mantle heat flow . . . . . . . . . . . . . . . . . . . . . . . 1756.5.3 Alternative solutions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 175

6.5.3.1 Tectonic emplacement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1756.5.3.2 Syn-thrusting heating . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 177

6.6 Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1796.7 Chapter Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 187

7 Discussion and Conclusions 1887.1 Exhumation of the Eclogite Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 188

7.1.1 Buoyancy-driven exhumation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1907.1.2 Return flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 194

7.2 Barrovian metamorphism: where’s the heat? . . . . . . . . . . . . . . . . . . . . . . . . . . . 1957.3 Tectonic Implications . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1977.4 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2037.5 Future Work . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 204

x

A Appendix 205A.1 Published manuscripts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 205A.2 Electron Microprobe analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 207A.3 Whole–rock REE ICPMS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 218A.4 U–Th–Pb and 40Ar/39Ar single grain fusion data . . . . . . . . . . . . . . . . . . . . . . . . . 218

References 256

xi

Chapter 1

Introduction

1.1 Project Aims

High-pressure (HP ) metamorphic rocks are a ubiquitous feature of Phanerozoic mountain belts.Their occurrence provides important constraints on both the transfer of material between the crustand mantle, and also on the timing of continental collision within zones of active and past con-vergence (De Sigoyer et al., 2000; Rubatto & Hermann, 2001; Leech et al., 2005; Parrish et al.,2006). Whilst subduction provides a plausible mechanism by which tracts of oceanic and conti-nental lithosphere are conveyed to upper mantle depths, the processes responsible for exhumingsuch rocks back to Earth’s surface are less well understood (e.g. Agard et al., 2009, and referencestherein). Within the typical orogenic cycle (ultra)high-pressure ((U)HP ) metamorphism occursearly, linked to subduction of intervening oceanic realms, before the inception of continental colli-sion which thickens the crust and drives subsequent Barrovian metamorphism. This accounts forcommonly observed amphibolite–greenschist facies overprinting of blueschist–eclogite facies meta-morphic rocks (e.g. England & Richardson, 1977; De Sigoyer et al., 1997, 2004). High–precisionradioisotope geochronology of both (U)HP and Barrovian metamorphic events allows calculationof average exhumation rates of (U)HP bodies. Recently, this approach has shown that exhumationrates are of the same order of magnitude (cm.a�1) as plate velocities (e.g. Rubatto & Hermann,2001; Baldwin et al., 2004; Me↵an-Main et al., 2004; Parrish et al., 2006). Critically, radiogenicdating of both (U)HP and Barrovian-style metamorphism in exhumed poly-metamorphic terranesconstrains the timescales of thermal equilibration of orogenic crust following continental collision.Prevailing thought suggests that heating of tectonically thickened crust is driven by conductive re-laxation of perturbed isotherms which operates over timescales of tens of millions of years (Oxburgh& Turcotte, 1974; Bickle et al., 1975; England & Thompson, 1984).

Recently, the accepted interpretation of the timing of tectono-metamorphic events in the East-ern Alps has been challenged (Glodny et al., 2005, 2008) and little is understood about the ratesof Alpine metamorphism in the Eastern Alps, largely due to a paucity of reliable geochronologicaldata. The aim of this project is to combine thermodynamic modeling of East-Alpine metamorphism

1

with precise radiogenic isotope geochronology to understand the rates of tectono-metamorphic pro-cesses responsible for the evolution of an eclogite terrane pertaining to the Alpine orogen’s lowerplate—the Eclogite Zone of the Eastern Alps. Calculation of P–T–t paths of both the EclogiteZone and adjacent units, a↵ords construction of a spatio-temporal model of East Alpine orogenisis.

Figure 1.1: Relief map of the Alpine arc using SRTM data. White arrows are convergence vectors between Adriatic

microplate (shaded region) and stable Eurasia, calculated by Calais et al. (2002) for the time interval 1996–2001; numbers

are convergence rates in mm.a�1; circles are 95% confidence ellipse. Star denotes Calais et al. (2002)’s rotation pole for

the Adriatic microplate. Dashed line highlights the position of the Tauern Window.

1.2 The Alps

The Alpine arc is the surface expression of ca. 65 million year’s worth (>514 km total convergence;Schmid et al., 2004b) of northward directed convergence between the European and African plates.During convergence, the Apulian microcontinent was trapped and acted as an indenter relative toEurope(Tapponnier, 1977), which resulted in the closure of the mid-Jurassic Tethyan realm duringthe Upper Cretaceous–Eocene, and eventually, continental collision close to the Eocene–Oligoceneboundary (Coward & Dietrich, 1989; Dewey et al., 1989; Laubscher, 1991; Stampfli et al., 1998;

2

Schmid & Kissling, 2000; Schmid et al., 2004b). Collision facilitated syn-metamorphic stacking ofinternal crystalline units (Oxburgh, 1968; Ernst, 1973b; Droop et al., 1990; Duchene et al., 1997;Berger & Bousquet, 2008) on top of external zones along the compressive front (Butler et al., 1986;Choukroune et al., 1986; Fry, 1989; Gratier et al., 1989; Butler, 1992), which forms a fan-shapedgeometry spanning ⇠180� from the Ligurian margin in the west to Slovenia in the east (Fig.1.1).The mountain belt varies between 200 and 400 kilometers in width and is ⇠1000 km along-strike.Peripheral foreland basins bound the mountain chain to the north and south—the Molasse and Pobasins respectively, whilst Oligocene–Miocene rifts form the northwestern extent and the Ligurianocean forms the southwestern boundary.

6°E

8°E

10°

12°

14°

48°N

46°

Tauern Window

ZürichBasel

Geneve

GrenobleMilano

Bozen

Lubljana

Graz

Wien

Po plain

GranParadiso

DoraMaira

Adula

Monte Rosa

Molasse basin

Penninic unitsEuropean margin(U)HP complexesHelvetic units

Austrolalpine unitSouthalpine unitOligocene granitoidsBasin deposits

Major thrust faultMajor normal fault

Major strike-slip fault

Mediterranean

AdriaticSea

100 km

A

B

Figure 1.2: Geological map of the Alpine arc after Schmid et al. (2004a), showing the positions of major lithotectonic

units. Red transect shows position of East Alpine cross section displayed in Fig.1.2; white box details map area of Fig.2.1.

The Alps are tectonically active. Instrumental earthquake monitoring of the Alpine chainshows that intra-chain seismicity is between 3 and 5 ML (Pavoni, 1961; Frechet, 1978; Menard,1988; Thouvenot, 1996; Eva & Solarino, 1998; Sue et al., 1999; Baroux et al., 2001; Kastrup et al.,2004). Seismicity reflects the complexity of the present-day tectonic regime (Pavoni, 1980, 1986;Deichmann & Rybach, 1989; Thouvenot et al., 1990; Delacou et al., 2004): in the Western Alps, themaximum horizontal compression axis is perpendicular to orogen strike, whilst in the inner Alpinearc, the stress field indicates extensional deformation. GPS monitoring constrains present-dayconvergence vectors between Africa and Europe to 3–8 mm.a�1 in a north to northwest direction(Fig.1.1; Argus et al., 1989; DeMets et al., 1990, 1994; Albarello et al., 1995; Kreemer & Holt,2001); relative displacements within the internal zone are <2 mm.a�1 (Calais et al., 2002).

3

The deep structure of the Alps reflects along-strike complexity in the orogen’s architecture(Schmid et al., 2004b). In the Eastern Alps, seismic tomography shows that the European mohodips to the south at ⇠7� and reaches a maximum depth of ⇠60 km where it abuts the Adriatic mohodipping at gentler angles to the north (Fig.1.3; Kummerow et al., 2004; Bleibinhaus & Gebrande,2006; Ebbing et al., 2006). In the western Alps, the European moho underplates the Adriatic mohoat ⇠50 km (Schmid & Kissling, 2000).

A B

20 km

Tauern Window

Calcareous Alps

Austroalpine

basement

Dolomites

Figure 1.3: Receiver functions from Kummerow et al. (2004) and scaled geological cross-section (Fig.1.4) along the

TRANSALP geophysical traverse highlighted in Fig.1.2. Abbreviations used on seismic profile: EM – European Moho;

AM – Adriatic Moho; I – base of the the flysch or Helvetic nappes; II – Sub-Tauern ramp; III - Sub-Dolomites ramp; ACI

– Adriatic Crust Interface—a north dipping discontinuity in the middle and lower Adriatic crust. Green and blue circles

represent crustal thicknesses used by in a subsequent figure.

Although they are complex, lithotectonic relations in the core of the Alpine collisional belt arewell exposed. This, in addition to the fact that it remains the most densely studied mountainbelt on Earth, makes the Alps an excellent laboratory to investigate the driving processes behindorogenesis.

1.2.1 Lithotectonic units

The orogen-scale geological architecture of the Alps can be classified as a series of laterally continu-ous lithotectonic units, juxtaposed by roughly east–west-bearing tectonic discontinuities (Fig.1.2).

European margin—the European margin constitutes the northern and western foreland ofthe Alpine chain and comprises a pre-Mesozoic crystalline basement and a Mesozoic coverseries (Schmid et al., 1996, 2004b). Cover nappes have been variably deformed and detachedfrom the basement—the Helvetic nappes exposed in the Swiss Alps have been completelydetached, whilst in the Eastern Alps the equivalent units (Inner Schieferhulle) remain pa-rautochthonous (e.g. Kurz et al., 1998b) and variably a↵ected by Alpine deformation. The

4

crystalline basement and cover nappes are exposed in the orogen’s core as kilometer-scaleculminations: the Tauern Window in the Eastern Alps (Oxburgh, 1968) and the Lepontinedome of the Central Alps. The foreland molasse basin is largely una↵ected by Alpine thrust-ing (Ceriani et al., 2001) and lies unconformably on the para–autochthonous cover nappes(e.g. Roeder & Bachmann, 1996).

Penninic nappes—a collage of oceanic and microcontinental domains was caught betweenEurope and Africa during collision—collectively these units constitute the Penninic nappecomplex. In the Eastern–Central Alps, remnants of the Valais ocean and its margins arepresent as variably-metamorphosed calc-schists, quartzites and ophiolitic fragments of theBundnerschiefer (Trumpy, 1955; Steinmann, 1994). Valaisian nappes form the structurallylowest member of the Penninic nappes. Tectonic units derived from the Brianconnais micro-continent terrane form the Middle Penninic nappes Schmid et al. (1997); Stampfli & Marchant(1997); Stampfli (2000). They di↵er from the Bundnerschiefer in that they include sliversof pre-Mesozoic basement. Nappes derived from the Piemont–Liguria oceanic realm (AlpineTethys) constitute the structurally highest member of the Penninic nappe system and com-prise a mixture of monotonous calcareous metasediments, ophiolitic fragments, which includepieces of sub-continental mantle, and non-metamorphic cover nappes (i.e. Helminthoid flyschof the Western Alps). Penninic nappes underwent both (U)HP (e.g. Meyre et al., 1999) andHP (e.g. Janots et al., 2009) metamorphism prior and subsequent to continental collisionrespectively.

Pennninic-Austroalpine transition zone—rifted portions of the distal Apulian (i.e. southof Piemont-Liguria ocean) margin were incorporated into the Alpine accretionary wedge(Froitzheim & Manatschal, 1996; Schmid et al., 2004b). Evidence for this is shown by theMargna–Sesia fragment: a composite of basement and ophiolitic melange fragments exposedin the Grisons area of the Central Alps (Muntener et al., 2000).

Austroalpine nappes— basement and cover slices derived from remnants of the southernmargin of the Piemont-Liguria ocean form the Austroalpine nappes. The Austroalpine nappeswere a↵ected by an orogenic episode during the Cretaceous (Eoalpine Froitzheim et al., 1994),which resulted in eclogite formation associated with closure of the Meliata ocean and adjacentmargin (Frank, 1987; Haas et al., 1995). Proximal to the Tauern Window, the Austroalpinenappes are thrust over Valaisian Penninic units (Oxburgh, 1968).

1.3 The Problem

A wealth of geochronological data shows that interior crystalline units of the Alpine arc were af-fected by Barrovian-style metamorphism during the period between 27 and 34 Ma (see Bergeret al., 2011, for review). Multimineral Rb/Sr internal isochrons from eclogite facies rocks of the

5

Northern Calcareous Alps Tauern Southern Alps

20 km

BA

Penninic basement

complex

Peripheral

Schieferhülle

Lower Austroalpine

nappe complex

Upper Austroalpine

nappe complex

(Palaeozoic)

Upper Austroalpine

nappe complex

(Mesozoic)

Altkristallin Sheet

(Austroalpine basement)

MolasseFlysch &

Helvetic Zones

Upper Austroalpine

nappe complex

(Grauwackenzone)

Figure 1.4: Geological cross section through the Eastern Alps after Schmid et al. (2004a). No vertical exaggeration in

scale; see Fig.1.2 for transect position. Red circular symbols correspond to direction of strike–slip motion on fault systems

bounding Tauern Window.

Eclogite Zone, Tauern Window, Eastern Alps consistently yield an age of 31.5±0.7 Ma (Glodnyet al., 2005). Not only do these data imply rapid exhumation rates for East Alpine eclogites,but, more importantly, a maximum of ⇠6 Ma for attainment of Barrovian conditions. Previousmodels of East Alpine orogenesis, which form the basis of current understanding of heat trans-fer in overthickened crust (Oxburgh & Turcotte, 1974; Bickle et al., 1975; England & Thompson,1984), require an interval of ⇠30 Ma between the termination of subduction by continental colli-sion and attainment of peak temperatures. Furthermore, U–Pb, Lu–Hf and Sm–Nd isotope datafrom eclogitic parageneses of the Central (Brouwer et al., 2005; Hermann et al., 2006) and WesternAlps (Rubatto & Hermann, 2001) suggest that rapid attainment (<10 Ma post-subduction) ofBarrovian conditions occurred pene-contemporaneously along the Alpine chain.

This apparent paradox between heating rates calculated from thermal modeling of conductiveheat flow and those obtained by radiogenic isotope geochronology in the Eastern Alps forms thebasis of the dissertation. Either the Rb/Sr geochronology is pervasively a↵ected by isotopic dise-quilibrium or there is a profound misunderstanding of the mechanisms and processes which driveheating of orogenic crust to Barrovian conditions. These hypotheses can be tested by precise dat-ing of the Alpine eclogite facies metamorphism in the Eastern Alps using an isotopic system whoseparent and daughter nuclides are e↵ectively immobile to di↵usive transport at peak conditions.

1.4 Fieldwork

Field observations provide the motivation for much of this work. Field work in the Tauern Windowwas undertaken during summers 2007–2010—each year, for a period of 2–4 weeks. Studies focusedon the immediate area of the Eclogite Zone, which forms a ridge of extreme topography to thesouth of the Großvenediger massif. The area has been the focus of intense field-based study sincethe pioneering work of Cornelius et al. (1939). A network of Alpine Club huts over 3000 m.s.l.

6

provide accommodation within the field area.

1.5 Thesis structure

Chapter 2—introduces the regional geology of the Tauern Window, major fabrics and litho-tectonic units, in addition to presenting a synthesis of structural data.

Chapter 3—presents salient petrographical observations from the Pennine nappe stack.Bulk-composition specific phase diagrams and the average P–T method are applied to rep-resentative samples in order to constrain P–T paths for the Venediger, Eclogite Zone andGlockner nappes.

Chapter 4—focuses on determination of the age of HP metamorphism in the Eastern Alpsusing U-Th-Pb isotope geochronology.

Chapter 5—addresses the post-HP thermal history of the Penninic nappes using the40Ar/39Ar white-mica geochronometer. Apparent ages are compared with U-Th-Pb agesof HP metamorphism to quantify the extent of excess 40Ar contamination, which is used asa proxy for time-integrated metamorphic permeability.

Chapter 6—the thermal evolution of the Tauern Window is modeled using a one-dimensionalfinite-di↵erence solution to the heat equation. P–T–t data from the previous chapters providethermal constraints.

Chapter 7—implications of calculated P–T–t paths and thermal models for Alpine orogen-esis and exhumation mechanisms, followed by concluding remarks and suggestions for futurework.

Appendices—data repository, including: abstracts of published manuscripts, electron mi-croprobe dataset and U–Th–Pb and 40Ar/39Ar isotopic data.

7

Chapter 2

Geology of the Tauern Window

2.1 Introduction

This chapter builds on the introduction of Alpine geology presented in the preceding chapterby focusing on the geology of the Tauern Window. Following a brief discussion of the regionalgeological context with reference to previous studies, field observations are presented in a detaileddescription of each of the major litho-tectonic units exposed in the Window. This is followed by adescription of the region’s main fabrics and structures, with a specific emphasis on the relationshipbetween mineral growth and fabric development within the Pennine nappe pile. Data are compiledto form a working hypothesis for the tectonic evolution of the Tauern Window.

2.2 Geologic Setting

The Tauern Window is a tectonic and erosional window through the Austroalpine nappes, exposingthe Alpine orogen’s lower plate and constituent Penninic nappes, which are observed in a post-subduction geometry reflecting exhumation associated with closure of the Valaisan oceanic realm.The Tauern Window itself is the largest (⇠40⇥180 km) of several Penninic windows in the EasternAlps, including the Rechnitz and Engadine Windows (e.g. Oxburgh, 1968; Schmid et al., 2004a).During the Alpine collision, the Austroalpine nappes were thrust northward over the Penninicnappes of the Eastern Alps—pertaining to the leading edge of the European continent. Excellentexposure of the overthrust terrane led to the development of the classic model for understandingthe thermal evolution of orogenic belts—(Oxburgh & Turcotte, 1974; Bickle et al., 1975; England& Thompson, 1984).

Penninic rocks exposed in the Tauern Window form a series of distinct lithotectonic units, whichcollectively form a nappe-pile of 5–8 km structural thickness (Fig.2.1). The tectonostratigraphicallylowest unit comprises a tonalitic–granitic basement complex of Hercynian age (e.g. Petra et al.,2010) and a suite of parautochthonous (Frisch, 1980) metasedimentary cover nappes, collectively

8

termed the Venediger complex (Finger et al., 1993; Frisch et al., 1993; Kurz et al., 1998b). Struc-turally above the Venediger complex lies a series of basement slices, intercalated with Mesozoicpassive margin successions variably reworked during the Alpine orogenic cycle and now formingthe Storz and Rote Wand-Modereck (Rote Wand) nappes (Exner, 1971, 1980; Kurz et al., 1998b).Metasedimentary successions above the Zentralgneiss are termed the Inner Schieferhulle. The up-per portions of the Penninic pile comprise a dismembered ophiolite suite complete with remnantpillow structures and a voluminous calc-schist–greenstone succession—the Glockner nappe (Staub,1924), and a low-grade melange complex dominated by metamorphic flysch sediments—the MatreiZone (Frisch & Raab, 1987). These units constitute the Peripheral Schieferhulle—analagous tothe Bundnerschiefer of the Western Alps. Tectonically inserted between the Inner and Periph-eral Shieferhulle complexes, the Eclogite Zone comprises a thin (<3 km thick) slice of continentalmargin material with a diverse range of lithologies from metagabbroic bodies to layered quartzites(Miller, 1974; Frank, 1987); metasediments pertain to the European continent and have Meso-zoic depositional ages (Kurz et al., 1998b, and references therein). Collectively, the lithofacialassemblages of Penninic nappes exposed in the Tauern nappe-stack reflect progression from stablecratonic crust to rifting and formation of a neo-Tethyan oceanic basin.

Penninic unitsEuropean margin(U)HP complexes

Helvetic units

Austrolalpine unitSouthalpine unit

Oligocene granitoidsBasin deposits

Major thrust faultMajor normal faultMajor strike-slip fault

20 km

Peripheral Schieferhülle

Rote-Wand/Storz/Venediger cover nappes Eclogite Zone

Venediger nappe - crystalline basement

Alpine granitoid intrusions

Major antiformal axisAustroalpine units

SEMP

BF KF

PAL

DAV

Figure 2.1: Simplified tectonic sketch map of the Tauern Window showing both syb- and post-collisional faults, after

Glodny et al. (2008). Map area corresponds to white box in Fig.1.2. Notations: SEMP, Salzach-Ennstal-Mariazell-

Puchberg faults system; BF, Brenner Fault; KF, Katschberg fault; DAV, Defereggen-Antholz-Vals fault; PAL – Periadriatic

line. Dashed white rectangle shows position of field area—Fig.2.2

Rocks of the Peripheral Schieferhulle preserve a dominant petrographic signature of regionalgreenschist to amphibolite facies metamorphism—the Tauernkristallisation event (Sander, 1911).This metamorphism is interpreted as representing the thermal climax of conductive heating fol-

9

lowing emplacement of the Austroalpine nappes; conditions were in the range: 6–8 kbar and500–570�C (e.g. Dachs, 1990; Inger & Cli↵, 1997). However, within the Eclogite Zone and, to alesser extent, within the Glockner, Rote-Wand and Venediger nappes there is localised evidenceof early HP metamorphism. Peak HP conditions in the Eclogite Zone were close to ⇠25 kbarand 630�C Holland (1977, 1979), whereas in the overlying Rote-Wand and Glockner nappes relictblueschists (Holland & Ray, 1985) and eclogites (Dachs & Proyer, 2001) record burial to shallowerdepths, between 10–17 kbar and 350–570�C (Holland & Ray, 1985; Dachs & Proyer, 2001; Gleißneret al., 2007). The underlying Venediger nappe experienced maximum pressures of 10–13 kbar at550–630�C (Selverstone et al., 1984; Franz et al., 1991; Cesare et al., 2001; Smye et al., 2011). TheP–T evolution of Penninic units is discussed in more detail in chapter 3.

Structurally, the Tauern Window is dominated by an elongate antiformal core of Europeanbasement, above which the Schieferhulle nappes form a thin (<6 km) carapace beneath the over-thrust Austroalpine nappes. The fault system bounding the Tauern Window (Fig.2.1) comprisestwo major normal faults at the eastern (Katschberg normal fault; Genser & Neubauer, 1989)) andwestern margins (Brenner normal fault; Selverstone, 1988), a prominent sinistral strike-slip fault(Salzach fault; Behrmann & Frisch, 1990) along the western portion of the northern margin anda dextral strike-slip system (Pustertal fault of the Periadriatic Lineament; Schmid et al., 1989;Neubauer, 1995) in the eastern half of the Window (Genser & Neubauer, 1989; Ratschbacher et al.,1991b; Neubauer et al., 1999; Rosenberg et al., 2004; Glodny et al., 2008). During Alpine collision,the kinematic field of convergence in the Eastern Alps changed from west-northwest- to north-northeast-directed shortening, which facilitated continuous transpression between the Europeanand Adriatic plates (Ratschbacher, 1986; Ring et al., 1988; Glodny et al., 2008). Transpressivenorth-south convergence and east-west extension was partitioned into the prominent strike-slipsystems that formed both within the Tauern Window and also within the overlying Austroalpinenappes.

According to thermobarometric data obtained from Schieferhulle units, the Tauern Windowwas exhumed from depths of ⇠35 km since ⇠30 Ma at an average rate of ⇠1 mm.a�1, reaching upto 3.5 mm.a�1 in Miocene times Selverstone (1985); Blanckenburg et al. (1989); Zimmermann et al.(1994); Glodny et al. (2008). The formation and mechanism of uplift of the Tauern Window iscontroversial. Selverstone (1985, 1988) suggested that the Window was formed via isostatic uplift,driven by east-west extension, similar to metamorphic core complexes, whereas Ratschbacher et al.(1991b,a) focussed on the importance of continental-escape-driven uplift as a result of the oblique-transpressional tectonic regime. Lammerer & Weger (1998) noted that both these models requiretectonic thinning of the Austroalpine nappes on top of the Tauern Window; they also notedthat the strain distribution preserved in the Schieferhulle nappes did not contribute to uplift asductile north-south shortening was compensated by east-west extension. Rather, in accordancewith seismic data (e.g. Schmid et al., 2004b), the Tauern Window is interpreted as a hanging-wall

10

anticline structure forming as a result of a deep-seated duplex system.

2.3 Field Area

The south-central Tauern Window provides exceptional exposure of the Penninic nappes and over-lying Austroalpine thrust-sheet. This study focuses on the geology of the southern Venedigergruppeto the north of the Virgen river and west of the town of Matrei im Ostirrol. Figure 2.2 details thegeographic extent and salient geological features of the study area. Due to special preservation ofearly Alpine HP parageneses, the region has been the subject of numerous petrological and struc-tural studies—Cornelius et al. (1939); Schmidegg (1961); Abraham et al. (1974); Holland (1977);Frisch & Raab (1987); Inger & Cli↵ (1997); Glodny et al. (2005).

As shown by Fig.2.2, topographic relief within the field area is close to 2000 meters fromthe floor of the Virgental valley to the peaks of the Venedigergruppe. Consequently, the north-south-trending hanging-valleys of Dorfertal, Timmeltal and Frosnitztal, each provide across-striketransects through 5–8 kilometers of the Penninic nappe pile. Structural mapping and samplecollection were focused on these exposures within these valleys.

The field area measures ⇠120 km2 and is bounded by the Tauerntal in the west, the Dorfertal inthe west, the east-west-trending Virgental in the south and the watershed formed by the Venedigermassif, to the north (Fig.2.2). In general, there is greater than >50% bedrock exposure above thetree line, which occurs close to 2000 m.s.a.l.. Within the area, four main litho-tectonic units areexposed (Fig.2.3); from north to south these are: the Venediger complex, the Eclogite Zone, theRote-Wand and the Glockner nappes. The Rote-Wand nappe is tectonically thinned and di�cultto distinguish from the base of the Glockner nappe, whereas the Eclogite Zone is at its thickest atthe head of the Tauerntal and Frosnitztal valleys.

2.4 Litho-tectonic units

The following section describes the dominant lithologies, internal structures and metamorphicgrade of each of the Penninic nappes represented in the southern Venedigergruppe. Descriptionsare based on field observations and are supplemented by published data. The aim is not toprovide an exhaustive detail of petrolographical features, but rather, to provide a context in whichgeochronological (Chapters 4 and 5) and thermal modeling (Chapter 6) work is set.

2.4.1 Venediger nappe

2.4.1.1 Lithologies of the Venediger nappe

The Venediger nappe is exposed as a crystalline spine of peaks over 2800 m.a.s.l. toward thenorthern extent of the field area. The base of the Venediger nappe is formed by orthogneisses ofthe Zentralgneiss unit, which are overlaid by a 0.5–2 kilometere-thick succession of metasediments

11

Figure 2.2: Summary geological map and tectonic cross-section of the field area, southern flanks of the Venediger massif.

Contours and spot heights are in meters. Note the exaggerated horizontal scale used in the cross section. Foliations

are schematic; arrows show sense of vergence. Notation: BMH, Bonn-Matreier Hutte (47�2’20”N, 12�25’39”W); EH,

Eissee Hutte (47�3’10”N, 12�23’5”W); JH, Johanishutte (47�5’58”N, 12�33’45”W); BH, Badaner Hutte (47�4’51”N,

12�25’32”W).

pertaining to the Inner Schieferhulle. The descriptions that follow are based on observations fromthe Dorfertal, north of the Johanishutte where the cover nappes are ⇠2 kilometers thick andglacially polished bedrock o↵ers excellent exposure of the Zentralgneiss complex.

Throughout the studied outcrop, the Zentralgneiss is dominantly tonalitic in composition;phaneritic (<1cm) plagioclase (⇠50 modal%), quartz (⇠40 modal%), biotite (⇠10 modal%) and mi-nor K-feldspar constitute the original plutonic assemblage, which, at the outcrop-scale, shows littleevidence for Alpine reworking and is modally homogenous on lengthscales <400 meters (Fig.2.4a).

12

Figure 2.3: Panoramic photograph of field area taken looking west from Saulkopf (3209 m.a.s.l., 47�2’55”N, 12�25’8”W).

Note the extreme topographic relief and degrees of bedrock exposure. Sketch outline beneath details geological units;

colour scheme is as used in Fig.2.2, with the addition of the Matrei Zone (flesh colour); overlying Altrkristallin is purple.

However, close inspection and thin-section analysis reveals that sub-millimeter clots of garnet +clinozoisite + biotite + muscovite overprint plutonic plagioclase to form a localised domain assem-blage. The assemblage appears to have overgrown both massive portions of the tonalite body inaddition to apophyses and veins of aphanitic quartz + plagioclase, which regularly intrude the Zen-tralgniess. Metre–centimeter-scale mafic boudins and pods of amphibole + biotite are envelopedwithin the metatonalite; rim domains show evidence for hydration in the form of abundant chlorite.Toward the upper (southern) limit of the Zentralgneiss, K-feldspar becomes increasingly abundantand fingers of tonalite are observed intruding the overlying Schieferhulle.

The outcrop-scale chemical homogeneity observed in the Dorfertal is not characteristic of theZentralgneiss exposed elsewhere in the Tauern Window. Finger et al. (1993) showed that the Zen-tralgneiss varies between granite to grano-diorite to tonalite between the Ahorn, the Granatspitze,the Tux, the Zillertal and the Venediger batholiths.

Schmidegg (1961) proposed that the Zentralgneiss was of early Alpine age, resulting frommelt-migration as a result of the Barrovian event. However, more recent work has shown theZentralgneiss to be Hercynian in age—296–340 Ma Hawkesworth (1974a); Cli↵ & Cohen (1980);Eichhorn et al. (2000).

Within the Dorfertal section, the Zentralgneiss intrudes a sequence of metasediments (Fig.2.4b)and rare amphibolites, interpreted by Frisch & Raab (1987); Kurz et al. (1998b); Frisch et al. (1993)to represent a parautochthonous cover sequence to the Zentralgneiss. These cover units collectivelyform the Inner Schieferhulle within the field area.

Calcareous micaschists are the most common rock type occurring in the region adjacent to theZentralgneiss. Schists are tan to gray in colour and contain subsets of the assemblage: garnet +muscovite + biotite + plagioclase + dolomite + chlorite + quartz. Quartz and albite form sub-centimeter-scale porphyroclasts; muscovite is the most abundance mica and is often associated

13

withe the chrome-rich white mica, fuchsite. In the mid reaches of Dorfertal, a ⇠100 meter-thickhorizon of graphitic mica-schist is intercalated and interfolded with adjacent calcschists. This rockunit is typified by samples ASA–42a and b, both of which display the following mineralogy: garnet+ phengite + chlorite + epidote + quartz + calcite + graphite + biotite (± sulphide ± titanite).Garnets are <1 cm in diameter, notably euhedral and contain synkinematic quartz-inclusion trails.

Exposure in the southern reaches of the Dorfertal is poor. However, available outcrops showa distinct change in lithology from carbonate and muscovite-schists to more coherent biotite +K-feldspar-bearing schists and gneisses (Fig.2.4g). Gneisses are heavily foliated and contain 1–10centimeter-wide K-feldspar augen; matrix domains are formed from aggregates of quartz + plagio-clase + biotite. Muscovite is notably rare. Enveloped within the gneisses, which form pronouncedroche moutonees in the valley floor, are meter-scale horizons of biotite-schists. To the immediatenorthwest of the Johahnishutte, on the western flanks of the Dorfertal, pods of heavily-chloritisedgarnet amphibolite are wrapped by biotite schists. These pods are up to 10 meters in diameter andaligned with the south-dipping regional foliation. Gneisses within this region are often cross-cutby quartz + feldspar pegmatites, which can contain centimeter-scale books of muscovite.

The section of schistose lithologies which form the structurally deepest ⇠1 kilometer, imme-diately adjacent to the Zentralgneiss, is interpreted to correlate wth the Hochstegen and KasererGroups described by Kiessling (1992). Fossil evidence for the stratigraphic age of the Hochsetgenmember is given by a single occurrence of a perisphinctide ammonite and by assemblages of spongespicules and radiolarians, which suggest deposition in the Oxfordian stage of the Jurassic (155–161Ma). On paleontological grounds the Kaserer group is thought to have been dposeited in a clas-tic environment during the Cretaceous (Kurz et al., 1998b, and references therein). Augen-gneisslamellae, biotite schists and garnet amphibolites are interpreted to belong to the polymetamorphic-basement member of the Storz nappe (Exner, 1971, 1980).

Consequently, metamorphic assemblages preserved in the Zentralgneiss and schists of theHochstegen and Kaserer groups are domonstrably Alpine, whereas those preserved by the basementrocks of the Storz nappe could plausibly be pre-Alpine (Caledonian or Hercynian).

2.4.1.2 Structures of the Venediger nappe

In general, the intensity of strain increases with proximity to the thrust-sense contact between theInner Schieferhulle units and the Eclogite Zone.

The Zentralgneiss is largely underformed (Fig.2.4a); close to its intrusive contact with theoverlying sediments (Fig.2.4b), feldspar and biotite aggregates are aligned to form a weak, S>L(flattening), foliation which strikes towards the northeast (⇠030�) and dips between 80–90�; nolineation or isoclinal folding was observed. Mafic enclaves exhibit a weakly oblate geometry andtheir long axis are aligned with the strike of the foliation. Within the intrusive boundary zone,meter-scale blocks of quartzofeldspathic country rock, pertaining to the Inner Schieferhulle, are

14

Figure 2.4: Lithologies and structures of the Venediger nappe. a. Metatonalite of Zentralgniess (station AS–09–41).

Note weakly defined fabric and clots of pale-flesh coloured garnet in amongst plagioclase domains. Field of view is 4

centimeters; b. Intrusive contact between Zentralgneiss and host quartzofeldspathic gneiss of the Inner Schieferhulle.

Host rocks is migmatitic in places; c. Xenolith of country rock within Zentralgneiss. Note open-tight folding of quart-

zofeldspathic layers. Outcrop is < 10 meters from intrusive contact of Zentralgneiss displayed in b; d. Isoclinal folding

of biotite schist belonging to the upper-Inner Schieferhulle; folds are likely related to early deformation (D1

or older);

e. Quartz augen gneiss boulder ⇠400 m southwest of the Badaner Hutte, close to the thrust-sense boundary with the

Eclogite Zone. The matrix domain comprises mylonitised biotite + quartz; f. Quartz mylonite boulder from Eclogite

Zone-Venediger nappe boundary. Note string L2

stretching lineation; g. Biotite-albite gneiss of the Storz nappe (upper-

Inner Schieferhulle). Note that the primary compositional layering has been transposed in to parallelism with the regional,

south-dipping S2

fabric; h. Early (D1

) isoclinal folding of quartz seams within Inner Schieferhulle metasediments; i. C-S

fabric within K-feldspar bearing schist of the Storz nappe, directly beneath the Eclogite Zone boundary. Shear sense is

top-to-the-north—consistent with sense of rotation of refractory porphyroclasts.

15

entrained within fingers of metatonalite. Such xenoliths, record remnant protolith banding, in theform of quartz-biotite layers, which has been open to tightly folded (Fig.2.4c). Fold wavelengthsare less than 10 cm and interlimb angles are between 15 and 30�. This confirms that the InnerSchieferhulle experienced at least one phase of deformation prior to the Hercynian intrusion ofthe Zentralgneiss: the competency contrast between tonalite and xenolith precludes assessment ofwhether fabrics within the included blocks also a↵ected the metatonalite. Cli↵ et al. (1971) andBickle (1973) noted similar structural relations between the Zentralgneiss and Inner Schieferhullein the eastern Tauern Window.

The Inner Schieferhulle dip steeply (50–85�) to the south–southest. Inner Schieferhulle metased-iments are isoclinally folded on a range of di↵erent lengthscales Cli↵ et al. (1971). Figure2.4d showscentimeter-scale folding of quartz–biotite layers in a boulder from the Inner Schieferhulle; this isin contrast to meter-scale interfolding of amphibolite bodies (5–10 meters thick) and their hostquartzofeldspathic schists of the upper Inner Schieferhulle. Larger isoclines are most clearly ex-posed in the western flanks of the Dorfertal valley. Close to the boundary with the Eclogite Zoneaplitic veins are commonly isoclinally folded to form boudins and fold noses. The regional fabricis axial planar to isoclines (F1) and generally strikes 070–080� and dips by 40–80� to the south-southeast. The fabric (S2) is defined by alignment of sheet-silicates—predominantly muscoviteand biotite—and quartz layers in metasediments, and by the preferred alignment of amphibolesin mafic bodies. Associated stretching lineations (L2) defined by single amphibole needles andquartz seams generally strike to the southwest (220–250�), but, are also observed striking to thesoutheast (100–150�). K-feldspar augen in gneiss lamellae of the Storz nappe are aligned with L2.In schistose litholgies, such as sample ASA–42a, the S2-mica fabric is openly crenulated in theearly stages of S3 fabric development; it appears that synkinematic garnet inclusion trails reflectD3 rotation. In graphitic schists, crenulations are visible on the outcrop scale and have fold axeswhich consistently trend 080–110� with plunges between 30 and 50�.

Outcrops of basement gneisses and quartzofeldspathic schists in the Johanishutte region arelocally mylonitic (Fig.2.4e and f). Refractory quartz clasts form �-type porphyroclasts with atop-to-the-north sense of vergence; C-S shear bands in augen gneiss also show a top-to-the-northshear sense (Fig.2.4i). Importantly, both shear indicators and lineations are commonly defined bysecondary, Barrovian-style mineral assemblages, such as hornblende and muscovite in calc-schistsand chlorite and Ca-amphibole within mafic bodies.

2.4.2 Eclogite Zone

2.4.2.1 Lithologies of the Eclogite Zone

The Eclogite Zone comprises a ⇠1–2 kilometer-thick sequence of tightly interfolded mafic andsedimentary lithologies, interpreted to represent a volcano-sedimentary succession of a distal con-tinental slope (Frisch & Raab, 1987; Kurz et al., 1998b). The nappe tapers out completely tothe west of the Dorfertal valley and is separated from the underlying Venediger nappe by a zone

16

of mylonitic shear with top-to-the-north vergence; a band of mylonitised pillow basalts formsthe boundary with the overlying Glockner nappe. Metasediments are more common than maficprotoliths, occupying ⇠60% of the nappe. All lithologies record a polymetamorphic evolution char-acterised by early eclogite facies mineral growth, which is variably overprinted by blueschist andgreenschist–amphibolite facies metamorphism. The Barrovian overprint is strongest in the nappe’supper-structural levels which are exposed in the south.

Mafic eclogites are concentrated at the base of the lithostratigraphic sequence. They formdekameter- to hektometer-scale boudins and lenses (Fig.2.5), around which the metasedimentarymatrix anastomoses. The wide degree of compositional heterogeneity to eclogite protoliths is re-flected by the fact that at least five di↵erent parageneses are observed (Table 2.1). Metagabbroiceclogites (Fig.2.5a) are characterised by the assemblage: garnet + omphacite + kyanite + talc +quartz + epidote + barroisite + phengite, and are typified by boulders to the northeast of the Eis-see (ASA–08–18). Garnet forms euhedral, centimeter-scale blasts which are pyrope-rich; omphaciteoccurs as similarly coarse, apple-green smears, which formed after original augite. Hock & Miller(1980) showed that metagabbroic eclogites are chemically similar to oceanic tholeiites. Mylonitisedeclogites (Fig.2.5b) are common close to the Eclogite Zone’s boundaries and are defined by a ma-trix of fine-grained (<1 mm) omphacite which hosts euhedral garnet blasts. Omphacite is stronglylineated (L1) and shows evidence for dynamic grain boundary recrystallisation. Dolomite or mag-neisite are also observed as refractory porphyroblasts. Banded eclogites (Fig.2.5c) are the mostcommon sub-division of eclogite; millimeter- to centimeter-scale barroisite-rich and garnet-rich lay-ers give the eclogite a streaky appearance. Horizons of dolomite and zoisite weather preferentiallyand are easily recognizable by their orange colouration. Banded eclogites are often exhibit chevron-style folding, particularly in the region to the immediate south of the Eissee, in the Timmeltal.Zoisite eclogites have a medium grained omphacite-garnet matrix and often house dolomite-richlayers. However, they are distinguishable from banded eclogites in that they contain mats of zoisiteneedles, which are best seen on S1(+S2) foliation planes. Needles are ⇠0.5–2 millimeters in lengthand regularly define a lineation; white paragonite mica is closely associated with the zoisite ag-gregates. Boudin necks regularly display quartz-filled voids, within which megacrystic omphacite,kyanite, garnet, dolomite and rutile grow. Sample ASA–08–30c (Fig.2.5d) from the upper Frosnitz-tal shows laths of kyanite, up to 15 centimeters long, spanning one such void which contains lathsof zoistie and pennies of rutile. These voids are interpreted as representing collections of locallyderived Si- and Al-rich fluid, which formed close to the peak of HP metamorphism. All types ofeclogite display evidence for retrogressive amphibolite and chlorite growth, which primarily occursaround discordant quartz veins and boudin boundaries.

There are four principal types of metasediments exposed within the Eclogite Zone. Horizons ofquartzite (0.5–3 meters) form prominent features in the local topography due to their resistanceto weathering. Quartzites are impure (Fig.2.5g): lamellae of minor garnet, green phengite anddolomite give the units a dusty appearance and define the limbs of isoclinal folds. Metapelites(Fig.2.5h) are ubiquitous within the south-central portion of the Eclogite Zone, where they form

17

Type ParagenesesMetagabbroic eclogite omph + gt + ky + tc + + qtz + ep + barr + paEclogite mylonites omp + gt + dol\magn + pa + pheBanded eclogites omph + gt + tc + ky + dol + rt + qtz + paZoisite eclogites omph + ky + zo + ky + pa + qtz ± rtEclogite mobilisates qtz + gt + omph + dol + zo + ky ± rt

Table 2.1: Summary of mineral parageneses for eclogites of the Eclogite Zone.

a ⇠300 meter-wide band, which grades into adjacent quartzites. Metapelites are characterised bysubsets of the following eclogite-facies assemblage: garnet + chloritoid + kyanite + phengite +clinozoisite ± rutile ± pyrite ± graphite. Matrix domains are formed by alternating microlithonsof phengite and quartz; kyanite shards (<1 cm) are aligned with phengite-layers and displayretrogressive textures at grain boundaries. Garnet is poikiloblastic and contains rutile needles,quartz trails and lozenges of allanitic epidote. Blasts are commonly greater than 1 centimeter indiameter. Station ASA–08-36, on the southern shore of Eissee, shows honeycombe-syle quartz-garnet intergrowths which are rimmed by dark-green chloritoid. More commonly, chloritoid formsblack tablets (<1 cm), which are aligned with the micaceous foliation. Both garnet and chloritoidare concentrated on the boundaries between pelitic and quartzitic subdomains. In the Zopet Spitzeregion, metapelites are less garnetiferous and contain abundant platy chloritoid.

Calc-schists, defined here as rocks with subordinate quartz relative to calcium carbonate, arethe dominant metasedimentary lithology within the Eclogite Zone and house the majority of maficeclogite bodies. Schists comprise subsets of the following parageneses: zoisite + choritoid + garnet+ dolomite + white mica + tourmaline + quartz ± sphene ± tremolite ± kyanite. Calc-schistsare intercalated with quartz-mica schists and garnet-chloritoid-kyanite metapelites; rarely, theyform gradational boundaries with banded eclogite bodies of tu↵aceous composition (ASA–08–36).Garnet is often black in colour, indicative of a high grossular content and the presence of graphiteinclusions, and forms poikiloblasts 0.5–2 cm in diameter. Brown dolomite/calcite forms centimeter-scale refractory blasts around which the micaceous matrix flows. As with the metapelites, thecalc-schists show abundant evidence for tectonic shu✏ing and retrogressive mineral growth in theform of sparry zoisite and amphibole.

Yellow to white marbles are intercalated with quartzites, metapelites and calc-schists on centimeter-to dekameter-scales. Marbles outcrop extensively within the southern portion of the Eclogite Zone;in particular, a 10–20 meter-thick band of zoisite marble is exposed close to the southern limit of theEclogite Zone (Fig.2.5i), where the unit can be traced west from the Zopet subarea to the westernflank of the Dorfertal valley where it is <5 meters wide. In addition to calcite and dolomite, marblebands contain: tremolite, zoisite, white mica, diopside, epidote, talc and subordinate sphene, pyriteand chlorite. The southern boundary of Frosnitzkees glacier, close to the Weißspitze, exposes ameter-thick band of marble which can be traced ⇠2 kilometers along strike. Shards of zoisite (0.5–3cm) and tremolite (0.2–2 cm) form a distinctive mesh texture on foliation planes. The occurrence

18

of marble bands close to the hanging and footwall boundaries of the Eclogite Zone was interpretedby Holland (1977) as evidence that the Eclogite Zone was dominated by a synformal structure.

The cofacial nature of the HP assemblages and common retrogressive textures in both metased-iments and senso stricto eclogites strongly suggests that the Eclogite Zone behaved as an internallycoherent nappe during subduction and exhumation.

2.4.2.2 Structures of the Eclogite Zone

Holland (1977) reported bent micas and folded trails of rutile within poikiloblasts of garnet fromthe Eclogite Zone. This is augmented by samples ASA–09–36b and 36i, both garnet, chloritoid,kyanite bearing metapelites from the Eissee region of Timmeltal, which show discordant rutile trailsand openly folded clinozoisite needles in the cores of eclogite-facies garnet. Raith et al. (1980) andKurz et al. (1998a) document pseudomorphs after lawsonite and relict glaucophane in the coresof HP garnet. Collectively, these data show that a pre-HP deformation event (D0) a↵ected themargin-sequence of the Eclogite Zone, most likely during subduction.

Mafic eclogites show a demonstrably HP foliation (S1) defined by the alignment of matrixomphacite grains. In gabbroic eclogites, eclogite facies omphacite is elongated to form a shape-peferred lineation fabric; this suggests that D1 was syn-metamorphic. However, both S1 and L1

fabrics are di�cult to reconstruct due to the often intense character of retrogressive amphiboli-tization and chloritization. Where present, S1 is transposed sub-parallel with the pervasive S2

foliation. L1 is defined in finer grained eclogites by omphacite, glaucophane and garnet lineations.Perhaps the clearest example of L1 is observed at station ASA–09-83, on the western shore of Eis-see; here, garnet is stretched to form a trail which strikes 200� and dips 45� to the south-southwest.Throughout the Eclogite Zone, L1 strikes between 160� and 240�. Metasediments rarely record L1

due to Barrovian ductile recrystallisation. However, close to Zopet Scharte (ASA–09–67), mus-

Figure 2.5 (following page): Lithologies and structures of the Eclogite Zone. a. Metagabbroic eclogite; note cm-scale

omphacite in a garnet matrix; long axes of omphacite is subparallel with S2

; b. Mylonitic eclogite, Darker layers are rich in

amphibole, whereas lighter layers are garnet + zoisite. Note the millimeter-scale of garnets; c. Crenulated banded eclogite

from station ASA–08–36. Orange horizons are dolomite-rich; fold are F2

generation; d. Sample ASA–08–30c—an eclogite

facies mobilisate from the Weißspitze region. Kyanite, zoisite and rutile form the mobilisate assemblage; boundary regions

are heavily chloritised; e. Fold interference in a banded eclogite. Refolding of an F2

isocline by an open D2

fold; f. Eclogite

boudin in calc-schist matrix on path between Eissee and Weißspitze, upper Timmeltal. Note deflection of the composite

S2

+S1

fabric around the eclogite pod, within which the early omphacite fabric (S1

) is preserved; g. Graphitic quartzite

from the Dorfertal. Fine grained graphite and muscovite give the quartzite a dusty appearancel note the sugary texture

of matrix quartz; h. Crenulated garnet-chloritoid-kyanite bearing schist from Eissee region. Note the spatial association

between garnet poikiloblasts and pelitic horizons; folds are F2

generation and show axial planar muscovite growth (S2

); i.

Marble band marking the hanging-wall contact between the Eclogite Zone and Glockner nappe. Photograph taken looking

west towards Zopet Scharte; j. Discordant greenschist-facies vein assemblage developed in calc-schist of the Eclogite

Zone. Vein assemblage comprises: actinolite + albite + chlorite + magnetite; k. D1

isoclinal fold in calc-schist; note

that the fold axis has been transposed sub-parallel with S2

.

19

20

covite, in a calc-schist horizon, is aligned to form a lineation of 268�\78, which is discordant toL2 and sub-parallel to cogenetic F1 isoclinal folds from the same outcrop. D1 isoclinal folding isoften observed in quartzite and marble bands (Fig.2.5k). Interlimb angles are commonly <10�,and fold limbs are parallel to S2. Fold axes strike variably between 070–270� with plunge valuesin the range: 30–90�. Such a large range in fold axis geometries reflects the e↵ect of D2 and D3

cross-folding. F1 folds in banded eclogite bodies are manifest as open crenulations to chevron-stylefolding dependent on the rheology contrasts between layers.

Figure 2.6: D1

–D2

fabric relations at station ASA–09–45. a. Early F1

isocline is refolded by sub-parallel D2

open fold.

b. Resultant dome-and-basin interference pattern viewed on S2

foliation plane. Inset shows equal area, lower-hemisphere

projections of respective fold axes—F1

, blue; F2

, red.

The regionally pervasive D2 event resulted in rotation of the HP S1 fabric to form a compositefoliation: S1+S2. This was accompanied by axial planar foliation development and cross-foldingof F1 in complex interference patterns. In metapelites, S2 is defined most clearly by crenulationof S1. Close to the southern extent of the Eclogite Zone in Dorfertal (ASA–08–14), graphiticschists display abundant centimeter-scale crenulations; S2 is axial planar to the fold axis andtrends 041�, dipping 79� to the southeast. D2 refolding of F1 isoclines creates composite foldinterference patterns. Station ASA–09-45, south of Eissee shows isoclinal folds in a quartzite–mica-schist succession, which are folded subparallel to S2 by an open F2 fold; this creates a dome-and-basin interference pattern (Fig.2.6, Twiss & Moores, 1992). Eclogites that are retrogressedduring D2 are characterized by an intense mylonitic foliation (S1 +S2) which is defined by thealignment of green actinolite, biotite, epidote, muscovite and chlorite. South of Eissee, ⇠100 metersnorth of the Eclogite Zone–Glockner nappe contact, banded eclogites are heavily retrogressed and

21

mylonitised (ASA–09–79). The composite foliation strikes 031� and dips 59� to the southeast;these measurements typify the geometry of S2. An associated lineation is often present (L2) andis defined by the alignment of epidote and amphibole. L2 generally strikes between 100–220� anddips up to 70�. Macroscopic F2 folding occurs in the cli↵s to the east of Zopet Spitze; here, eclogiteboudins and host calc-schists are openly folded without an obvious sense of asymmmetry. F2 foldsstrike broadly east-west. Sense of shear indicators document top-to-the-north shear towards thebase of the Eclogite Zone; shear sense is variable at the hanging-wall contact with �-style quartzand dolomite porphyroblasts indicating both top-to-the-east and top-to-the-west senses of vergence(Fig.2.7). This implies a component of transpressive motion during emplacement of the EclogiteZone into the Schieferhulle pile. D2 is synchronous with the development of amphibolite–greenschistfacies hydrous vein assemblages which cut all lithologies of the Eclogite Zone.

3500

3000

2500

Elev

atio

n (m

.a.s

.l.)

S N

Glockner-Eclogite Zone boundary Eclogite Zone-Venediger boundary

Figure 2.7: Observed M2

stretching lineations at the hanging-wall and footwall contacts of the Eclogite Zone. The

footwall contact is characterised by thrust-sense, top-to-the-NNE deformation in carbonate-metapelitic mylonites of the

basal Eclogite Zone and upper Venediger nappe (see text). Kinematics of the hanging wall contact of the Eclogite Zone

are more complex: the tectonic boundary is marked by a steeply south-dipping mylonite belt with dominantly south-west

plunging M2

stretching lineations, although south-east plunging muscovite and hornblende lineations are also present.

The final episode of deformation (D3) to a↵ect the Eclogite Zone was minor in comparison toD1 and D2. Bull quartz veins between 1 and 10 centimeters in diameter cut the entire tectonos-tratigraphic succession. In the Dorfertal, such veins are > 10 meters in length, but are generallyshorter elswhere in the Eclogite Zone. The veins strike at high-angles to the S2 foliation, com-monly in a north–south geometry. They are interpreted as resulting from late-stage Si-rich fluids

22

percolated into brittle tension gashes.

2.4.3 Rote-Wande nappe

Outside of the field area, the Rote-Wande nappe separates the Eclogite Zone from the Glocknernappe and comprises basement gneiss lamellae with a quartzite–calc-schist cover sequence (Kurzet al., 1998b, and references therein). However, the unit is tectonically thinned such that it is onlyrecognizable in the Dorfertal, to the south of the Ochnershutte. Here, a ⇠40 meter-thick successionof calc- and albite-schist is distinguishable from Eclogite Zone and Glockner nappe lithologiesdue to its heavily annealed appearance, indicative of a high strain history prior to Barrovianmetamorphism. Albite, calcite and quartz are fine-grained (sub-millimeter); white mica definesa schistosity which is concordant with S2: 069�\80� (right-hand convention). Platy muscovite isweakly aligned to form a lineation, which is associated with top-to-the-north asymmetry of quartzclasts—the lineation strikes 164� and dips at 70�.

2.4.4 Glockner nappe

The Glockner nappe forms lies structurally above the Eclogite Zone and comprises a 2–5 kilometer-thick succession of meta-ultramafic and metasedimentary rocks of oceanic a�nity.

2.4.4.1 Lithologies of the Glockner nappe

The base of the Glockner nappe is dominated by slices of former oceanic basement rocks. Immedi-ately above the boundary with the Eclogite Zone, a unit of strongly-lineated epidote-amphiboliteoutcrops (Fig.2.8a, b) and can be traced along strike from close to Steinsteg in the Frosnitztal,to the western flanks of the Dorfertal. The unit comprises the following amphibolite facies as-semblage: epidote + garnet + actinolite + albite + chlorite + muscovite + dolomite + quartz.Garnet, where present, is small (<2 mm), and deep red in colour; the first occurrence of garnet inthe Timmeltal occurs ⇠300 m south of the Eisseehutte. Actinolite needles are commonly alignedto form L2. The unit shows characteristic pillow structures (Fig.2.8a, b), as described by Holland& Norris (1979), whose long-axes are sub-paralled with L2. The rims to the pillows are formed bya rind of epidote, giving them a pale yellow-green colour. Pillow matrices are formed from epidote,calcite, actinolite, chlorite, albite, biotite and rare garnet. Throughout the epidote amphiboliteband, veins of bull quartz and quartz-dolomite cross-cut the regional fabric.

Serpentinite bodies are intercalated with the epidote amphibolite band and horizons of calc-schist. They form an array which can be traced, in the same structural position, from Frosnitztalto west of the Dorfertal. In the Dorfertal, serpentinite is economically mined at Steinbruch. Withincore regions of such bodies there is little evidence for preservation of original minerals; chrysotileand antigorite are ubiquitous (Fig.2.8c). In the Timmeltal, a serpentinite body ⇠80 meters wideoccurs in the valley floor, close to the Wurgenegge. Here, laths of zoisite, actinolite and occasional

23

octahedra of magnetite occur within the boundary between host calc-schists and the serpentiniteitself.

Calcareous mica schists outcrop at all structural levels within the Glockner nappe (Fig.2.8d).Their mineralogy is dominated by calcite, dolomite, white mica, graphite, albite, quartz, zoisite andchlorite. Calc-schists are intercalated with the metamafic rocks exposed at the base of the nappeand increase in abundance towards the southern extent of the Glockner nappe where calc-schistsabut against the Matrei Zone. On the western flanks of Timmeltal (ASA–09–76), calc-schistsare interbedded with both prasinites and quartz-rich schists. Here, the recessive weathering ofcalcareous dominated rocks forms clear foliation planes in the quartz-rich lithologies. Occasionally,marbles bands are exposed within calc-schist bodies; one such example is in the Dorfertal, closeto Steinbruch, where a marble band displays the rodingite assemblage: grossular + diopside +actinolite + epidote + chlorite + albite.

A 500 meter-thick unit of epidote greenstone lies structurally above, and is intercalated with,calc-schists of the middle Glockner nappe. The unit spans the entire length of the Glocknernappe exposed in the field area and comprises a prasinite matrix of epidote + albite + chlorite +actinolite + quartz ± biotite ± quartz. At the base of the unit, schists host abundant rectangularaggregates (0.2–1 cm) of albite + epidote + calcite + white mica + chlorite, which are interpretedas pseudomorphs after lawsonite (Fig.2.8e; Fry, 1973). Pseudomorphs are best preserved in thevicinity of the Bonn Matreier Hutte, where they are interlayered with pseudomorph-absent epidote-and calc schists. Structurally above the psuedomorph unit, a sequence of epidote- and albite schistsoccur. Epidote schists are characterised by centimeter-scale knots of epidote and magnetite in aprasinite matrix which contains oblate lenses of quartz and calcite. Epidote-rich horizons areoften isoclinally folded. To the immediate east of the Bonn Matreier Hutte, epidote-schists hostcurious spongy grey clots with rare cores of jadeite–acmite pyroxene, in addition to sub-centimeterblue-grey needles of retrogressed glaucophane on mats of paragonite. Together with lawsonitepseudomorphs, these assemblages are clear evidence that the Glockner nappe experienced early-HP metamorphism prior to decompression and prasinite development (Holland & Ray, 1985).Specific field relations pertaining to these rocks are discussed in section 3.4.2. Spotted albiteschists outcrop throughout the prasinite unit. The schists lack early HP mineral growth andare characterised by their abundant millimetric albite clasts, around which the greenschist fabricanostamoses.

2.4.4.2 Structures of the Glockner nappe

The structural history of the Glockner nappe is clearly preserved in outcrops of the epidote green-stone unit close to the Bonn Matreier Hutte. Due to their high competency, lawsonite-pseudomorphschists record early fabric relations, whilst weaker epidote schists are more thoroughly equilibratedduring D2. The earliest recognisable fabric is defined by epidote banding; these layers are a↵ectedby all subsequent deformation episodes and likely reflect primary compositional heterogeneity.

24

Figure 2.8: Lithologies and structures of the Glockner nappe. a. Strongly lineated (L1

) relict pillow basalts exposed

at the base of the Glockner nappe (ASA–08–35b). Note the epidote-rich rind and prasinitic core domains; b. Lineated

pillow basalts ⇠10 meters from the Eclogite Zone contact. Note the prolate geometry; c. Chrysotile vein in serpentinite

of the upper Timmeltal; d. Highly strained calc-schist of the mid-Glockner nappe. Refractory quartz clasts exhibit a

top-to-the-northeast sense of vergence; e. Lawsonite-pseudomorph schist from the vicinity of the Bonn Matreier Hutte.

Note sub-millimeter-scale albite clots in the matrix; f. Macroscopic F2

fold exposed in the western walls of the Dorfertal.

White lithology is calc-schist, whereas the surrounding cli↵s comprise prasinite. Note that F2

is sub-parallel to S2

; g. F2

crenulation in S1

fabric defined by aligned epidote bands; h. Eclogite Zone–Glockner nappe boundary looking east from

Zopet Scharte. Note Eissee in the middleground; i. Refolding of an F1

isoclinal fold by open F2

fold in epidote banded

greenstone unit; j. Macroscopic F3

buckle-type fold within calc-schists of the eastern Timmeltal.

25

Pseudomorphs of lawsonite, blue amphibole and jadeite all overgrow the banding, which is as-signed to D0. During D1, an L>S fabric developed through alignment of blue amphibole needlesand transposition of muscovite–epidote layering. After rotation, S1 is now sub-parallel with theregional S2 fabric; L1 generally dips to the east at between 20–50�, however, its orientation isdependent on the F2 folding. F1 isoclinal folds are abundant throughout the Glockner nappe, butparticularly so in horizons of epidote- and calc-schist (Fig.2.8i). Fold axes are sub-parallel withthe amphibole stretching lineation. The long-axes of relict pillow basalts exposed at the base ofthe Glockner nappe are aligned with L1 (Fig.2.8a,b), which dips at ⇠30� to the southwest in theregion close to the Eisseehutte. Holland & Norris (1979) recognised that the deformed pillows havebeen subjected to a large constrictional strain during D1 in order that their geometries are stillrecognisable in sections perpendicular to L1.

Station ASA–10–19 is located at the base of the Saulkopf, to the immediate north of the BonnMatreier Hutte. Here, lawsonite-pseudomorph schists are heavily crenulated: S1 is defined bythe alignment of millimeter-scale albite clasts and is folded into open-crenulations which exhibitaxial planar formation of S2. Critically, lawsonite pseudomorphs are wrapped by S1+S0, implyingpeak pressures were attained prior to or contemporaneous with D1 (Fig.2.9). Although rotated,S1 strikes ⇠300� and dips at ⇠60� to the north; S2 is defined by alignment of prasinite mineralsand strikes ⇠058� with dips between 75–85�to the south. F2 crenulations have axes which trendtowards the southwest and dip at moderate (⇠40�) angles (Fig.2.8f, g).

F2 F2

S 1+S 0

Figure 2.9: Sketch depicting fabric relations observed at station ASA–10–19. Lawsonite pseudomorph (white rim is

albite-rich; green core is prasinitic) is wrapped by a composite early fabric (S1

+S0

; blue dashed line), which is defined by

phengite and albite clasts, and is crenulated during D2

to form open F2

crenulations (fold axes: red dashed line). Prasinite

minerals within the pseudomorph are aligned with S2

. This suggests that lawsonite was stable during or slightly before

D1

.

26

S2 forms pervasively throughout the Glockner nappe and is concordant with the same generationfabric within the Eclogite Zone and Venediger nappes. L2 is less well developed than L1 and isdefined by preferential alignment of muscovite and actinolitic hornblende in calc- and epidote-schists; the lineation dips at shallow angles, commonly less than 20�, to the east. Shear indicatorsdo not indicate a consistent sense of vergence during D2. �-type quartz and albite porphyroblastsshow both top-to-the-northeast (Fig.2.8d) and top-to-the-southwest asymmetry, indicating that acomponent of transpressive motion occurred during D2 (Fig.2.7).

D3 is locally developed in the form of cross-folding of S1+S2 . Axial planes to F3 buckle-stylefolds commonly strike south (150–180�) and have large interlimb angles (⇠100–150�). In the cli↵sto east of the Eissee Hutte in upper Timmeltal, macroscopic F3 buckles are visible in calc-schistbands (Fig.2.8j). Fold axes strike towards the northeast at gentle angles (⇠20�) and display atop-to-the-north sense of vergence.

2.4.5 Matrei Zone

The Matrei Zone forms the hanging wall to the Glockner nappe and is characterized by cal-careous flysch sediments of low metamorphic grade. The lower boundary of the Matrei Zone isdefined by Frisch & Raab (1987) and Kurz et al. (1998b) as the structurally lowest occurrenceof Austroalpine-derived metasediment. However, field observations from the southern flanks ofthe Virgental show that there is a continuous metasedimentary transition from prasinites of theGlockner nappe to low-grade metasediments of the Matrei Zone. The most abundant rock-type iscalc-schist (Fig.2.10a–d), which is characterised by the following assemblage: calcite + quartz +muscovite + chlorite + epidote ± biotite ± black amphibole ± magnetite ± pyrite. Calcite-freemica-schists are intercalated with calc-schists (Fig.2.10a).

At least two episodes of deformation are recognizable within rocks of the Matrei Zone. Sheetsilicate minerals and quartz\calcite layering are concordant with the regional trend of S2 (Fig.2.10a,b). In the Rain Alm, foliations consistently strike between 072� and 100� for ⇠1500 meters ofstratigraphic thickness. Outcrops perpendicular to the strike of S2 reveal early isoclinally foldedquartz seams (Fig.2.10a, c); fold limbs are now parallel with S2 and fold axes plunge at shallowdips (<30�) to the east and west, depending on the presence of D3 cross-folding. Close to thecontact with the overlying Altkristallin thrust sheet, a strong southwest-west trending lineation isexposed—the lineation is defined by alignment of chlorite seams and, where present, needles of darkamphibole. This is interpreted as correlating with L2 of the Glockner nappe. Crenulated quartz–calcite layering (F2) often displays fold axes which are sub-parallel to the lineation (trend:270–320�;55–80� plunge). Centimeter-scale clasts of refractory calcite and quartz show a consistent top-to-the-north northeast sense of vergence. With proximity to the base of the Altrkristallin, quartzand calcite grain size decreases to sub-millimeter diameters and the schistosity develops a markedplaty appearance. C-S shear bands (C3) occur in mica-rich portions of the schists (Fig.2.10a,d);rarely do the bands excess ⇠3 centimeter in diameter and their sense of vergence is consistent with

27

top-to-the-north asymmetry. Several generations of discordant quartz veins cross-cut S2 and areubiquitous throughout the upper Matrei Zone.

2.4.6 Altkristallin

The Altkristallin complex comprises pre-Mesozoic (e.g. Siegesmund et al., 2007) basement rocks andis the largest of the allochthonous Austroalpine thrust sheets, extending ⇠500 kilometers east-west,and up to 50 kilometers in width. Within the field area, the Altkristallin directly overlies calcareousflysch of the Matrei Zone; its base is best exposed in the the vicinity of the Lasorling Group, to thesouth of the Virgen river. Here, garnet-amphibolite gneisses are interlayered with mafic enclaves,biotite gneisses and garnet-mica schists. Amphibolites display gneissic layering (Fig.2.10e,f), whichis defined by alternating proportions of amphibole + biotite and plagioclase + quartz in the as-semblage: plagioclase + quartz + garnet + white mica + biotite + amphibole. Garnet is generallysmaller than 3 millimeters in diameter and shows evidence for rotation—occasionally, garnets aresurrounded by coronas of dark brown amphibole. There are two populations of amphiboles: matrixgrains, aligned with the mylonite fabric, are tablet-shaped (3–4 mm) and dark-brown to green incolour, whereas porphyroblasts are black and less intensely deformed. Mafic enclaves are between30 centimeters and 10 meters in diameter and contain abundant amphibole and garnet, both ofwhich are much larger than observed in host amphibolites (> 1 cm). Approximately 600 meterssouth from the Altkristallin–Matrei Zone contact, close to Pragrat Torl amphibolites grade into aheavily schistose unit of silver garnet mica schists. Garnet is sub-centimeter size and appears darkgrey–black; the schists are rich in phengite and biotite, and are heavily foliated.

Altkristallin gneisses record intense concentrations of strain—all rocks are classified as L>Smylonites. Gneissic layering is concordant with regional S1+S2 and strikes broadly east-northeast–south-southwest (055–073�); dips are close to vertical. Long axes of amphiboles, quartz and pla-gioclase seams, in addition to the alignment of sheet silicate minerals, define a strong lineation(Fig.2.10g) which generally dips between 20–40� to the west. Pre-D2 isoclinal folding of quartzveins and gneissic layering developed fold axes which have been transposed sub-parallel to the my-lonitic lineation (L2; Fig.2.10g,h). Tablet amphiboles have been folded during this event, whereasporphyroblastic amphibole + biotite clots are aligned with S2. This shows that the gneissic layeringwas present prior to D2 and is likely to be Hercynian in age. Waters (1976) and Hoke (1990) claimthat at least three pre-Alpine episodes of deformation are recognizable in the Altkristallin. Thelimbs of isoclinal folds are a↵ected by subsequent open-style folding (F2 of the Schieferhulle)—foldaxes trend towards the southwest (220–240�) and dip between 15–30�. �-style quartz blasts showa top-to-the-north sense of vergence, which is consistent with the sense of shear observed in theunderlying flysch of the Matrei Zone.

28

Figure 2.10: Lithologies and structures of the Matrei Zone (a–d) and the Altkristallin (e–h). a. Calcareous flysh of the

Matrei Zone exposed in the southern reaches of the Lasorling Valley. Note unconsolidated nature of metasediments and the

distinctive green colour defined by chlorite + muscovite. S2

is prevasive; F1

isoclines are transposed sub-parallel to S2

; late

kink bands (C3

) cross-cut the section; b. Close-up of a.; note top-to-the-north shear sense; field of view is 30 centimeters

wide; c. F1

isoclinal folding of an early (S0

) quartz seam in calcareous matrix schist; d. C3

shear band in mica-schist; e.

Garnet amphibolite of the Altkristallin basement complex. Leucocratic layers are plagioclase-rich, whereas melanocratic

layers are biotite- and amphibole-rich. Note that matrix amphiboles are aligned with strong lineation; similarly note the

small grain size as a result of dynamic recrystallisation; f. Outcrop-scale view of e.; g. L2

fabric in garnet amphibolite;

view is of an S1

+ S2

foliation plane. Here, amphibole is heavily chloritised, hence the green colour; h. Early (D0

or

before) isoclinal folding of original quartz band in biotite gneiss of the Altkristallin. Fold axis is sub-parallel with L2

and

axial plan is parallel with regional S1

+ S2

fabric.

29

2.5 Fabric correlation across units

Before proceeding with investigations concerning the absolute timing and physical conditions ofmetamorphic events in the Tauern Window, it is necessary to establish a correlated structuralevolution of the Tauern nappe pile. Herein the following convention is used to describe discretemetamorphic events: M1 = eclogite facies; M2 = blueschist facies; M3 Barrovian. The followingstructural evolution is illustrated in Fig.7.6.

Relicts of pre-Alpine deformation (<D0) are preserved in the lowermost cover of the Venedigernappe, in to which the Zentralgneiss intrudes, in addition to gneissic banding within Altkristallinbasement lithologies. Both of these structural elements likely pertain to Variscan deformationevents (Waters, 1976; Hoke, 1990; Siegesmund et al., 2007); in the case of the Inner Schieferhulle,isoclinal folding of wall-rock sediments could plausibly have been caused by intrusion of the Zen-tralgneiss plutonic suite. The pervasive nature of Alpine deformation and the large degrees ofconvergence precludes further correlation between pre-Alpine deformation.

Inclusion trails in garnet from eclogites of the Eclogite Zone and a layered-epidote fabric inGlockner-nappe blueschists show that the first extensive Alpine deformation event occurred sub-sequently to a weak foliation (D0), which was almost entirely obliterated by Alpine deformationand likely developed during the early stages of subduction. This is consistent with preservation ofaligned gas vesicles in meta-pillow basalts of the Eclogite Zone Behrmann & Ratschbacher (1989)and pseudomorphs after lawsonite in both Inner and Upper Schieferhulle garnets. Therefore, D0

predated M1 and was the result of a flattening strain regime associated with subduction(i.e. Kurz,2005).

D1 occurred contemporaneously with M1 HP conditions within the Eclogite Zone and just priorto blueschist facies conditions in the Glockner nappe. In the Eclogite Zone, it is characterised byan omphacite foliation and associated lineation in mafic eclogites, and a phengite foliation withinmetasedimentary lithologies. D1 in the Glockner nappe is manifest as aligned trails of flattened(S>L) albite clots. Intense isoclinal folding and boudinage of S0 banding occurred throughout allunits. Pre-Alpine fabrics within the Altkristallin and Inner Schieferhulle were tightened to formsuch isoclinal folds. The D1 event marks the transition from a flattening (S>L) to a constrictional(L>S) strain regime. D1 occurred under conditions close to 25 kbar and 600� in the Eclogite Zone,whereas it occurred prior to lawsonite growth in the Glockner nappe, at pressures less than 13 kbar.Such a large disparity in metamorphic conditions (>10 kbar and 200�), shows that the EclogiteZone and Glockner nappe were vertically separated by more than 30 kilometers at the onset of D1.Therefore, it is unlikely that D1 occurred contemporaneously throughout the Penninic nappes.

D2 is characterised by transposition of the S1 fabric to form the regional south-dipping compos-ite foliation (S1+ S2) in all units. Isoclinal folds were rotated, and further tightened, sub-parallelto S2; D2 crenulation of S1 shows axial planar foliation development in metasediments of the Innerand Upper Schieferhulle. D2 initiated whilst the Penninic nappes were under blueschist facies con-ditions and ceased under Barrovian conditions: in the Eclogite Zone, glaucophane growth occurred

30

Deformation Metamorphism Regional Events

D1

D2

D3

D0

pre-D0

M3M2M1

M0

pre-M0N

-S c

onve

rgen

ceE

-W e

xten

sion

imbrication

Her

cyni

an

eclogitisation

uplift and dome formation

continental collision

exhumation

onset of subduction

EoAlpine subduction

D1Glockner

EZ

Olig

ocen

eM

ioce

nePliocene

PleistoceneE

ocen

e

Paleocene

Figure 2.11: Relationship between metamorphic and deformation events in Tauern Window nappes as determined by

field studies presented in the preceeding sections. Timings of events are not absolute; grey band represents the timing of

the flysch–molasse transition as determined by Sinclair (1997). Note the postulated relative positions of D1

in the Eclogite

Zone and Glockner nappes; non-continuous time scale.

concomitant with initiation of F2 crenulations, whereas, in the Glockner nappe, D2 occurred syn-chronously with decompression and pseudomorph development of lawsonite-bearing assemblages.S2 + S1 is concordant throughout the Tauern nappe stack, which shows that the nappes wereassembled prior to, or during the onset of D2. L2 mineral alignment occurred under Barrovianconditions and dips at shallow angles to the east and west throughout the nappe pile. Shear senseindicators are consistently top-to-the-north, with the exception of the Eclogite Zone–Glocknernappe contact. F2 fold axes trend towards the south-southwest and confirm that the Penninicnappes were assembled under a D2 transpressive strain regime. The change from north-south con-

31

vergence during D1 to east-west transpression in D2 is likely associated with collision between theAdriatic and European continental blocks.

D3 is locally developed throughout the nappe pile in the form of quartz-filled fissures and veins,together with F3 buckle folds, which are best exposed in calc-schists of the Glockner nappe; F3

axial planes strike north-south. D3 occurred as a result of continued transpressive exhumation anddoming of the Tauern nappes.

In summary, the structural record of Alpine deformation preserved in Penninic units of theTauern Window documents a transition from subduction-related flattening strain to exhumation-related constrictional strain. The strain-regime inflexion occurred at eclogite and blueschist-faciesconditions within the Eclogite Zone and Schieferh¨lle nappes, respectively, and likely correspondsto continental collision between the European and Adriatic continents.

32

Chapter 3

Conditions of Metamorphism

Pseudosection modeling of garnet–chloritoid–kyanite assemblages presented in this chapterforms the basis of the published manuscript: Smye et al. (2010).

3.1 Introduction

Advances in the understanding of the thermodynamic behavior of metamorphic phases, both rock-forming and accessory constituents (e.g. Helgeson et al., 1978; Holland & Powell, 1985; Powell &Holland, 1985; Berman, 1988; Gottschalk, 1997; Holland & Powell, 1998), permit the calculation ofaverage conditions of crystallization and quantitative phase diagrams for mineral systems thoughtto record chemical equilibrium. This approach provides a powerful tool for establishing P–T pathsof exhumed metamorphic rocks. This chapter presents petrographical and mineral chemistry datafrom representative samples collected from the Venediger, Eclogite Zone and Glockner nappes,which are used to derive detailed P–T trajectories. Collectively, this confirms that Inner andUpper Schieferhulle units experienced similar decompressive P–T loops, characterised by earlyblueschist conditions and subsequent decompression on Barrovian geotherms. A metapelite fromthe Eclogite Zone records peak pressures indicative of subduction to mantle depths, followed bybroadly isothermal decompression to mid-crustal levels.

3.1.1 Application of equilibrium thermodynamics

To calculate geologically accurate P–T–X relations from metamorphic rocks, it is critical thatthe system’s (i.e. mineral assemblage or sub–assemblage) phases are shown to represent a stateof thermodynamic equilibrium such that gradients in chemical potential (�µ) between reactantsand products of a reaction equal zero (in the beneath, R=molar gas constant, T=temperature,K=equilibrium constant and ai=activity of phase i):

�µ = 0 = �G� + RTlnK (3.1)

33

where,

µi = Goi + RTlnai (3.2)

Given that metamorphism is a dynamic process, with P , T and X changing throughout theduration of a metamorphic rock’s evolution, it is uncertain whether chemical equilibrium is a validassumption. The kinetics of intergranular di↵usion exert a dominant control on time and lengthscales of metamorphic equilibration (Carlson, 2002). Quantitative knowledge of the rates of inter-granular di↵usion for major and trace elements remain largely unknown. However, the presenceof an intergranular fluid phase is known to promote higher rates of elemental exchange and hence,metamorphic equilibration (Carlson, 2010). Accordingly, prograde (dehydration) metamorphic re-actions proceed more rapidly than retrograde (hydration) reactions—accounting for the commonpreservation of mineral assemblages formed under peak T conditions. This is supported by con-firmation of predictions made from phase equilibria modeling and observed metamorphic mineralassemblages. Despite theoretical validation of the equilibrium condition, commonly observed coro-nal microstructures (e.g. Whitney & McLelland, 1973; Ashworth & Birdi, 1990; Johnson & Carlson,1990), linked segregations (Carmichael, 1969) and pseudomorphous growth structures (Foster Jr,1986) provide petrographic evidence for chemical disequilibrium, where reactions ceased prior tothe full consumption of reactants. Such evidence is most commonly found in H2O under-saturatedconditions and suggests sluggish kinetics of intergranular di↵usion. Accordingly, it is fundamentallyimportant to screen samples for petrographic evidence of disequilibrium before applying equilib-rium thermodynamic principles. The equilibrium model of metamorphism cannot be ‘proved’ inthe same way as disequilibrium; instead, the absence of disequilibrium features is generally takenas supporting evidence. Furthermore, equilibrium assemblages must obey the phase rule (F = C

+ 2� P , where F=variance, C=the number of components and P=number of phases present) andshow consistent element partitioning between cogenetic phases.

Accepting the valid application of equilibrium thermodynamics to metamorphic rocks, thermo-dynamic descriptions of relevant phases (e.g. Helgeson et al., 1978) are combined with activity–composition models (a–x) describing the energetics of end-member phase interaction, to permitcalculation of the assemblage’s peak P–T . Originally, this approach was limited to specificallycalibrated reactions (e.g. Ferry & Spear, 1978), followed by consideration of numerous reactions,which allowed calculation of average P–T conditions (Powell, 1985; Powell & Holland, 1988, 1994).However, such thermobarometry is limited by a dependence on mineral chemistry—inherent to themethod’s inverse nature. Improvements in the quality and breadth of thermodynamic data anda–x relations based on the same phase end-member compositions (internally consistent datasets– Holland & Powell, 1998; Powell & Holland, 1988; Holland & Powell, 1990), underpinned thedevelopment of forward phase equilibria modeling, i.e. given a reactive composition, P–T , P–X,or T–X relations (pseudosections) can be calculated for all phases in the system of interest.

34

3.1.2 Pseudosection modeling

Pseudosection modeling is currently the most powerful technique employed to derive P–T datafrom exhumed metamorphic rocks (e.g. Powell et al., 1998; White et al., 2000; Tinkham et al.,2001; Powell & Holland, 2010). Pseudosections arose from the construction of petrogenetic grids,which show all the stable invariant points and univariant equilibria for all phases and bulk com-positions in a chemical system (e.g. Hess, 1969; Harte & Hudson, 1979a; Spear & Cheney, 1989;Powell & Holland, 1990a). Such grids provide information regarding the absolute stability of as-semblages, however, they do not provide information regarding the composition and abundance ofphases. Pseudosections provide a means by which the system’s petrogenetic grid is modified for aspecific bulk composition—e↵ectively forming a mineral assemblage map in P–T–X space (Powellet al., 1998).

In a chemical system (e.g. KFMASH), with n components, each phase (k) has ek end–membersand is thermodynamically described by ek�1 compositional variables. Therefore, a reaction in-volving p phases will have ⌃p

k=1ek = s end–members and ⌃pk=1ek-1 = ⌃p

k=1ek�⌃pk=11 = s�p

compositional variables. There will be s-n independent reactions in a system containing n chemi-cal components and s end–members (Powell & Holland, 1988). Solving these non–linear equationsrequires a number of variables (from P , T and s � p compositional variables) equal to the as-semblage’s variance must be fixed. This is because the number of fixed variables equals the totalnumber of variables minus the number of equations: s� p + 2� (s� n), and the phase rule showsthat p + v = n + 2, where v = number of variables. Calculated equilibria can then be assembledaccording to Schreinemakers Rules to form a petrogenetic grid.

As the bulk composition is specified in pseudosections, the non–linear equilibrium equationsdescribed above are augmented by an additional n mass balance equations in the form: Z =⌃i(propi[⌃jc

jZXi

j ]), where Z represents the amount of a component in the bulk composition, propi

is the modal proportion of phase i, cjZ is the amount of Z in the end–member j of phase i and

Xij is the quantity of j in phase i. This means that the total number of equations to be solved

will be s � n + n = s with s � p + 2 + p=s + 2 variables. Hence, if two variables are fixed theequations can be solved. This means that mineral modes and composition contours (isopleths) canbe calculated for all variance fields, constraining sections of P–T path. A more detailed treatmentof phase diagram thermodynamics is given by Powell et al. (1998).

The validity of the pseudosection approach relies on an accurate determination of the reac-tive bulk composition. As discussed above, the length scales of di↵usive equilibration are poorlyconstrained in metamorphic rocks, making this the greatest source of uncertainty inherent to thetechnique. Rocks containing zoned porphyroblasts are particularly challenging to model as the ef-fective bulk composition changes with porphyroblast growth, meaning that pseudosection topologywill vary as a function of time. This e↵ect is typified by the growth of garnet which preferentiallysequesters Mn during early stages of growth (Mahar et al., 1997a). It is possible to mitigate

35

against reactive volume fractionation by selecting a suitable composition of garnet—i.e. from core(initiation of growth), to rim (cessation of growth) portions, to be used in calculation of the bulkcomposition. Pseudosections presented in this chapter are based on bulk compositions calculatedfrom combining mineral modes and electron microprobe mineral analyses, unless otherwise stated.

A further limitation to the accuracy of pseudosection calculations is the size of chemical sys-tem in which equilibira are calculated. If modeling is undertaken in a system smaller than thatcontrolling the natural mineral assemblage there will be discrepancy between observed and cal-culated results (White et al., 2007; Powell & Holland, 2008). The magnitude of this discrepancyis determined by the system’s sensitivity to missing components. For example, the stabilities ofstaurolite and spinel are heavily influenced by the activity of Zn, and yet relevant Zn end–memberthermodynamic data for metamorphic phases are lacking. This means that phase equilibria model-ing in Zn–absent systems must be treated with caution. In general, and depending on assemblagemineralogy, larger systems converge closer to natural systems than systems with fewer chemicalcomponents (e.g. Tinkham et al., 2001; Smye et al., 2010).

Pseudosection calculations were carried out using Thermocalc v.3.33 software. Electronmicroprobe analyses were converted into end-member activities using AX software—available fordownload at http://www.esc.cam.ac.uk/research/research–groups/holland/thermocalc.

3.1.3 Average P–T

In systems where a reactive bulk is di�cult to establish, such as re–worked metamorphic rocks,or in which the peak mineral assemblage contains a small number of phases, the pseudosectionapproach can be complimented by conventional thermobarometry.Conventional thermobarometry includes all methods in which balanced reactions are written be-tween phase end–members present in equilibrium. If there are n end–members in a system describedby m components, there will be a maximum nCm+1 reactions, of which n�m will be independentof each other. From 3.1, the sum of the chemical potentials for each reaction must equal zero:

⌃niµi = 0 = ⌃niGoi + ⌃niRTlnai = �Go + RTlnK (3.3)

Thermobarometric information from a reaction is based on the variation of �Go with temperatureand pressure:

�Go ⇡ �Ho1bar � T�So

1bar + P�V o (3.4)

If �V o is known from X-ray–based structural refinements, �So and �Ho of an equilibria may bedetermined from experimental petrology such that the following expressions provide estimates ofa reaction’s P at a specified T andvice versa:

P =��Ho

1bar + T�So1bar �RTlnK

�V o(3.5)

36

T =�Ho

1bar + P�V o

�So1bar �RlnK

(3.6)

Given mineral analyses, a–x models and values for �So1bar, �Ho

1bar and �V o, n-m independentreactions can be positioned in P–T space to yield an average P–T estimate. The average P–T method of Thermocalc (Powell & Holland, 1994), performs such calculations by varying theactivities of end–members in proportion to their uncertainties (from thermodynamic data, electronmicroprobe data and a–x models) such that the equilibria intersect at a point—the average P–T .In doing so, errors associated to individual equilibria become correlated, leading to representationof the average P–T by an error ellipse of the form:

P 0 =�Pp

2(acos✓ + bsin✓) + Pav) (3.7)

T 0 =�Tp

2(acos✓ � bsin✓) + Tav) (3.8)

Where a =p

1 + ⇢PT and b =p

1� ⇢PT . The quality of an average P–T calculation isassessed in terms of diagnostic quantities. �fit is a measure of the discrepancy between observedand calculated values of enthalpy and activities for individual end–members—a �fit of 1.0 confirmsrealistic input values; e* is a measure of the di↵erence between an end–member activity calculatedfrom a mineral analyses versus the value required for all equilibria to intersect at the average P–T—large values show poorly fitted activities. Thermocalc undertakes a �2 test to calculate themaximum �fit value in which the average P–T is determined to 95% confidence level. A furtherdiagnostic, hk, or hat values gives an indication of the degree of influence an end–member has onthe �2 result. These diagnostics can be used to screen equilibria for problematic activities, whichcan be removed, thus lowering the �fit of an average P–T estimate. In fluid–bearing systems, �fit

can often be improved by varying values of aH2

O and aCO2

to locate a �fit minima.Commonly, assemblages with few end–members will fail to generate su�cient equilibria to calculatean average P–T . In such a case, average values for P and T can be calculated over intervals of T

and P respectively (Powell & Holland, 1994).

3.2 Eclogite Zone P–T evolution

The Eclogite Zone represents the largest exposure of eclogite facies Penninic crust in the EasternAlps. Consequently, understanding its P–T history is vital for models of Alpine orogenesis. Eclog-ite facies rocks are exposed in a 1–3 km thick tectonic sliver, juxtaposed between the underlyingVenediger nappe complex and overlying Rote Wand\Glockner nappes. The Eclogite Zone spansapproximately 20 km along strike, between the southern flanks of the Großvenediger range andthe western extent of the Großglockner range—south central Tauern Window.As discussed previously, the nappe comprises a heterogeneous mix of Mesozoic mafic and metased-imentary lithologies (Kurz et al., 1998b). In order of volumetric significance, the three dominant

37

rock types are: calcareous metasediments, pelitic metasediments and sensu stricto mafic eclog-ites. It is clear that the Eclogite Zone is the only tectonic unit within the Tauern nappe stack, tohave experienced pervasive eclogite facies metamorphism prior to regional recrystallisation underBarrovian conditions. Therefore, it is fundamentally important to constrain the Eclogite Zone’sP–T history in order to understand the tectono–thermal history of the nappe stack.

3.2.1 Previous work

The Eclogite Zone has been the subject of extensive petrographic and thermobarometric studythroughout the past seventy years (Dachs et al., 2005, and references therein). Eclogites andglaucophane-bearing rocks in the south–central Tauern Window were first described in the lit-erature by Cornelius et al. (1939), Scharbert (1954) and Schmidegg (1961). Further study ledAbraham et al. (1974) and Raith et al. (1977) to conclude that the eclogites represented a culmi-nation of progressive metamorphism from south to north, increasing in grade from the calcareousphyllites of the Matrei Zone, through the greenschist–amphibolite facies prasinites of the Glocknernappe, to the Eclogite Zone. Early thermobarometric calculations on eclogite facies assemblagesyielded 400–500�C, 8–11 kbar at low vales of aH

2

O (Miller, 1974, 1977). However, P–T conditionswere re–evaluated following the pioneering work of Holland (1977) and Holland (1979), who usedphase equilibria constrained by the coexistence of paragonite–omphacite–kyanite with dolomite–omphacite–quartz assemblages in mafic eclogites to calculate conditions of 620±30�C at 19.5±2.5kbar and XCO

2

= 0.02±0.02. Subsequent thermobarometric work has corroborated these findingsto bracket the P–T conditions of mafic eclogite formation between 20–25 kbar at 590–650�C (Franket al., 1987; Hoschek, 2001, 2004, 2007).

Eclogite Zone metasediments have received comparatively little attention (Holland, 1977, 1979;Franz & Spear, 1983; Dachs, 1986; Spear, 1986) despite showing that both metapelitic and metacar-bonate lithologies experienced eclogite facies conditions similar (20–25 kbar at 550–620�C) tothose derived from mafic eclogites. Collectively these data have been interpreted to indicate thatthe Eclogite Zone behaved as a coherent unit during its exhumation from 60–80 km depth (e.g.Behrmann & Ratschbacher, 1989).

Holland & Richardson (1979) and Eremin (1994) showed that the Eclogite Zone experiencedpartial recrystallisation of eclogite-facies amphibole and omphacite to glaucophane under blueschistfacies conditions. Quantitative estimates of P and T for this metamorphism are thwarted bypartial assemblage development. Holland & Richardson (1979) use glaucophane cores of zonedamphiboles which overgrow an omphacitic fabric to calculate anomalously low temperatures of200�C at 15 kbar. Eremin (1994) calculated a range of average P–T values between 450–550�C and10–20 kbar, from retrograde assemblages present in both mafic eclogites and metasediments. Theconditions of greenschist facies metamorphism are better constrained. Hoernes (1974) used oxgenisotope exchange between quartz and biotite to show peak temperatures did not exceed 550�C and

38

garnet S0+S1+S2

chloritoid

quartzite

Figure 3.1: Sample

TH–680. Note garnet

poikiloblasts and chlori-

toid tablets wrapped by

the quartz-mica S1

+S2

schistosity. F2

crenu-

lations are visible to-

ward the far-raight of

the sample. Scale bar

is 10 mm.

Zimmermann et al. (1994) showed that pressures were close to 7 kbar for the same temperatureinterval. More recently, Hoschek (2001) used average P–T methods on retrogressed metasediementsof the Eclogite Zone to constrain greenschist–amphibolite facies conditions between: 9–10.2 kbarand 552–565�C.

In light of the paucity of recent thermobarometric work focused on Eclogite Zone metapelites,despite their common occurrence, the following sections present detailed petrography, mineralchemistry and P–T constraints obtained from a garnet + chloritoid + kyanite-bearing metapeliticsample (TH–680; Holland, 1977)) collected from the interior of the nappe.

3.2.2 Sample petrography

Sample TH–680 (Fig.3.1) was collected by T.J.B. Holland in the 1970’s (Holland, 1979) fromthe Eclogite Zone exposed in Timmeltal, on a plateau north-east of Eissee and directly south ofWeißspitze (47�03’59”N, 12�22’43”E). The sample is representative of Eclogite Zone metapelitescommonly exposed as millimetre- to metre-scale horizons intercalated with metacarbonate andquartzite layers (3.1).

Sample TH–680 comprises the following high-pressure assemblage, in order of abundance:phengite, quartz, rutile, allanitic epidote, chloritoid, garnet and kyanite. The pervasive foliation(S0+S1+S2), which wraps porphyroblastic garnet, chloritoid, kyanite and clinozoisite–allanite,is defined by the shape-preferred orientation of phengite grains and quartz seams, and is locally,openly crenulated (Fig.3.2(a), Fig.3.2(b) and Fig.3.2(c)). Evidence for retrogressive mineral growthis confined to restricted patches of chlorite growth around the edges of chloritoid and garnet blasts.Garnet exists as rounded, mm-scale poikiloblasts which contain an inclusion assemblage dominatedby phengite, quartz and rutile with subordinate amounts of chloritoid and kyanite. These inclusionscommonly define sigmoidal fabric traces which are discordant to the matrix foliation. Chloritoidforms both mm-scale tablets within the matrix and also a ragged overgrowth around kyanite shards,suggesting two stages of growth. Kyanite exists in smaller quantities than both chloritoid and gar-

39

net as euhedral blasts within the matrix. Interestingly, margarite occurs as a rare, late-stage phase,overgrowing the penetrative phengite fabric. Subordinate amounts of apatite, pyrite, rutile anddetrital zircon are interstitially present between matrix micas.

(a) Included chloritoid within garnet (b) Matrix kyanite rimmed by chloritoid

(c) Phengite–epidote crenulation (F2

) (d) Late-stage margarite (center)—distinguishable from muscovite by itshigher relief, growing on top of muscovite

Figure 3.2: Photomicrographs of TH–680 taken under crossed–polars; field of view 5mm.

3.2.3 Mineral chemistry

Garnet poiklioblasts are chemically defined by an almandine–grossular–pyrope solid solution anddisplay weakly developed core–rim chemical zoning: Xalmandine = 0.68–0.63; Xgrossular = 0.17–0.20;Xpyrope = 0.13–0.18; Xspessartine = 0.02–0.005, from core to rim respectively. Sharp disturbances inthe core–rim profiles displayed in Fig.3.3 are due to the presence of included phases. Notably, lowXgrossular regions border included quartz and chloritoid, suggesting retrograde re-equilibration.

Chloritoid exists as two chemically distinct populations: blasts included within garnet (Fig.3.2(a))have XMg = 0.28–0.29, contrasting with matrix chloritoid XMg values between 0.33–0.35. Chlori-toid replacing kyanite Fig.3.2(b) is chemically indistinguishable from matrix chloritoid tablets.

Phengite is the dominant matrix constituent and has a compositional range between 3.2–3.4

40

0.55

0.60

0.65

0.70

0.75

0 500 1000 1500 2000 25000.00

0.05

0.10

0.15

0.20

Xalmandine

Core-rim distance (µm)

Xpyrope

Xgrossular

Xspessartine

0.60

0.65

0.70

0.75

0 200 400 600 8000.00

0.05

0.10

0.15

0.20

Core-rim distance (µm)

a.) b.)

Figure 3.3: Core–rim compositional traverses of garnet. End–members Xalmandine = Fetotal/Fetotal+Mg+Mn+Ca;

Xgrossular = Ca/(Ca+Fetotal+Mg+Mn); Xpyrope = Mg/(Mg+Fetotal+Mn+Ca); Xspessartine =

Mn/(Mn+Ca+Mg+Fetotal)

2.0

2.2

2.4

2.6

2.8

3.0

3.2 3.3 3.40.0

0.2

0.4

0.6

0.8

1.0

Si c.p.f.u (11 O)

Al

K

Mg

Fe

Na Figure 3.4: Muscovite composition (c.p.f.u.) on 11 O basis.

Note Tschermak-style Mg+Si=2Al substitution

Si cations per formula unit (c.p.f.u.) per 11 oxygens, XMg = 0.72–0.77 and XNa = 0.049–0.081.Figure 3.4 shows the degree of Tschermak MgSiAl�2 substitution seen in matrix mica—the majorityof matrix phengite analyses plot between 3.3–3.4 Si c.p.f.u., beneath 2.4 Al c.p.f.u, suggesting acommon growth history. This corroborates a lack of evidence for late-stage muscovite formation.

Allanite in TH–680 is present as rare inclusions within garnet and, more commonly, as ma-trix porphyroblasts. Grains are zoned from core to rim by increasing clinozoisite content (3.5),commonly ascribed to prograde metamorphic growth (Janots et al., 2007, 2008, 2009). Allaniticcore regions typically contain 0.5 c.p.f.u REE + Th and are heavily enriched in LREE (La/Yb⇡ 250–900). Clinozoisite rims contain less than 0.1 c.p.f.u REE + Th and show less fractionatedREE patterns (La/Yb ⇡ 30–200). A more detailed discussion of allanite petrogenesis is given insection 4.9.

41

0.5

1.0

0.0

1.0 1.5 2.5Epidote Clinozoisite

AllaniteFerriallanite

core

rim

� R

EE +

Th

Al

0.2

0.4

0.6

0.8

Figure 3.5: REE + Th vs. Al plot (Petrik et al., 1995) of allanite–clinozoisite (c.p.f.u.); note the core to rim decrease in

REE content and increase in clinozoisite end–member. Red dashed lines are labeled for Fe3+/Fetotal. Allanite–clinozoisite

mineral analyses were recalculated according to the criterion: ⌃(Si+Al+Ti+Fe+Mn+Mg)=6 c.p.f.u. (Ercit, 2002).

Mineral: garnetrim garnetcore chloritoid phengite epidote

SiO2

37.48 35.89 24.77 50.48 37.86TiO

2

0.03 0.08 0.00 0.33 0.19Al

2

O3

21.60 21.20 40.46 28.30 27.12Cr

2

O3

0.00 0.00 0.02 0.02 0.01FeO 29.19 30.93 19.17 1.97 6.37MnO 0.07 0.76 0.27 0.02 0.04MgO 4.39 3.27 6.11 3.40 0.31CaO 7.23 6.41 0.02 0.02 21.74Na

2

O 0.03 0.06 0.01 0.36 0.04K

2

O 0.00 0.00 0.00 10.81 0.07Total 100.02 98.60 90.83 95.76 93.75

No. of oxygen 12 12 6 11 12.5Si 2.95 2.90 1.11 3.36 3.04Ti 0.00 0.00 0.00 0.02 0.01Al 2.00 2.00 1.87 2.22 2.57Fe

tot

1.87 1.98 0.65 0.11 0.26Mn 0.01 0.05 0.01 0.00 0.00Mg 0.51 0.39 0.32 0.34 0.04Ca 0.61 0.55 0.00 0.00 1.87Na 0.01 0.01 0.00 0.05 0.01K 0.00 0.00 0.00 0.92 0.01

Table 3.1: Representative electron microprobe analyses of peak assemblage phases in TH–680. SeeAppendix A.2 for electron microprobe operating conditions.

3.2.4 Metamorphic modeling of TH-680

A P–T pseudosection for sample TH–680 was constructed in the expanded pelitic system Mn-NCKFMASHO, between 5–30 kbar and 400–650�C. TiO2 was ignored due to the low modal per-centage of rutile, the only Ti bearing phase present (<0.3 %). The reactive bulk composition wascalculated by combining mineral modes obtained via point counting with representative mineral

42

analyses (3.1). As the modes of garnet and epidote (the only phases in which rim compositionsdi↵er by �5% of core compositions) are small, TH–680 can be represented accurately via a singlebulk composition without the need to consider porphyroblast-growth-controlled bulk fractionation.In order to account for early growth of epidote/clinozoisite in the rock’s history, ferric iron was re-duced from the point counting value of O=0.7 (FeO:Fe2O3 = 7.33), to 0.188 (FeO:Fe2O3 = 27.94).Calculations involve garnet, epidote, chloritoid, chlorite, carpholite, kyanite/sillimanite, margarite,jadeite, paragonite, muscovite and biotite, in accordance with the observed mineral assemblage inTH–680 and also with pelitic mineralogy in general. Quartz/coesite, H2O and muscovite are forcedto be in excess throughout the investigated P–T range, which is consistent with the high modalabundance of muscovite and quartz.

3.2.4.1 Activity–composition models

Epidote–clinozoisite is modeled as a ternary solid solution, which incorporates non-ideal mixing(Holland & Powell, 1998). Garnet activities are calculated using the symmetric formalism modelof White et al. (2005). Chloritoid activity–composition relations are updated from those in Whiteet al. (2000). Recognising that chloritoid is characterised by two types of layer in which Al andFe3+ can only mix on one M1A site, its end-member formulae may be written as: fctd – Fe2Al3Al Si2O8(OH)4; mctd – Mg2Al3 Al Si2O8(OH)4; mnctd – Mn2Al3 Al Si2O8(OH)4; ctdo – Fe2Al3Fe3+ Si2O8(OH)4. Activities are given by:

afctd = X2Fe M1BXAlM1A�fctd

amctd = X2Mg M1BXAlM1A�mctd

amnctd = X2MnM1BXAlM1A�mnctd

actdo = X2Fe M1BXFe3+ M1A�ctdo

and all non-ideality taken as symmetric (regular solution) with Wfctd mctd = Wfctd mnctd = Wmctd mnctd =1.0 kJ and all W ’s involving ctdo being set to zero. The Gibbs energy of end–member ctdo wasmade from 2Gfctd + (Ghem �Gcor)/2 + 25.0 kJ.

For carpholite the activity composition relations are updated from those in Coggon & Holland(2002), with its end-member formulae written as: fcar – Fe Al Al2Si2O10(OH)4; mcar – Mg AlAl2Si2O10(OH)4; mncar – Mn Al Al2Si2O10(OH)4; caro – Fe Fe3+ Al2Si2O10(OH)4.

In order to model the white micas in systems with calcium and ferric iron, the mixing parametersfor margarite (ma) and ferrimuscovite (fmu) endmembers need to be added to the simpler whitemica model of Coggon & Holland (2002). The brittle mica margarite also requires consideration as aseparate phase because it occurs in rock TH-680 as late cross-cutting laths grown during subsequentdecompression and later Alpine metamorphism. Modification of the white mica model was doneby L.Greenwood and T.J.B.Holland (Greenwood, 2009), and is only summarized here—see Smyeet al. (2010) for a more complete discussion. Coggon & Holland (2002) built a white mica model

43

for phengites and paragonites in the simple KNCFMASH system, and included the endmembersmuscovite (mu), paragonite (pa), celadonite (cel) and ferroceladonite (fcel) to describe white micasrelated by the muscovite–paragonite solvus. Paucity of experimental data makes addition of themargarite (ma) end–member (CaAl2[Al2Si2]O10OH2) a di�cult task. Data from natural coexistingmica + margarite + paragonite triplets from greenschist–amphibolite facies rocks (Hock, 1974;Okuyama-Kusunose, 1985; Feenstra, 1996) are used to constrain mixing parameters (Fig.1–Smyeet al. (2010)). The basic model of Coggon & Holland (2002) provides the interaction energiesamong the first four endmembers detailed beneath; those involving ma and fmu were determinedvia best–fit methods augmented by assumptions regarding the Ca and K contents of muscoviteand margarite respectively, to replicate compositions of the natural coexisting mica triplets.

A M2A M2B T1 T2mu K Al Al SiAl Si2cel K Mg Al Si2 Si2fcel K Fe Al Si2 Si2pa Na Al Al SiAl Si2ma Ca Al Al Al2 Si2fmu K Al Fe3+ SiAl Si2

Addition of the fmu endmember is more approximate as there are virtually no data at hand,and we use the simplest model which will allow incorporation of a small ferric iron content into theM2B site. Thus the fmu endmember is taken as mixing ideally with all other endmembers, andany nonideality present is assumed to be absorbed into the Gibbs free energy, which was adjustedto produce similar ferric iron contents in micas to those coexisting with ilmenite and magnetite(Dyar et al., 2002). White mica activities are given by:

amu = XK AXAlM2AXAlM2BXSi T1XAlT1�mu

acel = XK AXMg M2AXAlM2BX2Si T1�cel

afcel = XK AXFe M2AXAlM2BX2Si T1�fcel

apa = XNa AXAlM2AXAlM2BXSi T1XAlT1�mu

ama = XCa AXAlM2AXAlM2BX2AlT1�ma

afmu = XK AXAlM2AXFe3+ M2BXSi T1XAlT1�fmu

Activities for chlorite and staurolite use the models of (Holland et al., 1998) and (Mahar et al.,1997b) respectively. Finally, jadeite is assumed to be of end–member composition.

3.2.4.2 TH–680 P–T pseudosection

The pseudosection (Fig.3.6) is characterized by the presence of large high variance (F = 4–6) garnetbearing fields. High �P/�T regions are defined by coexisting carpholite and jadeite, whereas

44

low �P/�T assemblages contain subsets of biotite–chlorite–margarite-bearing assemblages. Thepeak, garnet–chloritoid–kyanite assemblage is restricted to a narrow, temperature controlled fieldbetween 550–600�C at 17–25 kbar. Up-temperature, the stability field is defined by the breakdownof chloritoid, whereas the lower temperature boundary is defined by the first appearance of kyaniteat the expense of garnet + chloritoid + epidote. Epidote is the dominant ferric iron bearing phasein the system and is predicted to be present beneath ⇠600�C. Contours of fep (Fe3+:Al; Holland &Powell, 1998) proportion in epidote–clinozoisite (Fig.3.6) suggest that clinozoisite rims on allanitecores in TH–680 equilibrated between 425–500�C and 8–14 kbar—see section 4.9.3 for a moredetailed discussion. This corroborates with the microstructural setting of the clinozoiste–allaniteporphyroblasts as an early-growth phase in the P–T evolution (Fig.3.2). Comparison betweenpredicted and observed (Fig.3.3) garnet compositions suggest equilibration over a small P–T range,between 15.5–17.5 kbar and 560–580�C, using the symmetric formalism MnCFMASO garnet modelof White et al. (2005). The composition of chloritoid included within garnet poikiloblasts (Table3.1) is calculated to be stable between ⇡500–550�C within the large pentavariant garnet–chloritoid–chlorite–epidote field. Matrix chloritoid which, importantly, rims kyanite is slightly more Fe–richthan included grains (Xctd = 0.65–0.67) and is predicted to have equilibrated at 10–25�C higherthan the early chloritoid generation. The presence of kyanite indicates peak conditions �575�C.Si in matrix phengite varies between 3.25–3.4 c.p.f.u., consistent with equilibration above 16 kbar(3.4). The retrograde P–T path is constrained by the formation of margarite (Fig.3.2(d)) andlow-Mg chloritoid at the expense of kyanite, which collectively suggest (i.) a decompressive P–T evolution, and (ii.) approximate isothermal cooling from peak conditions within the garnet–chloritoid–kyanite/garnet–chloritoid fields to margarite–chlorite bearing stability fields—see P–T

loop overlaid on pseudosection (Fig.3.6). There is no petrographic evidence to support the stabilityof jadeite and/or carpholite during the rock’s P–T evolution, which is consistent with the extentof their stability fields on the pseudosection.

Observed mineral modes are in close agreement with those predicted by pseudosection calcu-lations (Fig.3.7) for coexisting garnet, chloritoid, kyanite, muscovite and quartz at 25 kbar and555�C, indicating the validity of chemical system, bulk composition and activity–composition mod-els employed.

3.2.4.3 Average P–T calculations

Further thermobarometric information pertaining to the peak conditions experienced by rock TH–680 was collected via average P–T calculations assuming chemical equilibrium between garnet rimportions, matrix chloritoid, kyanite, muscovite, quartz and an H2O-rich (XH

2

O > 0.5) fluid phase.The e↵ect of including epidote–clinozoisite in equilibria was also considered. Calculations betweenend-members of these phases used the following independent reactions:

45

400 450 500 550 600 6505

7

9

11

13

15

17

19

21

23

25

27

29

15

30

45

60

75

T(ºC)

P (k

bar)

Depth (km)

T(ºC)

P (k

bar)

400 450 500 550 600 6505

7

9

11

13

15

17

19

21

23

25

27

29

g ep car jdqtzcoe

g ky jd

g ky

g ctd ep car jd

g ky car jd

ctd chl ep

ctd chlep pa

g ky chl ma

g chl ma

g chl ma ep g chl ma bi

g ma bi st

g ky bi ma

g ctd ep

g ep car ctd

g ctd ep pa chl car

g ky car

g ctd ky

TH-680 - MnNCKFMASHOH2O + mu + SiO2

kysill

5

7

9

11

13

15

17

19

21

23

25

27

29

0.24

0.23

0.22

0.21

0.20

0.19

0.18

0.17

0.5

0.6

0.65

g ctd chl ma ep

g ky chl

3.243.26

3.283.30

3.323.343.36

0.23

0.24

0.25

0.26

0.27 0.28 0.29 0.30

0.8

0.7

0.6

0.5

g ctd chl ep

a.

b.

c.

clinozoisite growth

g gr

owth

Figure 3.6: a MnNCKFMASHO P–T pseudosection for TH–680. Reactive bulk is MnO: 0.091; Na2

O:0.243; CaO:1.308;

K2

O:3.580; FeO:5.629; MgO:3.692; Al2

O3

:13.956; SiO2

:71.312; O: 0.188—mineral modes are displayed in Table 3.2.

Yellow arrow denotes P–T path as constrained by phase relations and isopleths; oval transparencies indicate equilibration

conditions of clinozoisite rim portions and garnet porphyroblasts. b Inset shows Xctd (Fe2+/(Fe2++Mg)) isopleths in

green, fep (Fe3+

M13

/(Fe3+

M13

+ AlM13

)) isopleths in white and Si (c.p.f.u.) in muscovite isopleths in black. c Inset shows

Xg (Fe2+/(Fe2++Mg+Ca+Mn)) isopleths in yellow and zg (Ca/(Fe2++Mg+Ca+Mn)) isopleths in black.

0%

20%

40%

60%

80%

100%

kyanite

quartz

chloritoid

garnet

muscovite

CalculatedObserved

Figure 3.7: Comparison between calculated and observed

modes at 25 kbar and 555�C for MnNCKFMASHO TH–680

bulk composition. Mineral modes calculated by Thermo-

calc were converted into volume % modes using the follow-

ing molar volumes obtained from Holland & Powell (1998):

muscovite 14.05 J.bar�1, garnet (50:50 almandine:pyrope so-

lution) 11.3 J.bar�1, chloritoid 6.87 J.bar�1, kyanite 4.41

J.bar�1 and quartz 2.27 J.bar�1.

46

1. 3mctd + 2q = py + 2ky + 3H2O2. 3fctd + 2q = alm + 2ky + 3H2O3. 3cel + 4ky = py + 3mu + 4q4. 3fcel + 4ky = alm + 3mu + 4q5. 6cz = 4gr + q + 5ky + 3H2O

Due to the small �V of pressure-dependent equilibria, 1 and 4, uncertainty associated withpressure estimates is larger than that associated with temperature. This is a common problemencountered with eclogite-facies thermobarometry, where plagioclase is absent (Powell & Holland,2008). Combining equilibria 1–5 with aH

2

O=1 gives an average P–T of 26.3±2.4 kbar at 539±15�C(�fit=1.45; Fig.3.8). Excluding epidote–clinozoisite from calculations reduces the number of equi-libria to four (equilibrium 5 is omitted) and produces an average P–T of 26.7±2.7 kbar at 538±17�C(�fit=1.60). As epidote–clinozoisite grew early in the P–T path, before garnet, this P–T estimatebest reflects peak conditions.

Water activity is taken as unity for all thermobarometric and phase diagram calculations. Atthese pressures, although dissolved silicate material is expected to be significant, the solvus-likebehaviour of silicate–water systems (e.g. Shen & Keppler (1997)) implies extremely non-idealsolutions characterised by large activity coe�cients such that even for high silicate concentrationsthe water activity will remain much larger than mole fraction. Holland (1979) showed that highwater activity is valid in the Eclogite Zone of the Eastern Alps by measuring the composition offluid inclusions in eclogitic omphacite and zoisite, and concluding that XCO

2

in the fluid was lessthan 0.1. This is also supported petrographically by the lack of a common carbonate phase instudied garnet–chloritoid–kyanite schists. Average P–T conditions for TH–680 are insensitive toaH

2

O—lowering aH2

O from unity to 0.5 yields 26±2.6 kbar at 521±14�C (�fit=1.52), i.e. withinerror of calculations using aH

2

O=1. The e↵ect of ferric iron in muscovite was investigated byvarying Fe2+/Fe3+ between 0.02–0.05, which produced changes in pressure and temperature lessthan 1�C and 0.1 kbar.

3.2.4.4 Influence of Garnet a–X models

Figure 3.8 shows that the epidote–clinozoisite-absent average P–T results (aH2

O=1.0) lie withinerror of the upper boundary of the garnet–chloritoid–kyanite bearing stability field, around 25kbar. However, the P–T conditions over which TH–680 garnet core–rim compositions are stable(Fig.3.6; ⇡15–18 kbar, 555–575�C) is around 7 kbar lower than the calculated average pressure.This inconsistency can be accounted for by changing the symmetric formalism of the MnCFMASOgarnet a–X model of White et al. (2005) to consider a modest degree of asymmetric interaction(Holland & Powell, 2003) between end–members. Using the following interaction energies (sizeparameters for alm., py., gr., spess., and kho., are 1,1,3,1 and 1 respectively) which express a morepronounced degree of asymmetry associated with grossular interactions, grossular content isopleths(zgarnet) plot ⇠1–2 kbar higher than those calculated using the symmetric formalism:

47

16

18

20

22

24

26

28

30

500 510 520 530 540 550 560 570 580 590 600

T (°C)

P (k

bar)

5

2

4

1

31. 6cz = 4gr + q + 5ky + 3H2O2. 3fctd + 2q = alm + 2ky + 3H2O3. 2mctd + cel = py + mu + 2H2O4. 3fcel + 4ky = alm + 3mu + 4q5. py + 3fcel = alm + 3cel

iiiiii

Figure 3.8: Average P–T calculations for rock TH–680. Line numbers correspond to equilibria numbering scheme;

positions of equilibria correspond to P=26.3±2.4 kbar at 539±15�C (see text); red ellipse i. P–T estimate generated

from epidote–clinozoisite bearing equilibria (1–5), aH2O = 1.0; ii.P–T estimate generated from epidote–clinozoisite absent

equilibria, aH2O = 1.0 ; iii. as ii. with aH2O=0.5. Grey region corresponds to garnet–chloritoid–kyanite stability field

(Fig.3.6.

Walm.py = 2.5Walm.gr = 9Walm.spss = 0Walm.kho = 22.5Wpy.gr = 38Wpy.spss = 0Wpy.kho = 0Wgr.kho = -8Wgr.spss = 10Wspss.kho = 20

3.2.4.5 P–T evolution of TH–680

Combining pseudosection and average P–T calculations with sample petrography shows that chlo-ritoid, chlorite and epidote–clinozoiste/allanite were stable throughout much of the prograde P–T

path recorded by TH–680. Despite the work of Spear (1986), who interpreted rectangular shapedaggregates of epidote + quartz + paragonite + phengite + opaques + chloritoid + rutile as pseudo-morphs after prograde lawsonite growth in quartz mica schists from a similar location to TH–680,there is no evidence preserved by TH–680 which suggests either carpholite or lawsonite were sta-

48

ble along the prograde P–T path. These findings are supported by pseudosection calculationsperformed on a similar garnet–chloritoid–kyanite bearing metapelite from the Eclogite Zone byHoschek et al. (2010). The lack of significant (>10% variation from core values) major elementzonation in garnet suggests that poiklioblasts grew rapidly during the prograde path, over a smallregion of P–T space between 550–580�C, before the breakdown of epidote and appearance of kyan-ite. Garnet formation resulted from the breakdown of chlorite and chloritoid. Therefore, theassemblage garnet–chloritoid–kyanite likely records a late stage of the prograde evolution, prior totemperatures exceeding the thermal stability of chloritoid + quartz at 575–600�C, which is bestrepresented by the epidote–clinozoisite absent average P–T calculated above—26.7±2.7 kbar at538±17�C. Interestingly, these conditions imply that the Eclogite Zone crossed the quartz poly-morph transition into UHP conditions during its burial history. However, coesite is yet to bedocumented within the region.

Included high-Mg chloritoid within garnet poikiloblasts in addition to large (millimeter-scale)kyanite blasts present in the matrix, suggest that peak metamorphic conditions lie within thegarnet+kyanite stability field. The small (<3%) modal volume of kyanite present suggests thatthe metamorphic peak occurred proximal to chloritoid+quartz breakdown. Hoschek et al. (2010)favour peak conditions within the garnet+chloritoid+kyanite(+paragonite) stability field due totemperatures calculated from zoned chloritoid+quartz in garnet. Similarly, Stockhert et al. (1997)calculate peak conditions of 25–27 kbar at 575–625�C for the Eclogite Zone paragenesis: gar-net+chloritoid+kyanite+rutile.

The retrograde P–T path is constrained by the presence of rare, cross-cutting laths of margarite(Fig.3.2) in addition to second generation growth of (low-Mg) chloritoid, which grew after kyanite.Collectively, this shows that early stages of the retrograde path were dominated by exhumation frompeak pressures (⇡26 kbar) to <15 kbar under close to isothermal conditions. This is in good agree-ment with the retrograde path determined by Stockhert et al. (1997), who delineate a decompres-sion of 10 kbar with a concomitant temperature decrease of ⇠50�C. Fluid influx likely played an im-portant role in facilitating the growth of retrograde chloritoid. Negulescu et al. (2009) documentedretrograde formation of chloritoid in garnet+kyanite+chloritoid+paragonite+clinopyroxene bear-ing UHP pelites within the southern Carpathians. Here, aH

2

O values of unity for the retrogradefluid were required to calculate measured chloritoid compositions. Similarly, Hoschek et al. (2010)introduce ⇠0.56wt% H2O to their e↵ective bulk in order to model formation of secondary chloritoidduring exhumation.

Sample TH–680 is dominated by the petrographical signature of (U)HP metamorphism; min-eral growth evidence for the locally recorded blueschist facies event is not observed. Similarly,re–equilibration during the subsequent Barrovian event is limited to rare margarite growth. Matrixmuscovite preserves its high–P composition. Notably, TH–680 does not contain staurolite, whichis predicted to be stable over a small portion of P–T space between <9 kbar, 575–620�C–similar toBarrovian conditions. Together, this suggests that fluid mobility (metamorphic permeability) mayhave been limited subsequent to the formation of retrograde chloritoid during initial exhumation,

49

thus hampering recrystallisation over short (<10 Ma) metamorphic timescales.

The P–T path determined from sample TH–680 is in close agreement with paths calculatedfrom both mafic eclogites (e.g Hoschek, 2007) and calc-schist lithologies (e.g Spear, 1986), and istherefore treated as representative for the Eclogite Zone as a whole.

3.3 Garnet–chloritoid–kyanite assemblages

Sample TH–680 provides a good example of how garnet–chloritoid–kyanite parageneses are pow-erful tracers of eclogite facies metamorphism within metapelitic lithologies. Since the work ofHarte (1975) and Harte & Hudson (1979b) in preparing a semi-quantitative petrogenetic grid formetapelites, and the first thermodynamic calibration of it in the KFMASH system by Powell & Hol-land (1990b), it has been known that the assemblage garnet + chloritoid + kyanite in assemblagescontaining muscovite and quartz imply pressures greater than about 15 kbar at around 600�C andis therefore probably restricted to eclogite facies conditions. Pelites containing this assemblagewere first recognised by Holland (1979) as belonging to the eclogite facies rocks of the TauernWindow and as such assigned to conditions of around 20 kbar and 600�C. Recently, understandingof phase relations in metapelites at elevated pressures has been greatly improved via experimentsat high-pressure conditions (e.g. Schreyer, 1988; Massonne, 2000; Hermann, 2002b). In addition,quantitative petrogenetic grids may now be constructed in model systems, such as KFMASH (Cog-gon & Holland, 2002; Wei & Powell, 2003) and NCKFMASH (Wei & Powell, 2006) which provideimportant insights into natural rock assemblages occurring in such high-P metamorphic domains.

Field occurrences

Despite previous studies claiming the rare nature of this assemblage (Gabriele et al., 2003; Neg-ulescu et al., 2009), a survey of the relevant literature suggests that garnet–chloritoid–kyanitebearing schists are a common feature of tectonic environments in which subduction of pelitic ma-terial has occurred. The assemblage is reported from the following localities:

1. The Eastern Alps: eclogitic micaschists of the Eclogite Zone, Tauern Window contain garnet–chloritoid–kyanite parageneses as described in section 3.2.2.

2. The Western Alps: Vuichard & Ballevre (1988) report the occurrence of garnet–chloritoid–kyanite and garnet–chloritoid–chlorite schists from within the high-pressure Sesia zone. Themicaschists are located within the zone’s central portion, the Eclogitic Micaschist unit (Stella,1894), and are associated with paragneisses, marbles and amphibolites. Collectively theserocks have experienced a complex polyphase tectonometamorphic evolution (Zucali et al.,2002), and peak metamorphic conditions were thought to have reached 17–21 kbar and 550–650�C (Tropper et al., 1999) as a result of Paleogene subduction.

50

3. The Carpathians: recent work by Negulescu et al. (2009) documents the presence of metased-iments containing garnet–chloritoid–kyanite bearing assemblages in association with retro-gressed eclogites in the Bughea Complex of the Leaota Massif, South Carpathians. P–T

calculations suggested that the assemblage formed around 18 kbar and 580�C prior to over-printing during exhumation.

4. The Andes: Gabriele et al. (2003) present data from one of the few eclogite terraines withinthe Andean massif. The Raspas Complex of Ecuador represents an exotic fragment of meta-morphosed oceanic lithosphere, which comprises a suite of metaperidotites, eclogites andmetapelites. Exhumation of the Raspas high-pressure rocks and the parent El Oro metamor-phic complex resulted from and occurred during subduction and accretion of the Amotape-Chaucha terrane to preexisting Mesozoic continental arc systems of the Eastern Cordillera(Feininger (1980); Jaillard et al. (1990); Arculus et al. (1999)) . Raspas metasediments con-tain the key garnet–chloritoid–kyanite assemblage along with the subset associations: garnet–chloritoid and garnet–chlorite. Peak metamorphic grade was estimated to be in the region of20 kbar and 550–600�C. The few existing geochronological constraints on the high-pressureevolution of these rocks suggest they cooled through K–Ar closure during the late-Cretaceousperiod (Feininger & Silberman, 1982).

5. The Bohemian Massif: micaschists exposed in the central part of the Erzgebirge mountains,Saxothuringian domain of the Bohemian Massif (Czech Republic) display garnet–chloritoid–kyanite parageneses (Konopasek, 2001). The schists, which are intimately associated withmafic eclogites, were estimated to have equilibrated at conditions around 22 kbar and 640�C.

6. The Betics: garnet–chloritoid–kyanite-staurolite schist ALM-45 was collected by J. M. Bakeras a loose block from the basement rocks of the Sierra Cabrera in South Eastern Spain, wherethe Carboneras fault zone outcrops at the coast near Macenas tower, between Mojaccar andCarboneras. This assemblage can also be observed in several other Betic localities (Agardpers.comm.). They comprise white, phengite-rich schists with apparently rotated garnetswith strong C-S fabrics, presumably indicating extension during exhumation of these rocksfrom deep eclogite facies conditions. Fine grained phengite and quartz-rich bands form aregionally pervasive foliation, which is crenulated at thin-section scale. Garnets commonlycontain ring-shaped quartz inclusions, which are suggestive of an intermittent growth history.The unusual presence of some staurolite, which is often present in the form of corroded relicgrains in close association with quartz and phengite, is enigmatic, as calculations (see later)suggest that the pressures were too high for this phase. It may be that staurolite persistedduring compression into the eclogite facies, or, alternatively that the staurolite grew laterduring decompression. Some kyanites are partially overgrown by staurolite, whereas somestaurolites entrap a folded fabric, making textures ambiguous. Kyanite itself forms mm-scaleporphyroblasts within the matrix and, as with chloritoid blasts, is often overgrown by a fine-

51

grained phengite fabric. Porphyroblastic tablets of chloritoid are common within the matrixindicating its growth early within the rock’s history. Analyses of the principal minerals areprovided in the appendix.

7. The Kokchetav Block: Udovkina et al. (1977) and Massonne & Schreyer (1989) describemetapelites from the Kokchetav Block of northern Kazakhstan as containing garnet poikilo-blasts enclosing chlorite and chloritoid inclusions within a matrix of talc and kyanite.

8. The Norwegian Caledonides: Chauvet et al. (1992) and Hacker et al. (2003) report subassem-blages of garnet–chloritoid–kyanite–staurolite–paragonite–muscovite–quartz within highly alu-minous pelitic horizons of the Gasetjœrn unit, which is interleaved with the Western GneissComplex. Here, staurolite is thought to be a member of the peak assemblage, which is re-flected in peak P–T estimates of 14–16 kbar at 575–600 �C (Hacker et al., 2003)—slightlylower in pressure than the staurolite-absent garnet–chloritoid–kyanite assemblages reportedabove.

Thus it appears that the diagnostic garnet + chloritoid + kyanite assemblage is relativelywidespread in its occurrence worldwide, and provides a useful indicator of eclogite facies conditionsfrom pelitic rocks, even if mafic rocks are not present.

3.3.1 Metamorphic modeling of garnet + chloritoid + kyanite assemblages

Coggon & Holland (2002) took an average Dalradian pelite composition and constructed a P–T

pseudeosection involving garnet, kyanite/sillimanite, carpholite, chloritoid, biotite, chlorite, stau-rolite, quartz/coesite and H2O between 5–27 kbar and 300–750�C. The calculated equilibria showthat the garnet–chloritoid–kyanite paragenesis is stable over a surprisingly limited region of P–T

space, approximately between and 23–25 kbar and 590–600�C. Therefore, garnet–chloritoid–kyanitephase relationships o↵er an opportunity to tightly constrain P–T conditions of peak eclogite faciesmetamorphism, as seen above in TH–680. Given the number of localities from which garnet–chloritoid–kyanite bearing metasediments have been reported, it is of major importance to inves-tigate the bulk compositional controls on this eclogite facies paragenesis.

A series of P–T pseudosections was constructed, starting with the basic KFMASH systemand progressively incorporating Na2O, CaO, Fe2O3, and MnO to form the extended pelitic modelsystem, MnNCKFMASHO, for type sample TH–680. This was done incrementally to examine bulkcompositional controls on the garnet–chloritoid–kyanite assemblage. Bulk compositions (Table3.2) were calculated following the same methodology as outlined above for the construction of theMnNCKFMASHO pseudosection for TH–680 (Fig.3.6).

52

MnO Na2

O CaO K2

O FeO MgO Al2

O3

SiO2

O

KFMASH – – – 3.66 5.36 3.77 14.28 72.93 –NCKFMASH – 0.24 1.32 3.60 5.27 3.71 14.06 71.80 –NCKFMASHO – 0.24 1.29 3.53 6.53 3.64 13.76 70.31 0.70MnNCKFMASHO 0.09 0.24 1.29 3.52 6.53 3.63 13.75 70.25 0.70

Weight % 0.10 0.22 1.10 5.05 5.64 2.23 21.32 64.2 –

Table 3.2: Reactive bulk compositions used for TH–680 pseudosection calculations using the following modal abundances:

35% quartz, 50% muscovite, 8% garnet, 3% kyanite, 4% chloritoid, <2% epidote; molar % unless otherwise stated.

3.3.1.1 The KFMASH subsystem

The simple KFMASH pseudosection is characterized by the prevalence of trivariant and divariantfields, with only two quadrivariant fields seen by the TH–680 bulk composition. The garnet–chloritoid–kyanite stability field occupies a thin sliver of P–T space in the region, 18–27 kbar and540–610�C. The assemblage’s lower and upper temperature stability is defined by the kyanite-inand by chloritoid-out boundaries respectively. In turn, upper and lower pressure maxima are delin-eated by univariant reactions involving the growth of carpholite and chlorite, both at the expense ofchloritoid. Of particular interest is the expansive, trivariant garnet–kyanite field at higher temper-atures than the garnet-chloritoid-kyanite assemblage. This trivariant (in KFMASH) assemblage,garnet–kyanite–phengite–quartz, is observed in metapelites of the central Alpine, Adula nappe(Meyre et al., 1999; Dale & Holland, 2003). Bulk composition TH-680 is characterised by a con-siderably larger P–T field for the garnet–kyanite–chloritoid assemblage than seen in the averageDalradian pelite of Coggon & Holland (2002)

500 550 60014

16

18

20

22

24

26

28

30TH-680 g ctd ky schist

NCKFMASH + mu + qtz/coe + H2O

g ctd ky

g ky

g ky jdg ky jd car

g ky chl

g ky chl ctd

g chl ctd

g ctd

g ctd car

g car g car kyg car ky ctd

g chl ctd car

g chl ctd car pa

g ctd car pa

g jd car

0.9

0.8 0.7

3.3

3.275

3.25

coeq

a(H20) = 1.0

T C

XFe gSi. c.p.f.u mu

500 550 60014

16

18

20

22

24

26

28

30

T C

P kb

ar

g car

ctd car

car

ctd

g car ctd

g ky car

g ky

g ctd

ctd chl

ctd car chl

g ctd chl

3.3

g ky chl

g ctd ky

coeq

TH-680 g ctd ky schistKFMASH + mu + qtz/coe + H2O

0.75

0.70.8

a(H20) = 1.0XFe gSi mu c.p.f.u.

3.325

3.35

3.375

3.34

Figure 3.9: KFMASH and NCKFMASH pseudosections for rock TH–680.

53

3.3.1.2 The NCKFMASH subsystem

The incorporation of Na2O and CaO into the model system introduces the new phases, jadeiteand paragonite into pseudosection calculations, as well as the use of extended activity-compositionmodels for sodium and calcium bearing phases. The pseudosection di↵ers from that calculatedwithin KFMASH in that it is dominated by quadrivariant fields. The foundations of the KFMASHtopology remain, with a strongly temperature-sensitive garnet–chloritoid–kyanite field situateddown temperature of the large garnet–kyanite domain. However, it can be seen that univariantreactions involving the growth of chlorite and carpholite at the expense of chloritoid, enlarge toform trivariant fields. Paragonite only occurs at low temperatures between 20 and 23 kbar. Thegarnet–chloritoid–kyanite association covers a wider temperature range but smaller pressure rangeas a result of the larger chemical system, but still remains a relatively small stability field in theregion 19–24.5 kbar and 560–590�C. Jadeite is restricted to the highest pressures but occurs insuch low modal amounts that it would probably be undetectable in thin section even if it were tosurvive decompression.

3.3.1.3 The NCKFMASHO subsystem

Expanding the chemical system to include ferric iron and the extra phase epidote in calculationsrequires the use of extended activity models for micas, garnet, chloritoid, chlorite and jadeite.To investigate the e↵ect of ferric iron on pseudosection topology, we use the point-counting valueof O = 0.7, which di↵ers from the value used in Fig.3.6. Major di↵erences to the NCKFMASHtopology comprise the appearance of a low-T garnet stability boundary, the presence of epidote inall fields, except at the highest pressures and temperatures, and further extension of the importantgarnet–chloritoid–kyanite (± epidote) field to slightly lower pressures. The epidote occurs in smallamounts, less than 2 percent modally, in agreement with petrographic observations. Ferric ironin the bulk composition reduces paragonite stability to lower temperatures than shown on thediagram.

3.3.1.4 The MnNCKFMASHO system

The pseudosection was constructed because it is well known that manganese fractionates particu-larly strongly into garnet, chloritoid and carpholite, stabilising these phases in rocks (Mahar et al.,1997b). The amount of manganese added takes into account the fact that garnet growth initi-ated well before attaining peak conditions of interest. Therefore, garnet-rim chemistry provides amore realistic estimate of the reactive bulk composition than core regions. Manganese in epidotewas ignored, because only trace amounts of Mn are found in electron-microprobe analyses of theTH-680 epidotes (< 0.04 wt.%). The principal e↵ect of considering manganese in phase equilibriacalculations is to extend the garnet stability field down-temperature as described by Mahar et al.(1997b) and thus the pseudosection is only rigorously applicable to the stages post-dating initialgarnet growth. Manganese addition to NCKFMASHO can be seen to extend the pressure range

54

500 550 60014

16

18

20

22

24

26

28

30

g ctd ep jd pa

a(H20) = 1.0XFe gSi. c.p.f.u mu

g ctd chl ep

g ctd ep

ctd chl ep

g ky chl ep

g ky ep

g ky

g ky jd

g kyjd carg ky jd car epg jd car ep

g ky jd ctd car ep

g ctd ep jd car

g ctd ep jd

g ctd ky chl ep

ctd chl ep pa

ctd ep jd

jd ctd ep car

jd car ep

g ctd ky ep

TH-680 g ctd ky schistNCKFMASHO + mu + qtz/coe + H2O

0.8

0.75 0.7

3.25

3.275

3.3

3.325

3.35

3.375 coeq

T C

P kb

ar g ctd ky ep jdctd ep jd

g ctd epg ctd ep pa

g ctd ep pa chl

ctd ep chlctd ep chl pa

ctd ep pa

ctd ep pa jd

g ctd ep jd

g ky

g ky jd

g ky ep

g ky ep jd

g ctd ky ep car

g ctd ky ep

g ky ep car

g ky car jd

g ky ep car jd

Figure 3.10: NCKFMASHO pseudosection for rock TH–680.

of the garnet–chloritoid–kyanite(± epidote) field to occupy the region 14–25 kbar and 550–610�C.Di↵erences in topology between Fig.3.12 and Fig.3.6 reflect di↵erent normalized bulk compositionsresulting from slightly di↵erent ferric iron values.

500 550 60014

16

18

20

22

24

26

28

30

0.6

0.65

3.25

3.275

3.3

3.325

3.35

g ctd ep

g ctd ep chl

g ctd ky ep

g ky ep

g ky

g ky jdg ky ep jd

g ctd ky ep jd

g ky ep jd car

g ctd ky ep jd car

g ctd ky ep jd

g ctd ep jd

g ctd ep jd car

g ep jd car

g ctd ep jd pa

g ctd ep chl pa

g ctd ep pa

g ky ep chlg ky ctd ep chl

TH-680 g ctd ky schistMnNCKFMASHO + mu + qtz/coe + H2O

a(H20) = 1.0

T C

P kb

ar

coeq

XFe gSi. c.p.f.u mu

Figure 3.11: MnNCKFMASHO pseudosection for rock TH–680.

55

3.3.2 Bulk compositional controls

The series of P–T pseudesections for TH–680 show that sodium, calcium, ferric iron and man-ganese collectively expand the garnet–chloritoid–kyanite stability field from a thin (⇡15�C wide)field covering 8 kbar of P–T space within KFMASH, to a significant region of eclogite facies P–T ,within the full MnNCKFMASHO system. Although addition of extra components to KFMASH en-larges the field of garnet–chloritoid–kyanite, there are two further bulk compositional controls: theAl2O3:K2O and FeO:MgO ratios of the rocks. Figure 3.12 shows the e↵ect of varying Al2O3:K2Oon the size and position of TH–680’s garnet–chloritoid–kyanite field whilst keeping other composi-tional parameters constant. If this ratio is small, there will be little Al2O3 left over, after formingphengite, to produce much free kyanite. Likewise, larger values of this ratio will produce largeramounts of kyanite. Interestingly, at higher values of Al2O3:K2O, the lower temperature boundaryof the stability field is defined by the appearance of garnet at the expense of chloritoid and kyanite,as opposed to the reaction: chloritoid = garnet + kyanite + H2O at lower values (Al2O3:K2O<5).The typical Dalradian pelite composition has a smaller ratio (4.33) and hence a very small P–T

field for this assemblage, whereas the Betics rock ALM–45 has a large ratio (6.15) and a muchlarger stability field. Importantly, in the Dalradian bulk composition used in Coggon & Holland(2002) the Al2O3 content was adjusted downwards by removing Al2O3 assumed to be combinedwith Na2O and CaO in plagioclase, further reducing the apparent Al2O3:K2O ratio in KFMASHand hence in producing the unreasonably small field for the critical assemblage in that paper.Although Al2O3:K2O appears to be the dominant compositional control on stability-field size, therock’s FeO:MgO ratio also has considerable e↵ect on field position and topology. Figure 3.12a il-lustrates this graphically by varying TH–680’s bulk XFe number (FeO/(FeO + MgO)), with lowervalues shifting the field boundaries to higher temperatures and pressures. This is largely a resultof increasing the clinochlore component of chlorite and the magnesium component of chloritoid,which then stabilizes chlorite and chloritoid to higher pressures and temperatures respectively. TheAFM projection shown in Fig.3.13 shows that all of the garnet–chloritoid–kyanite schists studiedhere lie at lower values of FeO:MgO than the Dalradian pelite.

A comparison of the bulk-rock chemistry for several garnet–chloritoid–kyanite metapelites(Fig.3.13; Table 3.3) collected from the Tauern Window, the Raspas Complex, Ecuador, theCarpathians and the Betics respectively, shows that, relative to an average Dalradian pelite com-position (Mahar et al., 1997b), all have lower or similar manganese contents. This demonstratesthat the assemblage is not formed as a result of anomalously high-manganese bulk compositions,but that small amounts of manganese and ferric iron partitioning into garnet and chloritoid maystabilise the assemblage significantly. Similarly, these rocks are not particularly oxidised, havingferric iron contents comparable with those found in average pelitic compositions (White et al.,2000). However all of the schists have a higher Al2O3:K2O ratio than the average Dalradian pelite.Water activity is taken as unity for all pseudosection calculations—see 3.2.4.3 for discussion. Mod-eling with a reduced water activity of 0.8 shows that temperatures of critical boundaries and P–T

56

500 550 60014

16

18

20

22

24

26

28

30

5.5

3.9 (TH-680)

3.15

coe

q

T �  �  CP

(kba

r)

MnNCKFMASHO+ mu + qtz + H2O

500 550 60014

16

18

20

22

24

26

28

30

coe

q

0.65

0.58

(TH-680)

0.53

MnNCKFMASHO+ mu + qtz + H2O

T �  �  C

P (k

bar)

a b

Figure 3.12: a. E↵ect of variable FeO:MgO ratio on the garnet–chloritoid–kyanite stability field using the bulk compo-

sition of TH-680. Numbers represent bulk XFe (FeO/(FeO + MgO)) value; b. E↵ect of variable Al2

O3

:K2

O ratio on the

garnet–chloritoid–kyanite stability field using the bulk composition of TH–680. Numbers represent bulk Al2

O3

:K2

O ratio

value.

Al2O

3

FeO MgO

Raspas

Eastern Alps

Carpathians

Betics

Dalradian

Figure 3.13: Bulk rock AFM projection from muscovite,

quartz and H2

O, for a range of garnet–chloritoid–kyanite

schists and a Dalradian pelite. Data obtained from the fol-

lowing sources: Dalradian average pelite (triangle)—Mahar

et al. (1997b); the Eastern Alps (hexagon)—Smye et al.

(2010); the Carpathians (square)— Negulescu et al. (2009);

the Raspas complex (star)—Gabriele et al. (2003); the Betics

(ellipse)—this study.

estimates (Fig.3.3.3) only decrease by 15–20�C—a value within the likely error of the calculations.

57

TH � 6801 Raspas2 Carpathians3 Betics4 Dalradian5

SiO2 64.21 76.03 64.97 60.15 59.8TiO2 0.14 0.82 0.83 0.2 –Al2O3 21.33 13.05 17.49 24.23 16.57FeO 5.64 3.74 5.99 8.49 5.81MnO 0.08 0.03 0.07 0.29 0.10MgO 2.23 0.45 2.37 1.39 2.62CaO 1.10 0.18 0.97 0.92 1.09Na2O 0.22 0.68 1.57 0.46 1.73K2O 5.05 1.31 2.64 3.67 3.53

Table 3.3: Comparison of a suite of garnet–chloritoid–kyanite schist bulk compositions with a Dalradian average pelite.

Values are weight % oxide. Superscript notation: 1. TH–680—this study; 2. Raspas— Gabriele et al. (2003); 3.

Carpathians—Negulescu et al. (2009); 4. Betics—ALM-45 metapelite,Smye et al. (2010) ; 5. Dalradian average pelite—

Mahar et al. (1997b).

3.3.3 Comparison of peak P–T conditions

In order to assess the range and similarity in peak metamorphic conditions experienced by theglobal compilation of garnet–chloritoid–kyanite schists described above, a comparative average P–T study was performed using an identical approach to section 3.2.4.3. Mineral data from phaseconstituents of peak, eclogite facies, mineral assemblages were taken from both literature sourcesand freshly obtained electron microprobe datasets (Table 3.1; Table A.1). For consistency aH

2

O = 1and 5% of total iron in muscovite is Fe2O3 were assumed. All calculations used reactions 1–4 de-tailed in section 3.2.4.3.

Average P–T calculations were applied to all sample areas described above except for theKokchetav Massif due to a lack of suitable mineral data for comparison. P–T estimates andassociated error ellipses are presented in Table 3.4 and Fig.3.14 respectively.

Sample Assemblage P �P T �T �fit

Eastern Alps TH-6801 g ctd ky mu q 26.7 2.7 538 17 1.60Carpathians RU0072 g ctd ky mu q 24.6 1.8 590 12 0.67Western Alps unspecified3 g ctd ky mu q 29.9 2.7 544 18 1.63Andes 98RR11, 97Ce54 g ctd ky mu q 20.3 2.8 569 14 1.42Bohemian Massif HH-15 g ctd ky mu q 27.2 2.7 546 16 1.44Betics ALM-456 g ctd ky mu q 20.8 2.0 580 10 0.81

Table 3.4: Results of average P–T calculations. Superscript notation: 1. This study; 2. Negulescu et al. (2009); 3.

Zucali et al. (2002); 4. Gabriele et al. (2003); 5. Konopasek (2001); 6. This study.

Average pressures and temperatures cluster in two distinct groups separated by di↵erences inpressure. Critically, constraints associated with pressure estimates are poor relative to temperatureconstraints (see section 3.2.4.3). Fit parameters for all calculations lie within the 95% confidence

58

400 500 600 700 800

Temperature ᵒC

5

10

15

20

25

30

35

Pres

sure

kba

r

30

60

90

Depth km

7 ᵒC.km

-1

12 ᵒC.km-1

3

2

4

15

6

alm ky

fst

alm ky

fctd

jd qtz

ab

pa

ab ky

jd ky pafctd

jdalm

pa

ta alm chl

coeqtz

Figure 3.14: Comparsion of average P–

T estimates overlaid on pertinent region of

metapelitic NKFASH P–T grid of Wei & Pow-

ell (2006). Shaded field represents stability

region of garnet–chloritoid–kyanite assemblage

with Al2

O3

:K2

O = 5.5 for TH–680. Red dashed

lines represent the range of geothermal gradients

within which the target assemblage is formed.

Note the restricted temperature interval within

which all estimates reside. Error ellipses are 1�

width. Ellipses numbered as follows: 1. East-

ern Alps, 2. Carpathians, 3. Western Alps, 4.

Andes, 5. Bohemian Massif, 6. Betics.

interval (i.e. 1.73 for 4 reactions). Samples 98RR11 and ALM–45 from the Bohemian Massifand the Andes respectively, yield similar peak conditions around 20.5 kbar at 570–80�C and arethe lowest pressure rocks of all samples studied. Although similar in temperature, the pressureestimates for samples from both the Eastern and Western Alps and the Bohemian Massif aresignificantly higher than in the other localities. Interestingly, these estimates lie within the coesitestability field and yet coesite is yet to be reported from any of these localities. Calculations basedon mineral data presented by Zucali et al. (2002) for the Sesia zone eclogitic micaschists producethe highest pressures reported in this study. However, it is doubtful whether analyses presentedin Zucali et al. (2002) and employed in the calculation are from the same rock and hence theresultant estimate is approximate at best. Sample RU007 from the Carpathians yields an averageP–T estimate of 24.6 kbar and 590�C, between the clusters. Importantly, with the exception ofRU007, all P–T estimates lie within 40�C of each other, between 540–590�C. Despite the variationin pressures between samples, peak temperature estimates are better constrained and lie withina narrowly restricted range. This reflects the fact that the geothermometer reaction: fctd =alm + ky + H2O exerts a strong control on average T , and all rocks are rich in alm and fctdcomponents. Furthermore, this grouping of peak temperatures from a range of subduction zonelocalities indicates that temperature attained during subduction may be a critical factor a↵ectingthe depth of the subduction–exhumation transition.

59

3.3.4 Tectonic implications

It is very curious that such a large number of eclogite facies rocks from oceanic crust and passivemargin sequences, both pelitic garnet–chloritoid–kyanite bearing schists as shown here and alsomafic eclogites, seem to record maximum conditions of subduction within the restricted tempera-ture range 550–600�C at 23±3 kbar (see Agard et al. (2009) for review). The following discussionaddresses specifically the subduction of rheologically weak oceanic and shelf material and does notaddress the separate problem of ultra-high-pressure terrains.

Average P–T calculations presented above suggest that thermal grade is a more critical pa-rameter than pressure of subduction. The implication is that subduction tectonics require thedown-going margin slab-top, where both mafic and sedimentary protoliths may coexist, to passthrough a P–T window in the region of 60–75 km depth at temperatures of ⇠550–600�C, withinwhich eclogite facies units are decoupled from the slab and subsequently exhumed. Variations insubduction zone geothermal gradient mean that the slab-top will reach the critical temperaturerange at di↵erent depths, hence the spread of P values between 550–600�C in Fig.3.14. As evi-denced by the relatively uncommon nature of UHP terranes, rocks which are subducted to highergrade are seldom returned to the surface and if so, are commonly metamorphosed portions ofrheologically strong, buoyant continental basement e.g. Dora Maira massif, Western Alps (Chopinet al., 1991).

The question of what critically controls the subduction–exhumation transition in portions oftransitional crust has been approached in several petrographical and experimental studies. Stock-hert et al. (1997) investigated the microstructure of quartz in eclogite facies garnet–chloritoid–kyanite micaschists of the Tauern window and suggested that the mechanism of quartz deformationchanged from being controlled by grain boundary free energy during peak pressure conditions atlow deviatoric stress values, to the regime of dislocation creep during the early stages of exhuma-tion. Their well constrained P–T path shows that the critical assemblage was stable through latestages of subduction and early stages of exhumation i.e. through the period of detachment fromthe slab-top. This is also reported by Gabriele et al. (2003) in garnet–chloritoid–kyanite bearingschists of the Raspas complex, Andes. Together with the thermobarometric work of this study, thispoints towards garnet–chloritoid–kyanite assemblages in aluminous metapelites as representing anenhanced detachment state of subducted material.

Several authors point towards the importance of an interplate shear zone which migrates to-ward zones of lower rheological strength during subduction (Peacock, 1996, 1992; Stockhert et al.,1997). The lateral shift of a weak zone from the plate interface into footwall units would decoupleslab top material from the slab if buoyancy forces are greater than viscous-drag forces associatedwith subduction (e.g. England & Holland (1979)). More recently Jolivet et al. (2005) propose thatdecoupling of eclogite units from the slab occurs as a result of the progressive formation of shearzones synchronous with eclogitization and associated drop in rock strength. Based on observationsfrom the Western Gneiss Region, Norway, they show how a gradual succession from brittle fractur-

60

km0

50

100

kbar0

10

20

30

40

qtzcoe

continental crust

sediments

detachment of eclogitesfrom slab

600 C500 Cpassive margin

lithospheric mantle

exhumation of eclogitesvia return flow

Figure 3.15: Schematic showing

tectonic setting for the exhumation

of eclogites from a subducted con-

tinental margin. Isotherms and as-

sociated kinematic indicators taken

from Gerya & Stockhert (2006)

(model HYAA—15 Ma post subduc-

tion initiation) delineate the thermal

weakening threshold where detach-

ment of material from the slab oc-

curs.

ing, fluid infiltration and recrystallisation leads to localisation of strain in ductile shear zones andeventual decoupling of units along such zones. Rheologically weaker material would decouple fromthe slab-top at lower threshold values than those experienced by the Bergen eclogites. This hy-pothesis is consistent with the 2–D numerical modelling study of Carry et al. (2009), which pointsto development of two separate strength gradients evolving within the slab during subduction: oneparallel to the length of the slab where the ductile strength decreases with depth and a secondone normal to the slab where the ductile strength increases with depth as a result of the slab’stemperature distribution. They conclude that detachment of individual HP–LT units occurs whenthe subducted slab’s strength becomes lower than the applied net stress. Importantly, Carry et al.(2009), along with the earlier work of van den Beukel (1992), show that thermal weakening of theslab-top is a critical control on when detachment occurs.

Locality Setting References

Eastern Alps Passive margin sequence Holland (1979), Kurz et al.(1998), (1999)

Western Alps Austroalpine continental margin Compagnoni et al. (1997), Com-pagnoni (1997), Dal Piaz et al.(1972), Stella (1894), Tropper etal. (1999)

Carpathians Subduction melange complex Balintoni et al. (2009), Negulescuet al. (2009), Sabau (2000)

Andes Composite fragment of oceanic lithosphere Feininger (1980), Gabriele(2002), Gabriele et al. (2003)

Bohemian Massif Passive margin sequence Hofmann et al. (1988), Krohe(1996), Konopasek (2001)

Betics Passive margin sequence Akkerman et al. (1980), Weijer-mars (1991), Jolivet et al. (2003)

Table 3.5: Comparison of tectonic settings for garnet–chloritoid–kyanite schist localities.

61

The recent thermomechanical modeling studies of active continental margins conducted byStockhert & Gerya (2005) and Gerya & Stockhert (2006), produce decompressive P–T paths forslab–top material which support the common 550–600�C and 23±3kbar subduction maximumdiscussed above (Fig.3.15). Commonly such models predict peak temperatures which are coolerthan those deduced from thermobarometry. Importantly, peak temperature attained within theslab–top is largely a↵ected by convergence rate. The results of this study should help constrainsuch models by fixing the depths and temperatures where detachment of the crustal units occurs.

Thus, it appears that the strength of the subducted lithosphere is the critical factor in deter-mining the transition from subduction to exhumation of HP material. A comparison of protolithenvironments for the garnet–chloritoid–kyanite schists discussed above, shows that such rocks arecommonly associated with mafic eclogites and interpreted as representing supracrustal passivemargin sequences (Table 3.5). Given the coexistence of similar lithologies, it is likely that thenet strength, resultant from a number of variables including crustal thickness, subduction angle,convergence rate and thermal architecture, will be comparable across a wide range of collisionalsettings.

3.4 Glockner nappe

The Glockner nappe represents a dismembered ophiolite suite pertaining to the Valais ocean(Miller, 1977; Hock & Miller, 1980; Bousquet et al., 2002). It is dominated by the petrologicalsignature of Barrovian metamorphism and, in contrast to the Eclogite Zone, displays limited ev-idence for HP metamorphism. As the nappe directly overlies the Eclogite Zone, its P–T path isof direct tectonic importance—potentially constraining the conditions under which the Penninicnappe stack was assembled.

3.4.1 Previous work

The Glockner nappe has received less attention than the Eclogite Zone. Holland & Ray (1985)reported occurrences of retrogressed crossite and jadeitic pyroxene in banded epidote-schists fromthe northern slopes of the Virgental valley, close to the Bonn-Matreier Hutte. Jadietic pyroxenebreaks down to a zonal symplectite of albite + hematite + actinolite, whereas crossitic amphiboleforms a sequence of albite + actinolite + magnetite ± chlorite ± talc. Using the experimentallydetermined P–T position for the reaction: jadeite45 + quartz = albite (Holland, 1980) and thecalculated P–T position (Holland & Powell, 1985; Powell & Holland, 1985) for lawsonite + albite= zoisite + paragonite + quartz + H2O, the authors semi-quantitatively estimate peak pressuresof >8 kbar.

Dachs & Proyer (2001) documented exposure of retrogressed eclogites embedded in calcareousmicaschists and greenschists of the Glockner and Rote-Wande nappes in the Großglockner region,to the East of the Eclogite Zone. Pre-eclogite facies assemblages comprising: chlorite + actinolite

62

+ plagioclase, or glaucophane + paragonite + clinozoisite, are preserved in the cores of eclogiticgarnets. Conventional thermobarometry shows that the peak-metamorphic paragenesis of garnet +omphacite + paragonite + glaucophane + clinozoisite + quartz + rutile ± phengite records condi-tions of ⇠ 17 kbar and 570�C—close to those reported from the Eclogite Zone. High-P assemblageswere retrogressed via hydration and symplectite formation during the Barrovian overprint at 5–6kbar and 500–530�C.

Gleißner et al. (2007) dated glaucophane-bearing lawsonite pseudomorphs from the upperGlockner nappe. They semi-quantitatively interpreted the peak blueschist facies assemblage (gl +law + ep + chl+ sph + q) to represent conditions of 7–11 kbar at 400–500�C. The pseudomorphassemblage comprises albite + epidote + phengite + chlorite + calcite + actinolite ± quartz ±apatite. Subsets of this assemblage yield Rb–Sr isochron ages of 29.82±0.54 Ma, which is inter-preted to represent syntectonic decomposition of the lawsonite assemblage and crystallisation ofthe greenschist fabric.

Importantly, these studies show that HP metamorphism a↵ected the upper structural levelsof the Penninic nappe pile. Given the weak P–T constraints on HP metamorphism from theGlockner nappe immediately juxtaposed with the Eclogite Zone, the following section focuses oncalculating a HP–Barrovian P–T path from the crossite + jadeite breakdown textures reportedby Holland & Ray (1985).

3.4.2 Field relations

Pre-Barrovian HP symplectites outcrop in a⇠1000 meter-thick sequence of lawsonite-pseudomorphbearing greenschists, banded epidote-schists and calcareous mica-schists. Horizons are laterallycontinuous over 1–2 kilometers. The dominant foliation, S1+S2, dips moderately to steeply (50–80�) to the south, trends approximately east–west and is defined by alignment of sheet silicates.The three dominant lithologies are interfolded by isoclinal folds with axial planes parallel to S2.Crossite and jadeite pseudomorphs occur exclusively in the banded epidote-schists, which are de-void of lawsonite pseudomorphs and contain subsets of the assemblage: chlorite + epidote + albite+ actinolite + magnetite + phengite + calcite + quartz+ biotite.

Crossite breakdown textures are recognisable as ⇠5–20 millimetre deep-blue needles, whichare aligned to form a transposed L1 lineation on S1+S2 foliation planes of epidote-rich horizons.Pyroxene pseudomorphs are less common and form millimetre-scale spongy, grey patches, whichare best exposed on weathered foliation surfaces. The S1+S2 fabric anastomoses around bothcrossite and jadeite breakdown textures. Similarly, in the more competent epidote greenstonelayers, a crenulated S1 fabric, defined by albite clots, wraps lawsonite pseudomorphs. Collectively,this shows that HP mineral growth in the Glockner nappe occurred prior to or during the earlystages of D1. In the Eclogite Zone blueschist facies conditions prevailed after D1 and during theearly stages of D2, as evidenced by sample TH-672 (Holland, 1977): a banded zoisite eclogite,which shows glaucophane laths growing in the hinges of F2 generation crenulations (Holland &

63

Richardson, 1979).

Figure 3.16: Glockner nappe blueschists. a. Field context; south-dipping nappe pile of intercalated calc- (CS) and

epidote schists (ES). Photograph taken looking west from the Bonn Matreier Hutte; field of view is ⇠200m wide; b.

Ptygmatically folded (D2

) epidote rinds with horizons of epidote schist. Glaucophane needles are abundant on epidote

foliation planes. Field of view is 1 metre; c. Close-up of relict glaucophane needles on epidote foliation plane; note

their preferential alignment which forms a composite L1

+L2

lineation; d. Sample 221c (Holland & Ray, 1985); note the

cm-scale quartz veins; jadeite present as spongy grey patches; e. Sample N45b Ray (1986); note abundant glaucophane

needles and crenulated S1

+S2

fabric.

3.4.3 Petrology of breakdown textures

As the petrology of these rocks has been presented in detail by Holland & Ray (1985), the followingserves as a brief synopsis. Jadeite and crossite breakdown textures are typified by the samplesuite presented by Holland & Ray (1985); Ray (1986)—crossite symplectites, in various stages offormation, are present in all selected samples (N45b, B1, 221D, B3, B2 and 221C), whereas jadeitesymplectites occur in only three of the five samples (221C, B2 and B3).

Crossite textures are typified by sample 221D, which shows abundant needles of blue amphiboleembedded in rafts of paragonite. Crossite needles are ⇠5–10 mm long and have partially brokendown to a rim of actinolite and a broad (1–5 mm) shoulder of albite+chlorite, in that order, whichseparates the blue amphibole from matrix paragonite (Fig.3.17); chlorite is subordinate relative toalbite. Octahedra of magnetite and flakes of minor talc are present in the zones of albite+chlorite.Outside of the crossite symplectites, the matrix is formed by epidote, phengite, hematite, sphene

64

and minor calcite and quartz. These phases do not appear to have been involved in the breakdownreactions.

Jadeite breakdown textures are best preserved in sample 221C, which di↵ers from samples B2and B3 in that it contains pure sodic pyroxene as opposed to pseudomorphs of albite+hematite.Sample 221C contains the following assemblage: crossite + pyroxene + actinolite + phengite +albite + epidote + hematite + sphene + quartz + calcite. Pyroxene symplectites are cored byblocky jadeite (3–7 mm width), which is rimmed by a broad region of albite+hematite; fine aciclesof colourless actinolite are included in the outboard regions of the symplectite (Fig.3.17). Nomagnetite or chlorite are present in the breakdown textures and phengite is the dominant whitemica.

Figure 3.17: Photomicrographs of crossite (a) and jadeite (b) breakdown textures in rock 221C (Holland & Ray, 1985;

Ray, 1986). Fields of view are 1.5 and 2.6 mm for a and b respectively; photomicrograph a is taken using plane polarised

light wheras b is taken under crossed-polars.

3.4.4 Mineral chemistry

Pre-existing electron microprobe mineral analyses (Ray, 1986) are augmented by analyses of am-phibole, pyroxene, muscovite, albite and epidote obtained from sample N45b, performed at theUniversity of Cambridge (Table A.2). Amphibole, pyroxene and muscovite compositional trendsare presented in Fig.3.18. A brief description of the salient chemical trends follows—for a moredetailed presentation, see Ray (1986) and Holland & Ray (1985).

Na-amphibole occurs in all samples in various stages of decomposition and displays strongpleochroism from pale-mauve to dark blue. Fe2+ and Fe3+ concentrations (1.3–1.6 c.p.f.u. and0.3–0.7 c.p.f.u., respectively, using the ferric recalculation method of Holland & Blundy (1994) and23 oxygens) show that the amphibole is crossitic with little (<10%) chemical variation.

Ca-amphiboles form as breakdown products of both pyroxene and Na-amphibole. Figure 3.18ashows that most amphiboles are actinolitic in composition, although considerable variation to more

65

barroisitic and Al-rich (hornblende) compositions is present both within and between samples.Na-pyroxene is present as unaltered cores to symplectites of albite+hematite+actinolite in sam-

ples 221c and N45b. Electron microprobe analyses show show that the pyroxene is an intermediatejadeite–acmite solid-solution—16 spot analyses from 221c and N45b yield an average compositionof jadeite44, acmite45 and diopside11 (Fig.3.18c). Jadeite–acmite variation is <10%.

Paragonite occurs in crossite-bearing schists where it is the dominant mica (>35 volume per-cent) and shows mottled interference colours—microprobe data show that it contains no Mg2+.Minor paragonite (<2 volume percent) is present in pyroxene-bearing N45b.

Phengite is present in all samples; in pyroxene-bearing schists it is the dominant mica. Sicontents range between 6.6 and 6.8 c.p.f.u. (22 oxygens) for all samples. Fig.3.18b shows thatphengites generally have <50% celadonite component.

Epidote is ubiquitous in all rocks and lies between 91–94% epidote end-member. Texturally, itshows little evidence for involvement in the post-blueschist metamorphic evolution.

Pyroxene-bearing schists contain hematite, whilst crossite schists have magnetite as the domi-nant oxide phase, in addition to minor hematite (<0.5 vol.%). Ray (1986) reports that hematitecontains upto 7 wt.% TiO2, whilst magnetite is TiO2 free.

Albite is pure (>99%) NaAlSi3O8 end-member and is the dominant constituent of both crossiteand pyroxene breakdown symplectites.

3.4.5 Metamorphic modeling of gl+jd assemblages

Coexisting sodic amphibole and pyroxene are characteristic of blueschist facies metamorphism(Liou & Maruyama, 1987; Maruyama & Liou, 1988). The breakdown textures observed here areinterpreted to have formed during the transition from HP to Barrovian conditions, concomitantwith assembly of the Penninic pile (Holland & Ray, 1985). Therefore, these symplectites potentiallypreserve an important P–T record of the Glockner nappe during continental collision.

From petrographic observations alone, it is not clear to what extent the textures preservechemical equilibrium, making conventional thermobarometry di�cult. The following phases ap-pear to have been stable during peak HP conditions: glaucophane + jadeite + epidote + phengite+ paragonite/phengite + sphene (+ quartz + magnetite), before being variably reworked andoverprinted by the Barrovian assemblage: actinolite + albite + chlorite + epidote (+ biotite +sphene + calcite + hematite + magnetite). The principal aim of the pseudosection modeling pre-sented beneath is to use a newly calibrated NCFMASO clinopyroxene model to more preciselyconstrain peak conditions of HP metamorphism than the semi-quantitative estimate of Holland &Ray (1985): >8 kbar and 350–450�C.

Phase diagram calculations were performed with Thermocalc version 3.33 (Powell & Hol-land, 1988), using the November 2003 updated version of the Holland & Powell (1998) data set.All calculations were completed in the fully expanded mafic system: NCKFMASHTO. Phases con-

66

2.0

1.5

1.0

0.5

0.0

Na

M4

2.01.51.00.50.0

Al[IV]

Ca amphiboles

B1

B2

B3

221D

Na amphiboles

B1

B2

B3

N45b

221C

221D

Barroisite

HornblendeTremolite Tschermakite

Winchite

Glaucophane

NaM4 SiIV

CaM4 AlIV

2.0

1.5

1.0

0.5

0.0

2.01.51.00.50.0

gl schists

221d

gl+jd schists

B2

B3

N45b

221c

law schists

147b

160a

165g

MgS

iAl -2

NaFe3+Si2O

6

NaAlSi2O

6

Ca(Mg,Fe2+)Si2O

6

N45b

221c

jadeite

acmite diopside

omphacite

aegirine-augite

Al[VI]-2

a. b.

Al[I

V]

c.

Figure 3.18: Amphibole, pyroxene and mica chemistry using electron microprobe data obtained from sample N45b

and data from Ray (1986): a. Composition of Ca and Na amphiboles plotted as formula proportions (23 O) of NaM4

against Al[IV ]; b. Phengite solid solution in matrix muscovite (22 O); c. Acmite–jadeite–diopside ternary plot of pyroxene

compositions from rocks 221c and N45b.

sidered in the calculations and corresponding a–X models used are as follows: garnet and biotite(White et al., 2007), epidote and talc (Holland & Powell, 1998), chlorite (Holland et al., 1998),white mica (Coggon & Holland, 2002; Smye et al., 2010) and plagioclase-K-feldspar (Holland &Powell, 2003). Albite, lawsonite, rutile, sphene, hematite, spinel, magnetite, ilmenite, quartz andaqueous fluid (H2O) are treated as pure phases. Clinopyroxene and amphibole a–X models arediscussed beneath.

67

3.4.5.1 a–X models

The chains of Si–O tetrahedra, common to both amphibole and clinopyroxene structures, meansthat a wide variety of ionic substitutions can occur (Deer et al., 1992). This leads to complex a–X

composition relationships, which makes phase equilibrium calculations in mafic systems challeng-ing. Recent development of a–X models for both clinopyroxene (Green et al., 2007) and amphibole(Dale et al., 2005; Diener et al., 2007) means that pseudosections for mafic rocks can now be cal-culated in the extended NCKFMASHTO system (Stıpska & Powell, 2005; YANG et al., 2008;Phillips et al., 2008; Wei et al., 2009; Daczko & Halpin, 2009; Rebay et al., 2010; Wei & Clarke,2011). Despite the fact that these models have been calibrated from data over a limited window ofmineral compositional space, such calculations often accurately reproduce natural data (e.g. Rebayet al., 2010; Wei & Clarke, 2011).

However, the NCFMAS clinopyroxene model of Green et al. (2007) predicts that diopsidicclinopyroxene is stable under blueschist facies conditions (Diener et al., 2007). For a MORBcomposition (McDonough & Sun, 1995), Diener & Powell (In Press) show that diopside with XM2

Na

= 0.2–0.3 is stable over a P–T band ⇠1.5 kbar wide at P above 10 kbar and T above 425�C.Omphacite (XM2

Na = 0.4–0.5) occurs to higher P and lower T , but it becomes increasingly morediopsidic. The models also predict that with increasing Fe3+, diopside becomes more stable thanomphacite between 11–13.7 kbar at intermediate values of XFe3+ (<0.54; Diener & Powell, InPress). Under reduced conditions (XFe3+<0.4) diopside coexists with omphacite at low P and isstable without omphacite above 16 kbar; omphacite reaches a maximum XM2

Na of 0.47. Critically,the models do not predict jadeite to be stable anywhere between 4 and 20 kbar for a MORB withbetween 0 and 3.97 molar % O.

Clearly, these calculated clinopyroxene compositions do not adequately match natural blueschistphase relations, where omphacite is commonplace (Ernst, 1973a; Okay, 1980; Carpenter, 1982). Asa result, the clinopyroxene model has been recalibrated to make omphacite more ferric than diop-side under ferric conditions (Diener & Powell, In Press). This involves changing the interactionenergies involving acmite to be similar to those involving jadeite. There is a lack of natural datawhich can be used to constrain interaction energies between the Fe3+ (acmite and an ordered Al-Fe3+ end-member, jac) and the remaining end-members (diopside, jadeite, omphacite, hedenbergiteand an ordered Fe-Mg end-member, cfm).

The absolute stability of clinopyroxene and its propensity to sequester available Fe3+ over co-existing phases is determined by thermodynamic data of its end-members. Pertinently, the recali-bration of the clinopyroxene model involves an adjustment to the enthalpy of formation (�fH) ofacmite by -4 kJ.mol�1 from the value presented by Holland & Powell (1998). This adjustment repro-duces the relative order of Fe3+ partitioning documented by Holland & Ray (1985) in the Glocknernappe blueschists discussed here: jadeite>omphacite⇠glaucophane⇠epidote>diopside⇠actinolite(where, Fe3+ partitions as: Fe3+/(Fe3++Al)).

The amphibole model of Diener et al. (2007) has also been modified to incorporate Fe3+ via the

68

introduction of an additional end-member: magnesioriebeckite; its interaction energies are takenin proportion to those of glaucophane (Diener & Powell, In Press). Clinoamphibole end-member�fHs were adjusted from those adopted by Diener et al. (2007), by the following increments:gedrite, 20 instead of 23.5 kJ.mol�1; ortho-Mg-pargasite, 25 instead of 28.5 kJ.mol�1.

Jadeite and glaucophane phase relations are calculated beneath using the recalibrated a–X

models discussed here. The validity of the circular logic employed in calibrating �fHacmite usingthe jadeite + glaucophane Glockner assemblages relies heavily on an accurate interpretation of theorder of intensity for Fe3+ partitioning between coexisting mafic phases.

3.4.5.2 Estimating a reactive bulk composition

Millimetre-scale symplectite cores of unaltered jadeite and glaucophane (Fig.3.17) show that blueschist-breakdown reactions failed to reach completion and do not preserve a state of chemical equilibriumon lengthscales greater than the textures themselves (e.g. Whitney & McLelland, 1973; Ashworth& Birdi, 1990; Johnson & Carlson, 1990). An accurate bulk composition, responsible for genesis ofthe Glockner blueschists, must predict coexisting glaucophane and jadeite (0.45–0.5 jd; 0.45–0.5 ac)to be stable with muscovite, epidote, sphene and hematite/magnetite under blueschist conditions(ca. 8–15 kbar; 300–450. Blueschist pseudosection calculations presented here are not based on asingle sample; rather, they are based on the salient phase relations of the sample suite reported byHolland & Ray (1985) and Ray (1986).

Preliminary calculations detailed above show that the NCFMAS clinopyroxene model of Greenet al. (2007) incorrectly partitions Fe3+ between coexisting pyroxenes. Accordingly, calculationsherein employ the recalibrated a–X model of Diener & Powell (In Press).

In the full NCKFMASHTO system, required to describe all the phases of interest, mafic phaserelations are extremely sensitive to bulk compositional controls (e.g. Rebay et al., 2010). A series ofP–T pseudosections is calculated for a range of bulk compositions, starting with an oxidised MORB(Table 3.6; Rebay et al., 2010) and incrementally moving toward a whole-rock XRF compositionderived from sample 221c: a glaucophane + jadeite-bearing schist documented by Ray (1986). Cal-culations are performed between 3–20 kbar, 300–600�C under fluid- and SiO2-saturated conditions.Figure 3.19 shows the P–T stability fields of Ca-amphibole- and lawsonite-absent glaucophane +jadeite-bearing assemblages; bulk compositions are detailed in Table 3.6.

Mineral compositions are discussed in terms of compositional parameters used in thermocalc;clinopyroxene parameters justify explanation. According to the model of Green et al. (2007), Fe3+

and Al are permitted to order on M1a and M1m sites. Clinopyroxene Fe3+:Al is given by f(cpx),which is determined by: (XM1a

Fe+3

+XM1mFe+3

)/((XM1aFe+3

+XM1mFe+3

+(XM1aAl +XM1m

Al ). Similarly, Na and Caorder on the M2c and M2n sites; bulk Na content is defined by j(cpx): (XM2c

Na +XM2nNa )/2. Under

blueschist temperatures <450�, clinopyroxene will be highly ordered and proportions of acmite and

69

jadeite end-members are defined by f(cpx)·j(cpx) and j(cpx)·(1�f(cpx)) respectively. Clinopyroxenepresent in Glockner blueschists have values of f(cpx) between 0.45 and 0.50, and j(cpx) between0.85 and 1.

300 350 400 450 500 550 6003

4

5

6

7

8

9

10

11

12

13

14

15

16

17

18

19

20

0.48

a. MORB

o+gl+mu+ep+chl+ru/sph

0.36

0.49

0.47 + act

+ law

300 350 400 450 500 550 6003

4

5

6

7

8

9

10

11

12

13

14

15

16

17

18

19

20

0.9

0.68

300 350 400 450 500 550 6003

4

5

6

7

8

9

10

11

12

13

14

15

16

17

18

19

20

0.63

300 350 400 450 500 550 6003

4

5

6

7

8

9

10

11

12

13

14

15

16

17

18

19

20

0.88

b. Incremental bulk 1

o+gl+mu+ep+chl+ru/sph

c. Incremental bulk 4

gl+jd+mu+ep+hem+rud. 221c XRF

gl+jd+mu+ep+hem+ru

+ act

+ law

+ law

+ ab + ab

+ ab + ab

+ hb+ hb

+ law

Pre

ssur

e (k

bar)

Temperature�  (˚C)

0.37

0.34

0.62

0.61

0.60

0.59

0.530.52

0.51

0.5 0.49 0.48

0.47

0.86

0.84

0.82

0.80

0.78

0.76

0.74

0.720.70

0.68

0.86

0.84

0.82 0.

80 0.78 0.

76

0.64 0.88

0.86

0.84

0.82

0.80

0.78

0.76

0.74

0.720.70

0.680.66

0.86 0.

84 0.82 0.

80 0.78 0.

76 0.74 0.72

0.78 f(cpx)0.78 j(cpx)

NCKFMASHTO+q H2O

Figure 3.19: Calculated P–T stability fields for coexisting Na-pyroxene- and amphibole-bearing assemblages in the fully

expanded NCKFMASHTO system. Specific bulk details are detailed in Table 3.6. a Oxidised MORB bulk composition of

Rebay et al. (2010); b Incremental bulk 1; c Incremental bulk 4; d Whole-rock XRF bulk for sample 221c (Ray, 1986).

Fields are contoured for f(cpx) (red) and j(cpx) (blue) parameters—see text for discussion.

For a MORB composition, sodic amphibole and pyroxene coexist over an ⇠5 kilobar-wide

70

sliver of P–T space at 400�C, which tapers towards higher-T and -P , before omphacite reacts withglaucophane and actinolite to form garnet at ⇠550�C above 17 kbar. Lawsonite and actinoliteare stable throughout much of the blueschist facies, leaving only a thin (3 kbar at 400�) wedge ofquadrivariant P–T space (Fig.3.19a) for the assemblage: omphacite + glaucophane + muscovite +epidote + chlorite + rutile/sphene (+ quartz + H2O)—the most similar assemblage in the MORBsection to the Glockner blueschists.

Incremental bulk 1 (Fig.3.19b) yields a slightly expanded P–T field for the assemblage: om-phacite + glaucophane + muscovite + epidote + chlorite + rutile/sphene (+ quartz + H2O)assemblage, compared to the MORB section. This is because the increase in Fe3+ (O increasesfrom 1.5 to 2.08), results in a more acmite-rich pyroxene (f(cpx) = 0.51–0.63), which is stable atlower-P than ferric-poor compositions. Similarly, the decrease in bulk CaO content (12.43–11.22)reduces lawsonite and actinolite stability to higher- and lower-P respectively.

Incremental bulks 2 and 3 are characterised by the prevelance of quadrivariant and trivariantactinolite- and lawsonite-bearing fields. Respective topologies are not detailed on Fig.3.19 as sodic-pyroxene and glaucophane do not coexist in the absence of lawsonite and actinolite, as required.

Bulk composition 4 yields a large field of glaucophane + jadeite + muscovite + epidote +hematite + rutile, which spans 400–600� and 6–20 kbar (Fig.3.19c). The assemblage closelymatches those preserved in the Glockner nappe samples. Chlorite is no longer stable above 450�Cand 7 kbar; the assemblage is bounded by the occurrence of lawsonite and albite toward lower T

and lower P , respectively. Towards higher T , the Na-Ca amphibole solvus occurs between 500–600�

and 8–16 kbar. The high concentration of Fe3+ (O = 3.83) stabilizes 3–4 modal % hematite acrossthe section. Pyroxene is now a jadeite–acmite solution with values of j(cpx) and f(cpx) between0.86–0.64 and 0.88–0.68 respectively; compositions become more diopsidic up T .

The whole-rock XRF bulk composition of sample 221c yields an identical assemblage to bulk4, which is stable over a similar P–T interval, between 375–600�C and 5–20 kbar.

Despite the large range of investigated P–T–X space, the pseudosection calculations fail topredict clinopyroxene of similar composition to Glockner-nappe jadeite: j(cpx) is too low, spanning0.49–0.86 across the entire P–T–X range, whilst f(cpx) is highly variable between 0.34–0.9. Clearly,this indicates that the clinopyroxene model of (Diener et al., 2007) requires further calibration.

3.4.5.3 Breakdown reactions

The following section focuses on the reactions responsible for the formation of jadeite and glauco-phane breakdown symplectites during transition from blueschist–greenschist facies conditions.

Glaucophane breakdown reactionsIn the NCKFMASHTO system, the phases: glaucophane, actinolite, muscovite, paragonite, epi-dote, chlorite, rutile, sphene, albite, hematite, magnetite, quartz and H2O, yield the following

71

NaO CaO K2O FeO MgO Al2O3 SiO2 TiO2 O

MORB? 2.66 12.43 0.23 8.29 12.93 9.26 53.4 1.07 1.5bulk 1 2.84 11.22 0.73 8.73 11.19 9.71 53.44 1.47 2.08bulk 2 3.02 10.02 1.23 9.17 9.45 10.16 53.47 1.88 2.66bulk 3 3.19 8.81 1.74 9.61 7.71 10.62 53.51 2.28 3.25bulk 4 3.37 7.61 2.24 10.05 5.97 11.07 53.54 2.69 3.83221c† 3.55 6.4 2.74 10.49 4.23 11.52 53.58 3.09 4.41gl bkd. 5.28 8.61 0.35 8.35 2.89 8.68 53.60 0.65 2.70jd bkd. 6.06 7.98 0.28 11.52 2.72 14.10 54.28 0.87 4.18peak 3.79 6.48 2.89 8.28 4.59 13.87 57.59 0.29 2.19

Table 3.6: Molar-percent blueschist bulk compositions. ? Oxidised MORB composition of Rebay et al. (2010);† whole-

rock XRF composition from Ray (1986). Incremental bulks (1–4) separate MORB and 221c.

univariant (F=1) reactions, which account for the breakdown of glaucophane between 5–10 kbarand 400–550�C:

1.0 gl+0.36 pa+0.43 ep+0.01 mt = 0.98 act+0.02 mu 0.85 ab+0.15 hem+1.08 q+0.59 H2O [1]

1.0 gl+0.03 pa+0.22 chl+0.66 ep+0.15 mt = 1.38 act+0.002 mu+0.16 hem+1.57 q+ 0.88 H2O [2]

1.0 gl+1.06 pa+0.69 hem = 0.05 mu+0.38 chl+0.03 ep+3.01 ab+1.23 mt+1.36 q+2.07 H2O [3]

Mineral abbreviations follow Holland & Powell (1998). The chlorite-absent ([1]) univariantclosely matches the petrography of glaucophane breakdown symplectites, in which actinolite, albiteand hematite are formed at the expense of glaucophane. Paragonite and epidote are both requiredas reactants, which fits with their modal abundances in glaucophane schists (> 35% and >15%respectively). All reactions involve significant dehydration (release of greater than half the molarproportion of reactant glaucophane, H2O); this is contrary to the classic blueschist–greenschisttransition in NCMASH: 25 gl + 6 ep + 7 q + 14 H2O = 6 act + 50 ab + 9 chl (e.g. Will et al.,1998).

To model the breakdown of glaucophane via the chlorite-absent univariant recation, an NCKF-MASHTO pseudosection was calculated between 350–550�C and 3–18 kbar, for a bulk compositionconstructed using the following mineral proportions in conjunction with phase compositions fromthe univariant: 6 glaucophane, 10 paragonite, 7 epidote, 0.5 magnetite, 1, sphene, 1 quartz and0.5 muscovite. The resultant bulk composition is shown in Table 3.6 as gl bkd.

The pseudosection (Fig.3.20a) is dominated by trivariant assemblages. The blueschist–greenschist(lower amphibolite) facies transition is marked by a thin array of bivariant assemblages, whichmarks the appearance of albite and the expense of glaucophane, to lower P . Lawsonite is stableabove 10 kbar at 350�C and 18 kbar at 500�C; muscovite, paragonite, epidote, sphene, quartz andH2O are present throughout the P–T range. The chlorite-absent univariant runs from a full-system

72

invariant point at 7.3 kbar and 475�C to the limit of the examined P–T space at 550�C and 10.3kbar.

The topology of the pseudosection accurately explains the formation of albite, actinolite andmagnetite at the expense of glaucophane, hematite and paragonite. XCa

M4 isopleths in actinolite(dashed-green contours: Fig.3.20a) provide a strong constraint on the pressures of equilibrationduring Barrovian metamorphism. Average XCa

M4 values for actinolites proximal to glaucophanesymplectites are between 0.7–0.8 (1.4–1.6 c.p.f.u. Ca), which suggests pressures between 6–7 kbarat 550�C—the mid-point of Barrovian temperature estimates. Blueschist conditions are less wellconstrained. The bulk predicts Si-in-muscovite values between 3.12–3.2 c.p.f.u. for lawsonite-absent blueschist-facies assemblages (dashed-red contours: Fig.3.20a); however, the average phen-gite value measured from samples 221d and N45b—both glaucophane schists, is between 3.3–3.4c.p.f.u. This suggests that either the bulk composition is inaccurate, or, given the matchingpredicted and observed Ca-amphibole compositions, that the breakdown reactions were not iso-chemical. Furthermore, the model bulk composition does not account for the presence of chloritein reaction symplectites. Chlorite is stable down temperature of the full-system invariant pointwhere the actinolite-absent univariant reaction is responsible for glaucophane breakdown. Giventhe small quantities of chlorite (<5 volume %), this may be a result of sub-millimetric hetero-geneities in Fe2+, Mg and/or Al.

Jadeite breakdown reactionsJadeite decomposition was investigated independently from the reactions responsible for glauco-phane breakdown. The following univariant reactions were calculated in NCKFMASHTO:

1.0 jd+0.44 pa 0.02 mt+0.75 q = 0.07 act+0.03 mu+0.08 ep+1.08 ab+0.28 hem+0.3 H2O [1]

1.0 jd+0.42 pa+0.03 mt+0.82 q = 0.07 gl+0.03 mu+0.1 ep+1.04 ab+0.28 hem+0.27 H2O [2]

Reaction 1 matches jadeite-breakdown phase relations inferred from petrography and describedabove. However, pyroxene-bearing schists contain little or no paragonite (i.e. 221c and N45b),which suggests that either paragonite was originally present and has been fully consumed by thereaction, or, a di↵erent reaction, similar to reaction 1, occurred, but did not involve paragonite. Ifreaction 1 operated, the distribution of paragonite would account for the spatially variable extentof pyroxene decomposition. Similar to glaucophane-breakdown reactions, reactions 1 and 2 involvethe release of H2O. Furthermore, both reactions corroborate textural evidence that epidote andmuscovite were not involved in the post-blueschist P–T evolution.

A bulk composition was calculated such that it would ‘see’ the glaucophane-absent univariantreaction, using the following mineral modes combined with compositions from the univariant: 1glaucophane, 1 jadeite, 3 paragonite, 1 muscovite, 1 epidote, 0.3 hematite and 0.1 sphene. Thebulk composition (jd. bkd: Table 3.6) is more ferric- and aluminium-rich (Fe2+:Fe3+ = 0.75 versus

73

350

375

400

425

450

475

500

525

550

3456789101112131415161718

gl c

hl h

em la

w

gl m

t hem

chl h

em a

bch

l hem

mt a

bac

t mt a

bch

l mt a

b

act m

t ab

hem

NC

KFM

AS

HTO

+ m

u pa

ep

sph

q H

2O

0.90.

8

0.70.

6

gl c

hl h

em

3.2

3.19

3.18

3.17

3.16

3.15

3.14

3.13

3.12

Pressure (kbar)

Temperature�  (˚C)

400

425

450

475

500

525

550

3456789101112131415161718

gl jd

hem

law

gl jd

hem

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gl jd

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0.89

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0.810.800.79

NC

KFM

AS

HTO

+ m

u pa

ep

sph

q H

2O

Pressure (kbar)

Temperature�  (˚C)

0.9

0.8

0.7

0.6

ab

Fig

ure

3.20

:B

lues

chist–

gree

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ist

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sition

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ection

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for

bulk

:jd

bkd.

(Tab

le3.

6).

Das

hed-

blue

isop

leth

sar

ef(

jd)

((X

Fe3+

M1a

+X

Fe3+

M1m

)/(X

Al

M1a+

XA

lM

1m

+X

Fe3+

M1a

+X

Fe3+

M1m

);da

shed

-gre

enisop

leth

sar

eX

Ca

M4

inam

phib

ole.

Uni

varian

tre

action

sm

arke

dby

thic

ker

line

wei

ght.

74

1.09 for gl bkd.) compared to the glaucophane breakdown bulk. Resultant phase relations arecalculated in an NCKFMASHTO pseudosection between 400–550�C and 3–18 kbar.

The topology of the pseudosection (Fig.3.20b) is characterised by coexisting glaucophane andjadeite, which form bi- and trivariant assemblages up-P from the albite stability field. The assem-blge: actinolite + hematite + albite + magnetite + muuscovite + paragonite + epidote +sphene+ quartz + H2O is stable throughout the greenschist facies, between 400–550�C and 3–9 kbar.Jadeite and hematite are predicted to be stable across the T range, above ⇠9 kbar. Reaction 1occurs up-T from a full-system invariant point at 8.59 kbar and 505�C; the reaction has a gentle�P/�T and reaches ⇠9 kbar at 550�C. Down T from the invariant point, a jadeite-absent univariantreaction governs the breakdown of glaucophane. Chlorite is not predicted to be stable.

The pseudosection shows that symplectitic development of albite + hematite + actinolite atthe rims of jadeitic pyroxene, primarily occurs as a response to decompression from HP condi-tions, and not as as a result of heating. Isopleths of XCa

M4 in Ca-amphibole (green-dashed lines:Fig.3.20b) suggest that symplectite growth/equilibration occurred between 6–7.5 kbar, at 550�C,in agreement with values predicted by modeling of glaucophane-breakdown symplectites. However,isopleths of f(jd) (blue-dashed lines: Fig.3.20b) show that the calculated composition of pyroxeneis significantly more Fe3+-rich than observed compositions (f(jd) = 0.79–0.89 versus ⇠0.45, respec-tively); calculated values of j(jd) are between 0.8–0.9, which are close to measured values of ⇠0.9–1(Fig.3.18c). The discrepancy between measured and calculated pyroxene compositions precludesaccurately constraining peak-P conditions. Either the breakdown of jadeite is a non-isochemicalprocess, or there is an inconsistency in the calibration of Fe3+: Al partitioning in the a–X modelfor clinopyroxene.

3.4.5.4 Calibrating the clinopyroxene model

To investigate the energetics of jadeite–acmite stability in the clinopyroxene model, a syntheticbulk composition was constructed using the setbulk script of thermocalc. Mineral proportionsreflect the peak blueschist paragenesis (3 gl, 1 jd, 3.5 mu, 1 pa, 2 ep, 0.1 sph and 0.3 q; oxide phaseomitted) and representative mineral compositions were used from sample N45b. The resultantbulk is detailed in Table 3.6, as ‘peak’; the peak assemblage is predicted to be stable between360–420�C at 10 kbar.

At 10 kbar and 400�C, the Diener et al. (2007) pyroxene model yields a jadeite with f(jd)=0.652and j(jd)=0.864. Relative to the Glockner jadeites, j(jd) is within error of observed compositionswhilst f(jd) is ⇠20% enriched in Fe3+ relative to Al. As the P–T conditions of formation are broadlyconstrained to the interval 8–12 kbar and 350–450�C (Holland & Ray, 1985; Gleißner et al., 2007),and the bulk accurately reflects the observed mineralogy, this shows that acmite is too favorablypartitioned into model clinopyroxene. This accounts for the di�culty in calculating jadeite–acmitepyroxenes over a wide range of Fe3+:Al and Na2O:CaO compositional space (Fig.3.19). Figure3.21 (after Diener et al., 2007) shows the jadeite–acmite–diopside ternary system calculated at 10

75

acm

jac

jd om di

dac

Figure 3.21: Jadeite–diopside–acmite ternary diagram af-

ter Diener et al. (2007). Calculated at 500�C and 10 kbar.

Grey regions denote stable compositions whilst white areas

highlight miscibility gaps between coexisting pyroxenes. Red

star represents the Glockner pyroxene compositions whilst the

arrow shows the broad trajectory of sodic-pyroxene with in-

creasing/decreasing Fe3+:Al—this shows the need for a DQF

adjustment to calculate more jac-rich compositions.

kbar and 500�C. The red star shows the composition of Glockner jadeite whilst the arrow showsthe approximate trajectory of model clinopyroxene to higher and lower values of Fe3+:Fe2+ andFe3+:Al.

Diener et al. (2007) employ o↵sets, in the form of DQF parameters, to the �fHacm and �fHjac

of -4 and -3 kJ.mol�1 respectively. As the omphacite ordered end-member has a DQF of -2.9, thismeans that there is a strong preference for Fe3+ to be partitioned, at the expense of Al, into acmiteover the ordered jadeite–acmite end-member (Fig.3.21). The DQF corrections are based on theassumption that Fe3+ is preferentially incorporated, over Al, into jadeite relative to coexistingglaucophane, omphacite, epidote and oxide phases (Diener et al., 2007). Given that this is onlysemi-quantitative and based on the Glockner blueschists themselves, it is likely that the DQFs arethe source of error.

A DQF adjustment of +4 kJ.mol�1 to both acm and jac end-members is required to bringvalues of f(jd) into concordance with those of Glockner-nappe jadeite. This calibration yields thesame assemblage as before and a jadeite with f(jd)=0.465 and j(jd)=0.864 at 10 kbar and 400�C,in very reasonable agreement with the observed pyroxenes.

This work highlights the need for a careful re-calibration of the clinopyroxene model. By fixingP and T through independent estimates calculated from neighboring lawsonite schists (e.g. Gleißneret al., 2007), DQF parameters could be varied independently to produce the desired compositionsfrom a pertinent bulk composition.

3.4.5.5 Peak conditions

Figure 3.22, calculated using the recalibrated clinopyroxene model and the synthetic bulk compo-sition described above (Table 3.6), Fig.3.22 shows the P–T stability field for the peak assemblage:gl + jd + mu + pa + ep + sph (+ q + H2O). The assemblage is quadrivariant in NCKFMASHTOand is bounded by pentavariant lawsonite-bearing fields up-P and down-T , and albite-bearing as-semblages up-T and down-P . The ru–sph transition occurs between 11–13 kbar and 400–500�C

76

respectively. At 11 kbar and 400�C, calculated mineral modes (vol.%) of 30%gl, 9%jd, 28%mu,2%pa, 23%ep, 0.8%sph and 5%q are plausibly similar to natural modes, with the exception of gl,which should be <10%. Both hematite and magnetite were excluded from consideration becauseit is unclear as to whether they were present at peak P or whether they are solely related todecompression reactions.

The stability field is contoured for f(jd), j(jd) and z(gl) parameters (where z(gl)=XM4Na , Diener

et al., 2007). Clinopyroxene is an acmite–jadeite solid-solution with <10% diopside throughoutthe assemblage; jadeite becomes more acmite-rich to lower P , in agreement with the exsolution ofhematite during decompression (Fig.3.17), and increasingly sodic to lower T . Calculated valuesof f(jd) and j(jd) are within the range of Glockner jadeite (j(jd): 0.85–1; f(jd): 0.45–0.5), attemperatures lower than 450�C and pressures between 9 and 13 kbar. The combination of aquadrivariant assemblage and extreme sensitivity to bulk composition means that f(jd) isoplethsare widely spaced (⇠60�C) and yield little information regarding peak-P conditions. However,the fact that sphene is observed in place of rutile shows that peak P must have been <13 kbar.XM4

Na in glaucophane shows little variation throughout the field, but is consistent with observedvalues of ⇠0.90–0.98. Si in phengite values lie between 3.2–3.3 c.p.f.u., ⇠0.1–0.2 c.p.f.u.lower thanobserved values; this could be explained by partitioning more Fe3+ into the phengite structureduring tschermakite exchange between celadonite and muscovite end-members (Fig.3.18).

The agreement between calculated and observed assemblages shows that the e↵ective bulkcomposition responsible for the Glockner nappe blueschists is more Fe3+-, Na- and Al-rich thana typical MORB (Fe3+:Fe2+

MORB=0.56, Fe3+:Fe2+Peak=1.12; Na:CaMORB=0.21, Na:CaPeak=0.58) in

addition to being less calcic (Table 3.6). The ratio of Al to Fe3+ is broadly similar to a MORB(Al:Fe3+

MORB=6.17; Al:Fe3+Peak=6.33), which suggests that assemblage development is controlled

largely by Na:Ca and Fe3+:Fe2+ ratios.

3.4.5.6 Average P–T

Despite the non-equilibrium nature of the jadeite and glaucophane breakdown textures, phasescan be separated into those which belong to the peak and retrograde assemblages and utilized forconventional thermobarometry. Results augment the pseudosection modeling and are displayed inTable 3.7.

The following phases are considered from sample N45b to represent the peak assemblage: gl +jd + mu + ep + sph/ru + q (+mt + pa + H2O). Seventeen end-members generate 7 independentreactions which intersect at an average P of 7.8±3.66 kbar at 400�C (�fit=4.39) with XH

2

O = 1—an average P–T calculation is precluded by few geothermometer reactions. The large uncertaintyin this estimate is largely due to uncertainity over the activity if riebeckite in Na-amphibole. Onceexcluded, the estimate is refined to: 9.67±1.46 kbar (�fit=1.81; XH

2

O=1). Varying XH2

O between1 and 0.7 does not change the estimate outside of uncertainty. It is unclear whether paragonite

77

350 400 450 500 5508

9

10

11

12

13

14

15

16

17

0.46

0.88 0.

87 0.86 0.

85 0.84 0.

83 0.82 0.81 0.80 0.79 0.78

0.89

0.47

0.45

0.43

0.44

NCKFMASHTO+ q H2O

j(jd)

f(jd)

Temperature (°C)

Pre

ssur

e (k

bar)

0.99

0.98

z(gl)

+ ab

+law

gl+jd+mu+pa+ep+sph/ru

+ru

+sph

Figure 3.22: P–T pseudosection in NCKFMASHTO for peak-blueschist-facies assemblage: gl + jd + mu + pa + ep

+ sph (+ q + H2

O). Isopleths for j(jd) (blue), f(jd) (red) and z(gl) (green)—see text for definitions. Yellow line marks

peak conditions defined by j(jd)>0.85.

was stable at peak P , or whether it was generated during retrogression. Accordingly, the pressureestimate may represent a hybrid value between peak P and Barrovian values.

The assemblage associated with jadeite breakdown is: act + ab + mu + ep + bt + q + hem+ H2O. Using mineral data from sample 221c (Holland & Ray, 1985), and excluding the fcel end-member due to erroneously high uncertainty over its activity, yields an estimate of 5.5±2.5 kbarand 371±51�C (�fit=0.82; XH

2

O=1) from 4 independent reactions. The error correlation coe�cient(⇢) is 0.704, indicating that either P or T must be fixed to determine T or P respectively. Here,either an average P or T estimate is more suitable. Average P , without H2O included, within theBarrovian temperature interval of 450–550�C is 6.36±2.78 kbar (�fit=0.82); large errors in ln Kfor the geothermometer reactions shows that av.T estimates are poorly constrained, likely due tothe lack of strong Fe:Mg partitioning between phases.

78

The assemblage associated with glaucophane breakdown is: act + pa + ep + ab + chl + ta+ q + mt + H2O, as typified by sample 221d (Holland & Ray, 1985). Including all end-membersgenerates 9 independent reactions and an average P–T estimate of: 11.8±2.8 kbar and 539±82�C(�fit=3.16; XH

2

O=1). The poor �fit is largely controlled by clinochlore and fcel end-members.Additionally, including talc has a strong e↵ect on P . Given that it is unclear at what point duringthe decompressive P–T path talc formed it is omitted. These refinements yield an estimate of:5.4±1.6 kbar and 363±44�C (�fit=1.52; XH

2

O=1) with ⇢=0.704. Therefore, average P is favoured;for the Barrovian temperature interval (450–550�C) av.P is 8.34±0.75 (�fit=0.94). Average T ispoorly constrained due to weak Fe:Mg partitioning (high lnK errors) between phases.

Sample PM3 from Gleißner et al. (2007), is a lawsonite greenschist, collected ⇠500 metresup structural-section from samples 221c and 221d. Blueschist facies lawsonite has been totallyreplaced by: act + cz + pa + chl + ab + q + cc. Including H2O and CO2, whilst excludingceladonite and the ames end-member of chlorite yields: 7.1±0.9 kbar and 540±14�C (�fit=0.31;XH

2

O=0.8; ⇢=0.9). This estimate forms a �fit minima for values of XCO2

and XH2

O between 0–0.3and 0.7–1 respectively; however, the large error correlation coe�cient shows that T and P arebetter defined independently. CO2-absent average P between 450 and 550�C is 6.08±0.73 kbar(�fit=1.26); pressures increase up-temperature.

Relict glaucophane needles are present in the cores of lawsonite pseudomorphs in sample PM3Gleißner et al. (2007). Textural evidence suggests that the peak blueschist-facies assemblage was:gl + law + ep + mu + chl + sph/ru + q + H2O. These phases generate an average P estimatebetween 300–600�C of 14.31±0.68 (�fit=2.55; XH

2

O=1). Pressure is fixed by the exchange ofCa between clinozoisite and lawsonite. Omitting H2O, daphnite and the ames end-member ofchlorite due to large e* values, yields: 13.62±1.17 (�fit=2.13). Reducing the temperature intervalto between 350 and 450�C further refines the estimate to 12.74±1.01 (�fit=1.92).

Sample TH–661 is a garnet-bearing epidote-schist collected by T.Holland from base of theGlockner nappe, ⇠500 metres south of the Eisseehutte. The region displays intensely lineatedmetamorphosed pillow basalts as descried by Holland & Norris (1979). The main assemblage is:gt + hbl + chl + ep + ab + bt + pa + cc + q (H2O + CO2), which generate twelve independentreactions and a P–T estimate of 8.4±2.2 kbar at 519±32�C (�fit=3.84; XH

2

O=0.95). Omitting gland pargasite from hornblende (incorrect solvus positions) and also H2O and CO2, generates anestimate of 9.9±2.4 kbar and 547±80�C (�fit=1.14).

3.4.6 Discussion

Whole-rock bulk compositions do not accurately predict observed blueschist phase relations (Sec-tion 3.4.5.2). This suggests that the metamorphic evolution of jadeite- and glaucophane-bearingparageneses was controlled by sub-centimetre lengthscales of chemical equilibrium within a pro-tolith enriched in Fe3+, Al and Na. Intra-sample gradients in these components accounts for thelocal development of glaucophane and jadeite patches. The compositional characteristics are in-

79

Sample Assemblage P �⌦P T �⌦

T �fit

N45b† gl+jd+mu+ep+sph+q+mt+pa+H2O 9.67 1.46 350–450 - 1.81PM3† gl+law+mu+ep+sph/ru+chl+q+H2O 12.74 1.01 350–450 - 1.92221c? act+ab+mu+ep+bt+q+hem+H2O 6.36 2.78 450–550 - 0.82221d? act+pa+ep+ab+chl+q+mt+H2O 8.34 0.75 450-550 - 0.94PM3? act+cz+pa+chl+ab+q+cc+H2O+CO2 7.10 0.90 540 14 0.31PM3? act+cz+pa+chl+ab+q+H2O 6.08 0.73 450–550 - 1.26TH-661? g+hbl+chl+ep+ab+bt+pa+cc+q+H2O+CO2 9.90 2.40 547 80 1.14

Table 3.7: Average P and P–T calculations for Glockner nappe samples. ? Barrovian estimate; † Blueschist estimate;⌦ Error interval is 2�.

dicative of some sort of oxidised basalt–sediment succession, as to be expected in the upper levelsof an ophiolite complex.

The high variance (F=4) of peak assemblages, in addition to the extreme sensitivity of theclinopyroxene a–X model to bulk composition makes calculating peak blueschist facies conditionsdi�cult. The most reliable data are obtained from average P calculations (Table 3.7); samplePM3—a lawsonite+glaucophane schist from Gleißner et al. (2007)—shows that peak pressures werebetween ⇠11.7–13.7 kbar. The pressure estimate obtained from coexisting glaucophane and jadeiteis slightly lower and more uncertain: 9.67±1.46 kbar. Nevertheless both estimates corroborate one-another in showing that peak conditions did not exceed 14 kbar and were likely greater than 9 kbar.Pseudosection modeling of the gl+jd peak assemblage only loosely constrains peak conditions to9–13 kbar and 350–450�C.

Given that �fHacm and �fHjac used in the clinopyroxene a–X model have been calibratedagainst the jadeite-bearing assemblage of the Glockner nappe, it is important to note the concor-dancy between the pseudosection-based pressure estimate (Fig.3.22) and the independent pressureestimate obtained from sample PM3. In the absence of experimental data, this confirms that theinteraction energies between the acm and jac end-members, with the rest of the clinopyroxenesystem are a good approximation. Calibration of the parameters against the Glockner assemblageis made di�cult by the clinopyroxene model’s extreme sensitivity to minor changes Al:Fe3+ ratioof the bulk composition. Further work is required to accurately calibrate the model parameters.

Conventional thermorbarometry on retrogressive assemblages yields a series of compatible aver-age P estimates between ⇠5–8 kbar for the Barrovian temperature interval: 450–550�C (Fig.3.23).Temperature estimates are poor for all greenschist assemblages studied because of the absenceof a phase to strongly partition Fe:Mg between coexisting minerals, i.e. garnet. Barrovian tem-peratures in the upper Schieferhulle are constrained to between 425 and 475�C by oxygen isotopestudies of Hoernes (1974); Dachs (1990) use conventional geothermobarometry and calcite-dolomitethermometry to suggest a similar temperature range: 430–525�C.

Isopleths of XM4Ca in amphibole calculated for both jadeite and glaucophane breakdown textures

(Fig.3.20) are in agreement with Barrovian pressures between 5 and 8 kbar. However, the inability

80

of the univariant bulk compositions to predict the composition of peak-assemblage jadeite andmuscovite suggests that the breakdown reactions were non-isochemical. Taking the decompositionof jadeite, which forms albite and exsolves hematite, a generalised reaction can be written:

Na2AlFeSi4O12 ) NaAlSi3O8 + 0.5Fe2O3 + 0.5Na2O + SiO2

Therefore, jadite breakdown is non-isochemical and generates an excess of sodium and silica, whichare free to be removed from the symplectite and react elsewhere in the rock. Furthermore, bothjadeite and glaucophane univariant reactions (Section 3.4.5.3) require paragonite as a reactant.In glaucophane schists, this is consistent with the high modal abundance of paragonite present,however, jadeite schists such as N5b and 221c have only minor (<2%) paragonite, if present at all.Plausibly, if jadeite breakdown occurred via the univariant reaction, the exhaustion of paragoniteprovides an explanation for the halted reaction progress.

If the reaction textures are non-isochemical, application of equilibrium thermodynamics isredundant. However, Fig.3.23 plots a compilation of P–T data from this work, and existingdata, to show that the assemblages record ia consistent decompressive P–T trajectory from peakblueschist facies conditions of 9–14 kbar and 350–450�C, to Barrovian temperatures at ⇠5–8 kbar.

There is a correlation between structural position and peak metamorphic grade. Sample PM3from Gleißner et al. (2007) was collected ⇠400 metres from the base of the Glockner nappe andyields the highest pressure estimate for blueschist facies conditions (Fig.3.23). Sample TH–661comes from similar structural levels and yields the highest-grade estimate for Barrovian condi-tions. Whilst the thermobarometric data are insu�ciently accurate to calculate geotherms, theobservation suggests that the Glockner nappe was buried and exhumed as a coherent unit and thatprogressively deeper structural levels are preserved further to the north (e.g. Dachs, 1990).

3.4.7 Tectonic implications

The presence of coexisting jadeite and glaucophane assemblages, in addition to lawsonite-bearingparageneses, confirms that the Upper Schieferhulle was buried to between 30 and 45 kilometres,along a subduction-related geotherm of 8–15�C.km�1. Peak pressures are up to 10 kbar lessthan the underlying Eclogite Zone. Blueschist-facies metamorphism occurred prior to the sec-ond penetrative deformation event (D1) as the crenulated S1+S2 fabric anastomoses around HP

pseudomorphs. This relationship is also observed in the Eclogite Zone, where glaucophane in-fillsfractures formed by D1 crenulations (Holland & Richardson, 1979).

Considerable along-strike di↵erences in the depth of burial exist for the Glockner nappe. Fur-ther to the east, in the Großglockner region, peak depths are much deeper than elucidated fromthis work—around 50 kilometers (Dachs & Proyer, 2001). Together with the interpretation ofthe Eclogite Zone as remnant Glockner-basin material, this is consistent with the Glockner napperepresenting a slice of variably attenuated oceanic crust.

Subsequent to early blueschist facies metamorphism, the Glockner nappe was exhumed tomid-crustal levels (⇠18–28 km) where it experienced pervasive recrystallisation along a Barrovian

81

20

15

10

5

Pre

ssur

e (k

bar)

600500400300

cz+gl+law+ru

pa+clin+sph+q

jd 44+acm 40+q

ab

gl+ep+mu+q

cel+cz+ab+hem+H

2O

acm+pa+q

ab+hem+H 2O

gl+ep+q

tr+ab+pa+hem

+H 2O

Temperature (°C)

PM3

N45b

221c

221dPM3

TH-661

D&P

D&P

PS

a

b

Figure 3.23: Glockner P–T

paths. Boxes delineate average P

conditions across a specified tem-

perature (Table 3.7). Labels corre-

spond to sample code used in text;

D&P=Dachs & Proyer (2001). P–

T path a derived from this study;

b is that of Dachs & Proyer (2001)

from the Großglockner region. Re-

actions calculated using the same

a–X models as described in the

text and the following compositions:

law-out - sample PM3; jd-out -

221c; gl-out - 221c; acm-out - 221c.

Stippled area corresponds to the

oxygen isotope temperature interval

calculated by Hoernes (1974).

82

geotherm (⇠17–30�C.km�1). This occurred contemporaneously with development of the regionalfabric (S2). Barrovian grade is greater towards deeper structural levels, which is consistent with ex-humation of the Glockner nappe as a coherent body and conductive heating following emplacementof the Austroalpine nappes.

3.5 Venediger nappe

The continental basement of the Venediger nappe structurally underlies the Pennine nappe-stackand pertains to the European margin. The nappe comprises a pre-Variscan basement which isintruded by Variscan granitoids and overlain by a sequence of autochthonous Jurassic–Cretaceousmetasediments. Unlike the Glockner and Eclogite Zone nappes, it is unclear whether the Venedigernappe experienced HP metamorphism. Furthermore, despite being a fundamental constraint onthe thickness of the Austroalpine complex, peak Barrovian conditions attained by the nappe areloosely constrained. In the following, average P �T thermobarometry and pseudosections are usedto address these questions.

3.5.1 Previous work

Thermobarometric investigation has been focused on the metasedimentary cover nappes of theVenediger complex (Lower Schieferhulle); in contrast rocks of the Zentralgneiss plutonic suite havereceived little attention. There are few existing coupled estimates of both P and T for the Venedi-ger complex.

Bickle (1973) and Holland (1977) calculated pressures greater than 8 kbar using the assemblage:K-feldpar + albite + garnet + muscovite + biotite + epidote + quartz, which developed in mm-scale patches on plutonic albite of the Zentralgneiss.

Selverstone et al. (1984) investigated hornblende garbenschists of the Lower Schieferhulle, ex-posed in the southwestern Tauern, which contain the following assemblage: hornblende + kyanite+ staurolite + garnet + biotite + epidote + plagioclase + ankerite + quartz + rutile + ilmenite +chlorite + paragonite. Fe–Mg exchange thermometry between garnet and biotite yields tempera-tures of ⇠550�; pressures of 6–8 kbar were calculated from matrix garnet, plagioclase, kyanite andquartz, whereas plagioclase inclusions in garnet yield higher pressures between 9–10 kbar. Thesedata, along with major-element zoning in garnet, hornblende and plagioclase, are interpreted asindicating passage of the cover nappes along a decompressive P–T trajectory from ⇠530�C, 10kbar to ⇠550�C, 7 kbar.

Kruhl (1993) interpreted zoned phengites (Si increasing from core to rim domains) from Venedi-ger cover units of the northeastern Tauern Window in terms of burial of the nappe from 4–10 kbar.Similarly, pressures of 10–11 kbar (3.4 Si c.p.f.u.) were calculated by Franz et al. (1991), fromorthogneiss-phengites of the south-central Venediger nappe. Kurz et al. (1998b) used the limit-

83

ing assemblage: K-feldspar + biotite + quartz, in orthogneisses of the Lower Schieferhulle (Storznappe) to calculate pressures between 6–11 kbar.

Zimmermann et al. (1994) documented centimeter-long veins containing the assemblage: om-phacite + albite + actinolite + hornblende + calcite + apatite + titanite + quartz in a weaklyfoliated garnet amphibolite of the Lower Schieferhulle, exposed in the upper Frosnitztal. This is theonly reported occurrence of sodic-pyroxene from structural levels beneath the Eclogite Zone. Pyrox-ene is euhedral and displays sub-millimetric scale lammellae of diopside and omphacite; the jadeitecontent increases from ⇠25 mol.% in the core, towards ⇠45 mol.% at the rim. Diopsode–omphacitelamellae are interpreted as coexisting solvus compositions at a temperature of ⇠375±50�C. Pres-sures of 10±1.5 kbar are calculated using the reaction: albite = jadeite + quartz. Collectively,these data are interpreted as showing that blueschist conditions prevailed in the Venediger complex,prior to Barrovian overprinting.

Frisch & Raab (1987) provide further evidence for early HP metamorphism in the Venedi-ger complex, in the form of pseudomorphs after lawsonite within amphibolites of the LowerSchieferhulle.

Frisch & Raab (1987) and Zimmermann et al. (1994) provide the most convincing evidencethat basement units experienced HP metamorphism along an Alpine-subduction geotherm. How-ever, these studies target garnet amphibolites, located towards the structural base of the LowerSchieferhulle cover nappes. As these units are interpreted as having formed within a pre-Variscanisland-arc environment, it is unclear which metamorphic phases pertain to the Variscan or Alpineorogenic cycles.

In light of the above, thermobarometric calculations are performed on samples collected fromthe Zentralgneiss, a Hercynian plutonic body which was variably reworked during the Alpine event,and also on samples collected from the Jurassic–Cretaceous (Kurz et al., 1998b) cover sequence.

3.5.2 Sample petrography and mineral chemistry

Thermobarometric calculations are focused on samples from the Zentralgneiss. Accordingly, theirpetrography is described in more detail than samples from the cover units, which are used foraverage P–T calculations.

3.5.2.1 Samples TH–519 and ASA–41a

Both samples are coarse-grained, unfoliated metatonalites of the upper-Zentralgneiss, exposed inthe Dorfertal valley and have similar mineralogies and textural characteristics.

Metamorphic mineral growth is restricted to mm-scale patches of clinozoisite, garnet, bi-otite, muscovite and oxides which overgrow coarse-grained plutonic plagioclase. Garnet occursas <100 µm-diameter euhedral blasts and is an almandine–grossular solid solution, with val-ues of x(g) and z(g) around 0.41 and 0.56 respectively (x(g) = Fe2+/(Fe2++Fe3++Mg+Ca) andz(g)=Ca/(Fe2++Fe3++Mg+Ca)); these values deviate by less than 5% from core to rim. Biotite

84

Figure 3.24: Photomicrographs of Alpine-metamorphic mineral assemblage development in metatonalite TH–519. a.

Field of view is ⇠2 mm. Dashed box highlights photograph area of b; A and A* denote limits to chemical traverse of

host plagioclase grain displayed in .b Close-up photomicrograph of metamorphic phases; field of view is ⇠600 µm. Both

photomicrographs were taken using cross-polarised light.

flakes are between 300–800 µm diameter and display mottled birefringence, changing from bottlegreen in the core, to light brown at grain edges; XFe ⇡0.66 and the phlogopite component isminor: Na = 0.041 c.p.f.u. Small (<10 µm) inclusions of quartz and epidote are common andgrain boundaries are embayed, suggesting partial decomposition since growth. Muscovite containsbetween 3.32 and 3.35 c.p.f.u. Si and is volumetrically less significant than biotite. Where present,it occurs as <200 µm flakes proximal to clusters of garnet. Rafts of clinozoisite needles less than300 µm in length are ubiquitous (Fig.3.24a and b)—individual needles are aligned with the weakfoliation. M3-site Al:Fe3+ ratios are close to 1.51 and the total replacement of Ca by Fe2+, Mgand Mn is less than 0.089 c.p.f.u. Plagioclase grains host the metamorphic assemblage and reflectthe original plutonic texture. Grains are between 1–5 mm in diameter and show distinctive Carls-bad twinning (Fig.3.24a). Xalb varies between 0.93–0.99 across all grains (Fig.3.25), indicatingthat either the grains were formed chemically homogenous, or, that subsequent heating duringmetamorphism facilitated di↵usive smoothing of original compositional variations. In either case,the lack of compositional gradients proximal to metamorphic phases (Fig.3.25) shows that theplagioclase is in local equilibrium with the assemblage. Representative mineral compositions arepresented in the Table A.3.

As shown by Fig.3.24a, the metamorphic assemblage described above is spatially focused onhost plagioclase domains. The remainder of the rock comprises phaneritic (cm- to mm-scale)granoblastic quartz and plagioclase with minor (<5%) K-feldspar, hornblende and biotite. Due toits low modal abundance, K-feldspar was not analysed in target thin-sections of samples TH–519or ASA–41a. However, Bickle (1973) provides K-feldspar analyses from a petrologically similar

85

1.0

0.8

0.6

0.4

0.2

0.0

XalbK (c.p.f.u.)Al:Si

0 100 200 300 400 500 600 700 800Traverse distance (µm)A A*

Figure 3.25: Chemi-

cal traverse across host

plagioclase grain shown

in Fig.3.24. Xalb de-

fined as Na/(Na+Ca)

(c.p.f.u.). Data are fit-

ted by lines of visual

best fit. Note the high-

albite content and lack

of significant chemical

variation across the tra-

verse. Outliers show

the presence of included

phases. Measurement

errors are insignificant

relative to marker size.

metatonalite of the Zentralgneiss exposed in the Großglockner massif, and shows that it is puresanidine in composition. Alignment of millimetric hornblende laths defines a weak foliation, whichdips steeply to the south.

3.5.2.2 Sample SP614

This sample was collected by M.J. Bickle from exposure of the Zentralgneiss in the Granatspitzeregion, ⇠ 25 kilometres to the east of the Venediger massif. It is a metatonalite with similarpetrography to that of TH–519 and ASA–41a, comprising a domain assemblage of: garnet + biotite+ plagioclase + clinozoisite +k-feldspar + quartz (+ minor ilmenite). This sample is investigated toprovide an along-strike comparison in the P–T conditions preserved by the Zentralgneiss. Readersare referred to Bickle (1973) for a more detailed description of the sample’s chemistry.

3.5.2.3 Sample ASA–38a

Sample ASA–38a was collected ⇠400 m to the south of the Venediger nappe/Eclogite Zone bound-ary from an outcrop of biotite schist pertaining to the Venediger cover nappes. It is a stronglyfoliated garnet + biotite schist; augen of quartz are wrapped by a biotite + muscovite fabric, dip-ping steeply to the southeast (S2). The rock comprises: garnet + biotite + muscovite + plagioclase+ epidote + quartz with accessory rutile, detrital zircon and calcite. Garnet forms small (200–400µm), euhedral poikiloblasts with sub-microscopic inclusions of quartz and epidote. Chemical analy-sis shows that garnet is an almandine–grossular solid solution with Xalm ranging from ⇠0.30 in thecore to ⇠0.56 in rim domains. Grain boundaries of quartz and plagioclase are interlocked in ⇠500µm-wide microlithons, which separate micaceous domains. Plagioclase is pure albite (Xalb=0.99).Biotite is the dominant mica—flakes show a mottled appearance with green-pale brown birefrin-

86

gence. Phlogopite end-member is minor with typical Na c.p.f.u. values of ⇠ 0.012; XFe is between0.65 and 0.68. Muscovite flakes are commonly between 50–100µm in width and are aligned withthe biotite-defined schistosity. Larger grains show variation in celadonite content between 3.32to 3.16 c.p.f.u. Si, from core to rim, respectively. Epidote is clinozoisite-rich ( 0.37–0.41 c.p.f.u.Fe3+ and ⇠1.99 c.p.f.u. Ca) and forms lozenges (⇠300 µm in length) which are aligned with thefoliation. Representative mineral analyses are presented in Table A.3.

3.5.2.4 Sample ASA–42a

Sample ASA–42a is a garnet-mica schist collected from the Inner Schieferhulle exposed in the mid-Dorfertal, ⇠800 meters structurally above the basement–cover interface. The main assemblage is:garnet + muscovite + plagioclase + chlorite + quartz + calcite. Muscovite, quartz and chloriteare aligned to define an intense schistosity (S1), which is openly crenulated (S1+S2) and wrapspokiloblastic garnet. Garnet occurs as ⇠500–800 µm-wide, euhedral blasts with abundant quartzinclusions, which occasionally occur in high-enough concentrations to skeletal garnet textures.Compositionally, garnet is almandine-rich with Xalm values of ⇠0.72 in core regions and ⇠0.70in rim portions. Muscovite is present as fine (⇠50–100 µm-long) laths and shows variation in Sicontent between 3.08 and 3.23 c.p.f.u.. Plagioclase forms refractory, rounded blasts, which showevidence for decomposition at grain edges and deflect the schistosity. Xalb is ⇠0.96 and Ca contentsare <0.03 c.p.f.u.. Chlorite forms both fine (<300 µm) tablets in the matrix domains and also rimsto garnet poikiloblasts, which are largest in pressure shadows—values of XFe area close to 0.61.Calcite is a minor (<0.5 vol. %) phase constituent and is present as rounded clots in amongstquartz-rich domains. The kinematics of fabric relations in sample ASA–42a are discussed in moredetail in section 6.6. See Table A.3 for mineral composition data.

3.5.3 Average P–T

Multiple reaction thermobarometry was performed on each of the samples detailed above, usingThermocalc 3.33i and the November 2003 updated version of the Holland & Powell (1998) data set.Activity–composition models are detailed in section 3.4.5. Results are presented in Table 3.8 andFig.3.27.

3.5.3.1 Sample TH–519

The following end members yield a set of six independent reactions for XH2

O=1: pyrope, grossular,almandine, muscovite, celadonite, ferroceladonite, paragonite, anorthite, albite, clinozoisite, phlo-gopite, eastonite, annite and quartz. These equilibria intersect to estimate average P–T conditionsof 12.5±2.4 kbar and 594±64�C, with a �fit of 2.43. Omitting eastonite, paragonite and ferroce-ladonite end members due to low activities and incorrect solvus positions, yields an average P–T

estimate of 14.1±1.1 kbar and 603±27�C (�fit=0.99), from the following independent equilibria:

87

1. mu + 2phl + 6q = py + 3cel2. 12cz + 4phl + 6san = 3py + 8gr + 7mu + 3cel3. 4gr + alm + 4mu + 3q = 6cz + ann + 3san4. 5py + 12cz + 15ann + 33q = 8gr + 15alm + 15cel + 6H2O

Given the millimetre-scale development of the metamorphic assemblage and low values of per-meability common in metagranitoids, water activity (aH

2

O) is likely to be low. Omitting H2O fromconsideration means that reaction 4 is no longer used—the resultant P–T estimate is 13.7±1.6kbars and 571±76�C (�fit = 1.15). Average P–T conditions vary by less than 0.7 kbar and 46�Cfor values of aH

2

O between 1 and 0.5; a �fit minima is reached at aH2

O=0.7. Collectively, thesecalculations show that the calculation of an average P–T estimate is relatively insensitive to wateractivity. Using reactions 1–4 and aH

2

O=0.7 yields conditions of: 13.1±2.1 kbar and 601±38�C,which are interpreted as reflecting the best estimate of peak conditions.

3.5.3.2 Sample SP614

End members of anorthite, albite, sanidine, almandine, muscovite, celadonite, ferroceladonite,paragonite, phlogopite, annite, eastonite, quartz and H2O yield an average P–T estimate of15.9±3.2 kbar and 689±78�C (aH

2

O=1; �fit=2.39). The large uncertainty in T is due to thesmall activities of almandine and anorthite, and their involvement in geothermometer reactions(hat=0.46 and 0.53 respectively). As it is unreasonable to exclude these data, temperature in-formation is discarded and an average P of 13.3±2.2 kbar (�fit = 1.79) is calculated between550–650�C, without H2O.

3.5.3.3 Sample ASA–38a

End members of muscovite, celadonite, ferroceladonite, paragonite, phlogopite, annite, eastonite,albite, pyrope, grossular, almandine, clinochlore, daphnite, Al-chlorite, calcium carbonate, quartz,H2O and CO2 generate 6 independent equilibria which yield an average P–T of 12.6±2.6 kbar and613±57�C (�fit = 2.37; aH

2

O=1). Omitting the eastonite end-member from consideration due tolarge uncertainty associated with its activity refines the estimate to: 13.7±2.4 kbar and 610±48�C(�fit = 2.04; aH

2

O=1). aH2

O has a strong e↵ect on the temperature estimate—excluding H2O fromcalculations changes the temperature to 508±164�C, whilst P is una↵ected. A �fit minimum (1.55)is located under XCO

2

=0.5, which yields a best P–T estimate of: 13.7±2.2 kbar at 579±42�C.

3.5.3.4 Sample ASA–42a

The following end-members were considered: muscovite, celadonite, ferroceladonite, paragonite,phlogopite, annite, eastonite, albite, pyrope, grossular, almandine, clinozoisite, calcium carbonate,quartz, H2O and CO2. Using XCO

2

=0.05, 7 equilibria intersect to yield an average P–T of 11±1.7kbar and 531±25�C (�fit=2.71). Omitting the Al-chlorite end-member, on the basis that its

88

activity is loosely constrained, refines the estimate to 11.7±1.2 kbar and 538±17�C (�fit=1.79).Removing CO2, H2O and calcite from consideration on the basis that their activities are unknown,yields a best estimate of 10.2±1 kbar and 517±62�C (�fit=0.73).

Sample Assemblage P �?P T �?

T �fit

TH-519 g+cz+bi+mu+pl+kfs+q+H2O 13.1 2.1 601 38 1.66ASA-38a g+mu+bi+ep+pl+q+cc+H2O+CO2 13.7 2.2 579 42 1.55ASA-42a g+mu+bi+pl+chl+q+cc+H2O+CO2 10.2 1 517 62 0.73SP614 g+bi+mu+pl+kfs+q+H2O 13.3 2.2 - - 1.79

Table 3.8: Average P–T calculations for Zentralgneiss and Inner Schieferhulle samples. ? Error interval is 2�.

3.5.4 Metamorphic modeling

Metamorphism of granitoid protoliths is characterised by small changes in assemblage variancerelative to changes in P and T (e.g. Proyer, 2003). Therefore, growth of the assemblage: garnet +biotite + muscovite + clinozoisite + plagioclase + K-feldspar + quartz from a tonalitic compositioncan be investigated through construction of a P–T pseudosection.

Consequently, Fig.3.26 is a NCKFMASHTO pseudosection constructed for a whole-rock meta-tonalite (see Table 3.9 for bulk composition) collected from the Venediger massif and reported byWinkler (1996) and Frisch et al. (1993). Phase relations are calculated between 3 to 18 kbar, and400 to 650�C using a–X models discussed in section 3.4.5. Although it is unclear whether theassemblage was controlled at a whole-rock length scale of chemical equilibrium, consideration ofa representative tonalite bulk provides a starting point for such calculations. MnO will enhancethe stability field of garnet; however, it is omitted here due to the low Mn contents of garnet(<0.02 c.p.f.u.), in addition to the lack of Mn-bearing end-members in relevant phases. TiO2 isconsidered due to its e↵ect on stabilizing biotite and also the minor quantites of opaques presentin both TH–519 and sample ASA–41a. Tonalites typically contain between 1–4 molar % H2Owhich is principally hosted by biotite; prograde dehydration of biotite liberates H2O, which canbe sequestered by muscovite and epidote at higher P and T . Combining mineral modes from themetamorphic-mineral patches (5% biotite, 8% muscovite and 15% clinozoisite) with mineral H2Ocontents (calculated by di↵erence from ideal cation totals), yields an estimate of 1.08 weight %H2O for assemblage-bound water. FeO:Fe2O3 is set to 3.3 (O = 0.80) on the basis of clinozoisitecontaining ⇠0.36 c.p.f.u. Fe3+.

The pseudosection (Fig.3.26a) is characterised by a patchwork of small penta- to trivariantstability fields, which separate low-P , high-T H2O-present assemblages from high-P garnet +jadeite + kyanite assemblages. Plagioclase is predicted to be stable from 3 kbar, where it is ⇠40%anorthite, to ⇠11.3 kbar at 525�C, where it is close to pure albite (<6% anorthite). Up-P , theanorthite component of plutonic plagioclase reacts with biotite and its lattice-bound H2O to formgrossular-rich garnet, muscovite and quartz—accordingly, the biotite stability field mirrors that of

89

NaO CaO K2O FeO MgO Al2O3 SiO2 TiO2 O H2O

Winkler et al. 1996 3.24 6.48 1.37 5.32 4.47 10.56 63.36 0.61 0.79 3.77Nockolds 1954 3.98 5.24 0.95 4.09 3.04 9.65 69.60 0.49 0.54 2.42TH–519 4.91 6.79 1.15 3.89 1.06 11.34 66.30 0.14 0.70 3.72

Table 3.9: Compilation of metatonalite bulks. Winkler (1996) – XRF bulk of a Venediger-massif metatonalite from

the Zentralgneiss (used to construct Fig.3.26); Nockolds (1954) – average tonalite composition; TH–519: – domain

assemblage bulk based on 15% clinozoisite, 10% muscovite, 45% plagioclase, 5% garnet, 5% biotite and 20% quartz.

plagioclase. The plagioclase field is terminated up-P , by the fluid-absent reaction: albite = jadeite+ quartz (Holland, 1980). Kyanite is stabilised over high �P/�T regions by Al liberated fromanorthite and paragonite breakdown. Under such weakly oxidised conditions, ilmenite is calculatedto be stable across the entire P–T range except for a brief window between ⇠400–600�C, beneath⇠6 kbar, where rutile is the TiO2-bearing phase. H2O is mineral-bound beneath ⇠530�C; at highertemperatures, H2O is a free-phase—its modal abundance is <1% and varies inversely with that ofbiotite. The calculated solidus lies outside the investigated temperature range: between 680–700�Cand 3–18 kbar. Nevertheless, an experimentally determined wet-melting curve for a biotite granite(Johannes & Holtz, 1996) is superimposed on the grid to show the likely position of melt generationat higher H2O contents.

An important discrepancy between calculated and observed phase relations is that K-feldsparis not calculated to be stable across the specified P–T range. This is likely because the bulkcomposition contains su�cient H2O to stabilize biotite over K-feldspar, thus sequestering all avail-able K2O. If biotite in samples ASA–41a and TH–519 contain significant F, the contribution ofbioitite-bound H2O to the total H2O budget will be an overestimate.

The observed paragenesis of garnet + clinozoisite + muscovite + biotite + plagioclase (+ quartz+ ilmenite) is quadrivariant in NCKFMASHTO and forms a thin-wedge in P–T space, between610–650�C and 12–13 kbar (red field in Fig.3.26a). It is bounded by the occurrence of paragoniteto lower pressures, the appearance of kyanite to lower temperatures and the appearance of jadeiteto higher pressures. As the field is small, it provides a tight constraint on ambient P–T conditions.Interestingly, the average P–T estimate of 13.1±2.1 kbar and 601±38�C for sample TH–519, lieswithin error (2�) of the assemblage stability field, despite the absence of K-feldspar. Within thecalculated domain assemblage, garnet is an almandine–grossular solid-solution, with Xgr between0.22–0.24, muscovite is calculated to have 3.05–3.15 c.p.f.u. Si, and plagioclase is 83–85% albite(Fig.3.26b).

3.5.5 Discussion

Pseudosection and average P–T estimates corroborate that domain assemblages in metatonalitesof the Zentralgneiss suite record albite-quartz-amphibolite facies metamorphic conditions. As theaverage P–T calculations take into account the presence of K-feldspar, the best estimate of peak

90

400 425 450 475 500 525 550 575 600 625 6503

4

5

6

7

8

9

10

11

12

13

14

15

16

17

18

g mu jd ky ilm

g mu bi jd ky ilm

g mu bi jd ky pl mt ilm

g bi ky mt ilm

g pl bi pa ilm

pl bi ilm mt H2O

pl bi ru mt H2O

pl bi pa ru mt chl ilm ru pl mt bi pa

bi pa ky pl mt ilm

bi pa pl mt ilm

Pre

ssur

e (k

bar)

Temperature (°C)

NCKFMASHTO+ ep q

g mu bi ky pl mt ilm

g mu pl bi pa ky ilm

TH-519

-pl

+ky+jd

-bi

+mu

-H2O

+pa+mt-g

-pa

-bi

-pl+g

-mt+mu

-g

+g-jd

-ru

-pa

+ky-chl

+chl +H2O -pa -chl -ru +ilm

H2 O

-saturated solidus

g bi pa ky pl mt ilm g bi pa pl mt ilm

13

12.5

12600 625 650

0.22

0.23

0.24

0.84

0.853.10

Si (c.p.f.u.)XgrXalb

g jd ky mu bi pl ep ilm

g ky mu bi pl ep ilmg pa mu bi pl ep ilm

g mu bi pl ep ilm

g mu pl

bi pa ilm

a

b

Figure 3.26: NCKFMASHTO P–T pseudosection for a representative Zentralgneiss metatonalite (Table 3.9; Winkler,

1996). Red shaded area highlights stability field of domain assemblage present in samples TH–519 and ASA–41a: garnet

+ clinozoisite + muscovite + biotite + plagioclase (+ quartz + opaques). Italicized labels around the plot perimeter mark

phase appearance/disappearance up-grade. White dashed line is the experimentally determined wet-melting solidus of a

biotite granite (Johannes & Holtz, 1996). Box shows position of inset b, in which isopleths for Si in muscovite (c.p.f.u.),

Xgr in garnet (Xgr=Ca/(Ca+Fe2++Mg+Fe3+) and Xalb in plagioclase (Xalb=Na/(Na+Ca), are plotted.

conditions at ⇠500 meters below the cover, in the Zentralgneiss is 13.1±2.1 kbar and 601±38�C.The following presents a discussion of the reactions responsible for development of the domain

91

assemblage and the accuracy of the pseudosection.

The absence of K-feldspar and discrepancies between calculated and observed garnet, muscoviteand plagioclase compositions confirms that chemical equilibrium was not attained on whole-rocklengthscales, or, that the bulk composition of Winkler (1996) does not reflect the compositionof samples TH–519 and ASA–41a. Notably, calculated garnet is significantly less grossular-richthan observed (Xgr = 0.22–0.24, compared to ⇠0.56), muscovite is predicted to be less silica-richthan observed (⇠3.1 versus ⇠3.3 Si c.p.f.u.) and plagioclase has values of Xalb between 0.83–0.85,compared to 0.93–0.99 in TH–519.

In K-feldspar-bearing metagranitoids, the grossular component of garnet is formed via hy-dration (reaction 1 beneath), whereas the almandine component forms by dehydration of biotite(reaction 2 beneath) (Proyer, 2003, and references therein):

[1.] anorthite + K-feldspar + H2O = grossular + muscovite + quartz[2.] annite + muscovite + quartz = almandine + K-feldspar + H2O

If K-feldspar is not present, due to high-H2O conditions (as in Fig.3.26), K released by annitebreakdown is incorporated by muscovite at higher pressures. The pseudosection calculations showthat clinozoisite preferentially sequesters Ca liberated by anorthite breakdown over grossular, underamphibolite facies conditions. This suggests that the Fe2O3 content of the bulk composition hasa vital control on garnet composition: more oxidised bulk compositions will stabilize epidote–clinozoisite across a larger P–T range than lower values, leading to less grossular-rich garnet.This hypothesis is strengthened by the observations of Koons et al. (1987), who noted that inmetagranites of the Sesia Zone (W.Alps), K-feldspar was consumed during prograde re-equilibrationand zoisite was stabilized on the whole-rock scale.

In the Zentralgneiss samples, phengite formed in the domain assemblages as a result of thefollowing, pressure-dependent reaction (Massonne & Schreyer, 1987):

biotite + K-feldspar + quartz + H2O = phengite

If present, Fe3+ can substitute for octahedral Al in place of Si during tschermak exchange to formphengite. This would account for low calculated Si isopleths in Fig.3.26b.

A further di�culty in modeling equilibration of metagranitoid rocks is that the activity ofH2O plays a critical role in assemblage development and yet is very di�cult to constrain. Theproduction of phengite by the reaction detailed above requires hydration, either by external sourcesor by dehydration of biotite. Once arrested, kinetic e↵ects will arrest further mineral growth. Thismeans that biotite-absent assemblages in metagranitoid rocks most likely preserve the conditionsat which the P–T path becomes tangential to a specific H2O-mode isopleth, i.e. H2O is no longera free phase. The fact that relict igneous biotite is preserved in Zentralgneiss samples shows thatthe rocks did not experience full dehydration and that aH

2

O was non-zero. Furthermore, accordingto Fig.3.26, the presence of biotite shows that peak pressures were <14 kbar.

92

The fluid-absent breakdown of albite to jadeite and quartz provides a minimum/maxium es-timate of peak pressure in the absence of information on aH

2

O. The sole reported occurrence ofjadeite from the Venediger complex is detailed in section 3.5; sodic pyroxene is yet to be observedin the Zentralgneiss. Its absence in samples TH–519 and ASA–41a constrains peak pressures to beless than 12.5–13 kbar at 600–650�C (Fig.3.26a).

3.5.6 Tectonic implications

Figure 3.27 shows a compilation of P–T data, including those calculated above, derived from theZentralgneiss and the Inner Schieferhulle cover nappes.

With the exception of the P–T estimate generated by Zimmermann et al. (1994), using jadeiticpyroxene from a garnet amphibolite at the base of the Inner Schieferhulle pile, all available P–T

data for both the Zentralgneiss and cover nappes lie on Barrovian geotherms, between 16–25�C.km(Fig.3.27). Pressures calculated from Zentralgneiss samples lie in the range: 11–15 kbar (TH–519and SP614), whilst those calculated from cover-nappe assemblages lie between 6–11 kbar (ASA–42a and published data)—apart from sample ASA–38a, which yields considerably higher pressures:13.7±2.2 kbar. Collectively, this shows that metamorphic grade increases with structural depth inthe Venediger complex—as shown by the concentric geometry of oxygen-isotope isotherms, whichpeak at ⇠600–640�C in the core of the Zentralgneiss (Hoernes, 1974).

Tentative decompressive P–T trajectories are constructed for the both the Zentralgneiss andInner Schieferulle cover nappes by combining suites of P–T data displayed in Fig.3.27. The Zen-tralgneiss P–T path (blue arrow - Fig.3.27) is characterised by near-isothermal decompressionfrom peak temperatures of ⇠600�C at 11–15 kbar, to less than 7.5 kbar and 550–600�C, whereH2O- and CO2-rich fluid inclusions were trapped during emplecement of late-stage, discordantquartz+plagioclase+biotite veins. In contrast, the Inner Schieferhulle was exhumed (red arrow -Fig.3.27), whilst heating, from peak pressures close to 10 kbar, to ⇠7 kbar and ⇠550�C—based onthe calculations of Selverstone et al. (1984).

Collectively, these data show that the parautochthonous Inner Schieferhulle nappes were jux-taposed with the Zentralgneiss basement complex prior to attaining the thermal peak of Alpinemetamorphism. Also, the similarity in the geometries of decompressive P–T paths suggests thatthe unit behaved coherently during decompression.

However, it is unclear to what extent the Venediger complex was subjected to HP–LT con-ditions, prior to Barrovian metamorphism. This e↵ectively constrains whether a thermal-regimereflecting subduction was active when continental collision occurred in the Eastern Alps. In addi-tion to the findings of Zimmermann et al. (1994), the presence of kyanite and pseudomorphs afterlawsonite in hornblende micaschists of the Inner Schieferhulle, studied by Selverstone et al. (1984),is strong evidence for pre-Barrovian LT–HP metamorphism. According to the P–T trajectoriesdisplayed on Fig.3.27, early HP metamorphism would have operated at pressures greater than⇠10 kbar for the Inner Schieferhulle nappes and greater than ⇠12 kbar for the Zentralgneiss. Un-

93

EBS

EC

GS

EA

300 400 500 600 700Temperature (°C)

Pre

ssur

e (k

bar)

BS

A4

8

6

10

12

14

16

18

2

1

2

2

3

4

57

6

8

9

+ La

w jad + qtz

alb

Figure 3.27: P–T evolution of the Venediger complex, including average P–T estimates presented here and a compila-

tion of existing data. Red path shows the inferred P–T trajectory of Inner Schieferhulle garbenschists based on Selverstone

et al. (1984). Purple trajectory corresponds to P–T path inferred from sample ASA–42a. Blue path corresponds to P–T

trajectory inferred for the Zentralgneiss complex. Metamorphic facies based on Evans (1990) and uses the following nota-

tion: GS=greenschist; BS=blueschist; EBS=epidote blueschist; EC=eclogite; EA=epidote amphibolite; A=amphibolite.

Position of albite breakdown reaction taken from (Holland, 1980). P–T data as follows: 1. Jadeitic pyroxene from the

Inner Schieferhulle – (Zimmermann et al., 1994); 2. Hornblende garbenschists of Inner Schieferhulle – (Selverstone et al.,

1984); 3. Inner Schieferhulle P estimates from mu + qtz + kfs + bi – (Kurz et al., 1998b); 4. Sample ASA–42a; 5. Sam-

ple ASA–38a; 6. Sample SP614; 7. Sample TH–519; 8. Oxygen isotope temperature interval for Inner Schieferhulle and

Zentralgneiss – (Hoernes, 1974); 9. Trapping temperatures of H2

O–CO2

-bearing fluid inclusions from the Zentralgneiss –

Cesare et al. (2001). Colour scheme of P–T paths same as P–T estimates; dashed line = poorly constrained.

94

der such conditions, glaucophane and lawsonite would be stable in metasedimentary cover units,whereas jadeite would be stable in metagranitoid protoliths characteristic of the basement. Theirabsence may be accounted for by complete breakdown during high temperature equilibration as-sociated with the Barrovian event, or kinetically inhibited growth conditions, such as the lack ofan available grain boundary fluid, coupled with short (sub-Ma)-timescales of metamorphism.

3.6 P–T evolution of the Pennine nappes

During consumption of the Valais oceanic realm, each of the Penninic nappes experienced HP–LT

metamorphism. The Eclogite Zone was buried close to mantle depths (60–80 km), whilst the Innerand Upper Schieferhulle record burial to blueschist conditions, between 9 and 14 kilobars (30–45km). The Eclogite Zone was subsequently exhumed, under blueschist temperatures, to mid-crustallevels where it followed a common, decompressive, P–T trajectory through Barrovian conditionswith the Inner and Upper Schieferhulle units. A compilation of P–T paths for the Pennine nappesis presented in Fig.3.28.

A question of importance to East Alpine orogenesis is: when did juxtaposition of the Penninicnappe pile occur? Each of the units has been variably recrystallised during Barrovian metamor-phism, which occurred synchronously with D2 and regional fabric development (S2). Top-to-the-north shearing (D1) occurred throughout the nappe pile contemporaneous to, and postdating,blueschist-facies conditions. Furthermore, P–T trajectories of the Eclogite Zone and Schieferhullenappes converge within the blueschist facies, which suggests juxtaposition occurred during the wan-ing stages of subduction, prior to continental collision. Given that P–T conditions of blueschistfacies recrystallisation (Holland & Richardson, 1979; Eremin, 1994) in the Eclogite Zone span 10kilobars, between 20 and 10 kbar, the absolute depths of juxtaposition are uncertain.

The Glockner nappe represents an attenuated piece of oceanic lithosphere, whereas the Eclog-ite Zone is interpreted to represent a segment of transitional crust located close to the outermostcontinental shelf of the European continent (Kurz et al., 1998b, and references therein). To accom-modate the more distal position of the Glockner nappe relative to the Eclogite Zone, in addition tothe fact that the Glockner nappe only partially attained eclogite facies conditions (Großglocknereclogites, Dachs & Proyer, 2001), the Glockner nappe must have been accreted to the subduction-wedge during burial, whilst the Eclogite Zone was conveyed to mantle-depths.

The P–T paths of the Lower and Upper Schieferhulle are strikingly similar. Both units arecharacterised by early (pre- to syn-D1) burial to between 30–45 kilometres, along lawsonite-stablegeothermal gradients, followed by concomitant heating and decompression to 450–600�C and ⇠6–10kilobars. This alone does not prove that the units were juxtaposed prior to initiation of subduction.More convincing evidence is the increase in metamorphic grade with structural depth, both forBarrovian and HP metamorphic events (Figs.3.23 and 3.27). Furthermore, garnet growth pathscalculated from zoned garnets in the Lower and Upper Schieferhulle show concordant profiles of�P/�T (Selverstone, 1985). Early imbrication of the units could be explained by obduction of

95

the Glockner ophiolite suite on to the European margin prior to subduction, as observed in theWestern Alps, where the Lago di Cignana UHP terrain exposes preserved Jurassic litho-tectonicstratigraphy relating to a ocean-continent transition zone (Beltrando et al., 2010b). Such zonesare likely weak and provide a seed point for the initiation of subduction(e.g. Gerya & Stockhert,2006).

Correlation of P–T paths with major structural fabrics suggests the following sequence of tec-tonic events, which are illustrated in Fig.3.28. Prior to, or during, consumption of the TethyanValais ocean, the Glockner nappe was sutured with basement slices of the southern Europeanmargin. The composite margin sequence was then subducted to blueschist conditions, whilstseparate portions of the margin attained eclogite facies conditions. The Eclogite Zone was thenexhumed during active subduction, via a return-flow mechanism, and tectonically inserted intothe Schieferhulle block as an extrusion wedge, along a south-dipping thrust fault. Various base-ment slices were also involved in nappe stacking—the Rote Wande nappe was emplaced above theEclogite Zone along a thrust-sense contact. Entry of European continental crust into the subduc-tion zone caused a cessation in convergence and initiated isotherm relaxation within the thickenedcrustal pile, beneath the Austroalpine nappes.

Barrovian conditions were attained close to ⇠30 Ma Inger & Cli↵ (1997); Glodny et al. (2005),but estimates for the age of HP metamorphism range between 100–30 Ma (e.g. Bickle et al.,1975; Glodny et al., 2005). A precise age of the HP event in the Eclogite Zone is fundamentallyimportant to understanding the timing of nappe assembly and the onset of continental collision inthe Eastern Alps.

3.7 Chapter Summary

Conventional thermobarometry and pseudosection calculations on representative samples from theVenediger, Glockner and Eclogite Zone nappes yield similar P–T paths, characterised by early-HP

metamorphism followed by concomitant decompression and heating to Barrovian conditions.Garnet–chloritoid–kyanite equilibria from an eclogite-facies metapelite show that the Eclogite

Zone was subducted to the verge of UHP conditions: ⇠26 kbar and ⇠540�C. Such assemblagesare common throughout (U)HP terranes and consistently record peak temperatures in a narrowinterval between 520�C and 600�C at pressures between 20-26 kbar. This suggests that the as-sociation: garnet + chloritoid + kyanite is representative of a state of enhanced detachment forslab-top material.

Coexisting acmite–jadeite clinopyroxene and glaucophane are preserved in the cores of retro-gressive symplectites found in the upper levels of the Glockner nappe. Average P–T and pseudo-section calculations, which use a recalibrated a–X model for acmite–jadeite interaction in clinopy-roxene, show that the textures record decompression of the Glockner nappe from ⇠9–13 kbar and350–450�C to ⇠5–8 kbar at 450–550�C.

The Zentralgneiss preserves sub-millimetric garnet, biotite, muscovite and epidote in micro-

96

Europe Adria

Subduction

Pressure peak

Barrovian

EBS

EC

GS

EA

22

24

300 400 500 600 700Temperature (°C)

Pre

ssur

e (k

bar) BS

A4

8

6

10

12

14

16

18

2

+ law

jad + qtz

alb

20

26

28

30

coeqtz

CEC

LEC

Figure 3.28: Summary P–T paths and geodynamic evolution of the Venediger, Glockner and Eclogite Zone nappes.

Trajectories derived from the following: Eclogite Zone (blue): HP portion stems from this study, blueschist–greenschist

from Holland & Richardson (1979), Eremin (1994) and Hoschek (2001); Glockner nappe (green): lower grade path is

from this study, whereas high-P path is from Dachs & Proyer (2001); Venediger nappe: lower-grade path is for the

Inner Schieferhulle as determined by this study, whilst the higher grade loop is derived from sample TH–519—this study.

Facies definitions follow Fig.3.27. Circular markers correspond to tectonic cartoons in a clockwise (i.e. compressive) order.

During, or prior to, burial, the Inner and Upper Schieferhulle are juxtaposed; the Eclogite Zone occupies a more distal

position and attains mantle depths. Following the baric peak of metamorphism, the Eclogite Zone is detached from the

down-going slab and inserted within the Schieferhulle block. Subsequent emplacement of the Austroalpine nappes causes

conductive heating of the Penninic nappe pile and is contemporaneous with uplift and dome formation. Cartoon scales

are approximate.

domains developed at the expense of plutonic plagioclase. Modeling of the assemblage shows that

97

the unit experienced high-grade Barrovian metamoprhism, close to solidus conditions: 10–13 kbarand 550–620�C.

Correlating P–T information with regional deformation events shows that HP metamorphismoccurred prior to, or contemporaneous with assembly of the Pennine nappe stack. The nappeswere concomitantly exhumed and heated as a coherent pile from blueschist to greenschist faciesconditions following arrival of the European continent at the East Alpine subduction zone.

98

Chapter 4

U–Th–Pb Isotope Geochronology

U–Th–Pb work completed in collaboration with Prof. Randy Parrish (NIGL), Dr. DanCondon (NIGL), Dr. Matt Horstwood (NIGL), Dr. John Cottle (UC Santa Barbara) andDr. Steve Noble (NIGL). U–Th–Pb data discussed in this chapter have been published inSmye et al. (2011).

4.1 Introduction

Central to understanding the behavior of the Earth’s lithosphere in zones of convergence is the trac-ing of individual rock units through pressure-temperature space with time. Pressure–Temperature–time (P–T–t) loops are a powerful tool through which the rates of subduction related burial andsubsequent exhumation—processes which bracket mountain belt formation, can be constrained.They also provide important constraints for thermo-mechanical models of orogenesis (e.g. Beau-mont et al., 2001; Burov et al., 2001; Gerya et al., 2004; Burg & Gerya, 2005; Beaumont et al., 2006;Gerya & Stockhert, 2006; Yamato et al., 2007; Faccenda et al., 2008). Decay of radiogenic nuclides,often concentrated into accessory minerals by metamorphic equilibration, provides a measure ofabsolute time throughout an orogenic episode. However, this chronometric information and hencetectono-metamorphic rates can only be fully realized if a link between radiogenic isotope systemat-ics and P–T conditions is established. This linkage continues to be the subject of extensive study(e.g. Pyle & Spear, 1999; Foster et al., 2000; Hermann, 2002a; Rubatto, 2002; Foster & Parrish,2003; Hermann & Rubatto, 2003; Wing et al., 2003; Harley et al., 2007; Janots et al., 2007, 2008;Konrad-Schmolke et al., 2008). See Vance et al. (2003) for a comprehensive review.

Chapter 3 details the P -T trajectories of individual Pennine nappes throughout the Alpinecollisional cycle. In the following chapter, the U-Th-Pb system in allanite is used to addressthe fundamental question of the timing of HP metamorphism in the Tauern Window. This workforms a framework for the interpretation of 40Ar/39Ar data (chapter 5) in addition to underpinningmodeling of the Tauern Window’s thermal evolution, presented in chapter 6.

99

4.2 Aims

The principal aim of this chapter is to constrain the timing of peak eclogite facies metamorphismin the Eclogite Zone. Despite the Eastern Alps being amongst Earth’s best studied orogens,the age of HP metamorphism is weakly constrained. In the Western Alps, HP metamorphismoccurred between ⇠34–45 Ma along the mountain belt (see Berger & Bousquet, 2008, for review).Therefore, the age of eclogite formation in the Eastern Alps will provide an important controlon along-strike di↵erences in the timing of collision between Adria and the southern Europeanmargin. Furthermore, structural fabrics in the Eclogite Zone and Glockner nappe show that HP

metamorphism preceded assembly of the Penninic nappe stack and associated conductive relaxationof isotherms thought to drive the ca. 30 Ma Barrovian overprint. Accurate determination of theage of HP metamorphism would provide an upper-bound to the amount of time during which suchheating was active—key to understanding the driving mechanisms of heat transfer responsible forregional metamorphism.

4.3 Previous work

The age of eclogite facies metamorphism within the Eclogite Zone of the Tauern Window hasbeen the subject of discussion for over 30 years (e.g. Miller, 1974; Frank et al., 1987; Zimmermannet al., 1994; Glodny et al., 2005). Overthrusting of the Austroalpine nappes contemporaneouswith subduction of the Penninic margin was thought to have occurred during the late Cretaceouson the grounds of structural relations in the Gosau Beds (50–80 Ma) of the Northern CalcareousAlps (Austroalpine) (Oberhauser, 1968). Furthermore, detrital glaucophane thought to have beenderived from exhumed high pressure Penninic rocks is present in sections of the Campanian (ca. 80Ma) flysch (Oberhauser, 1968). Sixty five–100 Ma K–Ar mineral ages associated with amphibolitefacies strain markers at the base of the Altkristallin were thought to provide a maximum age limiton the timing of continental collision (Oxburgh et al., 1966; Brewer, 1969). Collectively thesedata were interpreted by classical studies of the Tauern Window’s thermal evolution to suggestthat high pressure metamorphism must have occurred ca.65–100 Ma (Oxburgh & Turcotte, 1974;Bickle et al., 1975; Bickle & Hawkesworth, 1978; England, 1978).

The first isotopic study of the age of peak pressure metamorphism within the Eclogite Zone wasundertaken by Van Breemen & Hawkesworth (1980), who applied the Sm–Nd method to garnetporphyroblasts, which are petrologically linked to eclogite facies conditions (i.e. Holland, 1979).Isotopic disequilibrium caused by microscopic inclusions within garnet cores, precluded calculationof geologically significant ages. Garnet growth rates in the Upper and Lower Schieferhlle from thewestern Tauern Window were calculated by Christensen et al. (1994). They used garnet–matrixfractionation of Rb–Sr isotopes to temporally calibrate the P–T path of Selverstone et al. (1984),showing that garnet growth, initiated ca. 65 Ma, prior to the ⇠10 kbar pressure peak being attained

100

ca. 45 Ma. This pressure peak is thought to be analogous to the locally observed blueschist faciesoverprint in the Eclogite Zone and also to the pressure peak in the Glockner nappe Zimmermannet al. (1994).

A wealth of 40Ar/39Ar data has been collected from K-rich phases across the nappe stack.Zimmermann et al. (1994) interpreted phengite ages between ca. 32–36 Ma, obtained from EclogiteZone lithologies, as representing cooling ages postdating eclogite facies metamorphism and thereforethe approximate age at which the Eclogite Zone was inserted between the Venediger and Glocknernappes. Ratschbacher et al. (2004) present 40Ar–39Ar ages from high-pressure amphibole, phengiteand phengite+paragonite mixtures, which are thought to document exhumation of the EclogiteZone through ⇠15 kbar and ⇠500�C at ca. 42 Ma to ⇠10kbar and ⇠400�C at ca. 39 Ma. Theysuggest that peak pressure conditions were attained in the Eclogite Zone 45 Ma. More recently,Kurz et al. (2008) combine phengite 40Ar–39Ar ages obtained via stepwise heating, with detailedmicro-structural analysis of host eclogites to propose that the pressure peak was reached 38Ma—see section5.2 for a more detailed discussion of previous 40Ar–39Ar study. Importantly, all40Ar–39Ar studies document varying degrees of excess radiogenic (40Ar) argon.

Glodny et al. (2005) applied the Rb-Sr method to white mica, omphacite, paragonite, apatite,carbonate, garnet, epidote and clinozoisite separated from pristine eclogite facies assemblages.Their multimineral isochrons consistently yield ages of 31.5±0.7 Ma, which are interpreted torecord peak conditions of 20–25 kbar and 600�C in the Eclogite Zone. However, the e↵ect ofisotopic resetting caused by subsequent Barrovian metamorphism is unknown (e.g. Thoni &Jagoutz, 1992). Glodny et al. (2005) show that Rb and Sr isotopes do not show evidence for opensystem behavior in their eclogite host rocks. However, this contradicts the findings of Inger & Cli↵(1997), who claim that Rb-Sr isochron ages between 26–30 Ma, obtained from whole-rock, epidote,carbonate and white mica from the Eclogite Zone, have to be interpreted as mixtures betweenpost-eclogite cooling and the greenschist–amphibolite overprint.

Prior to this study, the U–Th–Pb system was yet to be applied successfully to dating peakmetamorphism. Miller et al. (2007) attempted to date zircons from a layer of strongly retrogressedjadeite-bearing gneiss exposed in the central Eclogite Zone, but were unable to observe any evidenceof Alpine high pressure metamorphism in the zircons studied. Their ages spread between 288–691Ma, suggesting peri-Gondwanan and Variscan sources.

Therefore, despite there being a wealth of previous study, the timing of eclogite facies meta-morphism in the Eclogite Zone is only weakly constrained and awaits application of a single-graingeochronometer.

Unlike the age of high pressure metamorphism, the timing of Barrovian (re)crystallisation (⇠7kbar and 520–570�C, Zimmermann et al., 1994) in the Tauern Window is relatively well constrainedto between 27–32 Ma. Glodny et al. (2005) performed Rb–Sr analyses on cogenetic white mica,titanite, amphibole and plagioclase from a structurally discordant vein assemblage in a greenschist-facies retrogression aureole within mafic eclogite. They obtained an isochron age of 31±0.5 Ma(MSWD=0.63)—i.e. within error of their 31.5±0.7 Ma age calculated for the pressure peak. This

101

age is supported by the data of Gleißner et al. (2007), who carried out Rb–Sr isotopic analyses of apseudomorphic assemblage after blueschist facies lawsonite in the Glockner nappe. A multimineralinternal isochron involving white mica, apatite, epidote, clinozoisite, calcite and albite yields anage of 29.82±0.54 (MSWD=2.2), which is interpreted as representing syntectonic growth of thegreenschist facies matrix assemblage. Similarly, Inger & Cli↵ (1997) present white mica, epidoteand whole rock Rb–Sr isotopic data from eleven di↵erent localities from within the Eclogite Zoneand Glockner nappe. All of the ages lie between 27–31 Ma and are interpreted to record greenschistfacies metamorphism concomitant with penetrative deformation, influx of fluids and resetting ofSr isotopes. Low Si phengite, which formed during the Barrovian event (Ma2 of Zimmermannet al., 1994), 40Ar/39Ar ages obtained by Zimmermann et al. (1994) cluster around 27 Ma and arethought to represent cooling of the nappe pile following peak temperatures.

4.4 Geochronology of high–P metamorphism

Accurate and precise dating of high pressure metamorphism is a challenging exercise because theprocesses that drive formation of eclogite and blueschist facies rocks operate on plate velocitytimescales (i.e. cm.a�1); thus, requiring age resolution of generally 1–5 Ma (e.g. Rubatto &Hermann, 2001; Parrish et al., 2006). This is in contrast to temperature dominated events, suchas Barrovian metamorphism, which are thought to occur over timescales driven by the thermaldi↵usivity of crustal lithologies i.e. ten’s of millions of years (e.g. England & Thompson, 1984).

The geochronology of high pressure metamorphic events is best dated by decay schemes inhost minerals with a closure temperature (Tc; Dodson, 1973, 1976, 1979) in excess of the peaktemperature of metamorphism so as to avoid di↵usive loss of daughter nuclides after the meta-morphic peak. Furthermore, minerals which preferentially exclude daughter atoms of the decayscheme on formation allow more precise and accurate age determination, i.e. zircon U–Pb dat-ing. Geochronology using mineral chronometers with a high propensity for sequestering daughteratoms on crystallisation often incurs significant errors (>2%, 1�) associated with age corrections,i.e. common Pb corrections. Perhaps most importantly, the geochronometer of choice must be ableto be linked to the P -T conditions of interest (e.g. Hermann & Rubatto, 2003); without such alink, the age data are open to interpretation and are of limited value. Further issues which hamperthe application of high retentivity decay schemes (147Sm–143Nd, 176Lu–176Hf, 238U–206Pb, 235U–207Pb, 232Th–208Pb) to high pressure geochronology include: i. isotopic disequilibrium causedby the inclusion of parent or daughter nuclide rich sub-microscopic phases in host minerals (e.g.Van Breemen & Hawkesworth, 1980); ii. incomplete isotopic equilibration; inheritance of previousisotopic signatures (e.g. Luais et al., 2001); iii. interference of daughter nuclides from paralleldecay chains e.g. excess 206Pb from the decay of 230Th (e.g. Scharer, 1984).

Figure 4.1 shows that peak temperatures experienced by the Eclogite Zone during eclogitefacies metamorphism were beneath those required to facilitate di↵usive loss of radiogenic Pb fromzircon, monazite and allanite. This makes the U–Th–Pb system the system of choice for dating

102

U-Pb Zircon

U-Pb Monazite

Ar-Ar Hornblende

U-Pb Titanite

Rb-Sr Whole Rock

Sm-Nd Garnet

U-Pb Rutile

Rb-Sr Muscovite

Ar-Ar Muscovite

Rb-Sr Biotite

FT Titanite

Ar-Ar Biotite

FT Zircon

Ar-Ar K-feldspar

FT Apatite

(U-Th)/He Apatite

Surface exposure dating

T (°C)

Lu-Hf Garnet

U-Pb Allanite

200 400 600 800

Figure 4.1: Relative closure tem-

peratures to di↵usive loss of daugh-

ter nuclides for commonly used ra-

diogenic isotope systems. The green

band delineates the peak temper-

ature interval experienced by the

Eclogite Zone during eclogite facies

metamorphism. The width of boxes

indicate the relative degrees of un-

certainty associated with each de-

cay scheme—closure temperatures

are dependent on many factors in-

cluding i. cooling rate; ii. grain

size; iii. deformation; lattice de-

fects in the host mineral. Refer-

ences are as follows: U–Pb zircon,

Lee et al. (1997); U–Pb monazite,

Spear & Parrish (1996); U–Pb al-

lanite, Heaman & Parrish (1991);

Lu–Hf garnet, Scherer et al. (2000);

Ar–Ar hornblende, Harrison (1981);

McDougal & Harrison (1988); U–Pb

titanite, Heaman & Parrish (1991);

Cherniak (1993); Rb–Sr whole rock,

Jager et al. (1967); Purdy & Jager

(1976); Sm–Nd garnet, Ganguly

et al. (1998); U–Pb rutile, Kooij-

man et al. (2010); Ar–Ar muscovite,

Purdy & Jager (1976); Rb–Sr bi-

otite, Jenkin et al. (2001); Fission

track (FT) titanite, Coyle & Wagner

(1998); Ar–Ar biotite, Harrison &

McDougall (1985); FT zircon, Zaun

& Wagner (1985); Yamada et al.

(1995); Ar-Ar K-feldspar, Harrison

& McDougall (1982); FT apatite,

Naeser (1979); (U-Th)-He apatite,

Stockli et al. (2000).

103

Isotope Abundance(%) Half life(a) � Reference

238U 99.2743 4.468x109 1.55125x10�10 1235U 0.7200 0.7038x109 9.8485x10�10 1234U 0.0055 2.45x105 2.829x10�6 2

232Th 100.00 14.010x109 4.9475x10�11 1

Table 4.1: Abundances, halflives and decay constants (�) of principle naturally occurring isotopes of the U and Th.

Reference notation as follows: 1 = Steiger & Jager (1977); 2 = Lide & Frederikse (1995).

such high grade events.

4.5 U-Th-Pb geochronology

4.5.1 Principles and methodology

The decay of U and Th to stable isotopes of Pb is the basis for several important methods ofdating earth materials. Uranium (Z=92) and Th (Z=90) are both members of the actinide seriesand have similar electron configurations and atomic radii (U4+ = 1.05 A; Th4+ = 1.10 A), andtherefore display chemically coherent behaviour. Uranium has a series of three naturally occurringradioactive isotopes: U238, U235 and U234, whereas Th principally occurs as 232Th.

The decay of 238U to 206Pb can be described by the following expression, which shows thateach atom of 238U decays to an atom of 206Pb by emission of eight ↵ particles and six � particles,where Q = 47.4 MeV.atom�1 (Wetherill, 1966):

23892 U !206

82 Pb + 842He + 6�� + Q (4.1)

Each atom of 235U decays to form an atom of 207Pb after emission of seven ↵ particles and four� particles. Here, Q = 45.2 MeV.atom�1 (Wetherill, 1966):

23592 U !207

82 Pb + 742He + 4�� + Q (4.2)

The decay of 232Th gives rise to the emission of six ↵ particles and four � particles, result-ing in the formation of 208Pb. This is represented by the following expression, where Q = 39.8MeV.atom�1 (Wetherill, 1966):

23290 Th!208

82 Pb + 642He + 4�� + Q (4.3)

Accumulation of radiogenic Pb isotopes is expressed mathematically by equations derivablefrom the law of radioactivity (Rutherford & Soddy, 1902a,b). Radiogenic Pb isotopes are nor-malised to 204Pb, the only nonradiogenic Pb isotope:

104

206Pb204Pb

=✓206Pb

204Pb

◆

i

+238U204Pb

⇣e�

238

t � 1⌘

(4.4)

207Pb204Pb

=✓207Pb

204Pb

◆

i

+235U204Pb

⇣e�

235

t � 1⌘

(4.5)

208Pb204Pb

=✓208Pb

204Pb

◆

i

+232Th204Pb

⇣e�

232

t � 1⌘

(4.6)

Uranium–Pb and Th–Pb ages are determined by measuring isotopic ratios of U, Th and Pbvia mass spectrometry, and rearranging the respective equations above for t. It is important tonote that the ratios on the left of the equations presented above, correspond to present day values,i.e. after t years radiogenic ingrowth. Initial isotope ratios of Pb are taken from Pb-evolutionmodels (i.e. Stacey & Kramers, 1975). Ages are only sensitive to initial isotope ratios if sampleU/Pb of Th/Pb ratios are low, or if the rock\mineral is young (Tertiary–present). Generally, zirconexcludes non-radiogenic Pb during crystallisation, making it an extremely useful petrochronometer;Ca-bearing phases, such as titanite and allanite, however, have a propensity to sequester largeconcentrations of non-radiogenic Pb.

Age estimates from each of the decay schemes will yield the same result, known as concordancy,under the following conditions: (i.) the dated mineral has remained a closed system to U, Th andPb since its formation/equilibration, i.e. all U, Th and Pb residing in the mineral was incorporatedat the time of its formation/equilibration; (ii.) the initial Pb composition is well known; (iii.) thedecay constants are well known; (iv.) U has not been fractionated; v. the analytical measurementsare free from systematic error.

4.5.2 Accessory minerals in Eclogite Zone

An extensive sampling campaign (> 100 samples collected), carried out during summers 2007–09,focused on identifying suitable accessory phases for dating the eclogite facies event by U–Th–Pbgeochronology. The results of which are summarized in the following sections.

4.5.2.1 Techniques

Samples of potential interest were cut into slabs for thin section construction, undertaken atKingston University. Sections were then screened for microscopically visible accessory phases un-der plain- and cross-polarised light using a binocular microscope. Target slides were subsequentlycarbon coated and analysed in Back Scatter (BSE) mode on the Department of Earth Science’sJeol JSM 820 scanning electron microscope (SEM). Rare earth element rich accessory phases havehigh average atomic masses and are therefore bright and easily located under BSE. An energy dis-persive X-Ray detector helped distinguish U–Th accessory phases from other bright phases, suchas pyrite, magnetite and ilmenite. Digital images were taken of target grains, which were laterused for locating chemical domains and grain boundaries during laser ablation mass spectrometry.

105

[Zr]bulk [Zr]garnet [Zr]rutile [Zr]cpx [Zr]zircon Zircons (kg�1) Zircons (cc�1)

50? 3 200 4 481,202⇧ 1847 6100† – – – – 4640 16

Table 4.2: Mass balance calculations to estimate zircon density in a typical Tauern eclogite. Garnet, omphacite and

rutile are considered as the only volumetrically significant Zr–bearing phases. All [Zr] values are in ppm. Mineral modes

in volume % are as follows: garnet, 30%; clinopyroxene, 50%; rutile, 7%. Representative concentrations taken from

Bocchio et al. (2000) (garnet and clinopyroxene), and Zack & Luvizottow (2006) (rutile). Density for Tauern eclogite was

estimated to be 3.5 g.cc�1; mineral densities obtained from Deer et al. (1992). Zircon assumed to be 200µm long, and

deep, and 100µm wide. ? Lower bound of bulk eclogite [Zr] from Hock & Miller (1980). † Upper bound of bulk eclogite

[Zr] from Hock & Miller (1980). ⇧ Corresponds to 65 wt% ZrO2

(Deer et al., 1992).

4.5.2.2 Mafic lithologies

Accessory minerals present within Tauern eclogites datable via the U–Pb method are limited tozircon, rutile and sphene.Zircon can be an extremely powerful U–Th–Pb geochronometer due to its refractory nature andability to withstand reworking subsequent to peak metamorphism (e.g. Warren et al., 2003).Themajor and trace element geochemistry of mafic eclogites within the Eclogite Zone suggests a stronga�nity to ocean floor tholeiites (Hock & Miller, 1980). Correspondingly, Zr concentrations lie be-tween 50–150 ppm, with a graphical mean of ⇠100 ppm (Fig.5, Bickle & Pearce, 1975; Hock &Miller, 1980). Assuming new growth of zircon under eclogite facies temperatures, simple mass bal-ance calculations (Table.4.2) predict that for this range of Zr concentrations, between 6–16 zircongrains will be present per cubic centimeter of eclogite. Although these figures provide only a roughestimate of zircon density, they provide an explanation for its apparent paucity in Tauern eclogitesstudied using SEM. The best example of zircon in mafic eclogite is found in sample ASA–08–70a,a coarse grained (cm-scale garnet and omphacite), metagabbroic eclogite whose eclogitic assem-blage of, garnet–omphacite–kyanite–barroisite–rutile–epidote–quartz, pseudomorphs the previousgabbroic texture. Four zircon grains were observed in thin section analysis, which fall broadly intotwo textural settings: i Sub-rounded zircon: found as small grains (10µm diameter), in matrixomphacite and amphibole. Their small size and morphology is suggestive of metamorphic growthunder rapidly evolving conditions (Corfu et al., 2003); ii. Rounded zircon: dominant grain mor-phology with higher aspect ratio is indicative of magmatic zircon. They are found both in matrixand as inclusions in garnet. Because of the rare nature, small grain size and the probability thatmost grains are likely to be of magmatic origin (Miller et al., 2007, Rubatto pers. comm.), zirconU–Pb geochronology was not pursued.

Rutile and sphene are ubiquitous throughout all types of mafic eclogites observed. Typicallyboth phases occur as µm–mm scale flakes and grains in the matrix (Fig 4.2). However, rutile is of-ten observed as porphyroblasts located on the margins or within, quartz–kyanite–zoisite±dolomite

106

bearing eclogite facies mobilisates (i.e. ASA–08–30c; 4.2). The mechanisms driving formationof such aggregates under peak pressure conditions are uncertain, but must involve considerablechemical flux of SiO2, Al2O3 and TiO2 from the host eclogite into the selvage (Ague, 1995; Franzet al., 2001; Beitter et al., 2008). Given that the U–Pb system in rutile has a large range of ef-fective closure temperatures to di↵usive Pb loss (490–640�C, Kooijman et al., 2010) both matrixand mobilisate rutile were targeted for U–Pb geochronology in the hope that age data may o↵erconstraint on the timing of crystallisation or cooling.

4.5.2.3 Metasedimentary lithologies

Metasedimentary samples were screened for REE, U- and Th-bearing phosphate accessory mineralssuch as monazite, florencite, apatite, and xenotime, in addition to the REE-rich epidote end-member, allanite. Detrital zircon, identifiable by its diagnostic rounded grain morphology, waspresent in all samples studied; no evidence for metamorphic recrystallisation around grain rimswas observed. Generally, REE-bearing accessory phases are scarce in Eclogite Zone metasediments;allanite is the most common U–Th–Pb-bearing accessory phase and was found in only three (TH–680, HOS–10 and ASA–08–36h) of thirty five metasedimentary samples screened. Allanite occursin Eclogite Zone samples as euhedral–skeletal cores to clinozoisite porphyroblasts (100–1000µmlength; Fig.4.3) which are often aligned parallel with the dominant fabric (S0+S1). In sampleTH–680 and HOS–10, allanite is also found as inclusions within garnet poikiloblasts, indicative ofgrowth early in the rock’s metamorphic history.The paucity of monazite in Eclogite Zone metasediments is notable. Monazite was observed in asingle sample: HOS–10, obtained from Gert Hoschek (University of Innsbruck), who collected therock from the core region of the Eclogite Zone, approximately 1 km northwest of the Eissee lakeat 2900m. Samples TH–680 and HOS–10 are from closely spaced localities and contain identicalsilicate mineralogies comprising quartz, phengite, garnet, chloritoid and kyanite in decreasing orderof abundance (Smye et al., 2010; Hoschek et al., 2010). Monazite in HOS–10 is present as bothmatrix aggregates (10–200 µm diameter; Fig.4.3) and included growths within garnet where itappears to have grown as pseudomorphs after xenotime or allanite.There are two possible explanations for the rarity of monazite relative to allanite in Eclogite Zonepelites, as typified by the di↵erences in accessory phase mineral parageneses between TH–680 andHOS–10:

1. Bulk composition – rocks containing low concentrations of CaO, with su�cient LREE andPO4, generally form monazite within the greenschist facies, which then persists throughoutmetamorphism (Overstreet, 1967). Rocks with higher CaO are thought to form allanite underlow metamorphic conditions, which then breaks down close to the staurolite and/or kyaniteisograd, (e.g. Wing et al., 2003; Tomkins & Pattison, 2007), in some cases, to form monazite(Smith & Barreiro, 1990; Janots et al., 2008). Sample TH–680 contains 1.1 mol.% CaO,

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Figure 4.2: U–Pb accessory minerals in mafic eclogites. a. Metagabbroic eclogite ASA–08–70a. Note the apple green

omphacite aggregates after augite and garnet–rich matrix. b. Sub-rounded zircon, 3–4µm diameter in matrix omphacite

(ASA–08-70a). The grain’s morphology and matrix environment suggest a metamorphic origin. c. Rounded zircon grain

included within garnet rim region. The grain’s rounded geometry is typical of magmatic zircon (ASA–08-70a). d. Rutile

and sphene in matrix kyanite and omphacite respectively (ASA–08-70a). Note the embayed edges to the sphene grain

suggestive of late resorption. e. Rutile ‘penny’ in eclogite ASA-08–30c. The grain is located proximal to a peak eclogite

facies quartz–kyanite–zoisite bearing mobilisate.

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wheras monazite bearing HOS–10 has a much lower value of 0.24 mol.%, (Hoschek et al.,2010), in line with prevailing thought.

2. Temperature – in rocks of suitable CaO, LREE and PO4 composition to form allanite andmonazite, the decomposition of allanite to form monazite is highly temperature-sensitive inthe interval close to the chloritoid-out boundary, i.e. 550–600�C, (Janots et al., 2007, 2008).This temperature interval is comparable to the peak temperatures experienced by metasedi-ments within Tauern eclogite facies metamorphism (Smye et al., 2010). Therefore, completetransition from allanite to monazite in the Eclogite Zone may well have been hampered bylow temperatures.

Apatite is commonly present in garnet-bearing metasediments as micrometer-scale porphyroblastswithin phengite matrix domains (Fig.4.3). Sample HOS–10 shows apatite with embayed grainedges, suggestive of later reworking. Similarly, rutile forms between 0.1–3 vol.% of metasediments,occurring as µm–mm scale flakes within the matrix.

Due to its large grain size, textural evidence to suggest early growth and the sample’s well-constrained P–T path (see section 3.2.4.5) e↵orts to date peak eclogite facies metamorphism byU–Th–Pb geochronology were focussed on allanite grains from sample TH–680.

4.6 Allanite and REE-rich epidote

Allanite–CaREEAl2Fe2+Si3O11O(OH), is an epidote group disilicate mineral primarily found asan accessory mineral in felsic igneous (Exley, 1980; Schmidt & Thompson, 1996) and regionallymetamorphosed rocks (Smith & Barreiro, 1990; Wing et al., 2003). It is related to epidote by thefollowing coupled substitution: (Khvostova, 1963; Ploshko & Bogdanova, 1963):

REE3+ + Fe2+ $ Ca2+ + Fe3+ (4.7)

and to clinozoisite by:REE3+ + Fe2+ $ Ca2+ + Al3+ (4.8)

The allanite structure can accommodate weight per cent quantities of U4+ and Th4+ and minorto trace amounts of Sr2+, Mn2+, Mg2+, P5+, Ti4+, Zr4+, Cr3+, Pb2+ and Ba2+ (Deer et al.,1992; Giere & Sorensen, 2004). Th is preferentially incorporated over U through substitutionfor REE (Th4++Ca2+$2REE3+: Peter Gromet & Silver, 1983) or variations in oxidation state(Th4++Fe2+$REE3++Fe3+: Giere et al., 1999). Allanite is often defined apart from REE-richepidote on the basis of a REE+Th content � 0.5 c.p.f.u. or <20 wt.% CaO (Giere & Sorensen,2004).

Hermann (2002a) showed that allanite is the dominant LREE and Th carrier in subductedcrust; in granitic systems allanite’s controlling influence on trace element evolution is well docu-mented(e.g. Peter Gromet & Silver, 1983; Sawka et al., 1984; Oberli et al., 2004). In metamorphic

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Figure 4.3: U–Pb accessory minerals in Eclogite Zone metasediments. a. Allanite–clinozoisite shard in garnet–chloritoid

micaschist ASA–08–36h. The grain appears to have been reworked—much of the allanitic core has broken down. Note

also the presence of small (10–20µm detrital zircon in matrix. b. Apatite porphyroblast in sample ASA–08–36h. Embayed

and serrated grain edges imply subsequent reworking. Note the abundance of detrital zircon. c. Matrix monazite cluster

in sample HOS–10. The cluster overgrows a chloritoid blast, which appears to have partially broken down. d. Monazite,

TiO2

and florencite symplectite in sample HOS–10. Bright domains are Ce–rich monazite, which form a dendritic pattern

separated by domains of TiO2

and rare florencite, which is recognised by its intermediate brightness. Note the angular

lozenge of monazite, similar to the prismatic form of allanite observed in sample TH–680 for example.

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systems, allanite’s wide range of potential compositions means that it often crystallises in responseto a wider range of P , T , fO

2

and aH2

O than monazite and zircon(e.g. Janots et al., 2007, 2008).As detailed above, allanite forms at the expense of detrital or metamorphic monazite up-T fromthe chloritoid isograd and is replaced by monazite at temperatures �⇠600�C (Wing et al., 2003;Janots et al., 2008). Allanite’s greenschist–amphibolite facies stability window, coupled with itsretentivity of REE and trace element patterns which reflect evolving equilibrium mineral parage-neses, means that it has the potential to record prograde metamorphic events, unlike zircon andmonazite which generally record peak-T conditions (e.g. Rubatto, 2002; Hermann & Rubatto,2003).

Given allanite’s propensity to sequester Th and U, it is a target accessory phase for geochronol-ogy (Heaman & Parrish, 1991). However, as with other Ca-bearing U–Pb and Th–Pb geochronome-ters (i.e. sphene), allanite often incorporates significant quantities of non-radiogenic Pb (Pbc), thusimpairing precision and accuracy of age determination. Furthermore, the wide range of chemicalcompositions displayed by the allanite–epidote–clinozoisite solid-solution makes identificantion of amatrix-matched standard, suitable for in-situ (i.e. Laser Ablation ICPMS and Ion Probe) isotopicmeasurements, extremely di�cult (Catlos et al., 2000; Gregory et al., 2007). Strongly radiogenicallanite becomes metamict over short timescales relative to zircon and monazite (Giere & Sorensen,2004, and references therein).

Nevertheless, the use of allanite as a geochronometer of prograde metamorphism is an emergingfield, which, when coupled with peak metamorphic age constraints, has the power to estimate ratesof subduction and residence time of buried units at depth.

4.7 Techniques

U–Th–Pb geochronology was performed at NERC National Isotope Geoscience Laboratories (NIGL)in Keyworth, Nottinghamshire and was funded by NERC grant IP/1008/1108 (£21,100).

4.7.1 Grain selection

Hand sample TH–680 measures ⇠100⇥60⇥80mm (Fig.3.2.2). Slabs for thin section constructionwere taken across the width of the sample to obtain a spatially representative suite of sections.A total of three probe sections (dimensions: 28⇥48mm) with thicknesses between 40–50µm wereconstructed for in situ analysis. This higher than normal thickness permits detailed texturalanalysis using a polarizing microscope while making it possible to ablate for long enough to obtaina statistically valid number of isotope ratios. Prior to BSE imaging, a polarizing microscope wasused to identify possible allanitic cores to clinozoisite grains. Allanite is distinguishable fromother epidotes by its brownish yellow pleochroism and varying degrees of metamictization, whichproduces anastomosing cracks and a decrease in refractive index and birefringence (Deer et al.,1992).

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Following microscope and BSE imaging, the locations of several grains identified as allaniteor REE–rich clinozoisite were recorded for quantitative chemical analysis using a Cameca SX–100electron microprobe at the Department of Earth Sciences, University of Cambridge. Major elementanalyses (peaks measured: MgK↵, SiK↵, NaK↵, AlK↵, FeK↵, TiK↵, MnK↵, CaK↵, YL↵, ThM↵,SrL↵, LaL↵, CeL↵, PrL↵, NdL↵, SmL↵, EuL↵, GdL↵, DyL↵, YbL↵) were acquired using a 5µm spotdiameter, 15 KeV accelerating voltage, a beam current of 10 nA and 20 s data collection time.Rare Earth Element and trace element acquisition used 15 KeV, 200 nA and 100 s respectively.Natural and synthetic minerals and oxides were used as standards. Raw data were corrected on–line according to Cameca X-PHI procedure. Given that each of the 14 REE has a minimum of 12X–Ray lines in the L spectrum, peak interference can be problematic. Accordingly, peak positionswere chosen and signals were corrected using the methodology and correction factors outlined byReed & Buckley (1998). Several grains were also qualitatively X–Ray mapped for Y, Ce and Thconcentrations to resolve chemical morphology. Element maps were acquired on the large TAPand PET crystals for X–Ray peaks corresponding to YL↵, CeL↵ and ThM↵. Running conditionscomprised a focused 2µm beam diameter, an electron probe current of 100 nA, an acceleratingvoltage of 15 kV and a pixel dwell time of 30 ms/pixel. The high intensity of Y, Ce and Th signalsprecluded the need to subtract background energy. Quantitative spot analyses of ThO2 were usedto identify target grains and grain domains for laser ablation mass spectrometry. Due to the highlyvariable ThO2 content of allanite–clinozoisite in TH–680, detection limits lie between 0.5–300% ofmeasured values; ThO2 values lie beneath detection limits in clinozoisite. Typically, 1� errors are<2% for allanite cores >100% for clinozoisite rims.

4.7.2 LA–ICPMS

Prior to mass spectrometry, thin sections were washed in 2% HNO3 and milli-QTM water to min-imise potential surface contamination, such as common Pb. In situ, allanite U, Th and Pb isotopedata were collected using a 266 nm, Nd:YAG, New Wave Research UP193SS Laser Ablation systemcoupled to a Nu Instruments, Nu Plasma HR multi–collector mass spectrometer. Thin sectionswere ablated in a standard ca. 30 cm3 ablation cell with a ⇠8 s washout time. Smaller ablationcells permit faster washout times, more detailed time–resolved analyses and higher signal/noise ra-tios, but compromise fine beam focusing and sample mobility (Horstwood et al., 2003). Followingablation, sample aerosols were introduced into a He carrier gas; Ar is introduced to form a 50:50Ar–He mix midway between the cell and sample torch via a T-piece. The detection system inthe NIGL Nu Plasma instrument comprises 12 Faraday cups and 3 ETP discrete dynode electronmultipliers. Faraday detectors linked to 1011 ⌦ resistors were used to measure 238U, 235U, 232Th,208Pb, 205Tl, 203Tl and 202Hg, whereas 207Pb, 206Pb and 204Pb+204Hg were measured on the ioncounters. Due to the large range of masses, U–Pb isotopic data were acquired in a separate ablationsequence from Th–Pb data.

Mass bias fractionation is inherent to ICPMS and likely originates from mass and charge de-

112

pendent processes operating during sample ablation and introduction of sample aerosol into theplasma (e.g. Tanner et al., 1994). A Tl–235U monitor solution was aspirated into the plasma us-ing a Nu Instruments’ DSN–100 desolvating nebuliser, which was monitored to correct for bothinstrumental mass bias and plasma-induced inter–elemental fractionation. Ion counter gains rel-ative to the Faraday cups were calculated before sample data acquisition by firstly measuring205Tl- and203Tl in the Faraday array before peak switching the 205Tl beam through the ion coun-ters. The isobaric interference of 204Hg with the 204Pb signal was corrected by measuring 202Hgand assuming a natural ratio of 204Hg/20Hg = 0.22988, with a mass bias factor determined from205Tl/203Tl=2.38925.

Before analysis of thin section allanite, the grain surface was cleaned by rastering the laser beamover the target area (⇠50µm2), at a low fluence. Pre–ablation the laser was fired at operationalconditions behind the shutter for 10s to stabilize the energy profile of the beam. Analyticalconditions were constrained by the high concentrations of non radiogenic 206Pb present in allanitefrom TH–680, which often caused the ion counters to saturate over 2.5 J.cm�2. Accordingly,optimal signal strengths were attained using 45% power, a 2 Hz repetition rate and a 35µm diameterspot. Small amounts of available allanite precluded use of rasters, which are thought to limit down-hole laser induced inter–elemental fractionation(Kosler et al., 2001; Horstwood et al., 2003; Storeyet al., 2006). Correction for instrumental drift and laser induced elemental fractionation wasaddressed via analysis of zircon standard–91500 (Wiedenbeck et al., 1995), every three unknowns.Normalisation factors derived from repeat zircon standard analyses performed under the sameanalytical conditions as subsequent thin section measurements, were then applied to sample ratios–as described by Simonetti et al. (2006). Ideally, an allanite standard would be used to preventmatrix dependent fractionation, however, the isotopic systematics of available allanite standardsGregory et al. (2007) are insu�ciently constrained to justify their use. Manipulation of isotopicdata was performed o↵-line using an in-house Excel spreadsheet. Ages were calculated using Isoplotv.3 (Ludwig, 2003) and the decay constants of Ja↵ey et al. (1971).

4.7.3 ID–TIMS

Target allanite grains identified on the basis of spread in 238U/206Pb ratios via LA–MC–ICPMSanalysis were separated from thick sections via laser milling using a New Wave Research UP193SSlaser ablation system. Grains were then plucked via tweezers from the thin-section resin underethanol. U–Pb Isotope Dilution–Thermal Ionisation Mass Spectrometry (ID–TIMS) analyses wereperformed at NIGL and utilized the EARTHTIME 205Pb–233U–235U (ET535) tracer solution. Al-lanite shards were dissolved in Teflon R� capsules containing a 50:50 mix of HF:HNO3; capsuleswere housed in Parr-style bombs and kept at 220�C for ⇡48 hours. Uranium and Pb cuts wereseparated using U/Teva and AG–1 cation exchange resins, respectively, on standard-size columns(10 mm diameter). Thorium was separated from collected washes using Truspec resin and largecolumns (12 mm diameter). Uranium–Pb measurements were performed on a Thermo Triton

113

TIMS. Pb analyses for signal intensities 15 mV were measured in dynamic mode on a MassComSEM detector and corrected for 0.16 ± 0.04%/ a.m.u. mass fractionation. Linearity and dead-timecorrections on the SEM were monitored using repeated analyses of NBS 982, NBS 981 and U500.Pb signal intensities � 15 mV were measured in static Faraday mode on 1012 ⌦ resistors (204Pbsignal was � 3 mV). Uranium was measured in static Faraday mode on 1011 ⌦ resistors or forsignal intensities 15 mV, in dynamic mode on the SEM detector. Uranium was run as the oxideand corrected for isobaric interferences with an 18O/16O composition of 0.00205 (IUPAC value andalso determined through direct measurement at NIGL).Thorium was analyzed as a 2%HNO3 solution on NIGL’s Nu Plasma MC–ICPMS. Mass fraction-ation was constrained by analysis of a 10 ppb 230Th spike solution prepared by R. R. Parrish atthe Geological Survey of Canada—230Th/232Th=7.0724±0.007 (1�). Both 230Th and 232Th weremeasured on Faraday collectors (H1 and H3 respectively). Reported isotope ratios, which are av-erages of 30 individual analyses (1 block), were manipulated using an in-house Excel spreadsheet.U–Th–Pb ratios and uncertainties (corrected for fractionation, laboratory blank and tracer) werecalculated using the algorithms of Schmitz & Schoene (2007) and a 235U/205Pb ratio for ET535 of100.21± 0.1. Ages were calculated using Isoplot v.3 (Ludwig, 2003).

4.8 Results and discussion

4.8.1 Allanite LA–ICPMS

Uranium–Pb LA–MC–ICPMS data are presented in the Appendix: Table A.6. Reproducabilityof 207Pb/206Pb and 206Pb/238U ratios for zircon standard 91500 (Wiedenbeck et al., 1995) were1.4% (Std.error) throughout all analytical sessions. Normalised to zircon 91500, standard allanite,Siss (von Blanckenburg, 1992; Gregory et al., 2007), shows 238U/206Pb ratios between 10–100 and207Pb/206Pb ratios between 0.4–0.8, implying considerable enrichment of radiogenic Pb. These,uncorrected, data yield a Tera–Wasserburg intercept age of 37.33±0.98 Ma (MSWD=1.3; Fig.4.5)– within error of the corrected 238U–206Pb ID–TIMS age (36.54±0.51 Ma) determined by vonBlanckenburg (1992). However, the 232Th–208Pb ID–TIMS age for the same allanite fraction is31.5±0.35 Ma, which is corroborated by cogenetic zircons having a 238U–206Pb age of 31.9±0.1Ma. The di↵erence in ages is likely due to incorporation of excess 206Pb produced by the decay of230Th – a process common in high Th/U minerals (Scharer, 1984). 230Th correction calculationspredict a 238U–206Pb age of 34.05 Ma, which is still in excess of the 232Th–208Pb age. Furthermore,the reference ID–TIMS analyses of von Blanckenburg (1992) were performed on a bulk fraction of11 allanite grains, and therefore quite possibly mask out isotopic intra- and inter-grain U–Th–Pbisotope heterogeneity. Because of these concerns with the Siss allanite standard, sample allaniteanalyses were normalised to zircon 91500.

Allanite core regions in rock TH–680 typically contain 180–220 ppm U and have Th/U ratiosaround 20. Clinozoisite rim domains are much less radiogenic with U and Th concentrations of 2–

114

90 ppm and 6–200 ppm respectively. (Table A.6) Analyses of both core and rim portions of zonedgrains show that non-radiogenic Pb dominates the total Pb composition (Pb⇤/Pbc < 0.03). Inparticular, clinozoisite rim portions contain negligible radiogenic Pb (238U/206Pb < 0.15; Fig.4.4)and thus provide an estimate of the common Pb (Pbc) composition of the rock during allanitegrowth. The Tera-Wasserburg concordia (Tera & Wasserburg, 1972, 1973, 1974) is well suited tohigh common–Pb minerals such as allanite as it permits both age and initial Pb compositions to bederived from uncorrected data, where the errors involved in making a common-Pb correction areoften unfeasibly large. Analyses (n=14) of two allanite grains, which show characteristic allanite–clinozoisite zoning, (allanite 1 and 7) in thin section X1 provide a Tera-Wasserburg lower interceptage of 35±4.1 Ma (2�; MSWD = 15) and an initial 207Pb/206Pb composition of 0.8211±0.0015 (95%conf.). The free 207Pb/206Pb intercept value established here is slightly lower than the compositionof 35 Ma model Pb (207Pb/206Pb=0.8379; Stacey & Kramers, 1975). This possibly signifies that aminor component (⇠1.7% di↵erence in 207Pb/206Pb) of exotic, radiogenic Pb was incorporated intothe allanite during time of growth—a phenomenon common to low-µ phases (µ=238U/204Pb), suchas allanite (e.g. Romer & Siegesmund, 2003; Radulescu et al., 2009). However, such a small di↵er-ence in initial Pb composition is equally well explained by sample-scale Pb isotope heterogeneity,particularly given that the two grains analysed are separated by ⇠10mm. Furthermore, the Pb iso-topic evolution model of Stacey & Kramers (1975) is only an approximation to heterogeneous crust.

Isobaric interference of 204Hg precludes application of a 204Pb based common Pb (Pbc) cor-rection. However, assuming concordancy between 206Pb/238U and 207Pb/235U ages—valid forPhanerozoic samples where the probability of significant Pb loss and the range of radiogenic207Pb/206 are small, corrected ages can be calculated from:

206Pb?/238U = (1� f206)⇥206 Pb/238Umeasured (4.9)

where f206 is 206Pbc/206Pbtotal (Gregory et al., 2007). Single spot, corrected 238U–206Pb ages(Table.A.6; Fig.4.5) using a common Pb 207Pb/206Pb composition of 0.8211 (Fig.4.4), range be-tween 23–1935 Ma. In general, age and uncertainty are correlated with 238U concentration (Fig.4.5).Only 238U signals >⇠30 mV approach the uncorrected Tera–Wasserburg age of 35±4.1 Ma. Aweighted average of analyses in which the 238U signal is >15 mV yields an age of 40±16 Ma (2�;Fig.4.5). Collectively, this shows that for weakly radiogenic, Pbc enriched phases, the accuracy ofconcordancy-based, single spot corrections critically depends on both the precise determination ofthe sample’s initial Pb composition and its assumed age.

As allanite analyses are normalised against zircon 91500 (Th=17–32 ppm Wiedenbeck et al.,2004), Th–Pb data are likely to be Th-minimum values. Allanite cores typically have 208Pb/232Thratios 1–2 orders of magnitude greater than clinozoisite rims (TableA.6), which provide su�cientspread in Th–Pb isotopic space for the data to be plotted on 232Th–208Pb isochrons against a stable

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Figure 4.4: Tera-Wasserburg plot of uncorrected data normalised against zircon standard 91500. Inset graph shows

degree of extrapolation required to generate a lower concordia intercept age of 35.0±4.1 Ma (2�). Red and green error

elipses correspond to allanite grains 1 and 7 respectively, in thin section X1. Crossed-polar photomicrograph of allanite

grain 7, thin section X1, highlights radiogenic allanite core (light brown–yellow) and clinozoisite rim domains (dark brown).

Pb isotope—either 204Pb or 207Pb. Figure 4.6a. shows that uncorrected 232Th–208Pb–204Pb datafrom all allanite–clinozoisite analyses do not define a single isochron, but form a cluster, reflectedin the age uncertainty of 23±14 Ma (MSWD=22). Similarly, data scatter significantly above andbelow the 232Th–208Pb–207Pb isochron (Fig.4.6b) which yields an age of 31±17 Ma (MSWD). Ageuncertainty in both Th–Pb isochrons is controlled by scatter outside the magnitude of 2� measure-ment uncertainty, which is suggestive of variable concentrations of common Pb between analyses.Uncertainty in the 204Pb based isochron also includes variable levels of 204Hg interference. Never-theless, common Pb-rich clinozoisite provides a precise estimate of common Pb composition, with2� uncertainty <1% (Initial 208Pb/204Pb=39.02±0.30; Initial 208Pb/207Pb=2.530±0.13). There-fore, to minimise scatter, 232Th–208Pb data were plotted against 206Pbc, on a 232Th–208Pb–206Pbc

isochron (Fig.4.6c) in which 206Pbc was calculated using the f206 parameter (Gregory et al., 2007)and the intercept–defined, common Pb composition. 232Th/206Pbc is calculated from:

232Th/206Pbc =232Th/206Pbmeasured

f206(4.10)

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0 10 20 30 40 50 6020

40

60

80

100

120

140

160

U signal (mV)

238 U

–206 Pb

age

(M

a)

0.00 0.05 0.10 0.15 0.20 0.2520

40

60

80

100

232Th signal (V)

232 Th

–208 Pb

age

(Ma)

238

0.3

0.4

0.5

0.6

0.7

0.8

0 20 40 60 80 100 120

37.33±0.98 Ma MSWD = 1.3

data-point error ellipses are 2

0

20

40

60

80

100

120

140

160

180 box heights are 2

238U/206Pb

207 P

b/20

6 Pb

238 U

-206 P

b ag

e (M

a)

Wtd. av. 40±16 Ma (2σ)T-W 35±4.1 Ma (2σ)

a. b.

c. d.

Figure 4.5: a. Tera-Wasserburg concordia of Siss allanite (von Blanckenburg, 1992). Analyses are normalised to zircon

91500 and show spread from non-radiogenic (238U/206Pb⇡20) to radiogenic (238U/206Pb⇡20) compositions. Intercept age

of 37.33±0.98 Ma is significantly di↵erent from published ID–TIMS ages. b. Weighted average of common Pb corrected238U–206Pb, single-spot ages for 238U signals >15 mV. Light shading corresponds to 2� uncertainty on the average

age; darker shading signifies 2� Tera-Wasserburg age interval. c. 238U signal strength versus 207Pb-based, common Pb

corrected, single-spot age. Errors are 2�. d. 232Th signal strength versus 207Pb–based, common Pb corrected, single-spot

age. Errors are 2�.

117

The corrected isochron yields a precise age of 31.7±8.2 Ma, with minimal scatter in the data(MSWD=0.098). The age error reflects uncertainty in calculation of f206, and more significantly,the weakly radiogenic composition of the allanites. More spread in 232Th/206Pb is required toincrease accuracy. Similarly, single-spot 232Th–208Pb corrected ages (Fig.4.6d) yield a poorlyconstrained, weighted average age of 48±13 Ma (2�; for analyses in which 232Th signal strengthexceeded 0.02 V). Collectively, Th–Pb isochron and spot ages all lie within 2� uncertainty of theTera-Wasserburg age, which suggests that the U–Pb and Th–Pb systems were not decoupled duringor subsequent to allanite–clinozoisite growth.

37.5

38.5

39.5

40.5

41.5

42.5

0 400 800 1200 16002.48

2.52

2.56

2.60

2.64

2.68

2.72

0 20 40 60 80 100 120

2.02

2.06

2.10

2.14

2.18

2.22

2.26

2.30

0 20 40 60 80 1005

15

25

35

45

55

65

75

85

208 P

b/20

4 Pb

232Th/204Pb

data-point error crosses are 2σ

Age=23±14 MaInitial 208Pb/204Pb=39.02±0.30

MSWD=22

Age=31±17 MaInitial 208Pb/207Pb=2.530±0.013

MSWD=55

232Th/207Pb

208 P

b/20

7 Pb

data-point error crosses are 2σ

208 P

b/20

6 Pb c

Age=31.7±8.2 MaInitial 208Pb/206Pbc=2.11±0.011

MSWD=0.098

232Th/206Pbc

data-point error crosses are 2σ

box heights are 2σ

232 T

h-20

8 Pb

Age

(Ma)

Wtd.av. 48±13 Ma (2σ)

T-W 35±4.1 Ma (2σ)

a. b.

c. d.

Figure 4.6: a. 232Th–208Pb–204Pb isochron. Note well constrained initial Pb composition and scatter in more radiogenic

analyses. b. 232Th–208Pb–207Pb isochron. c. 232Th–208Pb–206Pbc isochron. Note that corrected analyses define a tightly

constrained regression; remaining uncertainty reflects common Pb correction errors and the weakly radiogenic nature of

the grains. d. 232Th–208Pb common Pb corrected, single–spot ages for analyses in which 232Th signal strength exceeded

0.02 V.

118

However, as allanite analyses are normalised against a zircon standard, the extent of matrix-induced isotopic fractionation remains unquantified. Therefore, LA–MC–ICPMS data cannot beused for robust geological interpretation, but rather a first–order indicator of expected allaniteage. In order to circumvent this, three allanite grains—two of which were previously used forLA–MC–ICPMS, were separated for U–Th–Pb ID–TIMS analyses.

4.8.2 Allanite ID–TIMS

The increased ionization e�ciency of U and Pb isotopes, higher signal to noise ratios and lackof reliance on mineral standard calibration promoted via ID–TIMS, permits U–Pb isotope ratiosto be measured at higher analytical precision (<0.5%1�) than LA–MC–ICPMS. ID–TIMS U andPb isotope data (Table A.7) obtained from both allanite and clinozoisite domains corroborateobservations on U–Pb systematics made from LA–MC–ICPMS measurements. Despite the limitedspread in 238U/206Pb compositional space (0.5–5), analyses of eight allanite-clinozoisite shardsdefine a Tera-Wasserburg regression with a lower concordia intercept age of 34.2±3.6 Ma (2�;MSWD = 3.8) and an initial 207Pb/206Pb composition of 0.8330±0.0014 (95 % conf.; Fig.4b).This age is within 1� error of the LA–MC–ICPMS Tera-Wasserburg age and has an initial Pbcomposition which is, again, slightly more radiogenic than the 34 Ma average upper crustal leadcomposition predicted by Stacey & Kramers (1975) (0.8379). Due to low allanite Th/U valuesand high common Pb contents, 207Pb/206Pb and 238U/206Pb ratios corrected for ingrowth of 206Pbfrom the incorporation of 230Th (Scharer, 1984), yield an age di↵erence within the uncorrected 2�

age uncertainty. The e↵ect of excess 206Pb on corrected spot ages is modeled using the followingequation (Scharer, 1984):

206Pb =238 U [(e�238

T � 1) +�238

�230(f � 1)] (4.11)

where f=[Th/Umin]/[Th/Urock] and T is the age of crystallization. Figure4.8 shows that for therange of Th/U ratios observed in Tauern allanite, the e↵ect of excess 206Pb on common–Pb cor-rected ages is minor compared to analytical uncertainty.

The Th–Pb ID–TIMS data (Table A.8) are comparable to those obtained via LA–MC–ICPMS(Fig.4.6). Uncorrected data show considerable scatter above and below 204Pb and 207Pb nor-malised 232Th—208Pb isochrons (Fig.4.9a and b), with MSWD’s ranging between 8–15. As aresult, isochron ages have large errors (26±17 Ma for 204Pb based isochron; 25±15 Ma for 207Pbbased isochron). Intercept 208Pb/204Pb (39.12±0.44) and 206Pb/207Pb (2.496±0.025) are bothwithin error of the values predicted by Stacey & Kramers (1975) for upper–crustal Pb at 34 Ma(38.57±0.28 and 2.496±0.025 respectively). Plotting the Th–Pb data on a 232Th—208Pb—206Pbc

isochron (using the ID–TIMS Tera-Wasserburg value for 207Pb/206Pbi Gregory et al., 2007) fails tocorrect such scatter and yields an age of 25±15Ma (MSWD=8.9; Fig.4.9c)—identical to the 204Pb

119

0.79

0.81

0.83

0.85

0 1 2 3 4 5 6

207 P

b/20

6 Pb

data-point error ellipses are 2σ

34.2 ± 3.6 MaMSWD = 3.8

clinozoisite

allanite

0.0

0.2

0.4

0.6

0.8

0 40 80 120 160 200238U/206Pb 238U/206Pb

a. b.

Figure 4.7: a. Tera-Wasserburg regression of ID–TIMS allanite–clinozoisite data. Regression yields an initial207Pb/206Pb of 0.8330±0.0014 (95%conf.). Note the small degree of scatter compared to LA–MC–ICPMS data

(MSWD=3.8). b. Tera–Wasserburg regression of ID–TIMS allanite–clinozoisite data, showing the weakly radiogenic

nature of the allanite and the large degree of extrapolation required to generate a lower concordia intercept age. Concor-

dia markers are spaced by 400 Ma.

0 1 2 3 4 5 6 7 899.6592

99.6592

99.6593

99.6593

f 206 (M

ax. 20

7 Pb/

206 P

b m) (

%)

0 1 2 3 4 5 6 7 897.4624

97.4626

97.4628

97.463

[Th/Umin]/[Th/Urock]

Age=34 MaTh/Urock=3

207Pb/206Pbc=0.8330

f206 (Min. 207Pb/ 206Pb

m ) (%)

Figure 4.8: Modeled e↵ect of excess 206Pb on common-Pb corrected ages following the correction method of Scharer

(1984). Parameter f206

corresponds to the total fraction of common-Pb as described in the text (Gregory et al.,

2007). [Th/Umin]/[Th/Urock] values reflect variation in allanite Th/U ratios; 207Pb/206Pbc taken from ID–TIMS Tera-

Wasserburg intercept value. Blue line is calculated using the maximum 207Pb/206Pbm value (0.83018), whereas the purple

line is calculated using the minimum value of 207Pb/206Pbm (0.81200), which correspond to maximum and minimum frac-

tions of common-Pb respectively. Note that variation in f206

due to excess 206Pb is 3 orders of magnitude smaller than

typical 1� (⇠1%) uncertainty of f206

resulting from analytical uncertainty.

120

based isochron. The observed Th–Pb scatter can be explained by loss/gain of Th and/or vari-able common Pb compositions. As allanite–clinozoisite shards contain >90% common–Pb, localvariations in 208Pb/204Pb, 208Pb/206Pb and 208Pb/207Pb will place a strong control on the Th–Pbisochron age. However, the small amount of scatter (MSWD=3.8) of the uncorrected 207Pb/206PbID–TIMS data shows that common-Pb composition does not vary in the U–Pb system on the scaleshown via Th–Pb data. It is, therefore, possible that the Th–Pb system was decoupled from theU–Pb system in allanite, as deduced via ID–TIMS analyses. LA–MC–ICPMS analyses lack su�-cient precision to observe this. Further evidence for decoupling of the U–Pb and Th–Pb systems ispresent in the form of a weak correlation between Th/U and 208Pb/206Pb (R2=0.8678; Fig.4.9d).

38.4

38.8

39.2

39.6

40.0

40.4

40.8

0 200 400 600 800 1000

data-point error crosses are 1σ

2.04

2.06

2.08

2.10

2.12

2.14

2.16

0 10 20 30 40 50

2.46

2.48

2.50

2.52

2.54

2.56

2.58

2.60

0 20 40 60

208 P

b/20

4 Pb

232Th/204Pb

Age = 26±17 MaInitial 208Pb/204Pb = 39.12±0.44

MSWD = 15

data-point error crosses are 1σ

Age = 25±15 MaInitial 208Pb/207Pb = 2.496±0.025

MSWD = 7.9

232Th/207Pb

208 P

b/20

7 Pb

a b

Age = 25±15 MaInitial 208Pb/206Pbc= 2.080±0.021

MSWD = 8.9

data-point error crosses are 1σ

232Th/206Pbc

208 P

b/20

6 Pb c

c

208Pb/206Pb2.06 2.08 2.10 2.12

R2 = 0.8678

5

10

Th/U

15

20

25d

Figure 4.9: a. 232Th–208Pb–204Pb ID–TIMS isochron. Note the large degree of scatter (MSWD=15) in the regressed

data. b. 232Th–208Pb–207Pb ID–TIMS isochron. c. 232Th–208Pb–206Pbc isochron after Gregory et al. (2007). f206

calculated using 207Pb/206Pbc from ID–TIMS Tera-Wasserburg plot (0.8330), and an expected age of 34.2 Ma. d. Th/U

versus 208Pb/206Pb for all ID–TIMS analyses; 1� errors are less than width of marker. Note the weak degree of correation

(R2=0.8678).

The Tera-Wasserburg ID–TIMS age is the most accurate and precise age obtained on allanite–

121

clinozoisite fractions from sample TH–680. Performing both ID–TIMS and LA–MC–ICPMS analy-ses on the same allanites is a powerful geochronological approach, combining the spatial resolutionof laser ablation and the high precision of ID–TIMS. Although the ID–TIMS Tera-Wasserburgage is within error of the LA–MC–ICPMS age, common Pb compositions are subtly di↵erent(�(207Pb/206Pb) ⇡0.0119). This is likely a product of matrix matching allanite with zircon—i.e. alow-µ mineral normalised against a high–µ mineral. Identification of a suitable allanite standard isrequired to constrain matrix dependent isotopic fractionation for U(–Th)–Pb LA–ICPMS allanitegeochronology.

4.9 Allanite petrogenesis

To determine the tectono-metamorphic significance of the 34.2±3.6 Ma U–Th–Pb allanite age, itis critically important to establish both the P–T conditions of allanite–clinozoisite growth andthe phase relationships between allanite and other, rock-forming phases. This is achieved usingtextural observations, REE+Y mass balance modeling and isopleth calculations.

4.9.1 Textural observations

Zoned epidotes with allanitic cores and clinozoisite rims are present as elongated lozenges (150–900µm length; 50–200µm width) which are wrapped by the rock’s phengite fabric (S0 + S1 +S2).This fabric, which additionally wraps around garnet poikiloblasts, formed close to peak pressureconditions (>20 kbar; Fig.3.6) and was crenulated during D2, to form F2 generation micro–folds(Fig.3.2(c)). Proximal to fold hinges, allanite grains have been transposed into alignment withfold limbs. This deformation event produced a new axial planar foliation during imbrication of thenappe stack, which defines regional structure (Holland, 1977; Behrmann & Ratschbacher, 1989).Importantly, euhedral allanite grains are also present as rare inclusions within both core andrim portions of garnet poikiloblasts (Fig.4.10)—indicative of pre-garnet growth. Included allanitedisplays identical allanite–clinozoisite, core-rim zonation as observed in matrix grains. Allaniteinclusions within eclogite facies garnet poikiloblasts have been observed elsewhere in the EclogiteZone. Hoschek et al. (2010) describe virtually identical phase relations for a garnet–chloritoid–kyanite bearing micaschist collected nearby TH–680, within the Eclogite Zone.Therefore, texturally, it appears that allanite grew early on, prior to initiation of garnet growthbetween 15–20 kbar (section 3.2.4.4 Hoschek et al., 2010) and also prior to development of theregionally pervasive mica fabric (D2).

4.9.2 REE+Y modeling

Allanite ((Ca,REE,Y)2(Al,Fe)3(SiO4)3OH) can incorporate weight percent levels of the LREEinventory and Th in crustal rocks (Hermann, 2002a) and forms a solution with epidote and clino-zoisite by the coupled substitution of Ca2++ Fe3+/Al3+ respectively for REE3++ Fe2+. Allanite

122

Figure 4.10: High contrast BSE image of a

euhedral allanite–clinozoisite grain included in a

garnet poikiloblast, sample TH–680. Note the

characteristic zoning observed in matrix grains

and also, the radial fracturing of proximal gar-

net, implying pre-tectonic inclusion.

in TH–680 is zoned from core to rim by increasing clinozoisite content (Fig.3.5) which is commonlyascribed to prograde metamorphic growth (Janots et al., 2007, 2008, 2009). Chondrite normalisedREE patterns (Fig.4.11) obtained via LA–ICPMS show that the allanite is heavily enriched in theLREEs relative to HREEs (average La/Yb = 294). Even so, the allanite is enriched in HREEsrelative to garnet, which suggests that allanite grew prior to garnet, as garnet is the rock-formingmineral that partitions most strongly HREE, Y and the heaviest MREEs (Hickmott et al., 1987;Lanzirotti, 1995; Cherno↵ & Carlson, 1999; Otamendi et al., 2002). Allanite lacks a significantEu anomaly, suggesting either growth outside the stability limit of plagioclase or growth in aplagioclase depleted rock. Given the abundance of garnet, the former explanation is preferred.

Mass balance calculations (Fig.4.12), which combine mineral modes determined via high-precision point counting (n=1000) and representative mineral analyses of garnet, allanite, phengiteand rutile, show that allanite and garnet dominate the REE budget of sample TH–680. Phosphateaccessory phases monazite, xenotime and apatite (common sinks for REE elements in metapeliticrocks) are not observed. Allanite and surrounding clinozoisite hold the LREEs and >50% of theMREEs, whereas garnet sequesters ⇠65% of the HREE budget (Fig.4.12).

Whole rock REE and trace element concentrations were determined via ICPMS, at the Univer-sity of Cambridge (Table A.6). Although, mass balance calculations provide a first order estimateof REE distribution, partition coe�cients (DREE

mineral) determined via combining bulk rock and min-eral analyses, are considerably more accurate at predicting the evolution of the rock’s REE budget.Values of DHREE+Y

garnet�bulk for TH–680 and several comparative studies are presented in Fig.4.13. Sam-ples 44A and 47A of Corrie & Kohn (2008), provide an external comparison of DREE+Y

garnet�bulk, in

123

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

0.01

1

100

10,000

1,000,000

Sample/chondrite

allanite

garnet

rim

core

Figure 4.11: Chondrite normalised REE pat-

terns of allanite and garnet within TH–680.

Note core–rim HREE depletion in garnet and en-

riched HREE concentrations in allanite. Chon-

drite values taken from McDonough & Sun

(1995)

0%

20%

40%

60%

80%

100%

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

% C

ontr

ibution to R

EE

budget

garnet-rim

garnet-core

rutile

epidote

allanite

phengite

Figure 4.12: Individual contributions of REE bearing phases to the whole-rock budget in sample TH–680. Mineral

modes (volume%) used are as follows: 50.45% phengite, 27.62% quartz, 0.04% allanite, 2.85% epidote, 3.49% rutile,

1.79% kyanite, 2.69% chloritoid.

124

Dy Ho Er Tm Yb Lu Y

50

100

150

D gt-b

ulk

TH–680Corrie et al. 47AOtamendi et al. RSR18Corrie et al. 44A

Figure 4.13: HREE garnet–bulk distribution

coe�cients (DHREE+Ygarnet�bulk) for TH–680 deter-

mined from ICPMS whole-rock data. Reference

values are taken from monazite and allanite ab-

sent garnet schists (44A, 47A) described by Cor-

rie & Kohn (2008) and a zircon+apatite bearing,

quartz+garnet+plagioclase+cordierite granulite

(Otamendi et al., 2002). Note that literature D’s

are consistently higher than those calculated for

TH–680.

the absence of allanite and monazite. Trace apatite is present within these rocks, which werecollected from the western Blue Ridge terrane, Great Smokey Mountains, U.S.A. Similarly, sampleRSR18 (Otamendi et al., 2002) provides an example of DREE+Y

garnet�bulk at granulite facies conditions,where garnet compositions reflect contemporaneous zircon crystallisation. It is clear that values ofDHREE+Y

garnet�bulk for sample TH–680, are lower than the range of literature derived values. Such a dis-parity between typical garnet HREE+Y concentrations and TH–680, is consistent with depletionof the HREE+Y budget, prior to initiation of garnet growth. Undefined reservoirs of LREE insamples 44A and 47A, and a lack of data for sample RSR18, preclude the possibility of comparingDLREE

garnet�bulk.Yttrium is a trace element with a strong a�nity for both allanite and garnet (i.e. DY

allanite�matrix

and DYgarnet�matrix>>1; Pyle & Spear, 1999) and mineral Y profiles are therefore very sensitive to

the relative timing of their growth. Garnets in TH–680 possess the same general Y profile (Fig.??a),consisting of a high Y core (⇠500ppm), followed by a variable width zoning shoulder, which dropsto a lower Y outboard region (⇠125ppm). The zoning is broadly consistent with growth of garnetunder conditions of constant Y activity, followed by consumption of the Y bu↵er and initiation ofRayleigh–style fractionation of the remaining Y. Therefore, in order to test whether the observedY profiles are most consistent with pre- or post-allanite, garnet growth, synthetic Y zoning profilesin garnet were calculated using the following equation (Hollister, 1966):

CG = D ⇥ CO ⇥ (1� WG

WO)D�1 (4.12)

Here, CG is the concentration of an element in garnet, D is the distribution coe�cient between

125

garnet and matrix (DYgarnet�matrix), CO is the initial concentration of the element in the bulk–rock,

WG/WO is the weight fraction of garnet in the bulk rock. Parameters used in Rayleigh fractiona-tion modeling are described in Table4.3. In order to fit the measured profile, the product of D andCO was kept constant at 499 ppm—the garnet core concentration at the point at which Rayleighbehavior commences. Pre- and post-allanite garnet growth was then modeled using values for CO

obtained from ICPMS bulk rock data, with the allanite(+clinozosite) contribution (mineral modesdetailed in Fig.4.12) subtracted, respectively. Values of D were calculated subject to the constantproduct constraint (Table4.3).Yttrium concentrations in garnet growing prior to allanite were calculated to be an order of magni-tude higher than measured values (⇠2500 ppm versus ⇠500 ppm; Fig.4.14). As shown in Fig.4.13,literature values for DY

garnet�bulk in metapelitic rocks with or without further Y bu↵ering phases,commonly lie between 30–200 (e.g. Pyle & Spear, 1999; Otamendi et al., 2002; Corrie & Kohn,2008). Pre-allanite garnet growth requires that DY

garnet�bulk=16.3, whereas garnet growth follow-ing allanite formation requires a more realistic value of 83.8. Figure 4.14 shows the di↵erence inpredicted garnet concentrations using DY

garnet�bulk=83.8 and the relevant value of CO for pre- andpost-allanite formation garnet growth. Note that for pre-allanite growth, the Y profile is broaderthan the measured profile. Interestingly, post–allanite Y profiles do not account for the outboardshoulder of constant Y concentration (⇠130 ppm), which suggests that Y was made available togarnet mid-way through its growth. Integrating polynomial and power-law fits to both measuredand best-fit, pre-allanite Y (DY

garnet�matrix=83.8; CO=5.96 ppm) profiles, shows that observedgarnets contain ⇠2.5 times more Y than predicted via Rayleigh fractionation. The most likelysource from which Y is liberated is the breakdown of clinozoisite during garnet growth. Giventhat clinozoisite rims contain ⇠377 ppm Y, 27% of the total clinozoisite (2.85 vol.%) observed inTH–680 would have to break down in order liberate the required amount of Y. This is consistentwith the pseudosection topology presented in Fig.3.6, in which the mode of clinozoisite (denoted epin section) decreases up temperature and eventually goes to zero in the garnet–chloritoid–kyanitestability field.

4.9.3 Metamorphic modeling

Assuming that ⌃[Si+Al+Ti+Fe+Mn+Mg] = 6 c.p.f.u. (Ercit, 2002), structural formulae for al-lanite and clinozoisite were recalculated in order to generate estimates for Fe3+ concentrations. Aspresented in Fig.3.6, isopleths for constant ferric iron concentration, and hence constant epidote–clinozoisite ratio were calculated between 5–25 kbar and 400–600�C using the epidote activity–composition model of (Holland & Powell, 1998). Given the high concentration of REEs present inallanite, the two component epidote–clinozoisite model may not be appropriate. However, lowerREE concentrations of the rims surrounding allanite cores make such portions suitable for appli-cation of the epidote activity model. Rim epidote–clinozoisite ratios show a comparable range

126

[Y]garnet�core [Y]garnet�rim [Y]allanite [Y]cz�rim [Y]bulk�rock CO DYgarnet�matrix

499.5 131.2 3267.9 377.9 30.6 30.6† 16.3†

5.9? 83.8?

2.9⇧ 172.8⇧

Table 4.3: Details of parameters used in Rayleigh fractionation modeling of garnet growth from a metamorphic matrix

before and after allanite growth. All [Y] values are in ppm. Calculations use the following mineral modes obtained via

high precision (n=1000) point counting: garnet – 12.2 wt.%; allanite–clinozoisite – 3.3 wt.%.† Values used for scenario in which garnet grows from a matrix prior to allanite formation—D is calculated subject to the

constant product constraint described in the text.? Values used for garnet growth from a matrix depleted by allanite growth; CO is calculated from the core to rim variation

in [Y] within allanite–clinozoisite grains, assuming allanite core composition corresponds to matrix compositions, and also

3.3 weight % allanite–clinozoisite. D is calculated subject to the constant product constraint.⇧ Values used for garnet growth from a matrix depleted by allanite growth; CO is calculated via subtracting the combined

allanite–clinozoisite Y contribution to the ICPMS bulk rock value, assuming 3.3 weight % allanite–clinozoisite. D is

calculated subject to the constant product constraint. Note the degree of similarity between both post allanite growth

estimates of CO.

2500

2000

1500

1000

500

00 250 500 750 1000 1250

Y (p

pm)

Radius (µm)

garnet prior to allanitegarnet post-allanitemeasured profile

50 µm100 µmmuscovite

garnet

Post-allanite gt growth

Pre-allanite gt growth

Radius (µm)0 250 500 750 1000 1250

2500

2000

1500

1000

500

0

Y (p

pm)

a b

Figure 4.14: Modeled Y concentration profiles from core to rim, assuming Rayleigh fractionation between garnet and

metamorphic matrix (Hollister, 1966). a. Both before (red solid line; DYgarnet�matrix=83.8; CO=30.62 ppm) and after

(blue solid line; DYgarnet�matrix=83.8; CO=5.96 ppm) allanite growth. Measured profile (dashed line) ( constructed

from LA–ICPMS Y data; 1� errors are smaller than the width of an analysis point marker. b. Yttrium profiles fitted to

the garnet–core concentration (Table4.3) using constant product constraint (see text) for pre-allanite growth (light blue

curve—DYgarnet�matrix=16.3; CO=30.6 ppm) and post-allanite growth (orange and red curves—DY

garnet�matrix=83.8;

CO=5.96 ppm and DYgarnet�matrix=172.8; CO=2.90 ppm, respectively).

127

in values to those calculated, between 8–13 kbar and 425–500�C—before garnet growth at 18–21kbar and 550–580�C. The formation of allanite during regional metamorphism of pelitic protolithsis intimately associated with the first appearance of chloritoid (Janots et al., 2008, 2009, 2010).Although loosely constrained, the interval of clinozoisite stability deduced from isopleth modelingof the TH–680 bulk composition is consistent with this conclusion. Chloritoid first appears attemperatures between 400–450�C and pressures of 5–7 kbar in TH–680, which are within error ofthe 400–430�C temperature interval derived by Janots et al. (2008) in the Central Alps.

Collectively, the textural, chemical and metamorphic evidence presented above are consistentin showing that both allanite and clinozoisite fractions grew prior to garnet and likely close to theappearance of chloritoid along the prograde metamorphic path.

4.10 Metamorphic rates

The 34.2±3.6 Ma U–Pb ID–TIMS age represents the time at which allanitic cores to clinozoisitesgrew at pressures around ⇠8–13 kbar at 400–500�C during ongoing subduction of the Penninicmargin. The new geochronological data, combined with previously reported data (Fig.4.15), permitcalculation of exhumation rates and also the first calculation of timescales for the process of eclogiteformation in the Eastern Alps.

Taking the angle of subduction as 45� and a descending slab velocity of 5 cm.a�1, the earliestpeak pressures (⇠26 kbar) would be attained in the subducted margin is ⇠2 Ma. after allanitegrowth i.e. by ca. 32 Ma. However, given the 2� uncertainty on the ID–TIMS age, peak pressureconditions are constrained to the interval 29–37 Ma. This corroborates the 31.5±0.7 Ma Rb–Sr ageof Glodny et al. (2005), obtained from a range of eclogitic mineral assemblages which suggests thatRb–Sr multi-mineral geochronology is applicable under eclogite facies conditions. The EclogiteZone was subjected to regionally pervasive Barrovian metamorphism between 27–32 Ma (Rb-SrInger & Cli↵, 1997; Glodny et al., 2005; Gleißner et al., 2007), which places a limit on the timetaken to both form and exhume the eclogite facies nappe to mid-crustal levels. Therefore, theEclogite Zone must have experienced ⇠60 km (i.e. from ⇠25 kbar to ⇠7 kbar) of exhumationat rates between 0.6 and >6 cm.a�1. Plate velocity exhumation of the Eclogite Zone demandsslip rates between 0.85–8.5 cm.�1 on bounding faults. Such exhumation rates are similar to thosereported from other (ultra)high-pressure nappes in the Alps and Himalayas: Dora Maira eclogites(Western Alps) – 3.4 cm.a�1 (Rubatto & Hermann, 2001); Kaghan eclogites (Pakistan) – 3–8cm.a�1 (Parrish et al., 2006). Independent evidence for plate velocity exhumation of the EclogiteZone comes from Dachs & Proyer (2002), who model garnet intra-crystalline di↵usion patterns ina relict eclogite and show that a minimum of 1 Ma elapsed between eclogite formation and coolingthrough 450�C. They then calculate average exhumation rates between 4.6–7.4 cm.a�1: withinerror of rates calculated here.

Due to subsequent equilibration of geochronometers at higher temperatures, dating progrademetamorphic processes is extremely challenging. However, given the data presented above, it is

128

300 400 500 600T (°C)

10

20

30

0

35

70

P (kbar)Z (km

)

ctd-inqtzcoe

allanite

clinozoisite

34.2±3.6 Ma

initiation of garnet growth

Barrovian27-32 Ma

Eclogitisation29-37 Ma

0.6->6 cm.a-1Exhumation

ma-in

>10.3 °C.Ma-1SubductionD2

D1

D1

D0

margarite

Figure 4.15: P–T–t–Z summary sketch of sam-

ple TH–680 showing the relationship between mineral

growth, deformation events and geochronological con-

straints. Note the presence of a D0

event, defined

on the basis of aligned quartz inclusions within gar-

net poikiloblasts. D1

brackets garnet growth given the

synkinematic nature of the inclusion assemblage. D2

involves crenulation of the high-P phengite fabric, fur-

ther rotation of garnet and allanite blasts and growth

of a regional fabric (S2

). Chloritoid- and margarite-

in equilibria positions from pseudosection calculations

for TH–680. Geochronological constraints for Barro-

vian event taken from Inger & Cli↵ (1997); Glodny

et al. (2005); Gleißner et al. (2007).

possible to calculate an estimate for the range of prograde heating rates experienced by the Euro-pean margin during Alpine subduction. Assuming allanite cores grew concomitant with the firstappearance of chloritoid at⇠430�C (Fig.3.6 Janots et al., 2008) and a peak temperature of 538±17�

(Section 3.2.4.3) shows that heating rates (peak-T at 29 Ma) must have been >10.3�C.Ma�1. Thisis comparable to the rates calculated by Lapen et al. (2003) of >15�C.Ma�1 from Lago di Cignana,Western Alps.

4.11 Tectonic Implications

Given its fast exhumation, it is extraordinary that the Eclogite Zone is understood to have beensubjected to local recrystallisation in the blueschist facies, prior to Barrovian overprinting, asevidenced by the growth of glaucophane at the expense of peak eclogitic omphacite. The P–T

129

10

10

10

10

10

10

10

10

10

-6

-5

-4

-3

-2

-1

0

1

2

Ma

1086420Nappe thickness (km)

Thic

knes

s of

EZ

Exhumation duration

Figure 4.16: Thermal time con-

stant for a range of nappe thick-

nesses (blue line), plotted against

width of the Eclogite Zone (grey

box) and the nappe’s exhumation

duration (blue box). Calculations

use a thermal di↵usivity () of 30

km2.Ma�1). Note that the exhuma-

tion interval is greater, outside of

uncertainty, than the thermal time

constant of the Eclogite Zone.

conditions of bluechsist facies metamorphism are constrained between 7–11 kbar and 400–450�C(Holland & Ray, 1985; Eremin, 1994). The thermal time constant (⌧=a2/(⇡2., where a is nappethickness) of the Eclogite Zone is ⇠ 10�2 Ma (Fig.4.16), assuming a nappe thickness of ⇠1.7km and a thermal di↵usivity of = 30 km2.Ma�1. Therefore, the amount of time it takes theEclogite Zone to be exhumed from eclogite-facies depths to mid-crustal levels, is greater, outsideof uncertainty, than the time it takes a thermal perturbation to e↵ect the nappe’s core. In lightof these calclations, the partial blueschist overprint is strong evidence that subduction was stillactive after detachment and during exhumaton of the Eclogite Zone from the down-going Penninicslab, i.e. the Eclogite Zone re-equilibrated to cold (<450�C) ambient temperatures, indicative ofan active subduction regime. The P–T estimates are consistent with calculated thermal structuresof slowly convergent subduction zones, such as the Nankai trough of SW Japan, which has a 45mm.a�1 subduction rate (Peacock, 1996, 2003). Such slow velocities are diagnostic of the cessationof subduction, just prior to continental collision.

As the Eclogite Zone represents transitional crust located on the southern margin of the Euro-pean continent (Kurz et al., 1998b, and references therein), the 29–37 Ma age of eclogite formationcritically provides a maximum estimate for the timing of continental collision in the Eastern Alps.This means that a maximum of 10 Ma separated emplacement of the Austroalpine nappes andattainment of peak Barrovian conditions at 27–32 Ma (Fig.4.17). Importantly, this shows thatprevious models for East Alpine orogenesis, which rely on a 30 Ma interval between Cretaceouseclogite formation and Barrovian metamorphism, are inaccurate by up to an order of magnitude(Oxburgh & Turcotte, 1974; Bickle et al., 1975). This challenges current understanding of the ratesof conductive heating (England & Thompson, 1984), the implications of which are investigated inchapter 6.

130

These new data rule out a close temporal relationship with the Cretaceous (91–93 Ma; Milleret al., 2005; Thoni et al., 2008; Janak et al., 2009) eclogites exposed in the Middle-AustroalpineKoralm/Saualm region. Rather, this suggests that the Eastern Alps was the result of a two-stagecollision process, with two south-vergent subduction zones responsible for transport of Austroalpineoceanic and, subsequently, Penninic oceanic crust to eclogite facies depths prior to continentalcollision(Fig.4.17; Neubauer et al., 2000).

Eocene–Oligocene HP metamorphism in the Eastern Alps is similar to ages of (U)HP meta-morphism from a range of terranes in the Western and Central Alps. From east to west, thefollowing nappes experienced eclogite facies conditions: the Adula nappe - 34.9–49.6 Ma (Becker,1993; Gebauer et al., 1996); the Monte Rosa unit, 34.9±1.4 Ma(Rubatto & Gebauer, 1999); theGran Paradiso massif, 43±0.5 Ma (Me↵an-Main et al., 2004); the (U)HP Dora Maira nappe,31.6–36.4 Ma (Duchene et al., 1997; Gebauer et al., 1997; Rubatto & Hermann, 2001). As all ofthese (U)HP nappes are interpreted to pertain to the thinned European margin (Schmid et al.,2004a), continental collision with the Adriatic plate, responsible for cessation of subduction and,likely, detachment of (U)HP rock from the slab, occurred contemporaneously, between ⇠30–40Ma, along the Alpine arc.

Tauern eclogite formation is also contemporaneous with Alpine magmatism to the south of theTauern Window (Rensen pluton 31.1–31.7 Ma - Barth et al. (1989); Rieserferner pluton 31.4–32.8Ma - Romer & Siegesmund (2003)) - see section 6.4.2 for detailed discussion. In the central Alps,Tertiary magmatism is confined to a belt of largely calc-alkaline plutons (Bergell and Adamelloplutons), which show a sharp maximum of plutonic and subvolcanic activity between 29 and 33Ma Hansmann & Oberli (1991); Oberli et al. (2004); Berger et al. (2009). This temporal relation-ship between (U)HP metamorphism and magmatism has been interpreted by Von Blanckenburg& Huw Davies (1995) to represent buoyancy driven exhumation and asthenospheric upwellingfollowing slab breako↵.

4.12 Chapter Summary

Chemically zoned allanite–clinozoisite grains from an Eclogite Zone metapelite yield U–Pb agesby LA–ICPMS and ID–TIMS of 35±5.1 and 34.2±3.6 Ma, respectively. The Th–Pb system cor-roborates U-Pb ages, although significant scatter in the composition of thorogenic Pb precludesprecise age determination. Petrographic and REE+Y Rayleigh fractionation modeling of gar-net growth show that allanite and clinozoisite rim domains grew during subduction, prior to theinitiation of garnet porphyroblast growth ca.18–21 kbar and 550–580�C. Isopleth calculations ofepidote/clinozoisite content suggest that rim portions grew between 8–13 kbar and 425–500�C.Thus, these data show that the European margin was buried to ⇠35 km at 34.2±3.6 Ma, beforeattaining eclogite facies conditions at 27–32 Ma and being exhumed, at plate tectonic rates (0.6–6cm.a�1), to mid-crustal levels, where the Penninic nappes experienced Barrovian metamorphism.

These data show that a maximum of 10 Ma existed between emplacement of the Austroalpine

131

ValaisAustroalpine

South Alpine

units

Helstatt-Meliata ocean

European

ca.90 Ma

ca.40-50 Ma

ca.33 Ma

ca.30 Ma

Eclogite formation

Assembly of Penninc nappe stackand Barrovian metamorphism

Eclogite Zone

Glockner nappe

Venediger nappe

Initiation of Penninic subduction

Subduction of Valaisian crust to ~80 km

Position of future Eclogite Zone

Formation of Middle-Austroalpine eclogites

N S

Conductive heating??

Figure 4.17: Tectonic model for the evolution of the Eastern Alps from Cretaceous to Oligocene, inferred from eclogite

formation at 33 Ma.

nappes, during continental collision, and attainment of Barrovian conditions ca.27–32 Ma. This isup to an order of magnitude faster than previously thought. Furthermore, the Eocene–Oligoceneage of eclogitisation confirm that an active subduction regime was still operating ca.35–30 Ma in theEastern Alps, in addition to there being contemporaneous (U)HP metamorphism and magmatismalong ⇠400 km of the Alpine arc.

132

Chapter 5

40Ar/39Ar Geochronology

40Ar/39Ar data collection undertaken in collaboration with Dr. Clare Warren (Open Uni-versity). Data discussed in this chapter form the basis of the published manuscript: Warrenet al. (2011).

5.1 Introduction

The fast di↵usivity of Ar in white mica under common metamorphic conditions makes 40Ar/39Ardating of white mica, a powerful tool for constraining cooling and therefore exhumation rates ofmetamorphic terranes (e.g. Hacker & Wang, 1995; Sherlock et al., 1999; Von Eynatten et al., 1999;De Sigoyer et al., 2000; Wiederkehr et al., 2009).

Traditionally, 40Ar/39Ar white mica ages have been interpreted according to the elegant closuretemperature concept (Tc) of Dodson (1973), in which:

Tc =R

E · ln(A · ⌧ · D0/a2)(5.1)

where, R is the gas constant, E the activation energy, ⌧ the time constant over which the coe�cientof di↵usion (D) diminishes, a is a characteristic di↵usion size and A is a numerical constant definedby the host mineral geometry and rate of parent nuclide decay. However, equation 5.1 is onlystrictly valid for metamorphic histories in which T�1 increases linearly with time i.e. a constantcooling rate (Dodson, 1973). Clearly, this is not applicable to high @P/@T tectonic regimes such assubduction zones, where changes in both pressure and temperature can occur over rapid (i.e. platevelocity) timescales (e.g. Rubatto & Hermann, 2001; Baldwin et al., 2004; Parrish et al., 2006).Furthermore, the 40Ar/39Ar system in high-pressure white micas is known to be highly susceptibleto contamination by excess radiogenic 40Ar—that is 40Ar incorporated into the mica lattice, whichis not produced by the in situ decay of 40K, resulting in spuriously old 40Ar–39Ar ages(e.g. Arnaud& Kelley, 1995; Ru↵et et al., 1995; Reddy et al., 1996; Scaillet, 1996; Pickles et al., 1997; Ru↵etet al., 1997; Kelley, 2002; Sherlock & Kelley, 2002; Warren et al., 2010). Historically, excess

133

Ar has been identified via an isochron approach, where 36Ar/40Ar is plotted against 39Ar/40Arand deviation from atmospheric Ar (40Ar/36Ar=295.5; Steiger & Jager, 1977) is constrained byregression intercept. However, the accuracy of this method critically relies on there being su�cient36Ar (0.3364±0.0006 atom% Ar Nier, 1950) present to obtain a precise measurement, which isoften not the case (e.g. Sherlock & Kelley, 2002).

Therefore, prior knowledge of the P–T–t evolution of target white mica is of fundamentalimportance to assessing the true thermobarometric significance of 40Ar/39Ar ages. Di↵usion mod-eling of Ar in muscovite following the expected P–T–t path provides a control on the range ofexpected 40Ar/39Ar ages and a theoretical estimate of Tc. Similarly, pre-existing geochronologicalconstraints from highly-retentive radioisotope schemes such as U–Th–Pb, provide a further checkon expected ages and hence the extent of excess Ar contamination or inheritance.

5.2 40Ar/39Ar geochronology in the Tauern Window

As detailed in section 6.2, previous understanding of the rates of tectonometamorphic activity inthe Tauern Window relies heavily on 40Ar/39Ar and K–Ar based chronology. Figure5.1 provides avisual summary of existing 40Ar/39Ar age constraints on the Tauern P–T–t path.

Zimmermann et al. (1994) classified three di↵erent groups of white micas from the Venediger,Eclogite Zone and Glockner nappes, according to their 40Ar/39Ar step-heating plateau ages: (i.)pre-Alpine low-Si relic micas with ages close to ca. 292 Ma, which have been variably reset byAlpine metamorphism; (ii.) phengites of variable Si content with ages between 32–36 Ma; (iii.)low-Si Alpine phengites with a maximum age of 27 Ma. Ages of group ii. phengites of theUpper Schieferhulle are interpreted to represent blueschist facies crystallisation, beneath Tc for Ardi↵usion in phengite. Group iii. mica ages are interpreted as constraining cooling of the nappe pilethrough Tc during late orogenic exhumation. Phengites from Eclogite Zone lithologies pertain tothe peak, eclogite facies event and yield ages between 34–36 Ma. The authors interpret such agesas the time during which the Eclogite Zone cooled through Tc during exhumation.

Ratschbacher et al. (2004) obtained 40Ar/39Ar step–heating spectra from two amphibole sep-arates from the Eclogite Zone: post-eclogitic amphiboles associated with the Barrovian overprintyield a convoluted degassing spectra indicative of excess Ar, with a maximum age of 35±2 Ma;peak, eclogite facies amphibole yield comparably complex, U-shaped, spectra with a maximum ageof 42±4 Ma. Phengite + paragonite separates from Eclogite Zone schists yield 40Ar/36Ar isochronages between 32–41 Ma, all of which have demonstrable excess Ar contamination (MSWD’s be-tween 0.2–166). Nevertheless, the authors claim that their data have geological significance, andusing values for Tc of Ar di↵usion in hornblende and mica derived by Blanckenburg et al. (1989),state that high–P amphibole and mica ages document exhumation of the Eclogite Zone from 15kbar and >500�C at ⇠42 Ma, to ⇠10 kbar and 400�C at ⇠39 Ma.

More recently, Kurz et al. (2008) investigated white mica fractions from both mylonitic andunfoliated eclogites collected from the northern, most pristine, portion of the Eclogite Zone. Anal-

134

400 500 600300Temperature (˚C)

5

10

15

20

25

Pres

sure

(kba

r)

32-36 Ma

27 Ma

42 Ma

<38 Ma56-42 Ma

32 Ma

Figure 5.1: Compilation of existing 40Ar/39Ar geochronol-

ogy and its relationship to the Tauern P–T evolution: orange

data from Kurz et al. (2008); purple data from Ratschbacher

et al. (2004); red data from Zimmermann et al. (1994).

Dotted arrow denoted prograde P–T path of the Peripheral

Schieferhulle.

ysed micas all contain greater than 38% celadonite end-member, demonstrating their high-P origin.Interestingly, micas analysed from mylonitic eclogites yield a narrow range of apparent ages be-tween 31.6–33.45 Ma, whereas the unfoliated eclogite sample gave an older age of 38±0.55 Ma.36Ar/40Ar versus 39Ar/40Ar isotope correlation plots show minor degrees of excess Ar contami-nations with MSWD’s ranging between 0.66–12. This is supported by intercept 36Ar/40Ar valuesbetween 0.00289–0.00347—close to the atomspheric value of 0.00338. The authors go on to sug-gest that deformation under eclogite-facies conditions was responsible for resetting the 40Ar/39Arsystem at ⇡32 Ma, during exhumation of the Eclogite Zone. Their unfoliated eclogite age of ca.38 Ma is taken to represent a maximum age for the eclogite peak, due to possible influence ofhomogenously distributed excess Ar.

However, the thermotectonic significance of 40Ar/39Ar data presented by the studies detailedabove is limited due to: i. the implicit assumption of the value of Tc, independent of cooling rate,grain size and the pressure dependence of Ar di↵usion in white mica; ii. the assumption thatanalysed minerals behave as an open system in which Ar di↵using out of a mineral is e↵ectively re-moved from the grain boundary reservoir (i.e. fluid). Many of the existing phengite and amphibole40Ar/39Ar ages are older than the 34.2±3.6 Ma U–Pb allanite age presented in section 4.8.2 and the31.5±0.7 Ma Rb–Sr age of peak eclogite facies metamorphism obtained via Glodny et al. (2005).This suggests variable incorporation of excess 40Ar throughout Eclogite Zone white micas despiteflat degassing spectra and atmospheric 36Ar/40Ar compositions on isotope correlation plots.

135

5.3 Methodology

The central aim of this work is to establish the thermo-tectonic significance of apparnt 40Ar/39Arwhite mica ages collected from each of the Penninic nappes exposed in the Tauern Window, utilisingavailable U–Pb and Rb–Sr constraints on each of the unit’s P–T–t paths. Following numericalmodeling of Ar di↵usion throughout the Tauern P–T–t cycle provides insight into the magnitudeof expected 40Ar loss and, therefore, the extent of excess 40Ar contamination obsrved.

5.3.1 Sample description and mica chemistry

Samples for Ar isotope analysis were collected across the entire structural profile of the Penninenappe stack from both metasedimentary and metabasite units of the Venediger, Eclogite Zone,Rote-Wand and Glockner nappes—see Fig5.2. Photomicrographs detailing sample petrographyare displayed in Fig.5.3 and chemical relations are detailed in Fig.5.4. The electron microprobedataset is presented in the Appendix (Table A.4).

12˚20'

12˚20'

12˚25'

12˚25'

12˚30'

12˚30'

47˚00' 47˚00'

47˚05' 47˚05'

2500

2000

3000

25003000

2500 2500

1500

150020

00

2000

1500

2500

3000

2000

1500

2 km

Glockner Nappe

Rote Wand/Modereck NappeEclogite Zone - metasediments

Venediger Complex

Eclogite bodies

Thrust fault

Sample location

N45-B

ASA-08-28a

ASA-08-35bASA-08-05ASA-08-84bASA-08-06a-b

CWT 12,13, 15

CWT 17

CWT 8

CWT 5,7CWT 19

Figure 5.2: 40Ar/39Ar sample locations; elevation contours in metres.

136

5.3.1.1 Venediger nappe

Sample ASA–08–28a is a calc–schist collected from the upper-most structural levels of theVenediger nappe (Inner Schieferulle), approximately 200m down structural section from thebase of the Eclogite Zone. The rock comprises a granoblastic matrix of calcite, dolomite andquartz. Hematite is present as an accessory phase. Small amounts (<5%vol.) of muscoviteare aligned to form a weak foliation, aligned with the regional fabric (S2), which reflectsjuxtaposition of the Eclogite Zone and Venediger nappe under Barrovian conditions. Mus-covite is compositionally homogenous, except for the outer 50µm, which displays a decreasein celadonite component, reflecting post peak equilibration, and contains ⇠6.7 Si c.p.f.u. per22 oxygens. Micas were picked from the 250–420µ size fraction.

Sample CWT 17 was collected from the Inner Schieferulle exposed in the upper reaches ofthe Dorfertal and contains the following assemblage, in order of abundance: quartz, albite,muscovite, chlorite, rutile and apatite. White mica flakes define a platy schistosity (S2)and contain ⇠6.7 c.p.f.u. Si; core to rim variation is <0.1 c.p.f.u. Si (Fig.5.4c). Graphiteinclusions were observed in muscovite and were avoided during picking of target grains fromthe >420µ fraction.

5.3.1.2 Eclogite Zone

Sample CWT 8 is an albite mica schist containing quartz, albite, calcite, phengite, paragonite,clinozoisite, hematite and titanite. Refractory albite porphyroclasts contain rotated inclusiontrails and are wrapped by the pervasive muscovite + paragonite foliation (S1+S2). Muscoviteis phengitic (⇠6.65 c.p.f.u. Si) and displays little compositional heterogeneity: core to rimvariation is ⇠0.15 c.p.f.u. Si. Micas were picked from the >420 µm size fraction.

Sample CWT 12 is an Eclogite Zone metapelite, comprising the following mineral assem-blage: garnet, quartz, phengite, chlorite, albite, rutile and kyanite. Centimeter-sized garnetpoikiloblasts contain inclusions of quartz and kyanite; phengites have a high tschermak con-tent (6.73 c.p.f.u. Si), indicating the fabric’s high pressure origin. Phengite for Ar isotopicanalysis was picked from the >820µm fraction and did not contain optically visible inclusions.

Sample CWT 13 is a banded eclogite containing a peak eclogite facies assemblage of: garnet,clinopyroxene, quartz, glaucophane, phengite, epidote, calcite, muscovite, ankerite, rutileand apatite. Compositional banding is defined by millimetre-scale garnet and clinopyroxenedomains. Phengite cores contain the highest celadonite content of all samples: 6.9 c.p.f.u.Si; rim domains (30µm) are less enriched: 6.83 c.p.f.u. Phengites were picked from the >420µm size fraction.

Sample CWT 15 is a retrogressed mafic eclogite comprising the following amphibolite-faciesparagenesis: hornblende, garnet, remnant clinopyroxene, paragonite, chlorite, rutile and zir-

137

Figure 5.3: Photomicrographs detailing the petrology of samples collected for 40Ar/39Ar dating. Photomicrographs

taken under crossed polars; field of view: 2.5⇥3.3 mm for ASA samples; 5.5⇥4.1 mm for CWT samples.

138

con. Hornblende forms symplectite structures; paragonite grains are rich in inclusions (rutile,zircon and hornblende) and are compositionally homogenous, containing ⇠1.83 c.p.f.u. Na(per 22 oxygens). Grains were screened for during picking from the >420 µm size fraction.

Sample ASA–08–06a was collected from the upper 100 m of the Eclogite Zone in the Dorfertal,close to the nappe’s contact with the Rote-Wand nappe. The rock is a coarse–grained (mmscale) calc–mica schist containing dolomite, calcite, muscovite, quartz, epidote, chlorite andmagnetite. Carbonate porphyroblasts are wrapped by the crenulated phengite fabric (S1+S2),within which, constituent phengite grains contains inclusions of carbonate, quartz and minorepidote. Muscovite displays celadonite-rich core regions (6.87 c.p.f.u. Si), surrounded by⇠100µm-wide rim domain of Si-poor (6.55 c.p.f.u. Si) muscovite, indicative of post-peakequilibration at lower pressures, likely during the Barrovian event. The most inclusion-freemicas were picked from the >420 µm size fraction.

Sample ASA–08–06b was collected from the same outcrop as sample ASA–08–06a and is amica schist containing quartz, muscovite, garnet, chlorite and dolomite. The rock is stronglyfoliated with quartz and white-mica forming millimetre-scale microlithons, which form theS1+S2 composite fabric. Phengite is fine (<300µm), contains a variation in Si c.p.f.u. of6.75–6.8 from rim to core, and is free from large-scale inclusions other than quartz.

5.3.1.3 Rote-Wand nappe

Sample CWT 5 is an intensely foliated (S1+S2) mica-rich amphibolite, collected proximalto the transpressional contact with the Eclogite Zone. The rock contains the following min-eral paragenesis: green amphibole, chlorite, epidote, muscovite, albite, K-feldspar, calcite,ilmenite, titanite and rutile. Muscovite is celadonitic in composition with Si c.p.f.u. rangingbetween 6.78–6.9, from rim to core. Fabric-forming mica grains were picked for dating fromthe >420 µm size fraction. Inclusions were not observed optically.

Sample CWT 7 is a quartz-calcite schist containing quartz, calcite, muscovite, paragonite,zoisite, chlorite, rutile, hematite and minor tourmaline. White micas contain included rutileshards which are aligned with the fabric foliation. There are two generations of white mica:coarse (>420 µm) grains form the rock’s fabric (S1+S2) and contain between 6.68–6.8 c.p.f.u.Si, whereas finer (ca.150µm) grained micas define quartz grain boundaries.

Sample ASA–08–05 is a garnet bearing carbonate mica schist. Poikiloblasts of garnet anddolomite are wrapped by muscovite and quartz domains, which define a regionally-concordantschistosity (S2). Analysed micas were picked from the 250–420 µm mesh fraction and showminor compositional variation between 6.58–6.75 Si c.p.f.u.

139

5.3.1.4 Glockner nappe

Sample CWT 19 is a coarse grained marble comprising the following mineral assemblage:calcite, quartz, muscovite, chlorite, epidote, titanite, apatite and hematite. Calcite, quartzand muscovite are all aligned to form a strong, south-dipping foliation (S1+S2). Muscoviteshows phengite solid solution between 6.4–6.8 Si c.p.f.u. (rim–core). Micas for dating werepicked from the 420–850 µm size fraction and were inclusion free.

Sample N45b was collected by N. Ray in the 1980’s (Ray, 1986) from the Bonn-MatreierHutte region, within the central portion of the Glockner nappe. The rock is a mafic schistcontaining quartz, paragonite, epidote, glaucophane, jadeite, albite, actinolite, hematite, ru-tile, and phengite. Phengite (6.68–6.75 c.p.f.u. Si) and epidote define the schistosity (S1+S2),which anastamoses around retrogressed patches of glaucophane and jadeite; paragonite (1.87c.p.f.u. Na) is rare, relative to muscovite and located proximal to glaucophane and jadeitesymplectites. Analysed micas were picked from the >420 µm size fraction.

Sample ASA–08–35b is a banded epidote greenstone from the central Timmeltal, south ofthe Eissehutte. The sample contains epidote, amphibole, albite, chlorite, paragonite andferroan calcite and was collected from the core of a lineated pillow basalt (Holland & Norris,1979) approximately 500m structurally above the base of the Glockner nappe. Muscovite isrelatively rare and occupies selvages between amphibole and epidote aggregates. Analysedmicas were picked from the 250–420µm mesh fraction an show variation in Si content between6.79–6.67, from core to rim.

Sample ASA–08–84b is a banded garnet amphibolite from the base of the Glockner nappeexposed in the Frosnitztal. Epidote, albite, chlorite, garnet, amphibole and paragonite formmm-cm scale bands between more micaceous sub-domains. Paragonite (⇠1.82 c.p.f.u. Na)forms the main schistosity in the rock. Rare large (>450µm) porphyroblasts of paragoniteof similar composition grow obliquely to the foliation. Analysed micas were picked from the>420 µm fraction.

5.3.2 Step-heating versus single grain fusion data

Previous 40Ar/39Ar dating studies in the Tauern region have used stepwise-heating (see McDougal& Harrison, 1988, for review) to liberate trapped Ar gas from host micas and amphiboles (Zim-mermann et al., 1994; Ratschbacher et al., 2004; Kurz et al., 2008). The step-heating techniquegenerally uses multiple grains in a single experiment for which repeated analyses are rarely pre-sented. Step-heating is assumed to represent the degassing of Ar held in di↵erent reservoirs. Earlygas fractions released at low-temperature increments correspond to loosely held Ar, commonlylocated close to the grain surface, whereas gas fractions released at higher temperatures generallyhave higher 40Ar*/39Ar ratios because Ar is then removed from more retentive crystallographic

140

Na

KCa

K

90

80

Mus

covit

eK 2A

l 4[Si 6A

l 2O20

](OH,

F)4

Na

90

80Pa

rago

nite

Na2A

l 4[Si 6A

l 2O20

](OH,

F)4

CW

T 8,

12,1

3AS

A 6a

,6b

CW

T 17

ASA

28a

CW

T 5,

7AS

A 5

CW

T 19

ASA

35b,

N45

bC

WT

8,13

,15

CW

T 7

ASA

35b

2.0

1.5

1.0

0.5

0.0

Al[IV]

2.0

1.5

1.0

0.5

0.0

Al[V

I] -2

CW

T 5

CW

T 7

ASA

5

CW

T 8

CW

T 12

CW

T 13

ASA

6a

ASA

6b

C

WT

19 N

45b

ASA

35b

A

SA 2

8a C

WT

17

Cel

Phen

gite

Mu

a. b.7.

0

6.5

6.0

5.5

5.0

Si (c.pf.u)

1400

1200

1000

800

600

400

200

0Di

stan

ce (µ

m)

CW

T 8

CW

T 12

CW

T 13

CW

T 15

ASA

6a

ASA

6b

7.0

6.5

6.0

5.5

5.0

2000

1500

1000

500

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T 9

N45

b A

SA 3

5b

7.0

6.5

6.0

5.5

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1200

1000

800

600

400

200

0

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T 5

CW

T 7

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7.0

6.5

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141

sites. According to this understanding, with increasing temperature, the 40Ar/39Ar history of asuite of micas is revealed. However, step-heating does not permit in situ correlation of 40Ar/39Arcompositions with grain position, which is critical for determination of whether a mica grain hasexperienced an open or closed system during its evolution (e.g. Wiederkehr et al., 2009; Warrenet al., 2010).

The single grain fusion approach, developed from early work by Megrue (1971), utilises thespatial resolution of a UV laserprobe to release trapped Ar from individual ablation sites withina single mineral grain, thus permitting evaluation of spatial variability in Ar concentrations (andhence apparent ages)(e.g. Sherlock & Kelley, 2002; Putlitz et al., 2005). The spatial resolution oflaserprobe 40Ar/39Ar age data is limited by the volume of Ar gas released, meaning that youngergrains require larger gas volumes (i.e. larger laser spot sizes; larger grains). The age resolutionobtained from, in-situ analysis of Tauern (ca.30 Ma) micas is insu�cient to resolve within-graindi↵erences outside analytical and systematic uncertainty. Therefore, this study employs the spatialresolution of the UV laserprobe approach applied to suites of 10–20 individual micas grains persample, which permits determination of Ar isotopic variability across the sample volume: ⇠1 cm3.

5.4 40Ar/39Ar analytical technique

Analytical work was undertaken by Dr. Clare Warren at the Open University; the methodologypresented beneath is detailed in Warren et al. (2010).

Target samples were crushed, washed and sieved prior to screening the white mica fraction fordeformation e↵ects and inclusions. The largest suitable mica grains were the washed in acetone,methanol and water before packing into foil packets for irradiation. An additional cold 0.2MHCl wash was applied to micas picked from calcareous rocks to remove any carbonate dust. Allsamples were irradiated at McMaster University in Canada. Irradiation flux was monitored usingthe GA1550 biotite standard with an age of 98.79±0.54 Ma (Renne et al., 1998). Sample J valueswere calculated by linear interpolation between two bracketing standards; a standard was includedbetween every 8–10 samples in the irradiation tube. Results were corrected for blanks, 37Ar decayand neutron-induced interference reactions. Background measurements bracket every 1–2 samples.Muscovite analyses and standards were corrected for (40Ar/39Ar)K = 0.0085 based on analysis ofK salts but not corrected for Ca decay.

Due to the small sample size, Ar sample and blank measurements approached detection limits,and were commonly within error of each other. Correcting for atmospheric Ar magnifies errorson the 36Ar measurement and results in anomalously high analytical errors on the final 40Ar/39Arage (e.g. Sherlock et al., 2005; Warren et al., 2010). Atmospheric Ar contents were measurable onlarger samples and were generally <5% of the total Ar. Samples were therefore only corrected foratmospheric Ar where the 36Ar measurement was >2⇥ the 36Ar blank and outside the blank mea-surement uncertainty. Uncertainty on the calculated age for samples which remained uncorrectedfor atmospheric Ar was doubled (e.g. Sherlock et al., 2005, 2008; Warren et al., 2010). Samples

142

were loaded into an ultra-high-vacuum laser port and placed under a heat lamp for eight hoursto reduce atmospheric blank levels. Total fusion of single grains was achieved using a Nd-YAG1064 nm infra-red laser coupled to an automated gas handling vacuum system and admitted intoan MAP 215–50 noble gas mass spectrometer. Gases were gettered for 5 minutes using two SAESgetters (at 450�C and room temperature respectively), and a liquid nitrogen cold trap before inletinto the mass spectrometer. Peaks from 36Ar to 40Ar were scanned 10 times each and amountswere extrapolated back to the inlet time.

5.5 Results

Raw 40Ar/39Ar single grain data are presented in Table A.9. Apparent 40Ar/39Ar ages frominfra–red laser fusion of 218 mica grains hosted in 15 di↵erent samples spread between 90–23 Ma(Fig.5.5). Intra–sample age variation between grains ranges between 4.6 Ma (32.8–37.4 Ma; sampleCWT–8) and 57 Ma (33.4–90.4 Ma; sample ASA–35b). Paragonite from sample ASA–84b of theGlockner nappe yields the youngest 40Ar/39Ar age measured in this study (23.57 Ma), whereas theoldest single grain fusion age was found in a deformed pillow basalt, ASA–35b (90.40 Ma).

Figure 5.5 plots apparent 40Ar/39Ar grain ages against sample number. Interestingly, a sig-nificant proportion of single grain ages are older than the 34.2±3.6 Ma U–Pb ID-TIMS allaniteage (Fig.4.7), interpreted to represent the time during which the European margin was subductedthrough ⇠8–13 kbar. Only 3 samples (CWT–19, ASA–05 and CWT–8) do not yield 40Ar/39Arages older than the 2� variation on the U–Pb allanite age.

Samples containing paragonite (± muscovite) generally show greater spread in single grain agesthan those containing solely muscovite. In part, this is due to the larger errors associated withsmaller volumes of 40Ar and 39Ar released during fusion.

Whole rock weight% K2O was calculated by combining white mica modal volume estimates ob-tained via point counting with representative electron microprobe analyses. This method assumesthat white mica is the only significant K bearing phase present—a valid assumption given samplepetrography (Fig.5.3). Figure 5.6 shows that there is no apparent correlation between the range ofages observed and bulk rock K2O content, contrary to what might be expected in a closed system(i.e. Foland, 1979). Similarly, apparent age variation does not correlate with structural position inthe nappe stack, or regional-scale intensity of deformation (Fig.5.6).

Isotope correlation graphs plotting 36Ar/40Ar versus 39Ar/40Ar (Turner et al., 1971; Roddicket al., 1980) provide an estimate of the degree to which data can be described by mixing betweennon–radiogenic and radiogenic Ar reservoirs. 39Ar/40Ar intercept values close to the atmosphericvalue of 0.00338, are indicative of subordinate contamination by excess Ar (McDougal & Harrison,1988). Tauern white mica data form highly scattered clusters close to the radiogenic 39Ar/40

intercept (Fig.5.7). The low concentrations of measured 36Ar and incoherent scatter of the datameans that intercept values have large errors and cannot be used to estimate the compositionof contaminant excess Ar. Additionally, several samples, such as ASA–35b (Fig.5.7b), display

143

two di↵erent isotopic signals, suggestive of mixing between three, or more, isotopically unique Arreservoirs (i.e. non-radiogenic and two di↵erent radiogenic sources).

Collectively, the data show that the studied white micas are variably contaminated with 40Arseparated from its parent 40K—excess Ar (40ArE). The degree of excess Ar contamination variesat the grain-scale, i.e. mm–cm, as well as at the regional scale i.e. km.

90

80

70

60

50

40

30

20

Appa

rent

40Ar

/39Ar

Age

(Ma)

N45

b(mu)

CW

T-19

(mu)

ASA-

35b

(pg+mu)

ASA-

84b

(pg) AS

A-05

(pg)

ASA-

06a

(mu)

ASA-

06b

(mu)C

WT-

5(mu)

CW

T-7

(mu+pg)

CW

T-8

(mu+pg)

CW

T-12

(mu)

CW

T-13

(mu+pg)

CW

T-15

(pg) CW

T-17

(mu)

ASA-

28a

(mu)

Figure 5.5: Apparent single grain fusion 40Ar/39Ar ages plotted against sample number and mineralogy (italicised

subscript: mu = muscovite; pg = paragonite). Colour scheme is as follows: green = Glockner nappe; orange = Rote-

Wand nappe; blue = Eclogite Zone; red = Venediger nappe. Individual data points represent a single analysed grain;

errors are ± 1�. Horizontal band represents 2� age range of U–Pb ID-TIMS allanite age (Fig.4.7).

5.6 Numerical Modelling

In order to assess the magnitude of unsupported 40Ar contamination in reported white mica39Ar/40Ar ages, it is important to understand the e↵ect of volume di↵usion of Ar within Tauernmicas. Thermally activated volume di↵usion of Ar in white mica can be approximated by Fick’ssecond law (Fick, 1855):

@C

@t= D

@2C

@z2(5.2)

where C is the concentration of Ar, t is time, z is spatial position and D is the di↵usion coe�cient.As this is a thermally driven process, the di↵usion coe�cient is defined by the following expression:

144

45

40

35

30

25

4

3

2

1

Weig

ht %

K2O

3500

3000

2500

Ele

vation (

m.a

.s.l.)

Appare

nt

40A

r/3

9A

r age (

Ma)

S N

N45b

CW

T-1

9

AS

A-3

5b

AS

A-8

4b

AS

A-0

5

CW

T-5

CW

T-7

AS

A-0

6a

AS

A-0

6b

CW

T-8 CW

T-1

2

CW

T-1

3

CW

T-1

5

CW

T-1

7

AS

A-2

8a

Figure 5.6: Apparent 40Ar/39Ar ages (note:smaller scale than Fig.5.5) plotted against sample weight % K2

O and

structural position in the nappe stack. Age errors are ± 1�; weight % K2

O calculated via combining mineral modes

obtained via point counting with representative white mica probe analyses. The following color scheme is used: green

= Glockner nappe; orange = Rote-Wand nappe; blue = Eclogite Zone; red = Venediger nappe. Sample positions are

approximate given along-strike variability in nappe thickness. Width of cross section is ⇠5 km. Horizontal shaded region

represents the 2� range of the 34.2±3.6 Ma U–Pb ID–TIMS allanite age; vertical bands denote areas of intense strain at

the nappe boundaries.

145

0.000

0.001

0.002

0.003

0.0 0.2 0.4 0.60.000

0.002

0.004

0.006

0.008

0.010

0.0 0.2 0.4 0.6

0.000

0.001

0.002

0.003

0.004

0.005

0.0 0.2 0.4 0.60.000

0.001

0.002

0.003

0.0 0.2 0.4 0.6

36Ar

/40Ar

39Ar/40Ar

Sample: N45bAge = 36.5±2.6 Ma

Initial 40Ar/36Ar = 436±270MSWD = 5.6

data-point crosses are 2σ

36Ar

/40Ar

39Ar/40Ar

Sample: ASA-35bAge = 48.04±6.6 Ma

Initial 40Ar/36Ar = 195±290MSWD = 2.3

data-point crosses are 2σ

36Ar

/40Ar

Sample: CWT-8Age = 34.7±1.6 Ma

Initial 40Ar/36Ar = 234±110MSWD = 2.8

data-point crosses are 2σ

39Ar/40Ar

Sample: ASA-28aAge = 56±24 Ma

Initial 40Ar/36Ar = 517±670MSWD = 11.2

data-point crosses are 2σ

39Ar/40Ar

36Ar

/40Ar

a. b.

c. d.

Figure 5.7: 36Ar/40Ar vs. 39Ar/40Ar isotope correlation plots (Turner et al., 1971; Roddick et al., 1980) for samples:

N45b (a.), ASA–35b (b.), CWT–8 (c.) and ASA–28a (d.).

D = D0e�ERT (5.3)

where D0 is the pre-exponential di↵usion coe�cient, R is the gas constant, T is the temperatureand E is the experimentally derived activation energy.

Numerical modeling of Ar di↵usion was undertaken using the MATLABTM program DIFFARG(Wheeler, 1996), which uses a finite-di↵erence algorithm to calculate bulk ages of individual micagrains given a specified thermal history. The program was modified (DIFFARGP) by C. War-ren (Open University) and F. Hanke (University of Liverpool) to include the latest experimentaldi↵usion parameters for Ar di↵usion in muscovite (Harrison et al., 2009): E=63±7 kcal.mol�1;D0=2.3 +70

�2.2 cm2.s�1, with an activation volume of 14 cm3.mol�1.The di↵usion modeling attempts to calculate the expected range in apparent 40K/40Ar mus-

146

covite ages for an open system scenario (i.e. zero grain boundary 40Ar concentration) for theindividual thermal histories recorded by each of the Penninic nappes of the Tauern Window.

5.6.1 Numerical techniques

Fick’s second law is augmented by a source term which takes into account the production of 40Arwithin the mineral grain:

@C

@t= D

@2C

@z2+ �[40K]ie��t (5.4)

where � is the total decay constant (i.e. production of both 40Ar+40Ca) of 40K and [40K]i is theinitial concentration (t=0) of 40K. This equation is integrated over time in a single spatial dimen-sion using the Crank-Nicholson finite di↵erence scheme (p.141,144, Crank, 1975), which is stablefor all sizes of timesteps and permits fast calculation times. The DIFFARG algorithm outputcomprises single grain profiles of apparent 40K/40Ar age. To obtain bulk grain ages, DIFFARGPapplies Simpson’s rule over the mesh interval in question. As numerical error on the bulk grainage is sensitive to the size of the mesh separation used, calculations were run for three di↵erentmesh sizes (10, 50 and 100 seed points), after which, the bulk ages were plotted against meshseparation and regressed to calculate a hypothetical infinite mesh density age: this is the valuereported beneath. Warren et al. (2010) provide a more detailed discussion of DIFFARGP.Ranges of apparent ages were calculated for the Venediger, Eclogite Zone, Rote Wand and Glock-ner nappes, taking into account uncertainty in both P–T estimates, the timing of metamorphismand grain size. As the di↵usion data of (Harrison et al., 2009), utilised in DIFFARGP, were calcu-lated using a multi-di↵usion domain model with a spherical geometry, as opposed to a cylindricalgeometry, model ages were calculated for spherical muscovite grains. A reference cooling rate of30�C.Ma�1, as calculated by Blanckenburg et al. (1989) for the Schieferhulle nappes of the westernTauern Window, was employed; decompression rates were calculated such that surface pressurevalues were attained at the same time as temperature values.

5.6.2 Results

Model ages are presented beneath, according to structural position in the Tauern nappe stack.

5.6.3 Venediger nappe

The reference model for muscovite from the Venediger nappe is based on a 0.5 mm radius grain(see section 5.3.1.1) crystallising at 13 kbar, and 600�C (see chapter 3) at 30 Ma (Inger & Cli↵,1997), before undergoing cooling at 30�C.Ma�1 and 0.65 kbar.Ma�1. These conditions yield anage of 26.58 Ma. Results are presented in Table 5.1.

Uncertainty in grain size and therefore, e↵ective di↵usion volume, between 0.1 and 1 mm radius,propagates to error on the reference age of 26.58+0.93

�1.93 Ma (models: VN 1–2). Variations in the

147

40 35 30 25 20 15 10 5 0time (Ma)

0

100

200

300

400

500

600

tem

pera

ture

(˚C)

Eclogite Zone

Glockner nappe

Venediger nappe

30˚C.Ma-1

Figure 5.8: Temperature–time trajectories used in di↵usion

modeling. Note: the Rote Wand nappe is not displayed as

it is thought to have experienced a similar T–t path to the

Glockner nappe.

pressure of Barrovian metamorphism (+2 kbar; -4 kbar; model: VN 5–6), yield ages 0.24 Ma and-0.45 Ma from the reference model. Temperature uncertainty has a stronger e↵ect: ±50�C yieldserror of 27.76+1.58

�1.33 respectively. Post-Barrovian cooling rates of 20, 40 and 50�C.Ma�1 yield ages-2.10 Ma, 1.00 Ma and 1.57 Ma di↵erent from the reference age (model: VN 9–11). The strongestcontrol on model di↵usion age is the timing of peak metamorphism: an uncertainty of ±3 Ma(27–33 Ma; see section 6.2) translates to a ±3 Ma uncertainty on the reference age (model: VN3–4). Despite there being scant petrological evidence for the Venediger nappe experiencing earlyHP metamorphism, model VN 12 considers the e↵ect on model di↵usion ages of muscovite growingduring at 15 kbar, 450�C and 33 Ma, before following the reference trajectory. The resultant age isidentical to the reference model, which confirms that the Barrovian event was of su�ciently highgrade to erase potential Ar ’memory’ of muscovite.

5.6.4 Eclogite Zone

The Eclogite Zone reference model is based on a 0.5 mm radius muscovite crystallising duringeclogite facies metamorphism at 25 kbar, 550�C and 33 Ma, followed by blueschist facies meta-morphism at 12.5 kbar, 450�C and 31 Ma, and late-stage Barrovian metamorphism at 7.5 kbar,550�C and 30 Ma. These parameters produce a model di↵usion age of 27.76 Ma. Modeling resultsare presented in Table 5.2.

Uncertainty in the P–T conditions of both the eclogite and blueschist facies metamorphic eventsdoes not e↵ect the model age (models: EZ 1–4, and 7), indicating that Barrovian temperatures arehigh enough to facilitate di↵usive resetting of the grain’s 40Ar budget. Uncertainty in grain size(e↵ective di↵usion volume) between 0.1–1 mm yields uncertainty in the reference age of 27.76+1.03

�1.8

Ma. Variation in Barrovian conditions of ±50�C, ±1.5 kbar and -3,+2 Ma yield deviations fromthe reference age of 27.76+2.03

�1.33 Ma, 27.76+0.18�0.17 Ma, 27.76+2

�3 Ma, respectively (models: EZ 8–13). Cooling rates (models: EZ 14–16) from the peak of Barrovian metamorphism of 20, 40 and50�C.Ma�1 translate to ages which are -1.57, 0.65 and 1.06 Ma from the reference age. Running

148

Model Size1 P2 T3 t4 C.rt.5 D.rt.6 Age7 �age8

VN ref. 0.5 13 600 30 30 0.65 26.58 -VN 1 0.1 13 600 30 30 0.65 24.66 -1.93VN 2 1 13 600 30 30 0.65 27.52 0.93VN 3 0.5 13 600 33 30 0.65 29.58 3.00VN 4 0.5 13 600 27 30 0.65 23.58 -3.00VN 5 0.5 15 600 30 30 0.75 26.82 0.24VN 6 0.5 9 600 30 30 0.45 26.13 -0.45VN 7 0.5 13 550 30 27.5 0.65 28.16 1.58VN 8 0.5 13 650 30 32.5 0.65 25.26 -1.33VN 9 0.5 13 600 30 20 0.433 24.49 -2.10VN 10 0.5 13 600 30 40 0.86 27.58 1.00VN 11 0.5 13 600 30 50 1.08 28.15 1.57VN 12? 0.5 13 600 30 30 0.65 26.58 -0.02

Table 5.1: DIFFARGP modeling results for muscovite following the Venediger nappe P–T–t path. Bold font highlightsthe variable changed from reference model conditions. Superscript notation: 1 Grain radius in mm; 2 Barrovian pressurein kbar; 3 Barrovian temperature in �C; 4 Age of Barrovian in Ma; 5 Cooling rate in �C.Ma�1; 6 Decompression rate inkbar.Ma�1; 7 Di↵usion age in Ma; 8 Di↵erence to reference age; ? Prograde growth of muscovite at 33 Ma, 15 kbar and450�C, prior to Barrovian event.

the reference model with a cylindrical geometry exerts only a minor control on model age: 28.2 Ma(model: EZ 17). Finally, a model muscovite growing during the Eclogite Zone’s prograde P–T pathat 13 kbar, 450�C and 36 Ma (i.e. within uncertainty of the U–Th–Pb allanite age), is expectedto have an identical di↵usion age to the reference model (model: EZ 18).

5.6.5 Rote-Wand nappe

The Rote Wand nappe experienced a similar P–T–t trajectory to the Glockner nappe (Dachs &Proyer, 2001), which precludes the need to model muscovite di↵usion ages separately.

5.6.6 Glockner nappe

The reference model for the Glockner nappe assumes a muscovite grain radius of 0.5 mm, growing33 Ma, during early, subduction related, HP metamorphism at 450�C and 10 kbar, followed byregional Barrovian metamorphism 30 Ma at 7.5 kbar and 550�C. These conditions yield a referenceage of 27.76 Ma, which is identical to the Eclogite Zone reference age (Table 5.2). Results arepresented in Table 5.3.

Uncertainty in model muscovite grain radius, between 0.1–1 mm (consistent with observed vari-ation in grain radius: section 5.3.1.4), propagates to uncertainty on the reference age of 27.76+0.97

�1.8

Ma (models:GN 1–2). As reported for both Venediger and Glockner nappe models, uncertainty inthe conditions of HP metamorphism does not e↵ect the reference age (models: GN 3–8). A ±50�Cerror in the temperature and a ±1.5 kbar error in the pressure of Barrovian metamorphism resultsin reference age uncertainty of 27.76�1.32

+0.04 Ma and 27.76+0.18�0.17 Ma, respectively (models: GN 9–12).

149

Model Size1 PHP2 THP

3 tHP4 PMP

5 TMP6 tMP

7 PBar8 TBar

9 tBar10 C.rt.11 D.rt.12 Age13 �age14

EZ ref. 0.5 25 550 33 12.5 450 31 7.5 550 30 30 0.41 27.76 -EZ 1 0.5 25 550 33 � � � 7.5 550 30 30 0.41 27.76 0.00EZ 2 0.5 20 550 33 12.5 450 31 7.5 550 30 30 0.41 27.76 0.00EZ 3 0.5 25 600 33 12.5 450 31 7.5 550 30 30 0.41 27.76 0.00EZ 4 0.5 25 500 33 12.5 450 31 7.5 550 30 30 0.41 27.76 0.00EZ 5 1 25 550 33 12.5 450 31 7.5 550 30 30 0.41 28.79 1.03EZ 6 0.1 25 550 33 12.5 450 31 7.5 550 30 30 0.41 25.96 -1.80EZ 7 0.5 25 550 37 12.5 450 31 7.5 550 30 30 0.41 27.76 0.00EZ 8 0.5 25 550 33 12.5 450 31 7.5 500 30 30 0.45 29.79 2.03EZ 9 0.5 25 550 33 12.5 450 31 7.5 600 30 30 0.37 26.43 -1.33EZ 10 0.5 25 550 33 12.5 450 31 9 550 30 30 0.49 27.94 0.18EZ 11 0.5 25 550 33 12.5 450 31 6 550 30 30 0.32 27.59 -0.17EZ 12 0.5 25 550 33 12.5 450 31 7.5 550 27 30 0.41 25.08 -3.00EZ 13 0.5 25 550 33 12.5 450 31 7.5 550 32 30 0.41 29.44 2.00EZ 14 0.5 25 550 33 12.5 450 31 7.5 550 30 20 0.27 26.19 -1.57EZ 15 0.5 25 550 33 12.5 450 31 7.5 550 30 40 0.54 28.42 0.65EZ 16 0.5 25 550 33 12.5 450 31 7.5 550 30 50 0.68 28.82 1.06EZ 17? 0.5 25 550 33 12.5 450 31 7.5 550 30 30 0.41 28.20 0.43EZ 18† 0.5 25 550 33 12.5 450 31 7.5 550 30 30 0.41 27.76 0.00

Table 5.2: DIFFARGP modeling results for muscovite following the Eclogite Zone P–T–t path. Bold font highlights thevariable changed from reference model conditions. Superscript notation: 1 Grain radius in mm; 2 Pressure of HP eventin kbar; 3 Temperature of HP event in �C; 4 Age of HP event in Ma; 5 Pressure of blueschist facies (MP ) event inkbar; 6 Temperature of blueschist facies overprint in �C; 7 Age of blueschist facies event in Ma;8 Pressure of Barrovianevent in kbar; 9 Temperature of Barrovian event in �C; 10 Age of Barrovian event in Ma; 11 Cooling rate (post-Barrovianmetamorphism) in �C.Ma�1; 12 Decompression (post-Barrovian metamorphism) rate in kbar.Ma�1; 13 Di↵usion age inMa; 14 Di↵erence to reference age in Ma; ? Reference model age for muscovite with a cylindrical geometry; † Model agefor prograde growth of muscovite at 13 kbar, 450�C, 36 Ma, before experiencing the reference model P–T–t cycle.

150

Model Size1 PHP2 THP

3 tHP4 PBar

5 TBar6 tBar

7 C.rt.8 D.rt.9 Age10 �age11

GN ref. 0.5 10 450 33 7.5 550 30 30 0.41 27.76 -GN 1 0.1 10 450 33 7.5 550 30 30 0.41 25.96 -1.80GN 2 1 10 450 33 7.5 550 30 30 0.41 28.73 0.97GN 3 0.5 10 450 37 7.5 550 30 30 0.41 27.76 0.00GN 4 0.5 10 450 31 7.5 550 30 30 0.41 27.76 0.00GN 5 0.5 15 450 33 7.5 550 30 30 0.41 27.76 0.00GN 6 0.5 7 450 33 7.5 550 30 30 0.41 27.76 0.00GN 7 0.5 10 400 33 7.5 550 30 30 0.41 27.76 0.00GN 8 0.5 10 500 33 7.5 550 30 30 0.41 27.76 0.00GN 9 0.5 10 450 33 9 550 30 30 0.49 27.94 0.18GN 10 0.5 10 450 33 6 550 30 30 0.32 27.59 -0.17GN 11 0.5 10 450 33 7.5 500 30 30 0.41 27.80 0.04GN 12 0.5 10 450 33 7.5 600 30 30 0.41 26.44 -1.32GN 13 0.5 10 450 33 7.5 550 27 30 0.41 24.76 -3.00GN 14 0.5 10 450 33 7.5 550 32 30 0.41 29.76 2.00GN 15 0.5 10 450 33 7.5 550 30 20 0.27 26.19 -1.57GN 16 0.5 10 450 33 7.5 550 30 40 0.54 28.41 0.65GN 17 0.5 10 450 33 7.5 550 30 50 0.68 28.82 1.06

Table 5.3: DIFFARGP modeling results for muscovite following the Glockner nappe P–T–t path. Bold font highlightsthe variable changed from reference model conditions. Superscript notation: 1 Grain radius in mm; 2 Pressure of HPevent in kbar; 3 Temperature of HP event in �C; 4 Age of HP event in Ma; 5 Pressure of Barrovian event in kbar; 6

Temperature of Barrovian event in �C; 7 Age of Barrovian event in Ma; 8 Cooling rate (post-Barrovian metamorphism) in�C.Ma�1; 9 Decompression (post-Barrovian metamorphism) rate in kbar.Ma�1; 10 Di↵usion age in Ma; 11 Di↵erence toreference age in Ma.

Uncertainty over the age of Barrovian metamorphism exerts a strong control on the model age:27.76+2

�3 Ma (models: GN 13–14). Finally, post-Barrovian cooling rates of 20, 40 and 50 �C.Ma�1

yield the following devaitions from the reference age: -1.57, 0.65 and 1.06 Ma (models: GN 15–17).

5.6.7 Uncertainty associated with di↵usion parameters

Harrison et al. (2009) impose experimental error of ±11% on the value of E and an error envelopebetween 0.1–72.3 cm2.s�1 for D0. The large range in possible D0 arises from significant scatter inthe values of di↵usion coe�cients plotted on an Arrhenius plot of di↵usion coe�cient against recip-rocal absolute temperature. Experimental uncertainty of di↵usion parameters exerts the strongestinvestigated control on di↵usion age (see Table 5.4): the ±11% uncertainty of E, translates to anerror of 27.76+3.9

�2.8 Ma on the Eclogite Zone reference model (models EZ 19–20), whilst maximumand minimum values of D0 yield an di↵usion age error of 27.76�1.94

+3 Ma (models EZ 21–22).

5.6.8 Discussion

Primarily, model di↵usion ages predict that muscovite present within Venediger, Eclogite Zoneand Glockner nappes should yield 40Ar/39Ar ages independent of their pre-Barrovian history—i.e.Barrovian temperatures between 500–600� at 6–9 kbar are su�cient to e↵ectively remove radigenic

151

Model E1 D02 Age3 �age4

EZ ref. 263592 230 27.76 -EZ 19 270585 230 31.69 3.93EZ 20 234334 230 24.88 -2.88EZ 21 263592 7230 25.82 -1.94EZ 22 263592 10 30.76 3

Table 5.4: E↵ect of uncertainty of thermal parameters on DIFFARGP Eclogite Zone reference model. Superscriptnotation: 1 Activation energy in J.mol�1; 2 Pre-exponential di↵usion coe�cient in mm2.s�1; 3 Model age in Ma; 4

Di↵erence to reference age in Ma.

40Ar from the muscovite lattice. This rule is broken only for minimum values of D0 and maximumvalues of E, inside experimental uncertainty.The Eclogite Zone and Glockner nappe yield identical model ages (27.76 Ma) as they experience thesame strength of Barrovian overprint, whilst the Venediger nappe is predicted to have a marginallyyounger reference age (26.58 Ma) due to higher Barrovian temperatures at the base of the structuralpile.

In order of importance, the following parameters operate to control di↵usive transport of 40Arwithin muscovite: i. Temperature of Barrovian overprint; ii. Grain size – e↵ectively controls thedi↵usive volume; iii. Cooling rate – faster cooling rates retain more 40Ar; iv. Pressure of Barrovian– although a relatively small e↵ect, pressure serves to slow di↵usion rates.

Logically, the apparent 40Ar/39Ar mica age is a function of the absolute age of peak Barrovianconditions. Rb–Sr multimineral (including white mica) geochronology shows this to have occurredbetween 27–32 Ma throughout the nappe stack (Inger & Cli↵, 1997; Glodny et al., 2005; Gleißneret al., 2007). The 5 Ma range in values likely reflects Sr’s sensitivity to deformative and fluidresetting as opposed to protracted metamorphism (e.g. Thoni & Jagoutz, 1992). However, as theBarrovian event controls 40Ar/39Ar mica ages, this uncertainty imposes a +2, �3 Ma error on allmodel ages.

Consideration of the full experimental error associated with E and D0 shows that they havethe potential to drastically change the rates of 40Ar di↵usive transport within muscovite at peakBarrovian temperatures. However, these errors are likely conservative, given the limited number ofdata points from which utilized values of D0 were calculated from (n=14; Harrison et al., 2009).

In summary, the modeling shows that under conditions in which 40Ar produced in-situ, withinmuscovite, is e↵ectively removed from the grain (i.e. open system), all Tauern 40Ar/39Ar muscoviteages should record the timing of cooling on exhumation from peak Barrovian metamorphism.

5.7 Discussion

A total of 13 out of 15 investigated samples yield single grain 40Ar/39Ar ages which are olderthan: i. the 34.2±3.6 Ma U–Pb allanite age (section 4.8.2), ii. the 27–32 Ma Rb–Sr age interval

152

for Barrovian metamorphism (Inger & Cli↵, 1997; Glodny et al., 2005; Gleißner et al., 2007), andiii. the numerically calculated model di↵usion ages, outside of uncertainty (section 5.6). In theremaining samples, ASA–05 and CWT–19, the majority of single grain ages are older than 32 Ma.

Table 5.5 compares 40Ar/39Ar weighted average ages, growth conditions and mineralogy of thesample mica population with nappe-specific model di↵usion ages. There is no apparent correlationbetween structural position in the nappe pile and age; similarly, neither do conditions of micagrowth and average age seem to be linked. Irrespective of geological context, all samples showaverage inter-grain ages between 3.9–22.3 Ma older than their relevant di↵usion ages. Interestingly,mafic samples, in general, show a greater spread in single grain ages than pelitic samples.

The 40Ar/39Ar data, when coupled with numerically calculated di↵usion ages, show that atleast one of the central tenets (i. Ar concentration within mica is controlled by tempertaure-dependent volume di↵usion; ii. target white mica crystallizes without any 40Ar within its lattice;iii. mica grains act as an open system in which Ar di↵using across the grain boundary is removedinto an open rock voume with an e↵ective Ar concentration of zero) under which 40Ar/39Ar agesare traditionally interpreted, has been compromised. According to numerical di↵usion modeling,temperatures of the Barrovian overprint were high enough to erase pre-Barrovian Ar history byvolume di↵usion. Therefore, the range of apparent white mica ages reflect variable incorporation of40Ar decoupled from its parent 40K – termed excess Ar (40ArE ; McDougal & Harrison, 1988). Thepresence of 40ArE shows that the white mica grains did not behave in a sensu-stricto open systemfashion in which the 40Ar grain boundary network concentration was non-zero, either during ofsubsequent to mica formation.

Generally, the presence of 40ArE in rocks is considered to be detrimental to the interpretation of40Ar/39Ar age data. However, 40ArE concentrations can be used to provide insight into the degreeof metamorphic connectivity throughout a particular P–T path (e.g. Camacho et al., 2005). Thepartitioning of 40ArE into white mica can be explained by two end-member scenarios (Fig.5.9):i. Open system – where 40Ar is fluid-borne within an interconnected grain boundary reservoiri.e. shear zones with high fluid fluxes, (e.g. Allen & Stubbs, 1982) and ii. Closed system –fluid poor systems, where the 40Ar concentration of the grain boundary network is controlled byrock porosity and whole-rock K2O content (Foland, 1979). The following discussion focusses ondiscerning which of these models best accounts for the range of 40ArE concentrations measured inTauern white micas.

5.7.1 40ArE partitioning

Before discussing the relative merits of open versus closed systematics in Tauern micas, it is aworthwhile exercise to summarize the behavior of Ar between phases in metamorphic systems.

Ar is a highly incompatible trace element, preferring to reside in grain boundary fluids ratherthan host minerals. This satisfies the zero initial 40Ar concentration constraint, central to the

153

Sample Unit Lithology1 Mineral Event2 Range3 40Ar/39Ar4 ±5 n6 Di↵arg7 �age8

ASA28a VN ab-s mu M3 13.82 48.9 2.9 14 26.58 22.32CWT17 VN c-s mu M3 10.28 33.58 0.77 10 26.58 7ASA6a EZ m-s mu M1 15.76 35.2 1.9 12 27.76 12ASA6b EZ m-s mu M1 16.39 39.6 3.1 12 27.76 11.84CWT8 EZ m-s mu+pa M1/3 4.78 34.58 0.81 15 27.76 6.82CWT12 EZ m-s mu M1 5.99 39.2 1 17 27.76 11.44CWT13 EZ m mu+pa M1 22.80 42.4 5.3 10 27.76 14.64CWT15 EZ m pa M3 16.55 38.6 3.3 12 27.76 10.84ASA5 RW m-s mu M1/3 6.49 32.59 0.97 16 27.76 4.83CWT5 RW m mu M3 8.39 31.7 0.85 17 27.76 3.94CWT7 RW c-s mu+pa M1/3 43.95 33.5 2.8 18 27.76 5.74N45b GN m mu+pa M1 8.32 37.6 1.6 10 27.76 9.84ASA35b GN m mu+pa M3 56.98 40 7 18 27.76 12.24ASA84b GN m pa M3 25.35 41 4.9 12 27.76 13.24CWT19 GN ma mu M1/3 3.53 34.75 0.57 13 27.76 6.99

Table 5.5: Compilation of 40Ar/39Ar age data, conditions of growth, mineralogy and model di↵usion ages. Superscriptnotation: 1 Lithology nomenclature: ab-s = albite schist, c-s = calc-schist (>50:50 carbonate:mica), m-s = mica-schist(<50:50 carbonate:mica), m = mafic, and ma = marble ;2 Metamorphic event associated with mica growth, inferred fromSi c.p.f.u. and fabric relations;3 Inter-grain fusion age range in Ma; 4 Weighted average age in Ma; 5 95% confidenceinterval on average age; 6 number of whole-grain fusion ages grains used in average age calculation; 7 DIFFARGP modeldi↵usion age; 8 Average age minus DIFFARGP model age.

success of 40Ar/39Ar geochronology. The strength of Ar fluid–mineral partitioning means thatgrain boundary fluid networks can e↵ectively be treated as an infinite reservoir for 40Ar di↵usingout of a K-bearing phase. Numerous experimental workers have shown Ar solubility in water to bestrongly dependent on both temperature and salinity (e.g. Weiss, 1970; Crovetto et al., 1982; Smith& Kennedy, 1983)—spanning the full range of metamorphic temperatures and likely salinities, Arsolubility in hydrous fluids is estimated to be betweem 10 and 100 ppm.bar�1 (Kelley, 2002). Arsolubility in K-beaing minerals is less well constrained, largely due to Ar’s propensity to absorbonto mineral defects (e.g. Broadhurst et al., 1992). However, with the advent of in-situ analyticalmethods, estimates of Ar solubilty in a range of di↵erent minerals show that Ar–mineral solubilitiesare �3 orders of magnitude smaller than fluid values: quartz – 3.75 ppb.bar�1 (Roselieb et al.,1997), clinopyroxene – 0.09 ppb.bar�1 (Brooker et al., 1998), olivine – 1 ppb.bar�1 (Brooker et al.,1998), plagioclase – 1.8 ppb.bar�1 (Kelley, 2002) and K-feldspar – 0.7 ppb.bar�1 (Wartho et al.,1999).

There is a dearth of experimental data constraining the solubility of Ar in hydrous minerals usedfor 40Ar/39Ar geochronology. Onstott et al. (1991) undertook preliminary experiments to showthat Ar solubility in biotite is between 3.6–36 ppm.bar�1, however, the accuracy of the results wascompromised by inclusion of an atmospheric Ar component. Nevertheless, natural 40Ar/39Ar dataconfirm a relative solubility order for the micas as muscovite is often observed with lower 40ArE

concentrations than co-genetic biotite (Roddick et al., 1980). Measurements on phlogopite confirmthat Ar solubilities in micas are ⇠1 order of magnitude greater than K-feldpsar (1.8 ppb.bar�1:

154

Roselieb et al., 1999).Using these experimental and natural solubility data to calculate partition coe�cients a↵ords

the opportunity to model mineral-fluid Ar distribution in both open and closed grain boundarynetwork systems (e.g. Kelley, 2002; Camacho et al., 2005).

40Ar

Open system

40Ar40Ar

Closed system

40Ar

40Ar

40Ar

Figure 5.9: Open and closed system mica–fluid 40Ar exchange scenarios. Red lines highlight transport of 40Ar. In the

open system, 40Ar di↵uses across the grain boundary into an intergranular medium of infinite volume with respect to 40Ar

concentration. This allows e�cient removal of 40Ar from in-situ decay of 40K bound in mica; note that porosity is greatly

exaggerated in the sketch. In the closed system, the intergranular medium is either of finite volume (i.e. between 200–1000

A when confined to the grain boundary width (Walther & Orville, 1982)) or anhydrous, thus limiting 40Ar mobility and

generating non-zero 40Ar grain boundary concentrations. This leads to hinderance of 40Ar di↵usion out of the mica grain,

and, under extreme conditions, partitioning of 40Ar back into the grain. Note that significant porosity does not preclude

closed system behaviour if permeability is low/absent—see void space in sketch.

5.7.2 Open system

Figure 5.10 shows calculated 40ArE excess ages for a muscovite in exchange equilibrium with agrain boundary fluid of variable 40ArE concentration. The model uses a range of KAr

Dmuscovite�fluid

values, between 10�2 and 10�7; zero 40ArE excess age is taken as 27.76 Ma—the model di↵usionage pertinent to the Eclogite Zone, Rote Wand and Glockner nappes.

According to solubility data discussed above, KArDmuscovite�H2O

values likely lie within the range10�3–10�4, two to three orders of magnitude greater than experimentally constrained values for

155

KArDK�feldspar�H

2

O(Kelley, 2002), and lower than the values of KAr

Dbiotite�H2

Oobtained by Onstott

et al. (1991) of 10�1–10�2.Single grain 40Ar/39Ar ages for Eclogite Zone samples are between 31.77 Ma (CWT–15) and

58.95 Ma (CWT–13), i.e. between 4–31 Ma in excess of the expected di↵usion age of the Barrovianoverprint. Using KAr

Dmuscovite�H2

O= 10�3, minimum 40ArE muscovite concentrations are reproduced

if the muscovite is in equilibrium with a fluid reservoir of ⇠23 ppm 40ArE , whereas maximum agesrequire grain boundary 40ArE concentrations of ⇠200 ppm. Rote-Wand single grain ages lie inthe range: 29.55–75.71 Ma (CWT–7), which requires a fluid 40ArE concentration range of 9–312ppm. Glockner nappe samples show single grain age variations between 30.09 (ASA 05) and 90.4Ma (ASA–35b), which translates to grain boundary 40ArE concentrations between 12 and 407ppm. Provided that ages are o↵set by �1.18 Ma, to account for the di↵erences in model di↵usionages, the model can be applied to Venediger nappe samples: grain ages between 30.53–69.6 Ma(CWT–17 and ASA–28a) can be reproduced by 40ArE concentrations between 23–326 ppm. It isimportant to note that these 40ArE are only strictly applicable to muscovite and not paragonitedue to the absence of di↵usion data available for Na micas, precluding calculation of a di↵usionage.

Therefore, with KArDmuscovite�H

2

O= 10�3 and grain boundary 40ArE concentrations similar to

typical metamorphic fluids, between⇠20–200 ppm, open system muscovite-fluid Ar exchange ac-counts for the distribution of the majority of single grain ages between 3–30 Ma in excess of theexpected di↵usion age. However, single grain ages >⇠60 Ma (CWT–7, ASA–28a, 35b) requireextreme grain boundary fluid 40ArE concentrations, the oldest of which (ASA 35b), requires 40ArE

concentrations >⇥2 the highest reported fluid inclusion 40ArE concentration (Harrison & Mc-Dougall, 1981). Given that KAr

Dmuscovite�H2

O= 10�3 is likely towards an upper bound, these 40ArE

are treated as minimum estimates.As KAr

Dmuscovite�H2

Ovaries antithetically with temperature, late-stage infiltration of meteoric

waters is often cited as the source of 40ArE influx (Torgersen et al., 1989; Kelley, 2002). However,in order to account for the range of excess ages observed, values of KAr

Dmuscovite�H2

Owould have to

be >10�1, which is outside temperature-dependent variation.The large range of intra-sample, single grain 40Ar/39Ar ages, between 4.78 Ma (CWT–8) and

56.98 Ma (ASA–35b) precludes the conclusion that open system behaviour operated over kilometrelength-scales. Rather, the fact that single grain ages di↵er outside of analytical uncertainty oversample (cm) length-scales, suggests that grain boundary 40ArE concentrations were di↵erent notonly between nappes and sample locations, but that concentration gradients existed also betweengrains.

A pair of samples from the upper-most Eclogite Zone provide further insight into the length-scales 40ArE heterogeneity. Samples ASA–6a and 6b were collected from the same outcrop, ⇠10 macross-strike distance apart; sample ASA–6a shows a spread of single grain ages (n=12) between31.73±1.04 Ma and 47.49±0.71 Ma, whereas sample ASA–6b yields a range of apparent 40Ar/39Arages between 33.02±0.76 Ma and 49.41±0.62 Ma. The spread of grain ages is notably similar. The

156

outcrop lithology comprises a coherent body of intensely foliated and crenulated calc- (ASA–6b)mica-(ASA–6a) schist. These data imply that gradients in grain boundary 40ArE concentrationexisted in both samples, but that e↵ective 40ArE concentrations were broadly comparable across⇠10 m, i.e. apparent open system Ar exchange over the outcrop-scale and yet closed system Arbehaviour on the grain scale. This paradox suggests that the grain boundary fluid originated locallyand did not travel over length-scales greater than inter-grain distances, i.e. µm. Muscovite fromboth samples preserve phengitic cores, which are surrounded by lower-P rim domains. Growth orequilibration of these rims in a grain boundary network concentrated in 40ArE released with fluidsduring decompression explains the data.

Therefore, mineral-fluid 40ArE partitioning calculations show that the apparent 40Ar/39Arages are more readily explained by equilibration with locally (meter-scale) derived 40ArE , thankilometer-scale open system Ar exchange.

100

90

80

70

60

50

40

30

Age

(Ma)

10-1 100 101 102 103 104 105 106

[40ArE] grain boundary fluid (ppm)

KD=1x10-2

KD=1x10-3

KD=1x10-4

KD=1x10-5

KD=1x10-6

KD=1x10-7

Met

eoric

flui

ds

Flui

d in

clus

ions

Figure 5.10: Apparent age of a 27.6 Ma muscovite as a function of fluid-borne 40Ar concentration (40ArE) contoured

for muscovite–fluid Ar partition coe�cients between 10�7–10�2. Age axis scaled such that minimum value corresponds

to the model di↵usion age of 27.76 Ma (Eclogite Zone, Glockner and Rote-Wand nappes). Light shaded area represents

typical concentrations of atmospheric 40Ar found in surface waters (Smith & Kennedy, 1983; Torgersen et al., 1989;

Kelley, 2002), whereas darker shading shows the observed concentration range of 40ArE found in fluid inclusions (Foland,

1979; Harrison & McDougall, 1981; Kelley et al., 1986; Cumbest et al., 1994; Kelley, 2002). Calculations use a muscovite

composition of 10.9 wt.% K2

O, a J value of 1.18⇥10�2, ⇢muscovite = 2.81 g.cc�1, 40K�total = 5.54⇥10�10 a�1 and40Ar� 5.81⇥10�11 a�1.

157

5.7.3 Closed system

In a closed system, the grain boundary network is finite and controlled by rock porosity. Grainboundary 40ArE concentrations will reflect the distribution of K2O throughout the rock, i.e. the ageand quantity of protolith K-bearing phases—if concentrations are significant, 40ArE will partitioninto muscovite.

For closed system dynamics to explain the spread in Tauern 40Ar/39Ar ages, it is criticalto assess the required combinations of protolith 40ArE and rock porosity. Figure 5.11 presentsmuscovite 40ArE age-porosity curves for a range of KAr

Dmuscovite�H2

O. The model assumes muscovite

grew/equilibrated with a grain boundary reservoir reflecting the 40ArE signature of a 110 Mabasalt, containing 0.9 wt.% K2O; ages are in excess of 27.76 Ma – the model di↵usion age of theBarrovian event in the Eclogite Zone, Rote-Wand and Glockner nappes.

Model parameters are tailored for sample CWT–13, a layered eclogite collected from the base ofthe Eclogite Zone containing eclogite-facies phengite grains which yield ages between 36.11–58.95Ma. The distinct garnet-clinopyroxene layering have been interpreted to reflect compositionalheterogeneity in a tu↵-like protolith (Kurz et al., 1998b), pertaining to the Valaisian oceanicrealm, ca. 110 Ma (Bousquet et al., 2002). For a KAr

Dmuscovite�H2

Oof 10�3, model calculations shows

that CWT–13 phengite ages are reproduced after equilibration with a grain boundary reservoircontaining 40Ar produced during ⇠82 million year’s worth of radiogenic in-growth of bulk-rock 40Kand which occupies between ⇠3⇥10�3–7⇥10�4 volume % of the rock. Use of KAr

Dmuscovite�H2

O=10�4

yields porosity estimates of ⇠3⇥10�4–7⇥10�5.Due to the varied mineralogy of pelitic protoliths, the e↵ective duration of radiogenic in-growth

is less well constrained than mafic samples. Eclogite Zone pelites (CWT–8,12, ASA–6a, 6b) showsingle grain phengite age populations which span ⇠31–50 Ma, i.e. ⇠3.2–21.7 Ma in excess of thedi↵usion age. Assuming a typical pelitic bulk-rock K2O content of 3 wt.% and a bulk-protolithage of 300 Ma, consistent with the Eclogite Zone representing a segment of thinned-Variscan crust,model calculations using KAr

Dmuscovite�H2

O=10�3 (broad dashed line in Fig.5.11) show that pelitic

phengite ages require porosities between 1.5⇥10�2–1.5⇥10�1 vol.% - 2–4 orders of magnitudegreater than mafic samples.

Porosities required by mafic samples (e.g. CWT–13) to reproduce the extent of internal Ardiscordance are similar to the range of metamorphic grain-scale porosities measured in exhumedterranes: between 10�6–10�3 vol.% (grey shaded box Fig.5.11; Norton & Knapp, 1977; Skeltonet al., 2000). Therefore, growth/equilibration of phengite under closed system conditions, withoutexternal addition of fluid, accounts for mafic 40Ar/39Ar age spectra. The fact that mafic samplesshow intra-sample, single grain, age di↵erences up to 56.98 Ma (ASA 35b) highlights that Arisotopic gradients existed at the cm-µm scale. This may be due to heterogeneous 40K distributionthroughout the protolith, or, locally heterogeneous fluid availability, which is likely during thegrowth of white mica under HP conditions as aH

2

O will be bu↵ered (Scaillet, 1996). In eithercase, fluid mobility must have been limited to less than intra-grain distances, i.e. <1 cm.

158

The high K2O content (>2 wt.%) of pelitic metasediments necessitates that grain-scale porosi-ties are 2–4 orders of magnitude greater than mafic grain boundary reservoirs—smaller porosityvolumes would generate ages in excess of those observed. As metamorphic grain-scale porosities arethought to be <10�4 volume % (Norton & Knapp, 1977; Skelton, 2011), this shows that inheritedAr must have been lost from the system during the interval between muscovite formation (HP

conditions; ca. 31–38 Ma) and the model di↵usion age (⇠27 Ma).The closed system model for Ar muscovite-fluid exchange elegantly accounts for internal Ar

isotopic discordance in both metasediments and mafic lithologies. Fluid mobility (permeability)must have been several orders of magnitude lower in mafic rocks than metasediments throughoutthe period of muscovite growth/equilibration at HP conditions.

Grain-scale porosity

70

60

50

40

30

20

10

0

Exce

ss 4

0A

r a

ge

(M

a)

10-6

10-5

10-4

10-3

10-2

10-1

100

101

Porosity (vol.%)

KD=1x10-6

KD=1x10-5

KD=1x10-4

KD=1x10-3

KD=1x10-2

KD=1x10-1

KD=1x100

300 Ma, 3 wt.% K2O

80 Ma, 0.5 wt.% K2O

Figure 5.11: Excess 40ArE age in a finite volume grain boundary network, plotted as a function of KArDmuscovite�H2O

and

volume % porosity for a muscovite in equilibrium with a grain boundary reservoir reflecting the signature of a 110 Ma basalt

(i.e. 80 Ma in-growth) with a bulk-rock K2

O content of 0.9 wt.%. Shaded area represents range of typical metamorphic

grain-scale porosities (Norton & Knapp, 1977; Skelton et al., 2000). Coloured lines are for constant KArDmuscovite�H2O

between 10�6–100; fine dashed line represents porosity-age relationship for a muscovite in dynamic equilibrium with an 80

Ma protolith containing 0.5 wt.%K2

O, whereas broad dashed line represents a Variscan sediment protolith (300 Ma and

3 wt.%K2

O). Calculations use a muscovite composition of 10.9 wt.% K2

O, a J value of 1.18⇥10�2, ⇢muscovite = 2.81

g.cc�1, 40K�total = 5.54⇥10�10 a�1 and 40Ar� 5.81⇥10�11 a�1.

159

5.7.4 40ArE as a tracer for metamorphic porosity

Grain-scale porosity reflects rock fluid volume (e.g. Connolly & Thompson, 1989). Open versusclosed system calculations presented above show that grain boundary fluids, and hence porosity, arelikely to have been autochthonous or locally derived. This means that the di↵erence in predictedgrain boundary reservoir volumes between mafic and metapelitic samples should correlate with thedevolatisation history of each rock-type, during the Tauern P–T–t path.

In the extended mafic system, NCKFMASHTO, a typical MORB (SiO2: 53.4, Al2O3: 9.2, CaO:12.4, MgO: 12.9, FeO: 8.29, K2O: 0.23, Na2O: 2.6, TiO2: 1.1, O: 1.5; Rebay et al., 2010) requiresaddition of ⇠25% of the H2O component required to yield saturated conditions in the greenschistfacies (4 kbar, 400�C), to yield H2O in excess conditions at 25 kbar and 575�C. Growth of law-sonite and talc, following decomposition of prasinite phases (albite, chlorite, epidote and actinolite),bu↵er H2O saturation at HP conditions. As lawsonite is not observed in mafic eclogites of theEclogite Zone, this suggests aH

2

O < 1. Peak conditions down-P of the lawsonite+talc stability field(garnet+glaucophane+actinolite+epidote+omphacite+rutile+muscovite+quartz+H2O), yield freeH2O amounts >35% than greenschist values (Fig.5.12b).

Modeling sample TH-680 (section 3.2.2) in the full pelite system, MnNCKFMASHO, shows thata similar P -T path, from 4 kbar, 400�C, to 25 kbar, 575�C generates a ⇠47% (at T> 600�C, thisvalue is above 60%; Fig.5.12a) loss in mineral-bound H2O. This devolatisation occurs primarilythrough the breakdown of chlorite between 21–17 kbar and 500–600�C; chloritoid and epidotecontribute smaller quantities of H2O.

These calulations show that pelites devolatise to a greater extent than mafic lithologies duringburial to HP conditions. Therefore, eclogite facies muscovite in mafic rocks is likely to equilibrateunder closed-system conditions of limited fluid mobility, in which precursor 40Ar, if present inappreciable quantities, can be incorporated into the mica grain. By contrast, the pelitic systemwill generate grain boundary networks orders of magnitudes larger in volume, i.e. towards theopen-system, which will dilute and, depending on permeability, sweep concentrations of inherited40Ar from the system. Using such a model, the extent of 40ArE contamination can be used a proxyfor time-integrated metamorphic connectivity.

5.7.5 Tectono-metamorphic significance of 40Ar/39Ar ages

As all samples contain mica grains which yield 40Ar/39Ar ages > than the model di↵usion ages foreach of the nappes in question, the single grain fusion data cannot be interpreted as cooling ages.Rather, ages represent proxies for the time-integrated e↵ects of fluid mobility, protolith age andK2O content, in addition to deformation.

This contradicts the work of Zimmermann et al. (1994), Ratschbacher et al. (2004) and Kurzet al. (2008), who assign specific tectono-metamorphic significance to 40Ar/39Ar white-mica spectra(section 5.2). Each of the authors acknowledge potential 40ArE contamination, but are unableto quantify its magnitude given the lack of an independent constraint on the timing of mica

160

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growth. Kurz et al. (2008) assign 40Ar/39Ar age di↵erences between 32–38 Ma to variations indeformation intensity within Eclogite Zone mafic eclogites. However, these ages are older than theexpected 27.76 Ma di↵usion age presented above, which alludes to variable 40ArE concentrations.Interestingly, Zimmermann et al. (1994) reports 40Ar/39Ar from low-Si Tauern micas with ages<27 Ma. These ages are consistent with Schieferhulle di↵usion ages and can be interpreted torepresent cooling of the nappe pile following Barrovian metamorphism. Muscovite showing suchyoung 40Ar/39Ar ages were found only as alteration products of earlier HP mica generations andnot as a pure sample. Sample ASA-84b is the only example of similar behaviour in this study’sdataset. Ratschbacher et al. (2004) calculate cooling rates of the Eclogite Zone between 42–56Ma on the basis of step-heating and single-grain 40Ar/39Ar fusion data. These ages are nowre-interpreted to represent significant 40ArE contamination.

Samples from the Rote-Wand nappe show ages which are discernibly younger than samplesfrom both footwall and hanging wall units. As the Rote-Wand nappe is <200m thick in theGroßvenediger region, these young ages may represent more e�cient cycling of fluid through thenappe, therefore reducing 40ArE concentration.

5.7.6 Suitability of 40Ar/39Ar geochronology under HP conditions

The 40Ar/39Ar mica geochronometer has been applied to the majority of Phanerozoic (U)HP

terranes in attempt to constrain the terrane’s passage through Tc on exhumation (e.g. Hubbard& Harrison, 1989; Li et al., 1999; De Sigoyer et al., 2000; Kurz et al., 2008; Warren et al., 2010;Wiederkehr et al., 2009). Commonly, 40Ar/39Ar mica ages from (U)HP rock-units are pervasivelye↵ected by 40ArE , yielding ages ten’s of millions of years older than independent constraints onthe timing of mica growth (Li et al., 1994, e.g.). The ubiquitous presence of 40ArE under (U)HP

conditions shows either that at temperatures above Tc, 40Ar is unable to e�ciently di↵use acrossthe grain boundary, or, during equilibration/growth of the mica grain, 40ArE was present withinthe grain boundary network in su�cient concentrations to partition into the grain. Both conditionscompromise the assumption of open-system Ar exchange.

The degree to which fluid-absent conditions are attained during (U)HP metamorphism is de-pendent on rock composition and permeability. While pelitic metasediments devolatise a significantfraction of their water during burial, growth of lawsonite in mafic rocks consumes any free H2O(section 5.7.4; Sherlock & Kelley, 2002), which contains a molecule of H2O per formula unit.Connectivity of the grain boundary network (permeability) depends on P–T conditions, fluid com-position and strain gradient (Bruce Watson & Brenan, 1987). The experimental work of Holness(1993) and Holness & Graham (1995) shows that pure carbonates are permeable (✓<60�, where✓ is the dihedral angle), to H2O–CO2 fluids at pressures ⇠<4 kbar, whereas quartzites are per-meable under temperatures ⇠<400�C. These observations hold for all values of porosity; however,compaction will still yield a minimum value for porosity (McKenzie, 1984). At values of ✓ greaterthan 60�, porosity will be finite and fluid can escape, aided by deformation. In light of this, per-

162

meability in (U)HP metasediments will be limited. Philippot & Rumble (2000) use stable isotopeheterogeneities in Alpine mafic eclogites to show that fluid fluxes during eclogitisation are as lowas 1.6 cm3.cm�2. This is supported by the findings of Getty & Selverstone (1994), who reportdolomite layers in disequilibrium with respect to both O and C isotopes, separated by ⇠1cm, inTauern banded eclogites—integrated fluid fluxes are ⇠1.65 cm3.cm�2.

Under fluid absent conditions grain boundaries are non-wetted (e.g. Scaillet, 1996). Di↵usionof anions through a free aqueous phase is up to 16 orders of magnitude faster than along dry grainboundaries (dry = 10�24; wet = 10�8 m2.s�1; Rubie, 1986). Therefore, homogenisation of locallyinherited 40ArE gradients will be limited by aH

2

O. Scaillet (1996) measured 40Ar/39Ar ages as highas 320 Ma from Alpine phengites of the Dora Maira eclogites and showed that phengite growthimposed fluid absent conditions, which prevented uniform Ar di↵usion close to the baric peak ofmetamorphism.

Therefore, it is apparent that Ar mica-fluid systems reflect closed-system behaviour under(U)HP conditions. Coupled with the e↵ect of pressure, which increases Ar di↵usivity (Harrisonet al., 2009), this confirms that 40Ar/39Ar white mica ages from (U)HP rocks cannot be interpretedaccording to the classic Dodson closure temperature concept. Accurate interpretation can onlybe achieved with an independent constraint on the timing of mica growth or equilibration, inconjunction with modeling of the expected di↵usion age. A definitive check on open versus closedmica–fluid Ar partitioning can only be achieved by mapping the concentration of 40Ar withindi↵erent minerals in an individual sample (e.g. Sherlock & Kelley, 2002; Warren et al., 2010).Analytical resolution is not yet su�cient to resolve within-grain age di↵erences in 30 Ma micas.

5.8 Chapter Summary

Single grain muscovite and paragonite 40Ar/39Ar ages from both mafic and pelitic lithologies acrossthe Pennine nappe pile yield ages between 23–90 Ma. Textural and chemical evidence show thatthe majority of micas grew or equilibrated close to the peak of HP metamorphism ca.33 Ma.Despite numerical modeling of Ar di↵usion in muscovite showing that Barrovian conditions weresu�cient to erase HP 40Ar histories, mica ages from all nappes and lithologies are older than theU–Pb and Rb–Sr age constraints on the timing of HP metamorphism. This confirms the presenceof ubiquitous 40ArE .

Modeling of open and closed system mica–fluid 40Ar exchange shows that intra-sample 40Ar/39Arage di↵erences are best reproduced by a finite volume grain boundary network with centimeterconnectivity length-scales. Mafic rocks show larger intra-sample age ranges because they do notde-volatise during burial.

This chapter shows that in the absence of an external constraint on the age of mica growth,40Ar/39Ar white mica data from (U)HP rocks cannot be interpreted according to the Dodsonclosure concept. Ages calculated from previous 40Ar/39Ar studies of the Penninic units of theTauern Window are shown to be pervasively a↵ected by 40ArE contamination.

163

Chapter 6

Thermal modeling

The thermal modeling presented in this chapter forms the basis of the published manuscript:Smyeet al. (2011).

6.1 Introduction

Profound questions remain concerning the length- and timescales of mechanisms responsible forthe transfer of heat, and therefore, the formation of metamorphic rock, within the Earth’s crust(e.g. Chamberlain & Rumble, 1988; Baxter & DePaolo, 2000; Camacho et al., 2005; Bjornerud &Austrheim, 2006; Ague & Baxter, 2007). The thermal di↵usivity of continental lithologies deter-mines that conductive heating driven by relaxation of isotherms in regions of over-thickened crustoperates on timescales of the order of tens of millions of years (e.g. England & Thompson, 1984).However, high-precision isotope geochronology has recently shown that in some regional meta-morphic terranes, peak Barrovian conditions are attained within much shorter durations (Oliveret al., 2000; Baxter et al., 2002; Ague & Baxter, 2007). Disparity between predicted and observedtimescales of heating necessitates additional mechanisms to conductive heating. These include:elevated mantle heat flow (Oxburgh & Turcotte, 1974; Stuwe & Sandiford, 1995), increased radio-genic heat production (Jamieson et al., 1998; Huerta et al., 1999; Engi et al., 2001), shear heating(Grujic et al., 1996; Searle et al., 2008; Whittington et al., 2009), magmatism (Lyubetskaya &Ague, 2010), advective transport of heat by fluids (Bickle & McKenzie, 1987) and advection ofheat by tectonics (Burg & Gerya, 2005).

The Tauern Window is a classic locality for modeling the thermal evolution of overthrust ter-rains (Cornelius et al., 1939; Oxburgh & Turcotte, 1974; Bickle et al., 1975; England & Richardson,1977; England, 1978; Richardson & England, 1979; Oxburgh & England, 1980; England & Thomp-son, 1984). Tauern Barrovian metamorphism is understood to have been driven by conductiveheating following northward emplacement of allochthonous Austroalpine units during Alpine oro-genesis (Oxburgh, 1968). Previous conductive heating models call for a period of ca. 30 Mapost-emplacement of the Austroalpine nappes for the European basement to attain peak condi-

164

tions of 520–550�C at 7–8 kbar – thus invoking ages of Tauern eclogite formation > 60 Ma. Clearly,this is incompatible with the U–Th–Pb allanite age presented in chapter 4, which shows that amaximum of 10 Ma passed between eclogite formation and the peak of Barrovian metamorphism.Here, a one-dimensional thermal model (e.g. Oxburgh & Turcotte, 1974; Bickle et al., 1975; Eng-land & Richardson, 1977) is used to investigate thermo-tectonic scenarios which satisfy rapid (<10Ma) attainment of Tauern Barrovian conditions.

6.2 Previous work

The Barrovian metamorphic event of the Tauern Window has provided the impetus behind severalclassic studies on the thermal evolution and energy budget of thickened continental crust—themost pertinent of which are detailed below.

Oxburgh & Turcotte (1974) conducted the first thermal modeling study of the Eastern Alps, inwhich the authors showed that isotherm relaxation associated with emplacement of a 15 kilometer–thick overthrust sheet on to the European basement at 65 Ma, fails to account for Barroviantemperatures (⇠500�C) ca.25–30 Ma. They conclude that Tauern Barrovian metamorphism musthave involved a significant contribution of heat from the mantle, on the order of 75 mW.m�2.The authors also proposed that the high–P , low–T metamorphism observed in the PeripheralSchieferhulle could have been driven by the low temperatures experienced by the footwall closeafter overthrusting.

Bickle et al. (1975) used an analytical approach to model one-dimensional conductive relax-ation assuming a 30 Ma duration separates overthrusting and the peak of Barrovian metamor-phism. Calculated solutions constrain mantle heat flow between 25.1–41.8 mW.m�2 for overthrustsheet thicknesses between 16–26 kilometers and an exponential distribution of radiogenic matter(radiogenic heat flow between 29–50 mW.m�2). Critically, this study illustrated that Barrovianconditions could be attained in 30 Ma without abnormally high mantle heat flow, consistent withpresent day heat flow measurements in the Tauern window. Despite using similar approaches,Oxburgh & Turcotte (1974) and Bickle et al. (1975) came to disparate conclusions, largely becauseBickle et al. (1975) consider a much greater range of parameter space. In particular, Bickle et al.(1975) consider a range of overthrust-sheet thicknesses up to 30 km, whereas Oxburgh & Turcotte(1974) only considered thrust sheets up to 15 kilometers thick.

England (1978) addressed Tauern metamorphism using a numerical model, taking into accountthe temperature dependence of thermal parameters. Conductive heat transfer solutions to themodel are similar to those calculated by Bickle et al. (1975) (15–30 km overthrust sheet thickness;radiogenic heat flow between 20–49 mW.m�2; mantle heat flow between 16–75 mW.m�2). Theauthor also presents calculations which show that for Tauern metamorphism to have been drivenby advection of heat via igneous fluids, the quantities of fluid required are implausible (10 kilometerthick layer to yield a 200�C rise in temperature at 15 kilometers depth)—particularly given thescarcity of field evidence for Tertiary plutonic rocks in the Tauern region.

165

Oxburgh & England (1980) show that surface heat flow measurements collected across theaxial zone of the Eastern Alps vary between 75–90 mW.m�2 and support the notion that the plateinterface was at 7.5±2 kbar (16–29 km) and 550±50�C ca. 30Ma. They claim that the uncertaintyinherent in metamorphic pressure estimates of the plate interface during Barrovian metamorphismis insu�cient to resolve between metamorphism driven by enhanced heat flow beneath a thinthrust sheet (i.e. Oxburgh & Turcotte, 1974), or metamorphism by deeper burial under conditionsof average continental heat flow (i.e. Bickle et al., 1975).

The work of England & Thompson (1984) builds on the aforementioned studies and servesas the foundation for current understanding of the rates of conductive heat transfer in regions ofoverthickened crust. The authors investigate the one-dimensional thermal evolution of overthrustcontinental crust for a wide range of geologically relevant parameters and show that the principalfeatures of P–T–t paths in such tectonic settings comprise: a period of thermal relaxation, duringwhich temperatures rise towards the elevated geotherm supported by the overthickened crustalpile, followed by a period of cooling as the rock crosses cooler isotherms on its passage towards thesurface, driven by erosion. The study predicts that attainment of peak temperatures in overthrustterranes should span ⇠10 Ma or more. Model solutions for a similar set of input parameterscorroborate the predictions of Bickle et al. (1975).

More recently, Stuwe & Sandiford (1995) investigated the combined thermo-mechanical e↵ectsof mantle lithosphere deformation on the metamorphic evolution of the Tauern Window. Theysuggest that conductive relaxation models do not account for changes in the thickness of mantlelithosphere observed, and that Barrovian metamorphism in the Tauern Window could have beendriven by localised attenuation of mantle lithosphere, which would increase the potential energyof the region’s crust and drive lateral extrusion of the Eastern Alps (Ratschbacher et al., 1991b,a).Such a scenario would imply large values for mantle heat flow.

6.3 Aim

The presence of detrital glaucophane in Eastern Alpine flysch sediments as old as 80 Ma (Ober-hauser, 1968), in addition to structural evidence that the Austroalpine Gosau beds underwentsignificant tectonic disruption between 50–80 Ma, led the authors of previous modeling studies(section 6.2) to assume that 30 Ma passed between HP–LT metamorphism at ca. 60 Ma andconductively driven Barrovian (re)crystallisation at ca. 30 Ma.

Critically, the U–Th–Pb allanite geochronology presented in chapter 4 shows that this cannotbe correct and that Tauern eclogite formation occurred much later, between ⇠31–37 Ma, whichis corroborated by multimineral Rb–Sr geochronology (Glodny et al., 2005). This means thata maximum of 10 Ma separates the two episodes of Alpine metamorphism. Furthermore, thisdrastically limits the duration of time available for conductive heating following northward-directedemplacement of the Austroalpine nappes on top of the European margin.

The thermal modeling presented beneath seeks to constrain combinations of thermal parame-

166

0

10

20

30

40

50

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70

80

90

100

Dep

th (k

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0 100 200 300 400 500 600 700 800 900 1000

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England & Thompson, 1984). The model is calculated with an erosion rate (u) of 0.9 mm.a�1, a thermal di↵usivity

() of 31.54 km2.Ma�1, an exponential distribution of radiogenic matter with an e-spacing of 8 km and a surface value

(Az0

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(10.7�C.km�1. Note: a 30 Ma erosional hiatus is required for the basement-top to attain Barrovian conditions. Grey box

represents Tauern Barrovian conditions; black dots mark the passage of a rock unit located at the plate interface in 5 Ma

increments.

ters required to attain Barrovian conditions at the plate interface in 10 Ma following thrust sheetemplacement, and, more specifically, to confirm whether or not the Barrovian event requires anadvective component of heat transfer. As lateral heat gradients are considered to be minor withrespect to vertical gradients produced during overthrusting, the Tauern Window’s thermal evolu-

167

tion is well approximated via a one-dimensional approach. This also permits direct comparisonwith previous, one dimensional thermal modeling studies (section 6.2).

Figure 6.2: Photograph of the southern boundary of the Tauern Window, detailing the plate interface zone. Thrust sense

markers (3-D; Oxburgh, 1968; Bickle & Hawkesworth, 1978; Wallis, 1988, ; section 2.4.6) show northward emplacement

of allocthonous Austroalpine nappes. Note the steeply, south-dipping white-marble band in the foreground,belonging to

the Matrei zone, and the rugged topography of the overthrust Altrkristallin basement complex. Photograph taken looking

south from summit of Sailkopf (3,209 m.s.l.).

6.4 The model

The thermal development of the Tauern Window is modeled by a finite-di↵erence approximationto the heat equation following the approaches of Oxburgh & Turcotte (1974), Bickle et al. (1975)and England (1978) in which the thermal profile following overthrusting is calculated for di↵erentvalues of total heat generation (radiogenic heat production + mantle heat flow). Additionally, thee↵ect of thrust sheet thickness on the evolving thermal profile is considered.

6.4.1 Numerical techniques

In the absence of advective heat sources, the total change in temperature with time at a depthz within a crustal column is controlled by: i. conductive heat transport; ii. mass transport (i.e.erosion); iii. heat generated via radiogenic heating. Numerically, these factors are expressed in the(one-dimensional) conductive heat equation where T is temperature at any time t, is thermaldi↵usivity, u is velocity (� in the case of erosion as z is positive with increasing depth), A isradiogenic heat production per unit volume rock, ⇢ is crustal density and Cp is the specific heatcapacity:

168

@T

@t=

@2T

@z2� u

@T

@z+

A

⇢Cp(6.1)

The equation is solved numerically, using the explicit Euler method of finite di↵erence approx-imation (Carslaw & Jaeger, 1959), over a 150 kilometer-thick profile gridded into nodes at 0.5 kmspacing, corresponding to the lithospheric thickness. Both di↵usivity () and conductivity (K) arekept constant at 1 mm2.s�1 (31.54 km2.Ma�1) and 2.5 W.m.C�1 respectively. Temperature depen-dence of these variables and also specific heat capacity (Cp) are not included as test calculations(Fig. 6.10) show that di↵erences are negligible in the absence of shear heating or magmatism.Heating as a consequence of deformation is not taken into account. The product ⇢.Cp is keptconstant at 2.7⇥1015 J km�3C – corresponding to an average crustal density of 2700 kg.m�3 andheat capacity of 1000 J kg�1C. Explicit stability is maintained by calculating the timestep for eachsimulation pending on the stability coe�cient and di↵usivity. Radiogenic heat production (Az) isapproximated by an exponential distribution throughout the entire lithosphere, in accordance withLachenbruch (1968, 1970), where Az corresponds to the magnitude of radiogenic heat productionat depth z and D is the e-spacing value:

Az = Az0

e(�z/D) (6.2)

A fixed surface temperature for all time-steps of 20�C is assumed for the upper-boundarycondition whereas the lower-boundary condition is determined by a constant temperature gradient(@T/@z) at 150 km depth. The grid was seeded with an initial geotherm calculated using the steady-state heat equation for a single-layer lithosphere with constant surface temperature (Tz

0

=20�C), afixed temperature gradient at infinte depth (P = @T/@z at 1) and an exponential distribution ofradiogenic heat production (Eq.6.2):

Tz = Tz0

+D2A0

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�zD

K+ Pz (6.3)

Instantaneous thrusting is assumed and the geotherm is repeated above and below the thrustdepth to create a saw–tooth geotherm prior to conductive relaxation.

6.4.2 Model constraints

Assuming that heat is transfered only by conduction, the following parameters control the model’sthermal evolution:

1. Initial thermal profile – The initial thermal profile is divided into two units representing theallocthonous Austroalpine complex and lower plate Penninic basement and cover nappes,including the Eclogite Zone. The initial temperature profile of the overthrust sheet is in-significant to the thermal evolution of the system (Bickle et al., 1975) and therefore, identicalequilibrium geotherm relations are used for both hanging– and footwall units (Eq. 6.3). As

169

stated above, the basement–top initial temperature is taken as 20�C—higher values are pre-cluded by the thickness of its cover sequence (1000m Kurz et al., 1998b). Constraints on thethickness of the Austroalpine complex are limited due to internal imbrication and the absenceof palaeo–horizontal markers. However, an estimate of the sheet’s minimum thickness at thetime of overthrusting is obtained from peak pressures experienced within the Zentralgneissof the Venediger nappe, during Barrovian metamorphism: 6–14 kbar equating to 20–45 km(Chapter 3). These thicknesses are slightly greater than assumed by previous studies (15–30km: Bickle et al., 1975; England, 1978; England & Thompson, 1984)), reflecting advances inthermodynamic data used in calculating metamorphic pressures.

2. Di↵usivity (), Conductivity (K) – These parameters control the rate of conductive heattransfer. See section 6.6 for a detailed consideration of the e↵ect of temperature-dependentthermal parameters on the rate of conductive heating.

3. Radiogenic heat production (A) – An exponential distribution of radiogenic material (Eq.6.2) within both hanging and footwall units fits the observations made by Birch et al. (1968),Lachenbruch (1968), Roy et al. (1968) and Lachenbruch (1970) that surface heat flow ap-pears proportional to heat production despite variable erosion rates. The distribution andmagnitude of heat produced depend on the characteristic drop-o↵ length (D) and the surfacevalue (Az0) adopted. Despite the fact that the distribution of radiogenic elements will be sig-nificantly di↵erent between the Austroalpine basement nappes and the Venediger nappe, thelatter is considerably more important than the former given that conductive heating is focusedon the plate interface. As a result, we use identical heat production distributions, based onthe geology of the Venediger nappe, in both upper and lower plates. The East Alpine Penninicbasement (Venediger nappe) represents an exhumed, crystalline portion of the Variscan oro-genic belt and comprises a Late-Proterozoic–Early Paleozoic magmatic–sedimentary succes-sion (Habach-Storz group: Kurz et al., 1998b)), which is intruded by a Variscan calc–alkalinegranitoid association—the Zentralgneiss. Radiogenic elements are concentrated within theZentralgneiss and are consistent with D = 8 km and Az0 = 5.02 µW.m�3 (Hawkesworth,1974a). A maximum bound to the amount of heat production possible (Az0.D) from withinthe basement is calculated assuming that the entire lower-plate is composed of Zentralgneiss– 50.16 mW.m�2 (1.19 HFU)(Bickle et al., 1975).

4. Mantle heat flow (qm) – It is important to note that the contribution of heat from the mantle isthe single most important parameter a↵ecting the thermal evolution of the system (England,1978). Convective transport of radiogenic heat to the base of the conductive lithosphereis generally between 20–40 mW.m�2 (0.47–0.96 HFU) in regions of stable continental crust(Sclater et al., 1980, 1981) and as high as 100 mW.m�2 (2.39 HFU) in areas of rifting (Lysak,1992). There are few available constraints on the magnitude of mantle heat flow during thrustsheet emplacement: Helium isotopic ratios (3He/4He) within East Alpine groundwaters show

170

a virtual absence of mantle–derived 3He (Marty et al., 1992) and there is a virtual absence ofAlpine igneous activity proximal to the Tauern Window. Given the lack of constraint, mantleheat flow (qm) is allowed to vary between 20–100 mW.m�2 (0.47–2.39 HFU), encompassingvalues characteristic of stable continental crust to rifting margins (Sclater et al., 1980, 1981).

5. Erosion (u) – Erosion facilitates exhumation of metamorphic rock, transporting such rockacross isotherms, and depends on the thermal architecture of the crust. In the simple one-dimensional overthrust model, erosion concurrently acts to counter conductive heating as itdrives hot rock through cooler, shallower isotherms towards surface exposure (England &Richardson, 1977). A central tenet of the classical studies of Oxburgh & Turcotte (1974),Bickle et al. (1975), England (1978) and England & Thompson (1984) is that their modelscall on a delay between the assembly of the orogen and the inception of erosion. This isbased on overthrusting being complete by ⇠65±10 Ma, whilst little molasse sedimentationoccurred prior to the Oligocene, thus leaving an erosional hiatus of 35 Ma. However, the34.2±3.6 Ma U–Pb allanite age shows that overthrusting occurred much later, coincidentwith the baric peak in Pennine units. Therefore, there is no justification for an erosionalhiatus following thrust sheet emplacement. The models used here involve a standard 0.9mm.a�1 (0.9 km.Ma�1) erosion rate Bickle et al. (1975), active immediately post-thrusting.

6. Strain heating – The ability of rocks to generate heat during metamorphic deformation de-pends on their strength in the ductile regime. In the Tauern Window, the plate-interfacezone is dominated by carbonate-rich lithologies such as aragonite–quartz schists. Assumingductile failure of schists within the shear zone, experimental data are combined with a generalpower law expression for ductile deformation (Eq. 6.4; Nabelek et al., 2010) to examine thevolumetric heating potential of plate-interface schists.

⌧ = (✏

B)

1

n · eQ

nRT (6.4)

Symbols are as follows: ⌧ = deviatoric stress (MPa) required to deform the rock; ✏ = strainrate (s�1); Q = activation energy (J·mol�1); n = flow-law constant, R = gas constant; T =absolute temperature (K); B = pre-exponential factor. Figure 6.3 shows the shear strength(Eq.6.4) of power law rocks and rheologies pertinent to the Tauern Window, as a functionof temperature, calculated using a strain rate of 1⇥10�13 (✏ = v/w, where v is the rate oftranslation of the shear zone (10 cm·a�1), and w is the width of the zone (5 km)). Dashedlines represent the brittle failure strength at the base of a 30 km thick thrust sheet inclinedat 30�, with an average ⇢ = 2.5 g·cc�1 and pore pressures (�) of 0.4 and 0.8 dynes, calculatedusing the following expression (Hsu, 1969):

⌧ = ⌧0 + (1� �) · ⇢ · g · z · tan✓ (6.5)

171

where, ⇢ is the average density (g·cc�1) of the overthrust sheet, g is gravitational acceleration(m· s�2) , z is the thrust sheet thickness (m) and ✓ is the thrust angle (in �).

Both aragonite and quartz (lines 3 and 2: Fig. 6.3) are relatively strong at low tempera-tures (<300�C), but lose their strength towards higher temperatures. This is an importantobservation as aragonite–schists present in the plate-interface zone will readily deform above⇠200�C at the base of a 30 km thrust sheet, emplaced at 10 cm·a�1. The experimental dataof Shea & Kronenberg (1992)show that mica–schists (line 3, Fig.6.3) are weak at low temper-atures, but retain their strength at high temperatures. Clinopyroxenite (line 4) is includedto illustrate the behaviour of lithologies stronger than quartz; its strength is insensitive totemperature between 0 and 800�C.

Collectively, these calculations show that strain heating of mixed aragonite–quartz schistswill be negligible above ⇠300�C at high pore pressures (�0.8), and above ⇠200�C for lowervalues (0.4). Lower pore fluid pressures move the brittle–ductile transition to greater depths(Brace & Kohlstedt, 1980). Accordingly, the thermal model does not take into account thee↵ect of strain heating, in line with Bickle et al. (1975).

7. Magmatism – Evidence for Tertiary magmatism in the Eastern Alps is limited to the Rieser-ferner tonalite batholith (Borsi et al., 1978) and a suite of unfoliated, alkali–basalt dykes(Deutsch, 1984b), both exposed within the Austroalpine basement bordered to the north bythe Tauern Window and to the south by the Defereggen–Antholz–Vals (DAV) Line (Sander,1925)—a strand of the Periadriatic fault system.

Rocks of the Rieserferner Pluton (⇠30⇥7 km) range from dioritic to granitic in compo-sition and show structural evidence for two deformation events, the second of which wascontemporaneous with mylonitization along the DAV shear zone (Muller et al., 2000). U–Pbgeochronology on plutonic allanite (Romer & Siegesmund, 2003) of the central RieserfernerPluton yields an age of 32.4±0.4 Ma, which is interpreted as the age of magma crysallisation.This age range is comparable to that obtained from Rb–Sr analysis of white micas from aDAV mylonite (Muller et al., 2000). These data have been interpreted to suggest that thepluton was emplaced along the DAV shear zone during north–south directed shortening andeast–west directed extension (Steenken et al., 2000; Wagner et al., 2006).

Mafic dykes in the Spittal region of the Eastern Alps were first described by Teller (1889) andcan be grouped into three suites according to major, trace and rare earth element mineralchemistry (Deutsch, 1984b): i. calc–alkaline dykes; ii. shoshonites (K2O/Na2O > 1); iii.alkali–basalts (K2O/Na2O > 1). The dykes are concentrated in an area ⇠100 km2 directlyto the south–east of the Tauern Window and range between 2–10 m in thickness. K–Ar data(Deutsch, 1984b) obtained from each of the dyke groups show that the alkali–basalt dykesbelong to a single magmatic event at ca. 30 Ma, whereas rocks of the shoshonitic intrusivesuite intruded between 24–30 Ma ago. K–Ar amphibole data from two calc–alkaline dykes

172

0 100 200 300 400 500 600 700 80010−1

100

101

102

103

104

105

106

107

T (°C)

ԑ = 1*10-13 s-1

τ (M

Pa)

! = 0.4

! = 0.8

[1]

[2]

[3]

[4]

[5]

h = 3 km

h = 5 km

Figure 6.3: Shear strength (⌧) plotted as a function of temperature for pertinent minerals and lithologies at a strain

rate of 1⇥10�13 s�1: [1] Aragonite (Rybacki et al., 2003); [2] Quartz (Rutter & Brodie, 2004); [3] Mica schist (Shea

& Kronenberg, 1992); [4] Clinopyroxenite (Kirby & Kronenberg, 1984); [5] Eclogite (Jin et al., 2001). Fine-dash lines

represent values of shear stress at the base of a 30 km thick thrust sheet inclined at 30�, using pore pressures (�) of 0.4

and 0.8, after Hsu (1969). Horizontal shading represents shear strengths in a model subduction channel calculated using

equation 6 of England & Holland (1979), for a subduction angle of 45�, a subduction rate of ⇠10 cm.yr�1, a density

contrast (slab-mantle) of 300 kg.m�3 and variably subduction channel thickness between 3–5 km. Vertical shading

represents range of peak-T experienced by the Eclogite Zone during HP metamorphism.

show a spread in ages between 26–40 Ma.

Collectively, these data suggest that there was a short lived (⇠5 Ma) period of magmatismat about 30 Ma, which was focused in hanging wall units, during the development of a majortranspressional shear zone above the subducted Valaisian slab.

However, evidence for Alpine magmatic activity is notably restricted to regions close to theDAV, showing that the lineament probably acted as a conduit for melt fractions generatedduring slab devolatisation. There are no plutonic rocks of Alpine age exposed within theTauern Window. Recent gravimetric measurements (Zanolla et al., 2006) across the TauernWindow show that the crystalline antiformal-spine to the Window defines a Bouger anomalylow (⇠-120⇥10�5m·s�2) relative to external regions of the orogen, which precludes the pres-ence of a deep-seated, high-density, magmatic complex (Ebbing et al., 2001, 2006). Giventhese constraints, advection of heat by igneous fluids is not considered as significant to thethermal evolution of the Tauern Window.

173

Combinations of the critical parameters discussed above must satisfy the following model cri-teria:

1. Metamorphic conditions – Peak temperature of Tauern Barrovian metamorphism is wellconstrained by conventional thermodynamic methods to between 500–570�C (Bickle, 1973;Droop, 1985; Dachs, 1990; Dachs & Proyer, 2001; Hoschek, 2001) within the south–centralTauern. Variations in temperature estimates from the Schieferhulle nappes are generallywithin uncertainty of each other. Pressure estimates are less well constrained due to fewerrelevant equilibria and uncertainties in activity–composition relations. Average P -T calcula-tions (thermocalc version 3.33, Holland & Powell, 1998) on a metatonalite (TH–519—seeREF) collected from the upper–most levels of the Zentralgneiss yield an average pressure of13.1±2.1 kilobars at 601±38�C (�fit = 1.66). Although loosely constrained, the estimate iscorroborated by pressures calculated from overprinted calc–schists of the Eclogite Zone, 2–3kilometres structurally above the Zentralgneiss, (Hoschek, 2001) between 9–10.2 kbar.

2. Time between thrust sheet emplacement and Barrovian conditions –The peak of Barrovianmetamorphism occurred between 27–32 Ma (Section 6.2 Inger & Cli↵, 1997; Glodny et al.,2005; Gleißner et al., 2007), however, the timing of emplacement of the Austroalpine nappes isless certain. Previously, Oxburgh & Turcotte (1974), Bickle et al. (1975) and England (1978)assumed overthrusting was complete by 65±10 Ma. This assumption is based on tectonicactivity within the Austroalpine Gosau Beds between 80–50 Ma and the fact that depositionof detrital high-pressure minerals, thought to originate from exhumed Penninic lithologies, inthe northern Flysch, was replaced by deposition of minerals characteristic to the Austroalpineunits during the Campian (Oberhauser, 1968). More recently, Rb–Sr deformation dating ofBachmann et al. (2009) suggests that subduction-related deformation ceased at ca. 50 Mawithin the Engadine Window region. This conclusion is supported by Ar40–Ar39 dating at theplate-interface close to the Tauern Window’s eastern boundary(Liu et al., 2001). However,given the Eclogite Zone’s more distal protolith environment, the allanite U–Pb geochronologyimplies that the European basement was overthrust syn- or post-eclogitization, between 27–37.8 Ma. This is supported by Paleocene–Eocene radiolaria found in carpholite-bearingmetasediments of the Central Alps (Bousquet et al., 2002) and also Rb–Sr data of Wallis(1988) from Alpine deformational fabrics at the base of the Austroalpine basement. Together,these data suggest that subduction of the Valaisian oceanic realm, beneath the Austroalpinewedge, was still active into at least the mid-Eocene. Therefore, model solutions must accountfor heating of the upper-structural levels (⇠6 km) of the Penninic footwall within 10 Ma ofthrust sheet emplacement.

Consequently, values for overthrust sheet thickness, radiogenic heat production and mantle heatflow in which conditions of 9–13 kbar and 550�C are attained in the upper 6 km of the lower platewithin 10 Ma of thrust sheet emplacement are considered in the following modeling calculations.

174

6.5 Solutions

6.5.1 Overthrust sheet thickness

Figure 6.4 shows calculated Depth–T–t curves for the basement-top for a range of overthrust sheetthicknesses between 20–60 km and a constant mantle heat flow of 27 m.Wm�2 (0.64 HFU). TheP–T evolution of the upper-basement is characterized by an initial phase of rapid heating (2Ma), driven by equilibration of temperatures across the sawtooth-shaped temperature depressionof the plate interface. Subsequent heating is exponentially slower until the basement is exhumedvia erosion through colder isotherms. The basement-top only attains Tauern temperatures within10 Ma of overthrusting if the thrust sheet is greater than ⇠55 km thick. Such a scenario wouldrequire peak basement pressures of 14–16 kbar. Peak temperatures for overthrust sheets of 55 and60 km thickness of 548 �C at 11.2 kbar and 579 �C at 11 kbar respectively, are reached between20–25 Ma post thrusting if rapid exhumation begins after 5 Ma. Interestingly, emplacement of a30 km thick thrust sheet – the preferred value of Bickle et al. (1975), heats the upper-basementto a peak temperature of 339�C in 13 Ma – 200 �C short of Tauern conditions. For thrust sheetsgreater than 27 km thickness, erosion must operate more rapidly than 0.9 mm.a�1.

6.5.2 Crustal heat production and mantle heat flow

The temperature attained by the basement-top following 10 Ma conductive heating under a 30 kmthick overthrust sheet is plotted against total heat contribution (A0.D) from the basement (Fig.6.5).Erosion is not active and mantle heat flow varies between 0.2–2.0 HFU (8.36–83.68 m.Wm�2).Saliently, the calculations show that for Tauern radiogenic heat production values 0.55–1.2 HFU(Hawkesworth, 1974a,b), Barrovian conditions will only be attained within 10 Ma if mantle heatflow is > 62.76–37.66 m.Wm�2 (1.5–0.9 HFU) respectively, at the time of emplacement. Suchvalues of heat flow into the lithosphere correspond to conductive mantle temperature gradients of25.10–15.06 �C.km�1, implying lithospheric thinning and asthenospheric upwelling (Sclater et al.,1980). However, Helium isotopic ratios (3He/4He) within East Alpine groundwaters show a virtualabsence of mantle-derived 3He (Marty et al., 1992) and there is a notable lack of extensive Alpinemagmatism within the Eastern Alps, both of which are inconsistent with elevated mantle heat flows.In the absence of erosion, these results provide lower-bound estimates as erosion will shorten thetime interval over which conductive heating is operative. Erosion would displace the constantmantle heat flow curves of Fig. 6.5 towards higher temperatures.

6.5.3 Alternative solutions

6.5.3.1 Tectonic emplacement

Assuming that the Austroalpine nappes were emplaced at ca.65 Ma, we now consider the scenarioin which the Peripheral Schieferhulle units were tectonically inserted into a heated crustal pile

175

and

ky

sill

0 Ma

5 Ma

10 Ma

15 Ma

20 Ma

Temperature (ºC)

Dep

th (k

m)

60

50

40

30

20

10

0

0 100 200 300 400 500 600

25 Ma

2

4

6

8

10

12

14

16

18

Pressure (kbar)

30 Ma

35 Ma

Figure 6.4: Numerically calculated Depth–T–t curves for sample rock units located at the plate interface for variable

thicknesses of overthrust sheet. Circles represent 5 Ma increments after thrusting. Mantle heat flow constant at 27

mW.m�2 (0.95 HFU), crustal heat production is 5.0 µW.m�3 (11.95 HFU) and erosion rate is 0.9 mm.a�1. Grey box

constrains metamorphic conditions at the peak of Alpine Barrovian metamorphism. Red trajectories highlight thicknesses

of thrust sheet which satisfy attainment of Barrovian conditions in <10 Ma.

at ca. 35 Ma. This end-member tectonic scenario accounts for the lack of petrological evidencethat the basement experienced blueschist-facies metamorphism prior to Barrovian conditions, inaddition to accounting for the Barrovian metamorphism by conductive heating alone. Figure 6.6shows the thermal evolution of a crustal pile in which a 4 km thick slab is inserted at plate interfacedepths, into a geotherm produced after 30 Ma of conductive heating beneath a 30 km thick thrustsheet. Initial temperature of the slab is taken as a constant 400�C, corresponding to temperaturesof the blueschist-facies event (Holland & Richardson, 1979). All other thermal parameters are asin Fig. 6.1 except that erosion is not considered due to the small timescales investigated. Giventhat the time constant (⌧ = a2/(⇡2)) of a 4 km slab under such conditions is 0.16 Ma, the slab

176

0.2

0.4

0.6

0.8

1

1.2

1.4

1.6

1.8

200 300 400 500 600 700 800

T after 10 Ma. (˚C)

A 0.D

(H

.F.U

)

Not applicable

Not applicable

1.8

2.0

1.6

1.4

1.2

1.0

0.8

0.6

0.4

0.2

H.F

.U.

Figure 6.5: Temperature attained after 10 Ma conductive heating of the basement after emplacement of a 30 km thick

thrust sheet plotted against total heat generation within the basement complex. The plot is contoured for solutions of

constant mantle heat flow (HFU); red shaded band delineates peak Tauern metamorphic temperatures; stippled region

corresponds to the probable range in A0

·D for the Zentralgneiss, of 0.55–1.2 HFU (Hawkesworth, 1974a,b). For simplicity

these solutions omit the e↵ect of erosion. Thermal properties of the thrust sheet are identical to those of the basement.

attains Barrovian conditions rapidly, within ca. 1.2 Ma. However, in order to reconcile this witheclogitization ages between 28–38 Ma, this scenario requires subduction to be active for 30 Mapost-emplacement of the Austroapline nappes at ca. 65 Ma.

6.5.3.2 Syn-thrusting heating

Oxburgh (1968) states that roughly 100 kilometers of displacement has occurred between thePenninic nappes of the southern Tauern window and the overlying Austroalpine during Alpineconvergence. If the maximum duration of thrusting is 10 Ma (38–28 Ma), then thrust sheet em-placement must have occurred at rates greater than 1 cm.a�1, akin to plate convergence rates.At high rates of thrusting the upper-basement will experience rapid heating due to the constantreplenishment of heat by the allochthonous sheet – e↵ectively maintaining a constant thermal gra-dient within the upper plate, analogous to the action of a hot iron. However, for this scenario tobe plausible, the basement must be located proximal to the inception of thrusting so as to avoidconductive cooling of the thrust sheet. Figure 6.7 shows the thermal evolution of a crustal pile un-dergoing syn-thrusting heating for a thrusting rate of 1 cm.a�1. It is clear to see that syn-thrusting

177

0

10

20

30

40

50

60

70

80

90

1000 200 400 600 800

Temperature (˚C)

Dep

th (k

m)

4 Ma post emplacement0.4 Ma time steps

0.4 Ma 4 Ma0 Ma

400 500 600300Temperature (˚C)

5

10

15

20

25

Pres

sure

(kba

r)

tectonic emplacement

Figure 6.6: One-dimensional conductive relaxation profile of a 4 km thick slab collectively representing the Glockner

nappe and the Eclogite Zone, inserted into a crustal pile with a geotherm representing 30 Ma post-thrusting (doubling

of thickness) conductive heating. The model represents the Eclogite Zone’s insertion into a blueschist facies nappe stack

and subsequent heating and exhumation to the upper-greenschist facies, ⇠550 �C (highlighted on P–T path of insert).

Time steps are 0.4 Ma duration.

heating would form inverted isotherms within the footwall during early stages of thrusting – asobserved in the Himalayan mountain belt. With subsequent conductive relaxation and radiogenicheat production, footwall isotherms attain similar values to the base of the overthrust sheet within10 Ma.

178

0

10

20

30

40

50

60

70

80

90

1000 100 200 300 400 500 600 700 800 900

Temperature (°C)

Dep

th (k

m)

100

80

60

40

20

00 10 20 30 40 50 60

Thru

stin

g ra

te -

r (km

.Ma-1

)

Depth from plate interface - Lper (km)

Displacement = 100 kmκ = 31.54 km2.Ma-1

Lper�  =�  √(κ.τper)

Exposed basement

0 Ma 2 4 6 810

Figure 6.7: The ’Hot Iron’ scenario. One-dimensional conductive heating of a basement complex with qm = 27 mW.m�2,

Az0 = 5.0 µW.m�3 and D = 8 km continually overthrust by a 35 km thick thrust sheet; erosion is neglected; 100 km

displacement occurs at 10 km.Ma�1 (1 cm.a�1). Horizontal distance from inception of thrusting (Lbase) 39 km. Grey

box reflects Tauern metamorphic conditions. Inset plot shows the depth a thermal perturbation (i.e. a sawtooth geotherm

of an overthrust sheet) would penetrate against rate of thrusting over 100 km displacement. Graph calculated from thermal

time constant relations: Lper =p

.⌧per where Lper = depth of penetration, = di↵usivity and ⌧per is the time over

which the temperature pertubation exists (i.e. duration of thrusting). Grey region delimits the range of basement depths

exposed within the Tauern Window.

6.6 Discussion

A maximum interval of 10 Ma separating thrust sheet emplacement and Barrovian metamorphismin the Eastern Alps raises profound questions over the length scales of heating responsible forregional metamorphism. The classic conductive relaxation model shows that the length scale ofconductive heating required to attain peak conditions in <10 Ma is incompatible with Barro-vian pressure estimates and detrital sediment loads (see beneath). Elevated mantle heat flow

179

explains the rapid heating, but lacks supporting geochemical evidence. Tectonic insertion of theSchieferhulle nappes along the plate interface favors rapid (<2 Ma) heating from blueschist toBarrovian metamorphic conditions. However, this scenario requires a mechanism to exhume theEclogite Zone from 70–80 km to ⇠30 km depth at plate tectonic rates of 1–6 cm.a�1, between ca.38 Ma and before Barrovian metamorphism ca. 27–32 Ma. Similarly, concomitant thrusting andheating can explain the <10 Ma timescale of heating, but demand temperatures at the base of theoverthrust sheet close to those of peak metamorphism and also a pulse in detrital sediment supplyto circum-East-Alpine basins.

For Tauern metamorphism to be driven by conductive relaxation of perturbed isotherms re-quires overthrust sheet thicknesses >55 km, which translates to plate-interface pressures between 14and 16 kbar (section 6.5.1). Such pressures are approximately twice the calculated value (7–8 kbar)of the Barrovian overprint within the Glockner nappe (Droop, 1985; Zimmermann et al., 1994).Although pressures calculated from the Zentralgneiss are similar (⇠ 13 kbar) to those required forthick-skinned thrusting, temperatures are also higher, around 600�C, consistent with the presenceof a ⇠20 kilometer overburden (i.e. thrust sheet thicknesses of ⇠30 km). Furthermore, a 55 kmthick thrust sheet would supply vast quantities of detrital sediment to circum-East-Alpine basins.Figure 6.8 plots the total detrital sediment mass derived from a range of thrust sheet thicknessesand Barrovian pressures, using an average crustal density of 2700 kg.m�3, an overthrust area of1.8⇥104 km2 (Oxburgh, 1968). Overthrusting of allocthons >55 km thick would supply >2.6⇥1018

kg of detrital sediment to circum-East Alpine basins - >1.5⇥1018 kg greater than the mass of theEast Alpine, Oligocene–present day, detrital sediment load estimated by England (1981) (1.1⇥1018

kg). Clearly, such length scales of conductive relaxation are too large to be considered as a viablesolution to the Tauern Window’s thermal evolution.

In the absence of an advective component to heating, it is of interest to consider the e↵ect oftemperature-dependent thermal parameters on the Tauern Window’s thermal evolution. Thermaldi↵usivity (), conductivity (k), specific heat capacity (Cp) and density (⇢) are related by thefollowing expression:

k = · ⇢ · Cp (6.6)

Specific heat capacity has a positive temperature di↵erential (@Cp/@T ), whereas thermal di↵u-sivity and conductivity both have negative temperature derivatives (@/@T , @k/@T ). Whilst thetemperature dependence of the Cp of crustal rocks is well constrained (Clark, 1966; Fei, 1995), onlyrecently has the strength of the temperature dependence of and k, been measured accurately(±2%: Whittington et al., 2009). Within orogenic crust these temperature dependences can pro-mote heat retention and thermal insulation in the vicinity of a heat source, such as shear heatingand magmatism (Whittington et al., 2009; Nabelek et al., 2010). Laser-flash analysis of represen-

180

0 0.5 1 1.5 2 2.5 310

20

30

40

50

60

Sediment mass (x 1018 kg)

Thru

st s

heet

thic

knes

s (k

m)

5

10

15

20

Plate interface pressure (kbar)

Eng

land

(198

1)

Kuh

lem

ann

(200

0)

Figure 6.8: East Alpine detrital sediment load (kg) plotted against overthrust sheet thickness (km) and Altkristallin–

Pennine plate-interface pressure (kbar) assuming an average crustal density of 2700 kg.m�3. Orange shaded area represents

the likely range of plate interface pressures (i.e. Zimmermann et al., 1994); blue shaded region highlights values of thrust

sheet thickness which yield Barrovian conditions within 10 Ma; grey, vertical, shading shows the East Alpine sediment

budget masses as calculated by England (1981) and Kuhlemann (2000a).

tative crustal lithologies (garnet schist, leucogranite and a rhyolitic tu↵) presented in Whittingtonet al. (2009), suggest a 50% decrease in thermal conductivity and a 75% decrease in thermal dif-fusivity of the average crust between 0–1000�C (Fig. 6.9). Both parameters decrease rapidly withincreasing temperature beneath the ↵–� quartz transition, above which, values asymptotically ap-proach a higher temperature limit. Di↵usivity values measured by Nabelek et al. (2010) from twoBohemian massif granulites, are slightly lower than those obtained by Whittington et al. (2009),and are plotted on Fig. 6.9 to show the likely variation in di↵usivity with temperature caused byvariation in crustal lithology.

To consider the e↵ect of temperature-dependent thermal parameters on rates of conductiverelaxation, the heat equation (eq. 6.1) is modified:

@T

@t=

1⇢Cp

· @

@z

✓Cp

@T

@z

◆� u

@T

@z+

A

⇢Cp(6.7)

Figure 6.10 plots T–t curves calculated, for both constant (⇠ 1mm2.s�1) and temperature-dependent thermal parameters, for a sample node located at the plate interface, beneath a 30

181

0 200 400 600 800 1000 1200 14000.2

0.4

0.6

0.8

1

1.2

1.4

1.6

1.8

2

2.2

Temperature (˚C)

κ (m

m2 .s-1

)

α−β

qua

rtz tr

ansi

tion

Nabelek et al. 2010

Whittington et al. 2009

200 400 600 800 1000 1200 14000

2

4

0.5

1

1.5

0

Cp (J.g

-1.K-1)k

(W.m

-1.C

-1) Cp

k

Temperature (˚C)

Figure 6.9: Temperature dependence of thermal di↵usivity using the functions of Whittington et al. (2009) and Nabelek

et al. (2010). Inset shows dependence of Cp and k using Whittington et al. (2009)’s data. Relations assume an average

upper-crustal density of 2700 kg.m�3.

km thick thrust sheet (i.e. conductive reference model - Fig. 6.1), with erosion and radiogenicheating omitted. Importantly, the curves show that, in the absence of an internal heat source, thee↵ect of temperature-dependent di↵usivity, conductivity and heat capacity on conductive heatingis to slow the heating rate (14.7�C.Ma�1 versus 16.2�C.Ma�1 between 0–10 Ma for temperature-dependent and constant models respectively) and to lower peak temperatures attained, relative toa lithosphere with constant values (T at 30 Ma: 175�C for T -dep. and 210�C for constant ).

Therefore, in the absence of an advective contribution to heating, the temperature dependencyof thermal parameters serves to increase the time in which peak Barrovian conditions are attainedin the Tauern Window. In terranes where syn-metamorphic magmatism and shear heating areactive, temperature-dependent thermal parameters need to be considered.

Mechanisms which account for the plate velocity exhumation of the Eclogite Zone to mid-crustaldepths are discussed in section 7.1; details relevant to the tectonic insertion scenario are discussedbeneath.

182

0 10 20 30 40 500

50

100

150

200

250

Time (Ma)

Tem

pera

ture

(°C

)

T dependent

1 mm2.s-1

Figure 6.10: Calculated conductive T–t paths for constant (1 mm2.s�1) and temperature-dependent (average crustal

composition (Fig. 6.9): Whittington et al., 2009) thermal di↵usivities following a reference node located at the plate

interface beneath a 30 km thick overthrust sheet; erosion and radiogenic heating are omitted.

Detachment of the Eclogite Zone from the down-going slab was likely triggered by the strongtemperature dependence of aragonite shear strength. Following detachment, rapid exhumation ofthe Eclogite Zone could plausibly have been driven by either a buoyancy contrast between the sub-ducted crustal material and surrounding mantle (�⇢ = 300 kg.m�3: England & Holland, 1979),or by combination of the nappe’s low e↵ective viscosity and external forces during convergence(e.g. Cloos, 1982). During exhumation deformation was partitioned into weak calc-schist chan-nels, which preserve a thrust-sense motion at the footwall contact and a transpressive sense ofmotion at the hanging wall contact. However, despite plausible exhumation mechanics, insertionof the Eclogite Zone as an extrusion wedge into a heated crustal pile requires that continentalcollision occurred ⇠30 Ma prior to eclogite formation between 38–28 Ma, e↵ectively refrigeratingsubducted oceanic crust for up to 30 Ma. This is highly unlikely given that temperatures of theblueschist overprint (200–500�C) are indicative of an active subduction regime between 38–28 Ma.Furthermore, shear-sense indicators along the plate interface are strongly thrust-sense, implyinglate-stage, northward emplacement of the Austroalpine nappes nappes as a tectonic ‘lid’ (Oxburgh,1968; Bickle, 1973; Wallis, 1988; Wallis & Behrmann, 1996; Liu et al., 2001; Bachmann et al., 2009).Emplacement of the Eclogite Zone and Schieferhulle nappes along the plate-interface infers thrust-sense and normal-sense contacts along the nappe’s footwall and hanging wall contacts respectively.In light of these problems with the mechanism of tectonic insertion it is not considered a feasibleexplanation to the problem.

183

Concomitant heating and thrusting are often overlooked in thermal modeling studies. How-ever, the calculations presented above show that they provide the impetus for rapid heating offootwall units close to the inception of thrusting ( r.⌧ , where r = thrusting rate and ⌧ = ther-mal time constant of the thrust sheet; Fig. 6.7). U–Th–Pb allanite geochronology constrainsaverage thrusting rates for the 100 km of displacement (Oxburgh, 1968) between upper and lowerplates to 10 km.Ma�1 – similar to rates calculated by Dewey et al. (1989), Rosenbaum et al.(2002) and Schmid et al. (2004b) via plate tectonic, magnetic isochron and palinspastic restora-tion respectively (Fig. 6.11). Therefore a range of thrusting rates between 10–100 km.Ma�1 isrepresentative of Eocene-Oligocene convergence within the Eastern Alps. The model presentedabove shows that syn-thrusting heating initially a↵ects the upper-most levels of the basement,forming inverted isotherms which are subsequently smoothed by conductive heating. Figure 6.7shows that syn-thrusting heating beneath a 30 km thick thrust sheet displaced at rates between10–100 km.Ma�1 could feasibly penetrate the upper 20 km of basement – consistent with Barrovianpressure estimates between 9–13 kbar calculated from basement rocks.

Using the detrital sediment loads presented in Fig.6.8, overthrusting of a 30 km thick allochthonacross the Eastern Alps would provide roughly 1.45⇥1018 kg of sediment to circum-East Alpinebasins in the interval 40 Ma.–present day, assuming a 10 km present-day average thickness of theAustroalpine nappes and an average erosion rate of 1.1 mm.a�1. Compared to both England (1981)and Kuhlemann (2000b) estimates of 1.1⇥1018 kg and 0.3⇥1018 kg for the East Alpine detritalsediment load, this suggests inconsistency between the detrital sediment record and metamorphicpressures, which could be accounted for by a hidden Eocene–present age sediment trap i.e. missinghanging wall rock, or, an overestimate of the metamorphic pressures i.e. thinner allocthon. Thelarge di↵erence (0.8⇥1018 kg) between England (1981) and Kuhlemann (2000b) estimates reflectsthe inherent di�culty associated with the accurate determination of Alpine sediment contribu-tions to mixed sediment traps, such as the Pannonian Basin and Danube fan in the Black Sea(Kuhlemann et al., 2002). The flysch–molasse sediment transition within circum-Alpine basinsoccurred across the Eocene–Oligocene boundary (Sinclair, 1997; Kuhlemann, 2000b; Kuhlemannet al., 2006), which strongly supports contemporaneous thrust sheet emplacement. Within theEastern Alps, cumulative sediment discharge rates rapidly increased from ⇠2000 km3.Ma�1 to⇠6000 km3.Ma�1, between 27–30 Ma, according to Kuhlemann et al. (2002).

Syntectonic growth of Barrovian assemblages would provide strong evidence in support of foot-wall heating due to rapid overthrusting of hot rock, as noted in the Lesser Himalaya (Caddicket al., 2007). However, the relationship between Barrovian mineral growth and fabric formationin the Penninic units of the Eastern Alps is less clear. Kurz et al. (1996) state that westwarddirected shearing initiated synchronous to the Barrovian peak, whilst Droop (1985) and Inger& Cli↵ (1997) state that the metamorphic peak post–dated regional fabric formation. SampleASA-08-42b is a quartz-mica schist collected from the Venediger cover nappes exposed in theDorfertal and comprises the following assemblage: garnet-muscovite-carbonate-biotite-chlorite-

184

Europe

Iberia

Africa

Adria

175 Ma

120

4633

19

Valais

Pyrenees Austroalpine

100 km

10 km.Ma-1

20 40 60 80 100 120 140 160

20

40

0

-20

Time (Ma)

km.M

a-1

Figure 6.11: Convergence of Africa relative to stable Eurasia 40 Ma. Palaeogeographic reconstruction taken from

Bousquet et al. (2002). Dashed line represents the trajectory of Africa relative to a fixed point in Europe (blue circle:

45�N; 6�W); red circle represent absolute time in Ma calculated from magnetic isochrons - data taken from Rosenbaum

et al. (2002). Inset plot shows convergenec rate (km.Ma�1 as a function of time; shading represents thrusting interval.

quartz±ilmenite±rutile±epidote. Muscovite and quartz micro-lithons define an intense foliation,which is tightly crenulated. Pertinently, this fabric is included within sub-euhedral garnet blasts(⇠3 mm diameter) as sigmoidal quartz-rutile-mica trails, suggestive of syn-kinematic garnet growth(Fig.6.12). The age of garnet growth in the Venediger cover is uncertain; however, given that mus-covite analysed from a proximal locality yields Alpine 40Ar-39Ar ages (sample Fr 7–16: 34.9±1.4Ma; Zimmermann et al., 1994) and the elevated position of the sample in the structural pile(Carboniferous-Permian Kurz et al., 1998b), a Variscan age is unlikely.

Plate velocity overthrusting of the Austroalpine nappes provides an elegant explanation forequally rapid exhumation of the Eclogite Zone as the same basal thrust system would be responsiblefor both processes (Fig. 6.13; e.g. Ring & Glodny, 2010). Similarly, the scenario accounts forcontinued subduction post-eclogite formation, explaining the low-temperatures of blueschist faciesoverprinting prior to Barrovian conditions. However, this mechanism of rapid heating requires thattemperatures at the base of the overthrust sheet approximate peak values.

The Altkristallin basement exposed directly to the south of the Tauern Window comprises

185

Figure 6.12: Photomicrograph of garnet-bearing, quartz-mica schist ASA-08-42b taken under crossed polars; field of

view is 2 mm. The well-defined S1

fabric is laterally continuous with shallow, sigmoidal inclusion trails within the garnet,

indicative of syn-tectonic poiklioblast growth. Note that the apparent (dextral) sense of shear required to produce garnet

inclusion trail asymmetry is consistent with that required to form F2

crenulations (white arrow), which mark the initial

stages of regionally pervasive S2

fabric development.

strongly deformed biotite-muscovite gneisses, hornblende-plagioclase gneisses, garnet-muscoviteschists, orthogneisses and amphibolites of the Lasorling group. These amphibolite facies assem-blages are though to have formed during pre-Alpine metamorphism, during, or shortly after theVasriscan orogeny (Peccerillo et al., 1979; Schulz et al., 2001; Siegesmund et al., 2007). Protolithages (U–Pb zircon) from metaigneous rocks of the Altkristallin cluster close to 470 Ma (e.g. Schulz& Bombach, 2003). During the Alpine orogenic cycle, temperatures at the base of the Austroalpinesheet were high enough to facilitate Ar di↵usion in biotite, chloritisation and development of low-temperature quartz microstructures at ca. 30 Ma. (Waters, 1976; Wallis, 1988), i.e. �300–350�C.Under these conditions, syn-thrusting heating would not account for the rapid attainment of Bar-rovian conditions. However, it is likely that pervasive development of new Alpine mineralogy washampered by the dehydrated nature of the reworked Altkristallin basement.

Therefore, syn-thrusting heating provides a plausible explanation for the rapid (<10 Ma.)tectono-metamorphic evolution of the Eastern Alps, despite the paucity of directly supportive fieldevidence. Tectonic replenishment of heat to footwall units may be a fundamental mechanism toother overthrust terranes, such as the Himalaya.

186

km

0

50

100

kbar

0

10

20

30

40

qtzcoe

Austroalpine

detachment of eclogitesfrom slab

600�  C500�  C

passive marginlithospheric mantle

1-10 cm.a-1 thrusting

Penninic

Valaisian slab

Austroalpine wedge

Eclogite Zone

ca. 38-28 MaN S

Figure 6.13: Tectonic schematic of scenario in which concomitant syn-thrusting heating of the Penninic basement and

plate velocity exhumation of the Eclogite Zone occur. This is possible because the same basal thrust system is responsible

for the rates of thrust sheet emplacement and exhumation. Isotherms are approximate and represent an early thermal

architecture associated with oceanic subduction (Gerya & Stockhert, 2006). Continental collision would raise isotherms

to shallower crustal levels.

6.7 Chapter Summary

The 34.2±3.6 Ma U–Pb allanite age for subduction-related metamorphism of the European margindrastically shortens the duration of time available for conductive relaxation of isotherms followingemplacement of the Austroalpine nappes, from ⇠30 Ma to 10 Ma. One-dimensional modelingshows that the thermal evolution of the Tauern Window cannot be explained by conductive heatingof overthickened crust on lithosphere with normal continental thermal gradients. Either the thermalevolution took place with a significant thermal contribution from the mantle, for which there islittle supporting evidence, or heating occurred during thrusting with heat provided by an overlyingthrust sheet. It is possible that syn–thrusting heating by rapid emplacement of a hot overthrustsheet, overlooked in previous thermal models, could plausibly have played an influential role in thetectono–metamorphic evolution of the Eastern Alps and in overthrust terrains in general.

187

Chapter 7

Discussion and Conclusions

The preceding chapters have shown that East Alpine orogenesis occurred in under 10 Ma, across theEocene–Oligocene boundary—up to an order of magnitude faster than previously thought. Suchrates of thermal and tectonic transport question the dominance of conductive heat transfer overtimescales of tectono-metamorphic evolution in orogenesis. The following discussion concentratesfirst on pertinent exhumation mechanisms responsible for transporting the Eclogite Zone at platetectonic rates from mantle depths, followed by a brief consideration of further examples of rapidheating to Barrovian condition elsewhere in the geological record. Finally, the tectonic implicationsof rapid East Alpine orogenesis for the evolution of the Alpine chain as a whole are discussed andoutstanding issues are raised for future work.

7.1 Exhumation of the Eclogite Zone

Whilst subduction provides a plausible means of conveying tracts of crustal material to mantle-depths, the processes responsible for exhuming (U)HP terranes to mid-crustal levels are not un-derstood (e.g. Ernst et al., 1997; Agard et al., 2009; Guillot et al., 2009). The Eclogite Zone wasexhumed at plate-tectonic rates of 0.6–>6 cm.a�1, between ca.38 Ma and ca.27–32 Ma, over avertical distance of ⇠50–60 kilometers (assuming a purely lithostatic pressure gradient). Suchrates have been recognised elsewhere in the Phanerozoic orogenic record—see Fig.7.1 for a globalcompilation of exhumation velocities: the Dora Maira terrane of the Western Alps was exhumedfrom ⇠35 kbar to 4–5 kbar between 35.1±0.9 Ma and 31.8±0.5 Ma, which translates to tectonicrates of 1.6–3.4 cm.a�1 (Rubatto & Hermann, 2001); the Tso Morari unit of the Ladakh Himalayawas exhumed from coesite-stable conditions at vertical rates faster than 0.5 cm.a�1 between 55 and47 Ma (De Sigoyer et al., 2000); Earth’s youngest exposed eclogite terrane in Papua New Guineawas exhumed from ⇠75 kilometres depth at rates between 3.5–5.1 cm.a�1 (Baldwin et al., 2004).

Rates of exhumation for subducted bodies of continental crust (Fig.7.1) are up to an order ofmagnitude faster than those derived for exhumed oceanic crust (e.g. the Franciscan Agard et al.,2009). Furthermore, bodies of continental crust are exhumed from mantle-depths in excess of100 kilometres, whereas oceanic crustal material is not exhumed from depths deeper than ⇠70–80

188

kilometres. Exhumation of (U)HP rocks requires the following criteria to be fulfilled: a mecha-nism for detachment of slab-top material; a driving force for exhumation in the opposite sense tosubduction; erosion of overlying rocks in order to maintain isostatic equilibrium of thickened crustduring exhumation. The extreme buoyancy contrast between subducted felsic crust and proximalmantle is of the order 200–400 kg.m�3 at (U)HP conditions and provides a viable explanation forrapid exhumation of continental crust.

9080706050403020100Exhumation rate (mm.a-1)

Franciscan

Zagros

ZermattDora Maira

Kokchetav

Papua New Guinea

Kaghan

Monviso

Tso Morari

Western Gneiss Region

Dabie Shan

Tauern Window

Figure 7.1: Global compilation of (U)HP exhumation velocities. Colour scheme correlates with nature of subducted

material: light green = oceanic crust; dark green = back-arc crust; deep red = continental crust. Velocities for Tauern and

Tso Morari terranes are only accurately constrained as minima. Rates calculated from the following studies : Franciscan,

Cloos (1986); Wakabayashi (1990); Ernst & Liou (1995); Anczkiewicz et al. (2004); Zagros, Agard et al. (2005); Agard

(2006); Zermatt Saas, Amato et al. (1999); Dora Maira, Rubatto & Hermann (2001); Kokchetav, Hermann et al. (2001);

Papua New Guinea, Baldwin et al. (2004); Kaghan, Parrish et al. (2006); Monviso, Rubatto & Hermann (2003); Tso

Morari, De Sigoyer et al. (2000); Western Gneiss Region, Kylander-Clark et al. (2008); Dabie Shan, Ayers et al. (2002);

Hacker et al. (2004); Tauern Window, this study and Glodny et al. (2005).

However, exhumation of dense tracts of oceanic lithopsphere is not readily explained by buoy-ancy forces and requires contribution from externally applied forces. The relative roles of bodyand externally applied forces in exhuming (U)HP passive margin sequences such as the EclogiteZone remains a fundamental question to tectonics (e.g. Platt, 1993; Ernst et al., 1997; Kurz &Froitzheim, 2002; Agard et al., 2009). The following section addresses the validity of both forcesin explaining the plate velocity exhumation of the Eclogite Zone.

Structural data presented in section 2.4.2.2 are consistent with the interpretation that theEclogite Zone was inserted into the Penninic nappe pile as an oblique extrusion wedge, with a

189

thrust-sense lower boundary and a top-to-the-west transpressional upper boundary (Behrmann& Ratschbacher, 1989; Kurz et al., 1996; Kurz & Froitzheim, 2002; Glodny et al., 2008; Ring &Glodny, 2010). These high strain zones accommodated between 50–70 km of vertical displacementduring exhumation. However, given the oblique transpressional nature of the hanging wall contactit is likely that the transport direction during exhumation was steeper than the plunge of the shearzone’s lineation (Robin & Cruden, 1994; Neufeld et al., 2008). Accordingly, average exhumationrates of 3.6 cm.a�1 calculated by Glodny et al. (2005) and 0.6–>6 cm.a�1 calculated in section4.10 are minimum estimates.

7.1.1 Buoyancy-driven exhumation

Buoyancy is often invoked as the primary body force behind rapid exhumation of continental-derived eclogite facies rocks (e.g. Rubatto & Hermann, 2001; Parrish et al., 2006). Detailed map-ping within the Eclogite Zone (Eremin, 1994) shows that mafic eclogites constitute up to 35% of thenappe’s volume. Given that the eclogites and host metasediments have average densities of 3400kg.m�3 and 2800 kg.m�3 respectively, a first order estimate of the Eclogite Zone’s average densityis 3020 kg.m�3. Taking mantle density at 70–80 km—peak conditions for the Eclogite Zone, as3340 kg.m�3 (Lee, 2003), the density contrast (�⇢) between mantle and subducted margin willbe of the order of 300 kg.m�3. According to the classical work of England & Holland (1979), a 3kilometer-wide subduction channel (h) with �⇢ of 300 kg.m�3, a subducting slab velocity (u) of 10cm.yr�1 at an angle (✓) of 45� requires the e↵ective viscosity (⌘) of the subduction channel to be �9.7⇥1017 kg.ms�1 for shearing forces associated with subduction to dominate over buoyancy. Thistranslates to a threshold value of the slab-top’s gross shear strength (⌧ = (⌘.u)/h) of 9.7 bars (0.97MPa), above which subduction will proceed and beneath which buoyancy will be the dominantforce. Figure 6.3 shows the range of slab shear strengths required for subduction, for values of h be-tween 3 and 5 km, using the parameters presented above, against power law rheologies of aragonite,quartz and eclogite. Quartz, mica–schist and eclogite are all between 101–103 MPa stronger thanthe shear strength required to subduct a 5 kilometer-thick slab-top at 600�C—the upper-bound tothe temperature of eclogite facies metamorphism. However, the plot shows that aragonite deformsvia power-law deformation at shear strengths less than those required to overcome buoyancy forceswithin a channel of ⇠4 km thickness at peak temperatures. Importantly, this confirms that the netrheology of the Eclogite Zone and the P–T conditions of the subduction–exhumation transition islikely to have been controlled by the rheology of calc-schist. Furthermore, this provides an elegantexplanation for focussing of exhumation-related strain on calc-schist horizons, close to the EclogiteZone’s boundary faults (e.g. Fig.7.4b).

Stockhert et al. (1997) found that garnet-chloritoid-kyanite schists of the Eclogite Zone recorda transition from interfacial free energy to dislocation creep dominated deformation of quartz. Us-ing the experimentally derived quartz flow law of Paterson & Luan (1990), under strain rates of

190

Vex = h2.g.sinθ.(ρm-ρs)12η

cosθ.Vs

mantle wedge

subducting lithosphere

eclogite boudins

metasedimentary matrix

θsub

Vs = 2a2.(ρe-ρs).g

η

ρm

ρs

9η

ρe

Vtot = Vex-cosθ.Vs

Figure 7.2: Technical illustration

of the force balance exerted on an

exhuming eclogite boudin.

10�14s�1, at 550–600�C, this transition is thought to occur at 10–30 bars (1-3 MPa). Clearly, thisshows that extremely low di↵erential stresses operated during the pressure peak of subduction,which accounts for the initiation of exhumation as shearing forces decrease.

Collectively, these data corroborate the results of pseudosection modeling presented in section3.3, which show that garnet–chloritoid–kyanite bearing assemblages within aluminous EclogiteZone metapelites are indicators of a state of enhanced detachment ca. 550�C. Detachment ofthe Eclogite Zone from the subducting Pennine slab was controlled by the strong temperature-dependence of shear strength of aragonite-bearing lithologies.

Following detachment from the down-going slab, shearing forces decrease and buoyancy forcesact to drive the fault-bounded slice back up the subduction channel at a rate proportional to thedensity contrast between proximal mantle and channel material. The average velocity of buoyancydriven exhumation (Vex) can be approximated by the following expression (equation 4: England& Holland, 1979):

Vex =h2 · �⇢ · g · sin✓

12⌘(7.1)

,where, h is the width of the subduction channel in metres, �⇢ is the density contrast between prox-imal mantle and subduction channel sediments (kg.m�3), g is gravitational acceleration (m.s�2), ✓

is the angle of subduction (�) and ⌘ is e↵ective viscosity in kg.ms�1. However, as noted by England

191

Boudin radius (m)0 100 200 300 400 500

1

2

3

4

5

6

7

8

9

2

4

6

8

10

12

Vs (cm

.a-1)

η x

1017

(kg.

m-1.s

-1)

45

6

3

2

1

0.1

Figure 7.3: Plot showing relationship between

boudin radius (m) and the e↵ective viscos-

ity of surrounding metasediments (kg.ms�1),

contoured for values of constant sink velocity

(cm.a�1) according to Sokes’ Law. Calculations

performed using ⇢e-⇢s=700 kg.m�3 and g=10

m.s�2.

& Holland (1979), mafic eclogite boudins within the Eclogite Zone are denser than surroundingmetasediments by as much as 600–700 kg.m�3. Therefore, for buoyancy to explain exhumation ofthe Eclogite Zone, eclogite boudins enveloped within the metasedimentary matrix must be su�-ciently small for viscous forces to exceed their negative buoyancy. Provided that the boudin is smallrelative to the width of the subduction channel (i.e. very small Reynolds number), the velocity atwhich an eclogite body will sink through a viscous matrix can be calculated according to Stokes’law:

Vs =2a2 · (⇢e � ⇢s) · g

9⌘(7.2)

where, a is the radius of a spherical eclogite boudin (m), ⇢e is the density of eclogite and ⇢e is thedensity of channel metasediment (kg.m�3). It is important to note that the following calculationsare only strictly valid for spherical eclogite bodies. The net velocity at which a spherical eclogiteboudin would be exhumed is simply:

Vtot = Vex � cos✓ · Vs (7.3)

This force balance is illustrated in Fig.7.2. Figure 7.3 illustrates the relationship between boudinradius and matrix viscosity, contoured for values of constant sink velocity (Vs); calculations usea density contrast of 700 kg.m�3 between eclogites and sediments. For boudins less than ⇠300metres in diameter, as exposed in the Eclogite Zone, Vs is slow (<1 cm.a�1) over a large rangeof matrix viscosities between 1–9⇥1017 kg.ms�1. Larger boudins, between 150–500 metres radius,yield sink velocities which are increasingly controlled by viscosity, such that Vs>5 cm.a�1 occursfor boudins >325 m radius, in sediments of e↵ective viscosities beneath ⇠250 kg.ms�1.

Figure 7.4a plots contours of constant boudin radius as a function of net exhumation veloc-

192

ity (Vtot; eq.7.3) and e↵ective viscosity. Contour geometries reflect an inflection from viscositycontrolled exhumation velocities to boudin size controlled behaviour at high e↵ective viscosities.Exhumation velocities are greater for lower e↵ective viscosities, according to eq.7.1. Eclogiteboudins <200 metres radius display similar characteristic exhumation rates. The inset plot high-lights the range of observed Eclogite Zone boudin dimensions against the likely range of averageexhumation rates calculated from both U–Pb and Rb–Sr isotope geochronology (section 4.10; 3–6cm.a�1; Glodny et al., 2005). This shows that for Eclogite Zone boudins to be exhumed at suchplate velocity rates, the e↵ective viscosity of the subduction\exhumation channel must have been⇠1–2⇥1018 kg.ms�1.

10

8

6

4

2

04321

0.1 m 1 10 100 200 300 400 500 600 700 800

6.0

5.5

5.0

4.5

4.0

3.5

3.01.51.00.5V to

t (cm

.a-1)

Viscosity (kg.ms-1)5 x 1018

2 x 1018

a b

Figure 7.4: a. Contours of constant boudin radius (m), as a function of e↵ective viscosity (kg.ms�1), and net exhumation

velocity (cm.a�1); ⇢e-⇢s = 700 kg.m�3, h=3000m and g=10 m.s�2. Inset plot highlights parameter space applicable to

Eclogite Zone boudins - grey band represents boudins between 0.1–200m radius. Contour colour corresponds to boudin

radius as shown in legend; b. Photograph of calc-schist shear zone displaying dextral sense of asymmetry; geological

hammer for scale. Dominant foliation (S1

+S2

) dips steeply to the south. Note: boudinage of refractory quartz pods in

mylonitic ground mass, indicative of viscosity contrast. Foliation strikes ⇠080�.

However, the heterogenous composition of the Eclogite Zone means that contrasts in shearstrength and hence, e↵ective viscosity, will have existed during exhumation. As shown above,power-law deformation of aragonite occurs at lower shear stresses than for quartz and clinopyroxene(Fig.6.3), meaning that exhumation-related deformation will have been partitioned into calc-schisthorizons Fig.7.4b. A 10 metre-wide aragonite shear zone operating at 400�C under an exhumationrate of 6 cm.a�1, will have an e↵ective viscosity of 5⇥1016 kg.ms�1, given by:

⌘ =⌧

�(7.4)

where, � is shear rate (u/h) in s�1 in a Newtonian fluid. This is two orders of magnitude lessthan the estimate of the Eclogite Zone’s net e↵ective viscosity derived from Fig.fig:visc. These lowviscosity channels also account for kilometer-scale lateral continuity of interlayered quartzite andmarble horizons, best exposed at the base of the Eclogite Zone, shown by Fig.7.5.

193

Upon reaching mid-crustal depths (⇠30 km), the density contrast between metasediments andproximal accretionary wedge material will decrease and buoyancy-driven exhumation will cease.This appears to corroborate with insertion of the Eclogite Zone between the Venediger and Glocknernappes just before Barrovian metamoprhism at ca. 7–9 kbar (24–30 km). Subsequent exhuma-tion of the entire nappe stack occurred at slower rates (Blanckenburg et al., 1989), driven by acombination of erosion and extension.

Figure 7.5: Field photograph detailing kilometre-scale stratigraphic continuity of Eclogite Zone metaseiments. Pho-

tograph taken looking south from the head of Frosnitzkees; field of view ⇠1.3 km. Dashed black lines mark laterally

continuous metasedimentary–mafic eclogite banding; inset details metre-scale stratigraphic heterogeneity.

7.1.2 Return flow

Return flow of low-viscosity material within a confined subduction channel has been postulated toaccount for the exhumation of (U)HP rocks (Cloos, 1982; Gerya et al., 2002; Gerya & Stockhert,2006). This model relies upon a subduction channel which narrows with depth, a low viscositymatrix and active subduction. Circulation of subducted material is critically controlled by the pres-sure gradient which develops according to channel geometry and rheology, along the subductionchannel Gerya et al. (2002). Hydration of the mantle wedge is thought to be a critical parametera↵ecting the channel’s pressure gradient and thus in facilitating return flow (Hermann et al. (2000);Guillot et al. (2000); Guillot, S. and Hattori, K.H. and de Sigoyer, J. and Nagler, T. and Auzende,A.L. (2001); Bostock et al. (2002); Gerya & Stockhert (2006)). The commonly observed associationbetween (U)HP rocks and serpentinite bodies supports this concept (e.g. Guillot et al. (2000);Guillot, S. and Hattori, K.H. and de Sigoyer, J. and Nagler, T. and Auzende, A.L. (2001)). Small(100 meter-scale) serpentinite bodies do exist proximal to the Eclogite Zone (section 2.4.4; Kurz& Froitzheim, 2002), however, they are few and volumetrically insignificant relative to the size ofthe nappe itself. Furthermore, the Eclogite Zone appears to have been exhumed as a coherent unit

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and not as a chaotic melange as expected with low viscosity material i.e. the Franciscan complex(Cloos, 1983, 1986). The coherent nappe geometry of the exhumed Eclogite Zone could well beexplained by a higher e↵ective viscosity within the subduction channel (Gerya & Stockhert, 2006)as a result of a non-pervasively hydrated mantle wedge. Plate velocity rates of exhumation ofthe unit are consistent with a Newtonian deformation regime (⌧ = "⌘, where " is the strain rate)within the channel, where strain was concentrated into carbonate-rich shear zones. Interestingly,the models of Gerya & Stockhert (2006) predict that buoyancy forces do not significantly a↵ectthe flow pattern of a subduction channel due to the large viscosity di↵erence between channel andhanging wall lithologies. However, as higher e↵ective viscosities of channel material are requiredfor coherent exhumation, the role of buoyancy will become more significant.

In summary, both buoyancy- and viscosity-driven exhumation provide a plausible means ofinserting the Eclogite Zone into the Schieferhulle nappe pile from mantle depths. To discriminatebetween the mechanisms, more precisely constrained exhumation rates are required.

7.2 Barrovian metamorphism: where’s the heat?

Combined U–Pb allanite geochronology and thermal modeling show that the thermal evolution ofthe Tauern Window occurred up to a magnitude faster than previously thought. Such rapid heatingcannot be explained by conductive heating of overthickened crust on lithosphere with normal conti-nental thermal gradients; rather, these data infer that an advective heat transport must have beensignificant. The relative contributions of conductive and advective heat transfer and their opera-tional lengthscales throughout regions of thickened crust remains a profound limitation to currentunderstanding of metamorphism. Further examples of rapid attainment of Barrovian conditionsin thickened crust has been documented in the Barrovian metamorphic zones of Scotland (Oliveret al., 2000; Ague & Baxter, 2007), the Wepawaug Schist of south-central Connecticut (Lancasteret al., 2008), the Tso Morari (U)HP terrane of the Ladakh Himalaya (De Sigoyer et al., 2000) andthe Lepontine Dome of the Central Alps (Engi et al., 2001; Brouwer et al., 2004; Berger et al., 2011).

Using records of Sr di↵usion preserved in apatite and corroborating Fe–Mg–Ca–Mn di↵usionpreserved in garnet, Ague & Baxter (2007) show that the thermal peak of Grampian (ca.460–470Ma) metamorphism in the classic Barrovian zones of Scotland (Barrow, 1893) prevailed for 200–300thousand years. Thermal modeling shows that background conductive heat transfer due to crustalthickening contributed to temperatures <500�C throughout the 10–15 Ma duration of orogeny—anadvective component is required to account for peak temperatures between 535–660�C within thegarnet—sillimanite zones respectively (Oliver et al., 2000). Advection of heat by emplacement ofsynchronous mafic intrusions and associated high regional fluid fluxes are invoked as a plausiblemechanism of advective heat transfer (Baxter et al., 2002).

Lancaster et al. (2008) used Sm–Nd and U–Pb geochronology on garnet and zircon, respectively,

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to show that peak Barrovian conditions were attained pene-contemporaneously across the Bt–Gtand St–Ky zones, in two distinct episodes—the first at ca. 389 Ma, and the second at ca. 380Ma, within the Wepawaug Schist of Connecticut, U.S.A.. These data challenge the predictionmade by conductive heating models of regional metamorphism that peak conditions are attained atdi↵erent times following thrust-sheet emplacement (England & Thompson, 1984); this, again, infersadvective heat transfer. The thermal peak of metamorphism was concomitant with emplacementof felsic igneous bodies and large fluid fluxes (⇠103 m3

fluid.m�2rock; Ague, 2011).

The Tso Morari dome of the Ladakh Himalaya represents a distal crystalline block of thinnedIndian continental margin (Guillot et al., 2008). Using multichronometric studies (U–Pb, Lu–Hf, Sm–Nd, Rb–Sr and Ar–Ar) De Sigoyer et al. (2000) showed that the unit was subducted toeclogite facies conditions (20±3 kbar and 550±50�C; Guillot et al., 1997) by ca.55 Ma; during quasi-isothermal exhumation, the Tso Morari rocks recrystallised under blueschist facies conditions (11±2kbar and 580±50�C; De Sigoyer et al., 1997), before being partially a↵ected by an amphibolitefacies event at ca.47 Ma (9±3 kbar and 610±70�C; Guillot et al., 1997; De Sigoyer et al., 1997). TheP–T–t evolution is remarkably similar to that derived from the Eclogite Zone of the Tauern Windowand raises similar questions regarding the timescales of Barrovian heating. The low temperaturesof the Tso Morari eclogites suggest that exhumation occurred closely after burial, during an activesubduction regime. Accordingly, Barrovian temperatures must have been generated in under ⇠8Ma—too quickly to be explained by overthrusting of Eurasian crust with a typical continentalgeotherm.

The Lepontine Dome of the Central Alps occupies a similar structural position to the TauernWindow of the Eastern Alps. Peak Barrovian temperatures were attained diachronously acrossthe dome: in the south Tmax occurred at 30–28 Ma; in the central Lepontine between 22–24 Ma(Engi et al., 1995, and references therein); in the north at 18–19.1 Ma (Janots et al., 2009).These data are fully consistent with conductive relaxation of isotherms following emplacement ofthe Austroalpine nappes. However, tectonically higher structural units, such as the Alpe Araminappe, were a↵ected by pre-Barrovian eclogite-facies metamorphism as late as 33–36 Ma (Brouweret al., 2004, 2005; Hermann et al., 2006). Following peak pressure conditions of ⇠19 kbar and⇠830�C, Lepontine gneisses near Alpe Arami cooled to 630�C during rapid exhumation and weresubsequently reheated by ⇠110�C to 740�C at 6 kbar by ca. 32.4 Ma (Brouwer et al., 2004, andreferences therein). Such high values of �T/�t invoke advective heat transfer, as is shown by thisdissertation to apply to the Tauern Window’s thermal evolution.

The records of Barrovian metamorphism described above show that mid-crustal heating alonggeothermal gradients between 20–50�C.km�1 frequently occurs on timescales 10 Ma. Simpleisotherm relaxation following crustal thickening occurs on timescales greater than 10 Ma (England& Thompson, 1984). The problem is quantitatively summarized by the Peclet number (PeL), whichexpresses the relative importance of conductive and advective heat transport (L=characteristiclengthscale; V=velocity; =thermal di↵usivity):

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PeL =L.V

(7.5)

Where PeL>1, advection dominates over conduction. Taking a pre-orogenic crustal thicknessof 35 kilometres, a convergence velocity of 1 cm.a�1 and =1 mm2.s�1, PeL⇡11. This confirmsthat in a typical convergent orogen, material is transported through the system at rates quickenough to preclude significant conductive heating. Whereas the heat deficit is readily explainedby syn-metamorphic magmatism in the Barrovian terranes of Scotland and eastern U.S.A., in theEastern Alps and the Himalaya there is no evidence for a concomitant advective heat source. Var-ious, largely theoretical, mechanisms have been invoked to account for the enigmatic heat source:Jamieson et al. (1998) and Engi et al. (2001) suggest that accretion of radiogenic upper-crustalmaterial to the base of the accretionary wedge is required to explain Barrovian temperatures;Brouwer et al. (2004) shows that asthenospheric upwelling following slab-breako↵ (Huw Davies &von Blanckenburg, 1995) will elevate temperatures in the mid-crust; Burg & Gerya (2005) suggeststhat viscous heating associated with crustal deformation may result in orogen-scale production of0.1–1 µW.m�3, which accounts for temperatures up to 200�C higher than conductive backgroundtemperatures in terranes such as the Lepontine Dome.

However, these models are only weakly supported by field data: Engi et al. (2001) base theircalculations on the Southern Steep Belt of the Lepontine Dome, which appears to be a unique fea-ture to the Central Alps—there is no evidence for an enriched radiogenic layer capable of supplying>2µW.m�3 in the Eastern Alps or the Ladakh Himalaya; there is no evidence for asthenosphericupwelling in either the Eastern Alps or Himalaya, which is predicted to accompany slab breako↵(Von Blanckenburg & Huw Davies, 1995); viscous heating requires deformation of rheologicallystrong lithologies, whereas in the Eastern Alps deformation related to Alpine shortening is parti-tioned into horizons of weak calcareous lithology which is incapable of supporting significant shearstresses above ⇠300�C.

Therefore, the mechanisms responsible for rapid (<10 Ma) Barrovian heating remain a fun-damental problem to current understanding of metamorphism and, therefore, orogenesis. Theintimate feedbacks between thermal and mechanical processes in regions of deforming crust makethis question fundamentally important to continental tectonics and deserving of special attentionin future research.

7.3 Tectonic Implications

The following discussion is supported throughout by Fig.7.6—an orogen-scale compilation of Alpinegeochronological data.

Past studies have shown that the age of HP metamorphism preserved in the Western and theCentral Alps cluster into three main episodes: 66–70 Ma, 54–44 Ma and 34–38 Ma (e.g. Ducheneet al., 1997; Gebauer et al., 1997; Rubatto & Hermann, 2001; Lapen et al., 2003; Bachmann

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et al., 2009; Radulescu et al., 2009; Wiederkehr et al., 2009; Beltrando et al., 2010a; Rubattoet al., 2011). Older ages are preserved in structurally higher levels, such as the Sesia Zone, whichunderwent two episodes of eclogite facies metamorphism between 65 and 79 Ma (Rubatto et al.,2011), reflecting incorporation of the thinned Adriatic margin to the Alpine accretionary wedgeduring early stages of convergence(Froitzheim & Manatschal, 1996). Younger ages of (U)HP

metamorphism are located at lower structural levels of the Alpine nappe pile. In the WesternAlps, the Dora Maira UHP complex, which represents continental basement of the Brionconnaismicrocontinent, was subducted to ⇠35 kbar between 35 and 37 Ma (Gebauer et al., 1997; Rubatto& Hermann, 2001) and the Gran Paradiso Massif preserves eclogite facies mineral parageneseswhich formed between 34–38 Ma (Me↵an-Main et al., 2004; Gabudianu Radulescu et al., 2009). Inthe Central Alps gneisses of the Lepontine dome (Alpe Arami and Monte Duria) were subductedto (U)HP conditions between 34–36 Ma (Gebauer, 1999; Brouwer et al., 2005; Hermann et al.,2006). The distribution of (U)HP ages throughout the Western and Central Alps is consistentwith the interpretation that the Alps evolved through episodic accretion of continental ribbons andoceanic basins located progressively further north–northwest, during consumption of the WesternTethys (e.g. Trumpy, 1980; Rosenbaum et al., 2002; Beltrando et al., 2010a).

The new geochronological data presented in this dissertation indicate that HP metamorphismoccurred between ca. 31–38 Ma in the Eastern Alps, contemporaneous (at 2�) with the youngestgroup of (U)HP ages in the Central and Western Alps. This implies that continental collisionbetween the European and Adriatic continents also occurred contemporaneously along ⇠1200kilometres of the Alpine chain. Deformation associated with (U)HP metamorphism between 31–38Ma is characterised by isoclinal folding—indicative of a shortening strain regime—at all structurallevels.

Following (U)HP metamorphism, large tracts of the Alpine chain experienced Barrovian meta-morphism within a restricted time interval, between 27–33 Ma, coincident with a transition fromshortening- to extensional-related deformation. Attainment of the steepest geothermal gradientshas been constrained to between 32.2–33 Ma in Lepontine gneisses (Gruf complex, Monte Duriaand Alpe Arami) of the Central Alps (Gebauer, 1999; Liati & Gebauer, 2003; Hermann et al.,2006). In frontal units of the Central Alps, Barrovian temperatures persisted locally until ⇠18–19Ma (Janots et al., 2009). In the Western Alps, the Dora Maira (U)HP terrane recrystallised underBarrovian conditions at ca.32.9±0.9 Ma (Michard et al., 1993; Rubatto & Hermann, 2001), whilstthe neighboring Gran Paradiso unit attained similar conditions at ca. 33.2±0.4 Ma (Freeman et al.,1997). Collectively, these data, along with those displayed in Fig.7.6 show that Barrovian condi-tions were attained extremely quickly after (U)HP metamorphism, i.e. <10 Ma, along the entireAlpine chain. This thermal anomaly was most pronounced in the Central Alps where temperaturesreached in excess of 700�C (Rubatto et al., 2009, and references therein).

Attainment of Barrovian conditions occurred concomitantly with plate-tectonic velocity ex-humation of subducted (U)HP terranes, in addition to the transition from flysch to molasse de-position in peripheral basins (Sinclair, 1997; Kuhlemann et al., 2002, 2006). As shown in chapter

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2, exhumation of the Eclogite Zone was facilitated by a transition from north-south shorteningto orogen-parallel transpressive extension between ⇠30–33 Ma (Glodny et al., 2008). In the Cen-tral and Western Alps, syn-exhumation extensional structures have been shown to be operationalbetween 30–34 Ma—i.e. the Turba deformation phase (Froitzheim et al., 1994; Marquer et al.,1996; Nagel et al., 2002; Beltrando et al., 2010a). Therefore, it appears that a pulse of exhumationoccurred along the Alpine chain as a result of entry of the European basement into the subduc-tion zone. This would presumably decrease convergence velocity and cause associated relaxationof isotherms at (U)HP depths; thermal weakening has been shown by modeling studies to be acritical factor controlling detachment of material from the slab-top (Carry et al., 2009).

Thermochronological data from detrital sediments deposited in circum-Alpine basins and de-rived from internal crystalline units show clear evidence for a regionally significant Early Oligoceneexhumation event. Zircon fission track and 40Ar/39Ar ages define a maximum at 28–32 Ma and33.2±0.4 Ma, respectively, in the Miocene Swiss Molasse (Von Eynatten et al., 1999; Spiegel et al.,2000); the Pennine units of the Western Alps yield similar zircon fission track ages (e.g. Hurfordet al., 1991). Cumulative sediment discharge rates for internal units of the Alpine chain doubledfrom 5000 to 10000 km3.Ma�1, between 28 and 31 Ma according to the calculations of Kuhlemann(2000b). Collectively, these data show that an orogen-scale pulse of exhumation occurred acrossthe Eocene–Oligocene boundary and facilitated rapid cooling of exhumed units. Apatite and zirconfission track ages for Penninic units exposed in the axial zone of the Alpine chain commonly liebetween 7 and 28 Ma. This has been interpreted as representing late-orogenic unroofing associatedwith dome formation in the Eastern (Tauern Window) and Central Alps (Lepontine Dome)

There is a close temporal relationship between exhumation, Barrovian metamorphism andregional Alpine magmatism. The HP event in the Eclogite Zone is contemporaneous with em-placement of Alpine intrusions directly to the south of the Tauern Window (see section 6.4.2):the Rensen pluton, 31.1–31.7 Ma (Barth et al., 1989); the Rieserferner pluton, 32.4±0.4 Ma and31.8±0.4 Ma (Romer & Siegesmund, 2003). Magmatism was more pronounced in the Central Alpsas evidenced by intrusion of the larger Adamello and Bergell granitoid bodies. The Adamellobatholith was intruded over an area of ⇠670 km2, between 43 and 33 Ma (Del Moro et al., 1983;Schaltegger et al., 2009) and is the largest of the Periadriatic intrusions. Compositionally, thebatholith represents a calc-alkaline fractionation trend from mafic melts exposed in the south, totonalites and granodiorites in the north. The Bergell pluton intrudes the Penninic nappes north ofthe Insubric line in the eastern Central Alps and comprises a balloon-shaped suite of granodiorite–tonalite granitoids emplaced over ⇠5 Ma between 33 and 28 Ma (Oberli et al., 2004). Emplacementof the Bergell and later stages of the Adamello batholiths occurred contemporaneously with ex-humation of (U)HP units and subsequent heating of mid-crustal levels to Barrovian conditions.

In summary, across the Eocene–Oligocene boundary the Alpine chain experienced a pulse ofheightened tectono-metamorphic activity. Furthermore, the sequence of tectonic, magmatic andmetamorphic events preserved in the Penninic nappes of the Tauern Window are concordant with

199

those preserved in the Central and Western Alps. These data have been interpreted as evidenceto suggest that the time interval between 30–34 Ma represents a transition from a period oflithospheric thickening to extension along the Alpine chain (Von Blanckenburg & Huw Davies,1995; Beltrando et al., 2010a).

Huw Davies & von Blanckenburg (1995) and Von Blanckenburg & Huw Davies (1995) proposethat the close temporal relationship between metamorphism, exhumation and magmatism areaccounted for by catastrophic loss of the Alpine orogen’s lithospheric root. In this model ofslab breako↵, buoyancy-driven detachment of Tethyan oceanic lithosphere promotes rapid ascent(exhumation) of positively buoyant crustal material in addition to upwelling of the asthenospherealong the locus of slab detachment. Accordingly, this would account for the thermal perturbationresponsible for rapid heating to Barrovian conditions following contemporaneous exhumation alongthe orogen’s strike and also the pulse in magmatic activity. However, the model has several criticalflaws: firstly, evidence (i.e. high 3He; mantle melts) for widespread asthenospheric involvement inAlpine magmatism is not observed (i.e. Marty et al., 1992); secondly, detachment of the slab wouldgenerate orogen-scale isostatic rebound and not necessarily di↵erential exhumation of discretenappes during ongoing subduction (Gerya et al., 2004); finally, east–west extensional deformationcontradicts the radial pattern of extension predicted in response to slab breako↵ (Schellart & Lister,2004). In light of these objections, slab breako↵ is not deemed a viable model for Alpine orogenesis.

Lithospheric thinning during transpressive convergence between Europe and the Adriatic promon-tory could also be accounted for by subduction rollback (Beltrando et al., 2010a). In this model,retrograde (relative to the overriding plate; north–northwest in the Alps) motion of the hinge ofthe subduction zone is driven by gravitational instability of the subducting slab. Importantly, themodel predicts a decrease in orogen-parallel stresses, which accounts for the marked transitionfrom north–south convergence to east–west extension along the Alpine belt. Furthermore, retreatof the subduction hinge causes thinning of the hanging-wall lithosphere, accounting for elevatedBarrovian geothermal gradients. Slab rollback could plausibly have been driven by southward re-treat of the north-dipping subduction of the African slab (Lustrino et al., 2009), which is thoughtto have been responsible for rotation of the Sardinia–Corsica block, opening of the Balearic basinand widespread extension throughout the Aegean ca. 32 Ma (Jolivet et al., 2003).

Seismic tomography confirms the presence of two distinct Moho surfaces at depths between30 and 60 kilometres beneath the Alpine arc (e.g. Schmid et al., 2004b). Therefore, in order forlithospheric thinning, via a mechanism such as slab rollback, to account for the tectonic associationcharacteristic of the Early Oligocene Alps, the observed lithospheric structure of the Alps mustreflect re-working post-30 Ma.

In addition to the Tauern Window, Valaisian nappes are exposed in the Engadine Windowand, further to the east, in the Central Alps in the vicinity of Petit St. Bernard (Trumpy, 1955).Comparing di↵erences in the P–T–t records preserved at each of these localities allows conclu-sions to be formed regarding the paleogeography of the Valais oceanic realm. Firstly, the Valais

200

ocean spanned at least 1000 kilometres east–west (approximate distance from Petit St. Bernardto Tauern). Secondly, the P–T path of western Valaisian exposures are di↵erent from those inthe Central and Eastern Alps in that they are characterised by higher metamorphic conditions(⇠17 kbar; Bousquet et al., 2002) and experienced heating during decompression. Schieferhullemetasediments of the Engadine Window form a 10–15 kilometer-thick pile and preserve peak pres-sures of ⇠12 kbar. Whereas in the Western Alps, Valaisian units are intercalated with basementslices of the Brianncconnais terrane (Nicolas et al., 1990; Fugenschuh et al., 1999), in the Enga-dine and Tauern Windows there is no evidence for microcontinent involvement. Finally, the U–Pallanite geochronology presented in this contribution shows that subduction-related HP metamor-phism was contemporaneous throughout the Valais oceanic realm at ⇠33–38 Ma. Furthermore,the occurrence of fossils at the top of the Valais stratigraphy with the same age as the HP event(40–35 Ma) shows that continental collision along the Valais trench must have occurred post–35Ma (Bousquet et al., 2002). These data are consistent with the interpretation that the Valais oceanwas connected along the length of the Alpine arc and independent from the Piemonte Ligurianbasin (Bousquet et al., 2002). High-pressure metamorphism is accounted for by an accretionarywedge which abutted the Brianncconnais microcontinent in the west, and the northern margin ofthe Austroalpine domain in the east; metamorphic pressures between 12–17 kilobars would implya wedge thickness of ⇠35–55 kilometers. The higher pressures experienced by the Eclogite Zoneattest to the unit being an exotic slice which was inserted into the wedge during exhumation.

Figure 7.6 (following page): Orogen-scale compilation of Alpine geochronological data between present day and

90 Ma. Locations ordered in approximate east–west trend to highlight along-strike di↵erences. Numbering of intrusive

bodies: 1 Adamello pluton, 2 Bergell pluton, 3 Rieserferner pluton and 4 Rensen intusion. Solid grey line represents

approximate timing of transition from subduction to exhumation. Note: assimilated units refers to litho-tectonic units

whose provenance is uncertain and now form the plate boundary in the Central Alps(Berger & Bousquet, 2008). Data

sources are as follows: Rechnitz: Ratschbacher et al. (2004); Tauern: this study, Blanckenburg et al. (1989); Zimmermann

et al. (1994); Inger & Cli↵ (1997); Ratschbacher et al. (2004); Glodny et al. (2005, 2008); Engadine Window: Wiederkehr

et al. (2009); Reckner: Dingeldey et al. (1997); Platta: Deutsch (1984a); Err nappe: Handy et al. (1996); Malenco: Villa

et al. (2000); Suretta: Wagner et al. (1979); Challandes et al. (2003); Alpe Arami: Wagner et al. (1979); Becker (1993);

Gebauer et al. (1996); Brouwer et al. (2005); Janots et al. (2009); Repiano: Wagner et al. (1979); Brouwer et al. (2005);

Monte Rosa: Me↵an-Main et al. (2004); Lapen et al. (2003); Antrona: Liati et al. (2005); Tisch Alp: Rubatto et al.

(1998); Amato et al. (1999); Sesia: Hurford et al. (1991); Venturini (1995); Reddy et al. (1996); Duchene et al. (1997);

Rubatto et al. (1999, 2011); Lago Di Cignana: Rubatto et al. (1998); Amato et al. (1999); Lapen et al. (2003); Dente

Blanche: Reddy et al. (2003); Cortiana et al. (1998); Janots et al. (2009); Mt. Emilius: Dal Piaz et al. (2001); Valaisian

nappes: Freeman et al. (1997); Cannic et al. (1999); Liati et al. (2005); Liati & Froitzheim (2006); Schistes Lustres: Agard

et al. (2002); Bucher et al. (2003); Grivola: Dal Piaz et al. (2001); Bucher et al. (2003); Gran Paradiso: Me↵an-Main

et al. (2004); Gabudianu Radulescu et al. (2009); Monviso: Duchene et al. (1997); Cli↵ et al. (1998); Rubatto & Hermann

(2003); Dora Maira: Tilton et al. (1991); Duchene et al. (1997); Gebauer et al. (1997); Rubatto & Hermann (2001);

Voltri: Federico et al. (2005); Cap Corsica: Brunet et al. (2000). Intrusive ages: Adamello: Schaltegger et al. (2009);

Bergell: Oberli et al. (2004); Rieserferner: Romer & Siegesmund (2003); Rensen: Barth et al. (1989).

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7.4 Conclusions

1. This dissertation presents the first U–Pb geochronological evidence for the age of eclogite-facies metamorphism within the Eclogite Zone, Tauern Window. Allanite-clinozoisite grewduring the prograde portion of the Eclogite Zone’s P–T–t path, between 8–13 kbar at34.2±3.6 Ma, prior to initiation of garnet growth. Combined with previously reported Bar-rovian metamorphism at 27–32 Ma, this places a maximum 10 Ma period between ⇠28–38Ma during which the peak of East Alpine eclogite-facies metamorphism occurred. Eclogitesfacies rocks were exhumed at plate velocity rates, between 1–6 cma�1, over a vertical distanceof ⇠50 km.

2. These new geochronological data drastically shorten the duration of time available for con-ductive relaxation of isotherms following thrust sheet emplacement, from ⇠30 Ma to 10Ma. Thermal modeling shows that the thermal evolution of the Tauern Window cannot beexplained by conductive heating of overthickened crust on lithosphere with normal continen-tal thermal gradients. Either the thermal evolution took place with a significant thermalcontribution from the mantle, for which there is little supporting evidence or heating oc-curred during thrusting with heat provided by an overlying thrust sheet. It is possible thatsyn–thrusting heating by rapid emplacement of a hot overthrust sheet, overlooked in previousthermal models, could plausibly have played an influential role in the tectono–metamorphicevolution of the Eastern Alps and in overthrust terrains in general.

3. Single grain fusion 40Ar/39Ar data from individual muscovite and paragonite grains collectedfrom representative samples pertaining to each of the Penninic nappes show that 40ArE isvariable on millimetric lengthscales for all lithologies. This strongly suggests that the grainboundary network operated as a finite volume to radiogenic 40Ar. Calculations confirm thatthe observed range of apparent 40Ar/39Ar ages can be explained by autochthonous fluidproduction within Jurassic protoliths. Furthermore, time-integrated porosities corroboratewith predicted devolatisation histories for mafic and pelitic lithologies.

4. The Penninic nappes experienced a minimum of three discrete episodes of deformation: D0

pertains to syn-subduction flattening during southward consumption of the Valaisian basin;D1 occurred contemporaneously with peak pressures in the Eclogite Zone and Glocknernappe; D2 delineates the transition from north–south convergence to east–west transpres-sive extension throughout the Alpine chain. The Eclogite Zone was tectonically insertedbetween the Schieferhulle nappes under blueschist facies conditions.

5. The Eocene–Oligocene boundary (30–38 Ma) represents a period of heightened tectonic,metamorphic and magmatic activity. Formation of Valaisian- and European-derived (U)HP

terranes, subsequent plate-tectonic exhumation and heating to Barrovian conditions occurredcontemporaneously, between 33–38 Ma, along the Alpine arc. This was accompanied by

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syn-metamorphic emplacement of several granitoid–calc-alkaline plutons in the Central and,to a lesser extent, in the Eastern Alps. The close temporal association of these eventsare compatible with lithospheric thinning after thickening, plausibly driven by gravitationalretreat of the African slab.

7.5 Future Work

The most important question raised by this work is that concerning the mechanisms and processesresponsible for orogen-scale attainment of Barrovian conditions in less than ten million years af-ter subduction-related (U)HP metamorphism. Calculation of precise rates of heating are a keyconstraint in discriminating between advective heating mechanisms. In turn, precision is limitedby the sensitivity of radiogenic isotope metrology. Recent advances in Sm–Nd geochronology ofzoned garnets mean that geologically meaningful ages (<1–2% precision) can be obtained fromextremely small sample sizes (2–10 ng Nd; Harvey & Baxter, 2009). This allows calculation ofgarnet growth rates from individual growth zones (Pollington & Baxter, 2010). Applied to garnetsof the Tauern Window, this method has the potential to provide invaluable constraints on the ratesand duration of HP and Barrovian metamorphism. More specifically, the age and conditions ofHP metamorphism in the Peripheral Schieferhulle nappes remain outstanding and yet are crucialto understanding the region’s tectono-metamorphic evolution.

More precise geochronological data constraining the age of both Barrovian and HP metamor-phism in the Tauern Window would refine the range of plausible exhumation rates of the EclogiteZone. This would discriminate between the models of buoyancy- and return-flow-driven exhuma-tion.

The 40Ar/39Ar white mica geochronometer is used ubiquitously to date cooling of metamorphicterranes through T c. However, more work is required to constrain the likely conditions under whichthe assumptions of Dodson (1973) are valid. In particular, under HP conditions 40ArE is pervasive,commonly partitioning back into the source mineral. In such cases, the 40Ar/39 method becomesa hydrochronometer—this method merits calibration against other fluid-mobile elements such asLi (Penniston-Dorland et al., 2010)and B(Tenthorey & Hermann, 2004), and has the potential toprovide quantitative insight into fluid connectivity throughout a metamorphic cycle. A startingpoint for further work is measurement of 40Ar/39Ar ratios from cogenetic phases.

204

Appendix A

Appendix

A.1 Published manuscripts

The following peer-reviewed manuscripts, presented in reverse-date order, were published duringthe course of the thesis:

205

206

A.2 Electron Microprobe analysis

Quantitative chemical analyses of all mineral phases discussed were performed at the University ofCambridge using a Cameca SX–100 electron microprobe with 5 wavelength dispersive spectrome-ters and a single energy dispersive spectrometer. Major-element analyses were acquired using a 15KeV accelerating voltage, a beam current of 10 nA and 20 s data collection time. REE and traceelement acquisition used 15 KeV, 200 nA and 100 s respectively. Natural and synthetic mineralsand oxides were used as standards. Raw data were corrected on–line according to Cameca X-PHIprocedure. CWT-sample micas were analysed at the Open University using a JEOL SX-100 5wavelength-dispersive spectrometer Electron Microprobe with the following operating conditions:a defocused beam with spot size of 10 µm; 15kV accelerating voltage, a 20nA beam current and a30s collection time. Natural standards were used for calibration and analyses were corrected usinga ZAF matrix correction routine. Analyses were bracketed by analyses of secondary standards tocheck for major-element reproducibility of 1%.

Table A.1: Major-element electron microprobe analyses of phases present in Betic sample ALM–45.

Mineral: garnetrim garnetcore chloritoid phengite

SiO2

36.51 38.14 24.52 47.98TiO

2

0.03 0.09 0.00 0.30Al

2

O3

21.28 20.98 41.07 32.74Cr

2

O3

0.01 0.00 0.00 0.01FeO 37.57 31.52 21.01 1.28MnO 0.27 2.05 0.04 0.00MgO 3.85 2.05 4.93 0.18CaO 0.73 5.74 0.04 0.00Na

2

O 0.13 0.16 0.05 0.15K

2

O 0.00 0.00 0.00 0.77Total 100.43 99.94 91.69 94.47

No. of oxygen 12 12 6 11Si 2.92 3.06 1.00 3.20Ti 0.00 0.09 0.00 0.02Al 2.00 1.98 1.98 2.58Fe

tot

2.42 2.11 0.70 0.07Mn 0.02 0.14 0.00 0.00Mg 0.46 0.15 0.30 0.18Ca 0.06 0.49 0.00 0.00Na 0.02 0.03 0.00 0.15K 0.00 0.00 0.00 0.77

207

Table A.2: Electron microprobe analyses of blueschist-facies phases from sample N45b (Ray, 1986) and symplectite-

forming phases from sample 221c (Ray, 1986).

gl jd phe pa ep ab act mu ep ab bi

N45b 221cSiO

2

57.62 55.71 49.84 45.92 37.74 71.42 56.54 50.15 37.37 68.59 40.59TiO

2

0.02 0 0.2 0 0.04 0.01 0 0.16 0 0 0Al

2

O3

8.88 11.18 26.08 37.09 22.22 18.61 1.03 26.63 21.54 19.95 13.81Cr

2

O3

0.03 0 0.05 0 0.03 0 0 0 0 0 0Fe

2

O3

4.22 15.55 1.87 0 14.93 0.83 4.56 0.22 14.23 0.18 2FeO 10.46 1.94 3.25 1.29 0.14 0 3.97 4.79 0 0 10.19MnO 0.09 0 0.02 0 0.27 0 0.89 0 0.95 0 0.2MgO 9.22 0.86 3.2 0 0.01 0.6 18.21 2.19 0 0 18.46CaO 0.38 2.42 0.12 0 22.9 0.17 9.5 0 21.61 0.13 0.24Na

2

O 7.05 12.83 0.4 6.8 0.03 11.37 2.05 0 0 11.41 0K

2

O 0.01 0 10.62 1.17 0.01 0.07 0.08 10.73 0 0 9.56

Totals 97.56 100.5 95.48 92.27 98.32 103.08 96.83 94.86 95.72 100.27 95.06

Oxygens 23 6 11 11 12.5 8 23 11 12.5 8 11

Si 8.003 2.005 3.362 3.046 3.011 3.025 7.955 3.405 3.056 2.984 2.968Ti 0.002 0 0.01 0 0.002 0 0 0.008 0 0 0Al 1.454 0.474 2.074 2.9 2.09 0.929 0.171 2.132 2.077 1.023 1.191Cr 0.003 0 0.003 0 0.002 0 0 0 0 0 0

Fe3+ 0.441 0.421 0.096 0 0.897 0.027 0.483 0.012 0.876 0.006 0.11Fe2+ 1.215 0.058 0.183 0.072 0.009 0 0.467 0.271 0 0 0.623Mn 0.011 0 0.001 0 0.018 0 0.106 0 0.066 0 0.012Mg 1.909 0.046 0.322 0 0.001 0.038 3.818 0.222 0 0 2.012Ca 0.057 0.093 0.009 0 1.958 0.008 1.432 0 1.894 0.006 0.019Na 1.899 0.896 0.052 0.875 0.005 0.934 0.559 0 0 0.963 0K 0.002 0 0.914 0.099 0.001 0.004 0.014 0.929 0 0 0.892

Sum 15.141 3.995 7.026 6.991 7.995 4.965 15.005 6.98 7.968 4.983 7.828

208

Table A.3: Electron microprobe analyses of Venediger-nappe samples, TH–519, ASA–08–38a and ASA–08–42a.

g bi mu pl ep kfs mu mu bi fsp g g ep mu fsp g chl

TH–519 ASA–08–38a ASA–08–42aSiO

2

38.18 35.47 49.27 67.66 38.41 65.78 49.10 46.42 35.04 68.67 37.96 38.05 38.05 48.23 68.09 36.88 24.21TiO

2

0.13 1.82 0.52 0.00 0.11 0.00 0.44 0.39 1.88 0.00 0.12 0.20 0.18 0.28 0.00 0.08 0.07Al

2

O3

21.43 16.65 26.68 19.81 27.58 17.73 27.90 31.43 17.35 19.74 21.10 20.97 28.49 31.51 20.05 21.05 22.17Cr

2

O3

0.01 0.01 0.01 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.05 0.00 0.02 0.02Fe

2

O3

1.00 0.00 0.39 0.04 6.13 0.07 0.00 0.00 0.00 0.00 1.06 1.59 6.58 0.00 0.00 3.08 0.01FeO 18.64 24.53 4.32 0.00 1.24 0.00 4.02 3.20 24.12 0.00 26.04 14.26 0.33 2.55 0.00 33.91 30.75MnO 0.31 0.37 0.06 0.00 0.05 0.00 0.02 0.01 0.24 0.00 1.11 4.75 0.08 0.00 0.01 1.86 0.25MgO 0.30 6.97 2.65 0.00 0.04 0.00 2.50 1.55 6.95 0.00 0.47 0.12 0.04 1.47 0.00 1.48 11.37CaO 20.15 0.16 0.02 0.86 23.73 0.03 0.06 0.07 0.01 0.11 13.51 20.38 23.83 0.03 0.66 4.41 0.04Na

2

O 0.10 0.27 0.18 11.39 0.22 0.48 0.36 0.49 0.08 12.24 0.04 0.06 0.00 0.81 11.51 0.02 0.00K

2

O 0.00 9.54 11.13 0.12 0.12 15.92 10.84 10.84 9.79 0.10 0.00 0.00 0.00 10.00 0.09 0.00 0.00

Totals 100.15 95.80 95.20 99.88 97.02 100.02 95.26 94.41 95.47 100.86 101.30 100.22 96.95 94.94 100.41 102.48 88.89

Oxygens 12 11 11 8 12.5 8 11 11 11 8 12 12 12.5 11 8 12 14

Si 2.98 2.77 3.35 2.97 3.01 3.03 3.32 3.16 2.74 2.98 2.99 2.98 2.97 3.23 2.97 2.92 2.59Ti 0.01 0.11 0.03 0.00 0.01 0.00 0.02 0.02 0.11 0.00 0.01 0.01 0.01 0.01 0.00 0.01 0.01Al 1.97 1.53 2.14 1.02 2.55 0.96 2.23 2.52 1.60 1.01 1.96 1.94 2.62 2.49 1.03 1.97 2.80Cr 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Fe3+ 0.06 0.00 0.02 0.00 0.36 0.00 0.00 0.00 0.00 0.00 0.06 0.09 0.39 0.00 0.00 0.18 0.00Fe2+ 1.22 1.60 0.24 0.00 0.08 0.00 0.23 0.18 1.58 0.00 1.71 0.93 0.02 0.14 0.00 2.25 2.76Mn 0.02 0.02 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.07 0.32 0.01 0.00 0.00 0.13 0.02Mg 0.04 0.81 0.27 0.00 0.01 0.00 0.25 0.16 0.81 0.00 0.06 0.01 0.01 0.15 0.00 0.18 1.82Ca 1.69 0.01 0.00 0.04 1.99 0.00 0.00 0.01 0.00 0.01 1.14 1.71 1.99 0.00 0.03 0.37 0.01Na 0.02 0.04 0.02 0.97 0.03 0.04 0.05 0.07 0.01 1.03 0.01 0.01 0.00 0.11 0.97 0.00 0.00K 0.00 0.95 0.97 0.01 0.01 0.94 0.94 0.94 0.98 0.01 0.00 0.00 0.00 0.86 0.01 0.00 0.00

Sum 8.00 7.85 7.04 5.01 8.05 4.98 7.04 7.06 7.84 5.03 8.00 8.00 8.02 6.99 5.01 8.00 10.00

209

SiO2

TiO2

Al2

O3

Cr2

O3

Fe2

O3

FeO MnO MgO CaO Na2

O K2

O Totals

ASA-28a-mus-1 51.18 0.23 29.71 0.04 0.00 1.33 0.02 3.23 0.09 0.59 9.83 96.26

ASA-28a-mus-2 49.21 0.23 28.71 0.07 0.00 1.24 0.00 3.09 0.08 0.59 9.64 92.87

ASA-28a-mus-3 48.34 0.23 28.85 0.03 0.02 1.13 0.00 2.96 0.09 0.46 9.17 91.29

ASA-28a-mus-4 50.42 0.23 28.90 0.02 0.00 1.37 0.00 3.22 0.09 0.55 9.73 94.54

ASA-28a-mus-5 50.87 0.28 29.29 0.01 0.00 1.37 0.00 3.05 0.09 0.57 9.85 95.39

ASA-28a-mus-6 49.79 0.29 29.19 0.03 0.00 1.33 0.01 2.95 0.06 0.59 9.80 94.05

ASA-28a-mus-7 50.60 0.26 29.06 0.02 0.00 1.24 0.01 3.19 0.06 0.56 9.61 94.62

ASA-28a-mus-8 50.54 0.26 28.59 0.03 0.00 1.36 0.00 3.21 0.21 0.55 9.89 94.65

ASA-28a-mus-9 50.20 0.28 28.80 0.04 0.00 1.39 0.02 3.14 0.08 0.51 9.77 94.24

ASA-28a-mus-10 48.32 0.26 28.69 0.05 0.00 1.24 0.01 3.04 0.07 0.57 9.52 91.78

C.p.f.u.

ASA-28a-mus-1 6.70 0.02 4.58 0.00 0.00 0.15 0.00 0.63 0.01 0.15 1.64 13.89

ASA-28a-mus-2 6.68 0.02 4.59 0.01 0.00 0.14 0.00 0.63 0.01 0.16 1.67 13.91

ASA-28a-mus-3 6.65 0.02 4.68 0.00 0.00 0.13 0.00 0.61 0.01 0.12 1.61 13.85

ASA-28a-mus-4 6.72 0.02 4.54 0.00 0.00 0.15 0.00 0.64 0.01 0.14 1.65 13.89

ASA-28a-mus-5 6.72 0.03 4.56 0.00 0.00 0.15 0.00 0.60 0.01 0.15 1.66 13.88

ASA-28a-mus-6 6.68 0.03 4.61 0.00 0.00 0.15 0.00 0.59 0.01 0.15 1.68 13.90

ASA-28a-mus-7 6.72 0.03 4.55 0.00 0.00 0.14 0.00 0.63 0.01 0.14 1.63 13.86

ASA-28a-mus-8 6.74 0.03 4.49 0.00 0.00 0.15 0.00 0.64 0.03 0.14 1.68 13.90

ASA-28a-mus-9 6.71 0.03 4.54 0.00 0.00 0.16 0.00 0.63 0.01 0.13 1.67 13.88

ASA-28a-mus-10 6.64 0.03 4.65 0.01 0.00 0.14 0.00 0.62 0.01 0.15 1.67 13.92

CWT-17-mus-1 48.66 0.50 28.21 0.00 0.00 3.20 0.00 2.00 0.01 0.34 10.21 93.13

CWT-17-mus-2 48.88 0.51 28.01 0.00 0.00 3.36 0.01 2.11 0.01 0.38 9.93 93.21

CWT-17-mus-3 49.30 0.50 28.08 0.00 0.00 3.53 0.01 2.13 0.01 0.38 10.15 94.08

CWT-17-mus-4 48.92 0.53 27.88 0.00 0.00 3.66 0.01 2.08 0.00 0.41 10.08 93.58

CWT-17-mus-5 48.76 0.52 27.95 0.00 0.00 3.27 0.01 2.14 0.02 0.28 10.12 93.06

CWT-17-mus-6 48.98 0.52 28.11 0.00 0.00 3.43 0.01 2.09 0.00 0.42 9.99 93.55

CWT-17-mus-7 48.88 0.55 28.03 0.00 0.00 3.38 0.01 2.09 0.00 0.46 10.03 93.43

CWT-17-mus-8 49.23 0.53 28.04 0.00 0.00 3.46 0.01 2.03 0.00 0.41 10.00 93.71

CWT-17-mus-9 49.04 0.52 27.75 0.00 0.00 3.70 0.01 2.06 0.01 0.42 9.95 93.45

CWT-17-mus-10 49.43 0.54 26.83 0.00 0.00 3.97 0.00 2.28 0.00 0.31 10.13 93.49

C.p.f.u.

CWT-17-mus-1 6.67 0.05 4.56 0.00 0.00 0.37 0.00 0.41 0.00 0.09 1.79 13.94

CWT-17-mus-2 6.69 0.06 4.52 0.00 0.00 0.39 0.00 0.43 0.00 0.10 1.73 13.92

CWT-17-mus-3 6.69 0.05 4.49 0.00 0.00 0.40 0.00 0.43 0.00 0.10 1.76 13.94

CWT-17-mus-4 6.69 0.06 4.49 0.00 0.00 0.42 0.00 0.42 0.00 0.11 1.76 13.94

CWT-17-mus-5 6.69 0.06 4.52 0.00 0.00 0.37 0.00 0.43 0.00 0.07 1.77 13.92

CWT-17-mus-6 6.68 0.06 4.52 0.00 0.00 0.39 0.00 0.42 0.00 0.11 1.74 13.93

CWT-17-mus-7 6.68 0.06 4.52 0.00 0.00 0.39 0.00 0.42 0.00 0.12 1.75 13.94

CWT-17-mus-8 6.70 0.06 4.50 0.00 0.00 0.40 0.00 0.41 0.00 0.11 1.74 13.92

CWT-17-mus-9 6.70 0.06 4.47 0.00 0.00 0.42 0.00 0.42 0.00 0.11 1.74 13.93

CWT-17-mus-10 6.77 0.06 4.33 0.00 0.00 0.46 0.00 0.47 0.00 0.08 1.77 13.93

CWT-12-mus-1 49.17 0.29 29.30 0.00 0.00 2.33 0.00 2.31 0.00 0.66 10.09 94.17

CWT-12-mus-2 49.93 0.18 27.83 0.00 0.00 2.67 0.00 2.69 0.00 0.44 10.40 94.15

CWT-12-mus-3 49.95 0.23 28.25 0.00 0.00 2.58 0.00 2.53 0.01 0.52 9.98 94.05

CWT-12-mus-4 50.28 0.17 27.91 0.00 0.00 2.70 0.00 2.66 0.00 0.46 10.23 94.41

CWT-12-mus-5 50.32 0.17 27.64 0.00 0.00 3.00 0.02 2.65 0.01 0.38 10.22 94.39

CWT-12-mus-6 49.77 0.18 28.18 0.00 0.00 3.01 0.00 2.50 0.00 0.52 9.99 94.15

CWT-12-mus-7 49.70 0.19 28.19 0.00 0.00 3.08 0.00 2.45 0.00 0.44 10.23 94.28

CWT-12-mus-8 50.18 0.15 27.90 0.00 0.00 2.81 0.00 2.60 0.00 0.47 10.22 94.34

C.p.f.u.

210

SiO2

TiO2

Al2

O3

Cr2

O3

Fe2

O3

FeO MnO MgO CaO Na2

O K2

O Totals

CWT-12-mus-1 6.63 0.03 4.66 0.00 0.00 0.26 0.00 0.47 0.00 0.17 1.74 13.96

CWT-12-mus-2 6.75 0.02 4.44 0.00 0.00 0.30 0.00 0.54 0.00 0.12 1.79 13.96

CWT-12-mus-3 6.74 0.02 4.49 0.00 0.00 0.29 0.00 0.51 0.00 0.14 1.72 13.92

CWT-12-mus-4 6.77 0.02 4.43 0.00 0.00 0.30 0.00 0.53 0.00 0.12 1.76 13.93

CWT-12-mus-5 6.79 0.02 4.39 0.00 0.00 0.34 0.00 0.53 0.00 0.10 1.76 13.93

CWT-12-mus-6 6.73 0.02 4.49 0.00 0.00 0.34 0.00 0.50 0.00 0.14 1.72 13.93

CWT-12-mus-7 6.72 0.02 4.49 0.00 0.00 0.35 0.00 0.50 0.00 0.12 1.77 13.95

CWT-12-mus-8 6.77 0.02 4.43 0.00 0.00 0.32 0.00 0.52 0.00 0.12 1.76 13.94

CWT-8-mus-1 48.32 0.27 32.87 0.00 0.00 1.19 0.01 1.49 0.02 0.89 9.25 94.31

CWT-8-mus-2 48.70 0.27 32.33 0.00 0.00 1.32 0.01 1.77 0.07 0.78 9.26 94.51

CWT-8-mus-3 48.83 0.24 31.07 0.00 0.00 1.44 0.00 2.00 0.03 0.67 9.40 93.69

CWT-8-mus-4 50.29 0.26 29.21 0.00 0.00 1.62 0.00 2.75 0.03 0.56 9.41 94.13

CWT-8-mus-5 50.65 0.31 30.10 0.00 0.00 1.54 0.00 2.64 0.02 0.60 9.42 95.28

CWT-8-mus-6 49.79 0.31 29.97 0.00 0.00 1.41 0.00 2.41 0.00 0.66 9.42 93.98

CWT-8-mus-7 49.20 0.31 29.75 0.00 0.00 1.39 0.00 2.19 0.00 0.61 9.36 92.82

CWT-8-mus-8 49.87 0.29 30.30 0.00 0.00 1.42 0.02 2.26 0.01 0.62 9.44 94.23

CWT-8-mus-9 50.22 0.31 29.99 0.00 0.00 1.44 0.01 2.30 0.01 0.62 9.41 94.32

CWT-8-mus-10 49.88 0.30 30.06 0.00 0.00 1.47 0.01 2.27 0.01 0.66 9.45 94.10

CWT-8-mus-11 49.94 0.32 29.71 0.00 0.00 1.51 0.00 2.45 0.01 0.60 9.44 93.98

CWT-8-mus-12 50.52 0.33 29.36 0.00 0.00 1.55 0.01 2.74 0.02 0.60 9.40 94.52

CWT-8-mus-13 50.04 0.27 30.01 0.00 0.00 1.35 0.01 2.40 0.01 0.60 9.37 94.07

CWT-8-mus-14 49.82 0.31 29.68 0.00 0.00 1.44 0.01 2.38 0.00 0.60 9.42 93.65

CWT-8-mus-15 50.02 0.26 29.77 0.00 0.00 1.38 0.01 2.39 0.01 0.64 9.45 93.94

CWT-8-mus-16 49.59 0.29 29.54 0.00 0.00 1.35 0.00 2.39 0.01 0.63 9.48 93.27

CWT-8-mus-17 50.10 0.28 29.53 0.00 0.00 1.47 0.00 2.38 0.00 0.60 9.43 93.79

CWT-8-mus-18 49.61 0.28 29.86 0.00 0.00 1.50 0.00 2.32 0.00 0.60 9.41 93.58

CWT-8-mus-19 49.75 0.29 29.95 0.00 0.00 1.42 0.01 2.29 0.00 0.64 9.44 93.80

CWT-8-mus-20 49.79 0.30 29.30 0.00 0.00 1.57 0.00 2.72 0.00 0.54 9.49 93.72

CWT-8-mus-21 48.80 0.24 30.70 0.00 0.00 1.51 0.00 2.16 0.01 0.60 9.46 93.48

CWT-8-pa-22 47.16 0.07 38.68 0.00 0.00 0.17 0.01 0.12 0.32 7.08 1.02 94.62

CWT-8-pa-23 47.28 0.04 38.77 0.00 0.00 0.15 0.01 0.10 0.32 7.00 1.00 94.67

CWT-8-pa-24 46.21 0.08 39.35 0.00 0.00 0.18 0.00 0.10 0.39 6.78 0.86 93.94

C.p.f.u.

CWT-8-mus-1 6.44 0.03 5.16 0.00 0.00 0.13 0.00 0.30 0.01 0.23 1.57 13.85

CWT-8-mus-2 6.47 0.03 5.07 0.00 0.00 0.15 0.00 0.35 0.01 0.20 1.57 13.85

CWT-8-mus-3 6.56 0.02 4.92 0.00 0.00 0.16 0.00 0.40 0.01 0.18 1.61 13.85

CWT-8-mus-4 6.72 0.03 4.60 0.00 0.00 0.18 0.00 0.55 0.01 0.14 1.61 13.83

CWT-8-mus-5 6.68 0.03 4.68 0.00 0.00 0.17 0.00 0.52 0.01 0.15 1.58 13.82

CWT-8-mus-6 6.66 0.03 4.72 0.00 0.00 0.16 0.00 0.48 0.00 0.17 1.61 13.84

CWT-8-mus-7 6.66 0.03 4.75 0.00 0.00 0.16 0.00 0.45 0.00 0.16 1.62 13.82

CWT-8-mus-8 6.65 0.03 4.76 0.00 0.00 0.16 0.00 0.45 0.00 0.16 1.61 13.82

CWT-8-mus-9 6.69 0.03 4.71 0.00 0.00 0.16 0.00 0.46 0.00 0.16 1.60 13.81

CWT-8-mus-10 6.67 0.03 4.74 0.00 0.00 0.17 0.00 0.45 0.00 0.17 1.61 13.83

CWT-8-mus-11 6.68 0.03 4.69 0.00 0.00 0.17 0.00 0.49 0.00 0.15 1.61 13.83

CWT-8-mus-12 6.72 0.03 4.60 0.00 0.00 0.17 0.00 0.54 0.01 0.15 1.60 13.82

CWT-8-mus-13 6.68 0.03 4.72 0.00 0.00 0.15 0.00 0.48 0.00 0.15 1.60 13.81

CWT-8-mus-14 6.69 0.03 4.70 0.00 0.00 0.16 0.00 0.47 0.00 0.15 1.61 13.82

CWT-8-mus-15 6.69 0.03 4.69 0.00 0.00 0.15 0.00 0.48 0.00 0.17 1.61 13.83

CWT-8-mus-16 6.69 0.03 4.69 0.00 0.00 0.15 0.00 0.48 0.00 0.17 1.63 13.83

CWT-8-mus-17 6.72 0.03 4.66 0.00 0.00 0.17 0.00 0.47 0.00 0.15 1.61 13.81

CWT-8-mus-18 6.67 0.03 4.73 0.00 0.00 0.17 0.00 0.46 0.00 0.16 1.61 13.83

211

SiO2

TiO2

Al2

O3

Cr2

O3

Fe2

O3

FeO MnO MgO CaO Na2

O K2

O Totals

CWT-8-mus-19 6.67 0.03 4.73 0.00 0.00 0.16 0.00 0.46 0.00 0.17 1.62 13.83

CWT-8-mus-20 6.69 0.03 4.64 0.00 0.00 0.18 0.00 0.54 0.00 0.14 1.63 13.85

CWT-8-mus-21 6.57 0.02 4.87 0.00 0.00 0.17 0.00 0.43 0.00 0.16 1.62 13.86

CWT-8-pa-22 6.07 0.01 5.87 0.00 0.00 0.02 0.00 0.02 0.04 1.77 0.17 13.96

CWT-8-pa-23 6.08 0.01 5.87 0.00 0.00 0.02 0.00 0.02 0.04 1.74 0.17 13.94

CWT-8-pa-24 5.98 0.01 6.01 0.00 0.00 0.02 0.00 0.02 0.06 1.70 0.14 13.93

CWT-15-pa-1 47.02 0.09 38.43 0.00 0.00 0.66 0.00 0.21 0.19 7.30 0.94 94.84

CWT-15-pa-2 46.97 0.15 38.15 0.00 0.00 0.66 0.00 0.26 0.18 6.97 1.17 94.51

CWT-15-pa-3 46.71 0.10 38.96 0.00 0.00 0.51 0.00 0.14 0.25 7.41 0.61 94.69

CWT-15-pa-4 46.92 0.08 39.13 0.00 0.00 0.41 0.00 0.08 0.31 7.49 0.55 94.97

CWT-15-pa-5 47.05 0.07 38.73 0.00 0.00 0.48 0.00 0.18 0.23 7.23 0.80 94.78

CWT-15-pa-6 46.88 0.12 38.37 0.00 0.00 0.61 0.00 0.23 0.18 7.09 1.13 94.61

CWT-15-pa-7 46.82 0.08 37.75 0.00 0.00 0.81 0.00 0.78 0.21 6.65 0.98 94.09

CWT-15-pa-8 46.93 0.08 38.69 0.00 0.00 0.52 0.01 0.19 0.21 7.31 0.66 94.60

CWT-15-pa-9 46.84 0.10 39.03 0.00 0.00 0.49 0.00 0.14 0.24 7.48 0.53 94.85

C.p.f.u.

CWT-15-pa-1 6.06 0.01 5.83 0.00 0.00 0.07 0.00 0.04 0.03 1.82 0.15 14.01

CWT-15-pa-2 6.07 0.02 5.81 0.00 0.00 0.07 0.00 0.05 0.03 1.75 0.19 13.98

CWT-15-pa-3 6.01 0.01 5.91 0.00 0.00 0.06 0.00 0.03 0.03 1.85 0.10 14.00

CWT-15-pa-4 6.02 0.01 5.91 0.00 0.00 0.04 0.00 0.02 0.04 1.86 0.09 13.99

CWT-15-pa-5 6.05 0.01 5.87 0.00 0.00 0.05 0.00 0.03 0.03 1.80 0.13 13.98

CWT-15-pa-6 6.05 0.01 5.84 0.00 0.00 0.07 0.00 0.04 0.03 1.78 0.19 14.00

CWT-15-pa-7 6.07 0.01 5.77 0.00 0.00 0.09 0.00 0.15 0.03 1.67 0.17 13.95

CWT-15-pa-8 6.04 0.01 5.87 0.00 0.00 0.06 0.00 0.03 0.03 1.83 0.11 13.98

CWT-15-pa-9 6.02 0.01 5.91 0.00 0.00 0.06 0.00 0.03 0.03 1.86 0.09 14.00

CWT-13-mu-1 50.51 0.28 26.49 0.00 0.00 2.58 0.00 3.40 0.02 0.71 10.13 94.12

CWT-13-mu-2 50.96 0.26 25.83 0.00 0.00 2.40 0.00 3.71 0.02 0.60 10.43 94.22

CWT-13-mu-3 51.31 0.29 25.97 0.00 0.00 2.44 0.00 3.73 0.01 0.58 10.44 94.76

CWT-13-mu-4 51.55 0.28 25.86 0.00 0.00 2.43 0.00 3.69 0.02 0.57 10.35 94.75

CWT-13-mu-5 51.41 0.29 25.93 0.00 0.00 2.39 0.00 3.62 0.01 0.58 10.34 94.58

CWT-13-mu-6 51.50 0.25 25.93 0.00 0.00 2.43 0.01 3.70 0.01 0.59 10.45 94.87

CWT-13-mu-7 51.22 0.29 26.17 0.00 0.00 2.41 0.00 3.59 0.01 0.62 10.34 94.65

CWT-13-mu-8 50.86 0.28 26.21 0.00 0.00 2.40 0.00 3.55 0.01 0.64 10.38 94.32

CWT-13-mu-9 51.03 0.27 26.19 0.00 0.00 2.40 0.01 3.57 0.01 0.64 10.35 94.45

CWT-13-mu-10 50.76 0.30 26.24 0.00 0.00 2.38 0.01 3.55 0.00 0.62 10.39 94.25

CWT-13-mu-11 51.02 0.23 26.18 0.00 0.00 2.57 0.00 3.54 0.01 0.71 10.23 94.47

CWT-13-mu-12 51.36 0.27 26.16 0.00 0.00 2.62 0.00 3.55 0.02 0.70 10.22 94.91

CWT-13-pa-13 47.04 0.11 36.98 0.00 0.00 0.83 0.00 0.32 0.16 5.83 1.54 92.81

CWT-13-pa-14 47.44 0.12 37.72 0.00 0.00 0.81 0.00 0.25 0.17 5.80 1.25 93.57

CWT-13-pa-15 47.05 0.15 37.51 0.00 0.00 0.84 0.00 0.28 0.19 6.37 1.32 93.71

CWT-13-pa-16 47.62 0.12 37.43 0.00 0.00 0.89 0.01 0.32 0.31 6.22 1.54 94.47

CWT-13-pa-17 47.30 0.12 37.63 0.00 0.00 0.76 0.00 0.19 0.23 6.55 0.93 93.70

C.p.f.u.

CWT-13-mu-1 6.83 0.03 4.22 0.00 0.00 0.29 0.00 0.69 0.01 0.19 1.75 14.00

CWT-13-mu-2 6.89 0.03 4.11 0.00 0.00 0.27 0.00 0.75 0.01 0.16 1.80 14.01

CWT-13-mu-3 6.89 0.03 4.11 0.00 0.00 0.28 0.00 0.75 0.00 0.15 1.79 13.99

CWT-13-mu-4 6.92 0.03 4.09 0.00 0.00 0.28 0.00 0.74 0.00 0.15 1.77 13.97

CWT-13-mu-5 6.91 0.03 4.11 0.00 0.00 0.27 0.00 0.73 0.00 0.15 1.77 13.97

CWT-13-mu-6 6.91 0.03 4.10 0.00 0.00 0.28 0.00 0.74 0.00 0.15 1.79 13.99

CWT-13-mu-7 6.88 0.03 4.15 0.00 0.00 0.27 0.00 0.72 0.00 0.16 1.77 13.98

CWT-13-mu-8 6.86 0.03 4.17 0.00 0.00 0.27 0.00 0.72 0.00 0.17 1.79 14.00

212

SiO2

TiO2

Al2

O3

Cr2

O3

Fe2

O3

FeO MnO MgO CaO Na2

O K2

O Totals

CWT-13-mu-9 6.88 0.03 4.16 0.00 0.00 0.27 0.00 0.72 0.00 0.17 1.78 13.99

CWT-13-mu-10 6.86 0.03 4.18 0.00 0.00 0.27 0.00 0.72 0.00 0.16 1.79 14.00

CWT-13-mu-11 6.88 0.02 4.16 0.00 0.00 0.29 0.00 0.71 0.00 0.19 1.76 14.00

CWT-13-mu-12 6.89 0.03 4.14 0.00 0.00 0.29 0.00 0.71 0.01 0.18 1.75 13.99

CWT-13-pa-13 6.18 0.01 5.72 0.00 0.00 0.09 0.00 0.06 0.02 1.49 0.26 13.82

CWT-13-pa-14 6.16 0.01 5.78 0.00 0.00 0.09 0.00 0.05 0.02 1.46 0.21 13.77

CWT-13-pa-15 6.13 0.02 5.75 0.00 0.00 0.09 0.00 0.06 0.03 1.61 0.22 13.90

CWT-13-pa-16 6.15 0.01 5.70 0.00 0.00 0.10 0.00 0.06 0.04 1.56 0.25 13.89

CWT-13-pa-17 6.14 0.01 5.76 0.00 0.00 0.08 0.00 0.04 0.03 1.65 0.15 13.87

ASA-06a-mu-1 51.21 0.34 27.13 0.00 0.00 2.76 0.01 3.44 0.03 0.36 10.77 96.06

ASA-06a-mu-2 51.34 0.33 26.63 0.01 0.00 3.04 0.00 3.73 0.02 0.30 10.95 96.36

ASA-06a-mu-3 51.23 0.23 26.25 0.00 0.00 3.06 0.00 3.85 0.02 0.29 11.09 96.03

ASA-06a-mu-4 51.13 0.23 25.70 0.00 0.01 2.90 0.00 3.98 0.03 0.25 10.77 95.01

ASA-06a-mu-5 51.51 0.32 26.29 0.01 0.00 2.96 0.00 3.76 0.02 0.28 10.98 96.14

ASA-06a-mu-6 49.86 0.35 28.13 0.00 0.95 2.51 0.02 3.22 0.02 0.43 10.78 96.18

ASA-06a-mu-7 49.84 0.37 27.88 0.02 0.67 2.99 0.02 3.16 0.03 0.62 10.64 96.19

ASA-06a-mu-8 50.46 0.35 27.94 0.02 0.45 3.02 0.00 3.10 0.04 0.47 10.52 96.33

ASA-06a-mu-9 50.12 0.42 27.45 0.00 0.13 3.36 0.01 3.08 0.04 0.54 10.42 95.57

ASA-06a-mu-10 48.75 0.45 29.27 0.01 0.00 3.09 0.01 2.56 0.05 0.53 10.60 95.33

C.p.f.u.

ASA-06a-mu-1 6.80 0.03 4.25 0.00 0.00 0.31 0.00 0.68 0.00 0.09 1.83 14.00

ASA-06a-mu-2 6.82 0.03 4.17 0.00 0.00 0.34 0.00 0.74 0.00 0.08 1.86 14.03

ASA-06a-mu-3 6.83 0.02 4.13 0.00 0.00 0.34 0.00 0.77 0.00 0.08 1.89 14.06

ASA-06a-mu-4 6.88 0.02 4.07 0.00 0.00 0.32 0.00 0.80 0.00 0.07 1.85 14.02

ASA-06a-mu-5 6.85 0.03 4.12 0.00 0.00 0.33 0.00 0.75 0.00 0.07 1.86 14.02

ASA-06a-mu-6 6.64 0.04 4.41 0.00 0.10 0.28 0.00 0.64 0.00 0.11 1.83 14.05

ASA-06a-mu-7 6.65 0.04 4.38 0.00 0.07 0.33 0.00 0.63 0.00 0.16 1.81 14.08

ASA-06a-mu-8 6.70 0.03 4.37 0.00 0.05 0.33 0.00 0.61 0.01 0.12 1.78 14.01

ASA-06a-mu-9 6.72 0.04 4.34 0.00 0.02 0.37 0.00 0.62 0.01 0.14 1.78 14.03

ASA-06a-mu-10 6.55 0.05 4.64 0.00 0.00 0.35 0.00 0.51 0.01 0.14 1.82 14.06

ASA-06b-mu-1 51.35 0.20 27.49 0.02 0.00 2.26 0.02 3.39 0.03 0.43 10.46 95.66

ASA-06b-mu-2 51.56 0.26 27.99 0.02 0.00 2.38 0.00 3.25 0.02 0.43 10.64 96.56

ASA-06b-mu-3 50.66 0.23 27.30 0.03 0.00 2.07 0.02 3.33 0.03 0.48 10.39 94.55

ASA-06b-mu-4 51.34 0.22 27.51 0.02 0.00 1.97 0.01 3.36 0.02 0.44 10.52 95.42

ASA-06b-mu-5 50.75 0.21 27.22 0.03 0.00 2.08 0.01 3.36 0.04 0.40 10.41 94.52

ASA-06b-mu-6 49.64 0.28 28.11 0.00 0.00 2.37 0.01 2.81 0.02 0.41 10.58 94.24

ASA-06b-mu-7 50.87 0.25 27.44 0.03 0.00 2.12 0.01 3.30 0.03 0.49 10.77 95.32

ASA-06b-mu-8 50.92 0.26 27.64 0.04 0.00 2.35 0.00 3.25 0.03 0.44 10.47 95.41

ASA-06b-mu-9 50.60 0.30 27.57 0.03 0.00 2.37 0.03 2.94 0.05 0.38 10.52 94.80

ASA-06b-mu-10 48.84 0.26 27.00 0.03 0.00 2.49 0.02 2.76 0.03 0.34 10.36 92.14

C.p.f.u.

ASA-06b-mu-1 6.82 0.02 4.30 0.00 0.00 0.25 0.00 0.67 0.00 0.11 1.77 13.95

ASA-06b-mu-2 6.79 0.03 4.34 0.00 0.00 0.26 0.00 0.64 0.00 0.11 1.79 13.96

ASA-06b-mu-3 6.80 0.02 4.32 0.00 0.00 0.23 0.00 0.67 0.00 0.12 1.78 13.96

ASA-06b-mu-4 6.82 0.02 4.31 0.00 0.00 0.22 0.00 0.67 0.00 0.11 1.78 13.95

ASA-06b-mu-5 6.82 0.02 4.31 0.00 0.00 0.23 0.00 0.67 0.01 0.10 1.78 13.95

ASA-06b-mu-6 6.71 0.03 4.48 0.00 0.00 0.27 0.00 0.57 0.00 0.11 1.82 13.99

ASA-06b-mu-7 6.79 0.03 4.32 0.00 0.00 0.24 0.00 0.66 0.00 0.13 1.83 14.00

ASA-06b-mu-8 6.79 0.03 4.34 0.00 0.00 0.26 0.00 0.65 0.00 0.11 1.78 13.96

ASA-06b-mu-9 6.79 0.03 4.36 0.00 0.00 0.27 0.00 0.59 0.01 0.10 1.80 13.95

ASA-06b-mu-10 6.76 0.03 4.40 0.00 0.00 0.29 0.00 0.57 0.00 0.09 1.83 13.97

213

SiO2

TiO2

Al2

O3

Cr2

O3

Fe2

O3

FeO MnO MgO CaO Na2

O K2

O Totals

CWT-5-mu-1 50.19 0.24 26.67 0.00 0.00 2.67 0.00 3.18 0.02 0.44 10.54 93.96

CWT-5-mu-2 50.59 0.23 26.29 0.00 0.00 2.58 0.00 3.34 0.00 0.41 10.49 93.94

CWT-5-mu-3 50.46 0.23 26.26 0.00 0.00 2.66 0.00 3.31 0.01 0.36 10.62 93.92

CWT-5-mu-4 50.75 0.29 24.79 0.00 0.00 3.36 0.01 3.54 0.01 0.19 10.83 93.77

CWT-5-mu-5 50.66 0.20 24.76 0.00 0.00 3.42 0.00 3.50 0.00 0.19 10.89 93.62

CWT-5-mu-6 50.32 0.21 25.84 0.00 0.00 2.81 0.01 3.37 0.01 0.28 10.72 93.57

CWT-5-mu-7 50.47 0.21 26.24 0.00 0.00 2.63 0.00 3.25 0.00 0.38 10.59 93.78

CWT-5-mu-8 50.04 0.22 26.37 0.00 0.00 2.87 0.00 3.17 0.00 0.42 10.46 93.56

CWT-5-mu-9 49.82 0.29 26.70 0.00 0.00 3.11 0.00 3.07 0.01 0.45 10.48 93.94

C.p.f.u.

CWT-5-mu-1 6.81 0.02 4.27 0.00 0.00 0.30 0.00 0.64 0.00 0.12 1.83 14.00

CWT-5-mu-2 6.86 0.02 4.20 0.00 0.00 0.29 0.00 0.68 0.00 0.11 1.82 13.98

CWT-5-mu-3 6.85 0.02 4.20 0.00 0.00 0.30 0.00 0.67 0.00 0.09 1.84 13.99

CWT-5-mu-4 6.94 0.03 3.99 0.00 0.00 0.39 0.00 0.72 0.00 0.05 1.89 14.01

CWT-5-mu-5 6.94 0.02 4.00 0.00 0.00 0.39 0.00 0.72 0.00 0.05 1.90 14.02

CWT-5-mu-6 6.87 0.02 4.16 0.00 0.00 0.32 0.00 0.69 0.00 0.07 1.86 14.00

CWT-5-mu-7 6.86 0.02 4.21 0.00 0.00 0.30 0.00 0.66 0.00 0.10 1.84 13.99

CWT-5-mu-8 6.83 0.02 4.24 0.00 0.00 0.33 0.00 0.64 0.00 0.11 1.82 14.00

CWT-5-mu-9 6.78 0.03 4.28 0.00 0.00 0.35 0.00 0.62 0.00 0.12 1.82 14.01

CWT-7-mu-1 49.68 0.24 28.49 0.00 0.00 2.50 0.00 2.56 0.03 0.59 9.98 94.08

CWT-7-mu-2 50.25 0.33 27.84 0.00 0.00 2.58 0.02 2.70 0.03 0.45 10.11 94.31

CWT-7-mu-3 50.33 0.21 27.67 0.00 0.00 2.44 0.01 2.78 0.02 0.47 10.15 94.09

CWT-7-mu-4 50.48 0.23 27.51 0.00 0.00 2.36 0.00 2.78 0.02 0.45 10.17 94.01

CWT-7-mu-5 50.28 0.23 27.49 0.00 0.00 2.33 0.00 2.86 0.02 0.44 10.18 93.84

CWT-7-mu-6 50.45 0.21 27.52 0.00 0.00 2.31 0.00 2.89 0.02 0.43 10.18 94.01

CWT-7-mu-7 50.22 0.24 27.86 0.00 0.00 2.14 0.01 2.81 0.02 0.49 10.17 93.96

CWT-7-mu-8 49.92 0.25 28.52 0.00 0.00 2.32 0.01 2.52 0.03 0.56 10.11 94.23

CWT-7-mu-9 49.72 0.26 28.25 0.00 0.00 2.51 0.01 2.61 0.02 0.55 10.08 94.01

CWT-7-mu-10 49.66 0.24 28.37 0.00 0.00 2.45 0.01 2.47 0.02 0.57 10.09 93.88

CWT-7-mu-11 49.65 0.28 28.60 0.00 0.00 2.56 0.00 2.54 0.04 0.64 10.05 94.35

C.p.f.u.

CWT-7-mu-1 6.70 0.02 4.53 0.00 0.00 0.28 0.00 0.52 0.01 0.15 1.72 13.94

CWT-7-mu-2 6.77 0.03 4.42 0.00 0.00 0.29 0.00 0.54 0.01 0.12 1.74 13.92

CWT-7-mu-3 6.79 0.02 4.40 0.00 0.00 0.28 0.00 0.56 0.01 0.12 1.75 13.92

CWT-7-mu-4 6.81 0.02 4.37 0.00 0.00 0.26 0.00 0.56 0.01 0.12 1.75 13.91

CWT-7-mu-5 6.80 0.02 4.38 0.00 0.00 0.26 0.00 0.58 0.01 0.12 1.75 13.92

CWT-7-mu-6 6.81 0.02 4.38 0.00 0.00 0.26 0.00 0.58 0.01 0.12 1.75 13.92

CWT-7-mu-7 6.78 0.02 4.43 0.00 0.00 0.24 0.00 0.57 0.00 0.13 1.75 13.92

CWT-7-mu-8 6.72 0.03 4.53 0.00 0.00 0.26 0.00 0.51 0.01 0.15 1.74 13.93

CWT-7-mu-9 6.72 0.03 4.50 0.00 0.00 0.29 0.00 0.53 0.01 0.14 1.74 13.94

CWT-7-mu-10 6.72 0.02 4.53 0.00 0.00 0.28 0.00 0.50 0.01 0.15 1.74 13.94

CWT-7-mu-11 6.69 0.03 4.54 0.00 0.00 0.29 0.00 0.51 0.01 0.17 1.73 13.96

ASA-84b-pa-1 47.23 0.07 39.52 0.04 0.00 0.43 0.00 0.16 0.42 7.24 0.79 95.90

ASA-84b-pa-2 46.59 0.07 39.08 0.01 0.00 0.36 0.00 0.18 0.41 7.07 0.92 94.69

ASA-84b-pa-3 48.12 0.07 40.37 0.03 0.00 0.34 0.01 0.12 0.40 7.41 0.76 97.63

ASA-84b-pa-4 46.35 0.07 38.52 0.02 0.00 0.41 0.00 0.14 0.36 6.34 0.69 92.90

ASA-84b-pa-5 47.34 0.07 39.06 0.03 0.00 0.49 0.00 0.20 0.32 7.20 0.89 95.60

ASA-84b-pa-6 47.06 0.07 39.40 0.02 0.00 0.35 0.01 0.12 0.40 7.04 0.63 95.10

ASA-84b-pa-7 47.10 0.08 39.63 0.03 0.00 0.47 0.00 0.15 0.38 7.26 0.74 95.84

ASA-84b-pa-8 47.38 0.07 39.82 0.03 0.00 0.59 0.00 0.15 0.39 7.30 0.74 96.47

214

SiO2

TiO2

Al2

O3

Cr2

O3

Fe2

O3

FeO MnO MgO CaO Na2

O K2

O Totals

ASA-84b-pa-9 46.91 0.07 39.19 0.05 0.00 0.49 0.00 0.17 0.35 7.31 0.89 95.43

ASA-84b-pa-10 46.94 0.09 38.98 0.04 0.00 0.82 0.00 0.17 0.30 7.10 0.78 95.22

C.p.f.u.

ASA-84b-pa-1 6.00 0.01 5.92 0.00 0.00 0.05 0.00 0.03 0.06 1.78 0.13 13.98

ASA-84b-pa-2 6.00 0.01 5.93 0.00 0.00 0.04 0.00 0.03 0.06 1.77 0.15 13.99

ASA-84b-pa-3 6.00 0.01 5.94 0.00 0.00 0.04 0.00 0.02 0.05 1.79 0.12 13.98

ASA-84b-pa-4 6.05 0.01 5.93 0.00 0.00 0.04 0.00 0.03 0.05 1.61 0.11 13.83

ASA-84b-pa-5 6.04 0.01 5.87 0.00 0.00 0.05 0.00 0.04 0.04 1.78 0.14 13.98

ASA-84b-pa-6 6.02 0.01 5.94 0.00 0.00 0.04 0.00 0.02 0.05 1.75 0.10 13.93

ASA-84b-pa-7 5.99 0.01 5.94 0.00 0.00 0.05 0.00 0.03 0.05 1.79 0.12 13.98

ASA-84b-pa-8 5.99 0.01 5.94 0.00 0.00 0.06 0.00 0.03 0.05 1.79 0.12 13.99

ASA-84b-pa-9 6.00 0.01 5.91 0.01 0.00 0.05 0.00 0.03 0.05 1.81 0.15 14.01

ASA-84b-pa-10 6.02 0.01 5.89 0.00 0.00 0.09 0.00 0.03 0.04 1.76 0.13 13.97

ASA-05-mu-1 50.58 0.27 28.79 0.01 0.00 3.06 0.02 2.95 0.04 0.46 10.81 97.00

ASA-05-mu-2 50.65 0.27 27.24 0.02 0.12 3.54 0.01 3.16 0.05 0.49 10.59 96.14

ASA-05-mu-3 50.32 0.27 27.36 0.02 0.50 3.32 0.01 3.10 0.06 0.50 10.51 95.93

ASA-05-mu-4 50.86 0.28 27.66 0.01 0.15 3.60 0.00 3.08 0.02 0.50 10.71 96.86

ASA-05-mu-5 48.48 0.29 26.38 0.01 0.32 3.28 0.01 3.00 0.14 0.49 10.03 92.41

ASA-05-mu-6 50.40 0.36 27.25 0.01 0.00 3.77 0.00 3.04 0.05 0.57 10.52 95.98

ASA-05-mu-7 49.63 0.28 26.82 0.01 0.34 3.44 0.00 3.07 0.14 0.51 10.23 94.44

ASA-05-mu-8 50.76 0.28 27.57 0.00 0.24 3.43 0.02 3.10 0.05 0.54 10.49 96.47

ASA-05-mu-9 49.79 0.24 27.29 0.02 0.56 3.23 0.00 3.08 0.05 0.57 10.41 95.19

ASA-05-mu-10 49.10 0.39 28.82 0.04 0.64 2.49 0.03 2.89 0.04 0.36 10.73 95.48

ASA-05-mu-11 49.46 0.27 27.45 0.01 0.77 2.98 0.02 3.03 0.06 0.51 10.40 94.89

ASA-05-mu-12 49.40 0.30 28.33 0.02 0.81 2.94 0.00 2.91 0.04 0.72 10.37 95.77

ASA-05-mu-13 49.28 0.38 27.74 0.04 1.17 2.62 0.00 3.11 0.04 0.61 10.39 95.27

ASA-05-mu-14 49.10 0.33 28.32 0.01 0.43 2.99 0.00 2.79 0.08 0.58 10.27 94.87

ASA-05-mu-15 47.39 0.39 27.61 0.05 0.89 2.48 0.00 2.74 0.08 0.52 9.96 92.03

C.p.f.u.

ASA-05-mu-1 6.7 0.0 4.5 0.0 0.0 0.3 0.0 0.6 0.0 0.1 1.8 14.0

ASA-05-mu-2 6.8 0.0 4.3 0.0 0.0 0.4 0.0 0.6 0.0 0.1 1.8 14.0

ASA-05-mu-3 6.7 0.0 4.3 0.0 0.1 0.4 0.0 0.6 0.0 0.1 1.8 14.0

ASA-05-mu-4 6.7 0.0 4.3 0.0 0.0 0.4 0.0 0.6 0.0 0.1 1.8 14.0

ASA-05-mu-5 6.7 0.0 4.3 0.0 0.0 0.4 0.0 0.6 0.0 0.1 1.8 14.0

ASA-05-mu-6 6.7 0.0 4.3 0.0 0.0 0.4 0.0 0.6 0.0 0.1 1.8 14.1

ASA-05-mu-7 6.7 0.0 4.3 0.0 0.0 0.4 0.0 0.6 0.0 0.1 1.8 14.0

ASA-05-mu-8 6.7 0.0 4.3 0.0 0.0 0.4 0.0 0.6 0.0 0.1 1.8 14.0

ASA-05-mu-9 6.7 0.0 4.3 0.0 0.1 0.4 0.0 0.6 0.0 0.1 1.8 14.0

ASA-05-mu-10 6.6 0.0 4.6 0.0 0.1 0.3 0.0 0.6 0.0 0.1 1.8 14.0

ASA-05-mu-11 6.7 0.0 4.4 0.0 0.1 0.3 0.0 0.6 0.0 0.1 1.8 14.0

ASA-05-mu-12 6.6 0.0 4.5 0.0 0.1 0.3 0.0 0.6 0.0 0.2 1.8 14.1

ASA-05-mu-13 6.6 0.0 4.4 0.0 0.1 0.3 0.0 0.6 0.0 0.2 1.8 14.0

ASA-05-mu-14 6.6 0.0 4.5 0.0 0.0 0.3 0.0 0.6 0.0 0.2 1.8 14.0

ASA-05-mu-15 6.6 0.0 4.5 0.0 0.1 0.3 0.0 0.6 0.0 0.1 1.8 14.0

ASA-35b-mu-1 49.31 0.30 26.97 0.00 0.00 2.75 0.00 3.09 0.08 0.45 10.33 93.30

ASA-35b-mu-2 49.48 0.21 27.57 0.00 0.00 2.45 0.00 2.90 0.04 0.50 10.47 93.62

ASA-35b-mu-3 50.06 0.21 27.06 0.00 0.00 2.37 0.02 3.23 0.03 0.50 10.55 94.04

ASA-35b-mu-4 49.93 0.21 27.29 0.00 0.00 2.09 0.00 3.13 0.03 0.50 10.37 93.56

ASA-35b-mu-5 49.98 0.27 27.27 0.00 0.00 1.66 0.00 3.26 0.03 0.54 10.29 93.31

ASA-35b-mu-6 48.78 0.42 26.15 0.00 0.00 3.38 0.00 4.26 0.05 0.51 10.34 93.88

ASA-35b-pa-7 47.16 0.04 38.56 0.00 0.00 0.48 0.01 0.21 0.16 7.33 0.78 94.72

215

SiO2

TiO2

Al2

O3

Cr2

O3

Fe2

O3

FeO MnO MgO CaO Na2

O K2

O Totals

ASA-35b-pa-8 46.62 0.06 39.08 0.00 0.00 0.34 0.00 0.11 0.22 7.54 0.70 94.68

ASA-35b-pa-9 47.14 0.04 38.55 0.00 0.00 0.40 0.00 0.16 0.20 7.29 0.79 94.57

ASA-35b-pa-10 46.89 0.08 38.67 0.00 0.00 0.42 0.01 0.15 0.19 7.49 0.84 94.74

ASA-35b-pa-11 47.07 0.06 38.66 0.00 0.00 0.36 0.00 0.13 0.19 7.47 0.63 94.56

ASA-35b-pa-12 46.96 0.06 38.08 0.00 0.00 0.45 0.00 0.22 0.17 6.98 1.08 93.99

ASA-35b-pa-13 46.92 0.07 38.27 0.00 0.00 0.44 0.02 0.21 0.18 7.21 1.07 94.38

ASA-35b-pa-14 47.22 0.10 38.19 0.00 0.00 0.46 0.00 0.25 0.22 7.03 1.04 94.51

ASA-35b-pa-15 46.80 0.07 38.78 0.00 0.00 0.34 0.00 0.10 0.23 7.44 0.54 94.32

C.p.f.u.

ASA-35b-mu-1 6.75 0.03 4.35 0.00 0.00 0.31 0.00 0.63 0.01 0.12 1.80 14.01

ASA-35b-mu-2 6.74 0.02 4.42 0.00 0.00 0.28 0.00 0.59 0.01 0.13 1.82 14.01

ASA-35b-mu-3 6.78 0.02 4.32 0.00 0.00 0.27 0.00 0.65 0.01 0.13 1.83 14.01

ASA-35b-mu-4 6.78 0.02 4.37 0.00 0.00 0.24 0.00 0.63 0.01 0.13 1.80 13.98

ASA-35b-mu-5 6.79 0.03 4.37 0.00 0.00 0.19 0.00 0.66 0.01 0.14 1.78 13.96

ASA-35b-mu-6 6.68 0.04 4.22 0.00 0.00 0.39 0.00 0.87 0.01 0.13 1.80 14.14

ASA-35b-pa-7 6.07 0.01 5.85 0.00 0.00 0.05 0.00 0.04 0.02 1.83 0.13 13.99

ASA-35b-pa-8 6.00 0.01 5.93 0.00 0.00 0.04 0.00 0.02 0.03 1.88 0.12 14.03

ASA-35b-pa-9 6.07 0.01 5.85 0.00 0.00 0.04 0.00 0.03 0.03 1.82 0.13 13.98

ASA-35b-pa-10 6.04 0.01 5.87 0.00 0.00 0.04 0.00 0.03 0.03 1.87 0.14 14.03

ASA-35b-pa-11 6.06 0.01 5.86 0.00 0.00 0.04 0.00 0.03 0.03 1.86 0.10 13.99

ASA-35b-pa-12 6.09 0.01 5.82 0.00 0.00 0.05 0.00 0.04 0.02 1.75 0.18 13.96

ASA-35b-pa-13 6.07 0.01 5.83 0.00 0.00 0.05 0.00 0.04 0.03 1.81 0.18 14.00

ASA-35b-pa-14 6.09 0.01 5.80 0.00 0.00 0.05 0.00 0.05 0.03 1.75 0.17 13.96

ASA-35b-pa-15 6.04 0.01 5.90 0.00 0.00 0.04 0.00 0.02 0.03 1.86 0.09 13.98

N45b-mu-1 49.91 0.18 26.24 0.00 0.58 3.96 0.03 3.11 0.04 0.36 10.99 95.36

N45b-mu-2 50.25 0.18 26.13 0.02 0.32 4.33 0.00 3.05 0.02 0.34 11.04 95.65

N45b-mu-3 50.60 0.18 26.75 0.01 0.54 4.17 0.00 3.01 0.01 0.36 11.13 96.72

N45b-mu-4 49.92 0.19 26.21 0.00 0.68 4.14 0.02 2.94 0.01 0.37 10.89 95.31

N45b-mu-5 49.83 0.20 27.13 0.02 1.06 3.80 0.03 2.81 0.03 0.46 10.83 96.10

N45b-mu-6 49.21 0.23 26.40 0.03 0.95 4.22 0.02 2.68 0.01 0.51 10.69 94.87

N45b-mu-7 49.93 0.17 26.79 0.01 0.36 4.25 0.00 2.84 0.01 0.46 10.89 95.69

N45b-mu-8 49.80 0.22 26.56 0.00 0.19 4.35 0.04 2.80 0.01 0.39 10.89 95.24

N45b-mu-9 49.06 0.18 26.33 0.01 0.00 3.67 0.02 3.06 0.03 0.31 10.81 93.49

N45b-mu-10 47.22 0.06 38.14 0.00 0.00 1.56 0.00 0.16 0.04 7.51 0.44 95.13

C.p.f.u.

N45b-mu-1 6.76 0.02 4.19 0.00 0.06 0.45 0.00 0.63 0.01 0.09 1.90 14.10

N45b-mu-2 6.79 0.02 4.16 0.00 0.03 0.49 0.00 0.61 0.00 0.09 1.90 14.10

N45b-mu-3 6.75 0.02 4.21 0.00 0.06 0.46 0.00 0.60 0.00 0.09 1.90 14.09

N45b-mu-4 6.76 0.02 4.19 0.00 0.07 0.47 0.00 0.59 0.00 0.10 1.88 14.08

N45b-mu-5 6.69 0.02 4.29 0.00 0.11 0.42 0.00 0.56 0.00 0.12 1.85 14.08

N45b-mu-6 6.71 0.02 4.24 0.00 0.10 0.48 0.00 0.54 0.00 0.13 1.86 14.10

N45b-mu-7 6.73 0.02 4.26 0.00 0.04 0.48 0.00 0.57 0.00 0.12 1.87 14.10

N45b-mu-8 6.75 0.02 4.24 0.00 0.02 0.49 0.00 0.57 0.00 0.10 1.88 14.09

N45b-mu-9 6.75 0.02 4.27 0.00 0.00 0.42 0.00 0.63 0.00 0.08 1.90 14.08

N45b-mu-10 6.07 0.01 5.78 0.00 0.00 0.17 0.00 0.03 0.01 1.87 0.07 14.01

CWT-19-mu-1 47.72 0.43 31.93 0.00 0.00 2.11 0.00 1.59 0.03 0.84 9.91 94.55

CWT-19-mu-2 49.31 0.27 29.47 0.00 0.00 2.18 0.00 2.58 0.03 0.63 10.09 94.56

CWT-19-mu-3 49.33 0.21 29.69 0.00 0.00 2.11 0.00 2.46 0.03 0.64 10.12 94.59

CWT-19-mu-4 50.83 0.20 28.06 0.00 0.00 1.87 0.00 3.10 0.02 0.51 10.02 94.61

CWT-19-mu-5 50.96 0.21 27.77 0.00 0.00 1.80 0.01 3.13 0.02 0.47 9.97 94.34

CWT-19-mu-6 51.12 0.21 27.90 0.00 0.00 1.83 0.00 3.12 0.03 0.47 10.01 94.70

216

SiO2

TiO2

Al2

O3

Cr2

O3

Fe2

O3

FeO MnO MgO CaO Na2

O K2

O Totals

CWT-19-mu-7 51.32 0.22 27.83 0.00 0.00 1.81 0.01 3.08 0.02 0.45 9.99 94.73

CWT-19-mu-8 51.02 0.19 27.67 0.00 0.00 1.83 0.01 3.11 0.01 0.45 10.07 94.36

CWT-19-mu-9 51.06 0.22 27.77 0.00 0.00 1.84 0.00 3.08 0.02 0.45 10.01 94.45

CWT-19-mu-10 50.92 0.24 27.81 0.00 0.00 1.82 0.00 3.03 0.03 0.46 10.04 94.35

CWT-19-mu-11 50.76 0.23 27.70 0.00 0.00 1.86 0.01 3.10 0.02 0.45 10.10 94.23

CWT-19-mu-12 50.93 0.24 27.92 0.00 0.00 1.85 0.01 3.04 0.02 0.51 9.98 94.50

CWT-19-mu-13 51.11 0.24 27.70 0.00 0.00 1.89 0.00 3.15 0.02 0.52 10.07 94.69

CWT-19-mu-14 50.90 0.23 27.79 0.00 0.00 1.85 0.01 3.09 0.02 0.52 9.99 94.39

CWT-19-mu-15 51.14 0.23 27.54 0.00 0.00 1.86 0.01 3.18 0.02 0.46 10.06 94.48

CWT-19-mu-16 51.02 0.22 27.90 0.00 0.00 1.89 0.01 3.13 0.02 0.48 10.00 94.67

CWT-19-mu-17 50.86 0.23 27.83 0.00 0.00 1.83 0.00 3.16 0.02 0.50 10.06 94.49

CWT-19-mu-18 51.16 0.21 27.73 0.00 0.00 1.57 0.01 3.41 0.00 0.51 10.05 94.65

CWT-19-mu-19 49.51 0.21 28.99 0.00 0.00 1.66 0.00 2.65 0.02 0.52 10.09 93.64

CWT-19-mu-20 50.87 0.20 27.52 0.00 0.00 1.94 0.02 3.32 0.00 0.48 10.06 94.39

C.p.f.u.

CWT-19-mu-1 6.41 0.04 5.05 0.00 0.00 0.24 0.00 0.32 0.01 0.22 1.70 13.98

CWT-19-mu-2 6.62 0.03 4.66 0.00 0.00 0.24 0.00 0.52 0.01 0.17 1.73 13.97

CWT-19-mu-3 6.62 0.02 4.69 0.00 0.00 0.24 0.00 0.49 0.01 0.17 1.73 13.96

CWT-19-mu-4 6.79 0.02 4.42 0.00 0.00 0.21 0.00 0.62 0.01 0.13 1.71 13.90

CWT-19-mu-5 6.82 0.02 4.38 0.00 0.00 0.20 0.00 0.62 0.01 0.12 1.71 13.88

CWT-19-mu-6 6.82 0.02 4.38 0.00 0.00 0.20 0.00 0.62 0.01 0.12 1.71 13.88

CWT-19-mu-7 6.84 0.02 4.37 0.00 0.00 0.20 0.00 0.61 0.00 0.12 1.70 13.86

CWT-19-mu-8 6.83 0.02 4.37 0.00 0.00 0.20 0.00 0.62 0.00 0.12 1.72 13.88

CWT-19-mu-9 6.83 0.02 4.38 0.00 0.00 0.20 0.00 0.62 0.00 0.12 1.71 13.87

CWT-19-mu-10 6.82 0.02 4.39 0.00 0.00 0.20 0.00 0.61 0.01 0.12 1.72 13.88

CWT-19-mu-11 6.81 0.02 4.38 0.00 0.00 0.21 0.00 0.62 0.00 0.12 1.73 13.90

CWT-19-mu-12 6.81 0.02 4.40 0.00 0.00 0.21 0.00 0.61 0.01 0.13 1.70 13.88

CWT-19-mu-13 6.83 0.02 4.36 0.00 0.00 0.21 0.00 0.63 0.00 0.14 1.72 13.90

CWT-19-mu-14 6.81 0.02 4.38 0.00 0.00 0.21 0.00 0.62 0.00 0.14 1.71 13.89

CWT-19-mu-15 6.84 0.02 4.34 0.00 0.00 0.21 0.00 0.63 0.01 0.12 1.72 13.88

CWT-19-mu-16 6.81 0.02 4.39 0.00 0.00 0.21 0.00 0.62 0.01 0.13 1.71 13.89

CWT-19-mu-17 6.80 0.02 4.39 0.00 0.00 0.20 0.00 0.63 0.01 0.13 1.72 13.90

CWT-19-mu-18 6.82 0.02 4.36 0.00 0.00 0.18 0.00 0.68 0.00 0.13 1.71 13.90

CWT-19-mu-19 6.69 0.02 4.61 0.00 0.00 0.19 0.00 0.53 0.01 0.14 1.74 13.92

CWT-19-mu-20 6.82 0.02 4.35 0.00 0.00 0.21 0.00 0.67 0.00 0.13 1.72 13.91

Table A.4: 40Ar/39Ar-sample mica chemistry. Data divided into weight% and cations per formula unit (calculated on

basis of 22 O).

217

A.3 Whole–rock REE ICPMS

Whole rock REE analyses were performed using a Perkin Elmer Elan DRCII ICPMS at the Uni-versity of Cambridge. After weighing and drying, approximately 0.1 grams of sample powder wasdigested in a mixture of HF and HNO3 before evaporation. Residual salts were taken up in 2%HNO3. Machine drift was monitored via repeat analysis of the USGS international standard –BCR–2 (Wilson, 1997). Percentage recovery against known standard values was between 97–103%for the REEs.

TH–680

Y 30.62Zr† 8.02La 32.09Ce 102.41Pr 10.37Nd 46.88Sm 13.95Eu 3.32Gd 13.53Tb 1.78Dy 7.88Ho 1.13Er 2.42Tm 0.26Yb 1.39Lu 0.17Hf 0.20Pb 40.40Th 12.76U 6.83

Table A.5: TH–680 bulk rock ICPMS values for REEs and selected trace elements. Analyses presented are corrected for

detrital zircon component assuming that all Zr (†) measured derives from detrital zircon and that the detrital zircon is of

the same composition as that reported by Miller et al. (2007; zircon C1-3 core).

A.4 U–Th–Pb and 40Ar/39Ar single grain fusion data

218

Sam

ple

saU

bP

bb

Th

U

206P

b

204P

b

1�%

207P

b

206P

b

1�%

206P

b

238U

1�%

207P

b

235U

1�%

⇢c

208P

b

232T

h

1�%

f 206

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all1

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311

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0.00

012

11.0

60.

970.

7044

67.8

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7078

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1.52

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all1

-222

111

626

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0.15

0.81

060.

030.

4098

1.97

45.8

01.

971

0.00

025

9.25

0.98

0.62

4467

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18.8

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9.62

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4.67

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0.75

50.4

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0.68

153.

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8.38

7.29

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all1

-781

110

29.3

18.7

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0.06

5.67

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3764

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85.6

276

29.6

114

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8579

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3.77

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41.

47

Tab

leA

.6:

U–T

h–Pb

isot

ope

data

for

alla

nite

–clin

ozoi

site

from

TH

–680

,an

alyz

edby

Nu

Pla

sma

LA

-MC–I

CPM

S

aLab

elle

dfo

rth

inse

ctio

n(X

1),al

lanit

e–cl

inoz

oisi

tegr

ain

and

spot

num

ber

.bA

ppro

xim

ate

conce

ntr

atio

ns

to±

10%

-bas

edon

calibra

tion

agai

nst

zirc

onst

andar

d91

500

(Wie

den

bec

ket

al.,

1995

).c206

Pb/2

38

U-2

07

Pb/2

35

Uco

rrel

atio

nco

e�ci

ent

d206

Pb

c/2

06

Pb

-ca

lcula

ted

follow

ing

met

hod

ofG

rego

ryet

al.(2

007)

eC

alcu

late

das

sum

ing

conco

rdan

cybet

wee

n235

U-2

38

Usy

stem

s

219

Fra

ctio

na

Weig

htb

Uc

Pb

cT

hU

dP

b⇤

Pbc

eP

bce

206P

b204P

b

f208P

b206P

b

g207P

b206P

b

h2�

i204P

b206P

b

h2�

i238U

204P

b

j2�

i238U

206P

b

h2�

i207P

b235U

j2�

i⇢

k 8�

7

⇢k 8�

4

⇢k 7�

4

X1

all.2-

10.

005

32.2

79.6

1.76

0.01

0637

1.23

18.9

281.

699

0.82

780.

0837

0.05

280.

1658

26.1

60.

341.

380.

2982

.58

0.39

-0.3

50.

23-0

.91

X1

all.2-

20.

005

31.7

30.6

3.54

0.02

5414

0.62

19.1

742.

474

0.81

780.

0884

0.05

210.

1696

68.3

90.

733.

570.

7131

.62

0.77

-0.4

4-0

.04

-0.7

8X

1al

l.2-

40.

002

139.

421

6.4

21.1

50.

0264

331.

3818

.969

5.05

70.

8245

0.08

830.

0527

0.18

7542

.33

0.37

2.23

0.32

50.9

30.

43-0

.33

0.18

-0.7

6X

1al

l.4-

10.

003

2.2

15.3

4.71

0.00

7847

.88

18.8

322.

086

0.83

020.

1161

0.05

310.

1861

9.52

2.58

0.51

2.57

226.

412.

62-0

.13

-0.1

3-0

.33

X3

all.3-

10.

004

113.

912

9.9

3.69

0.02

3844

8.64

19.1

282.

579

0.81

960.

0804

0.05

230.

1748

57.4

50.

303.

000.

2537

.62

0.36

-0.4

20.

26-0

.90

X3

all.3-

20.

001

774.

752

5.8

3.15

0.02

9845

1.4

19.3

292.

151

0.81

210.

0798

0.05

170.

1667

97.0

70.

295.

020.

2322

.30

0.35

-0.4

70.

28-0

.95

X3

all.3-

30.

008

30.0

28.7

2.90

0.02

2323

8.11

19.1

482.

199

0.81

910.

0851

0.05

220.

1990

67.6

50.

473.

530.

4331

.97

0.53

-0.3

70.

09-0

.74

X3

all.3-

40.

003

96.0

912.

460.

0279

250.

419

.218

2.47

30.

8177

0.08

200.

0520

0.18

6469

.80

0.44

3.63

0.40

31.0

50.

50-0

.40

0.10

-0.8

1

Tab

leA

.7:

aLab

elle

dfo

rth

inse

ctio

n(X

1),al

lani

tegr

ain

and

shar

dnu

mber

.b

Fra

ctio

nwei

ghts

calc

ulat

edus

ing

phot

mic

rogr

aph

grai

ndi

men

sion

mea

sure

men

tsan

dan

alla

nite

dens

ity

of3.

93g.

cc�

1

(obt

aine

dfrom

Web

min

eral

-

http

://w

ebm

iner

al.c

om/d

ata/

Alla

nite

-(Ce)

.sht

ml).

cN

omin

alU

and

tota

lPb

conc

entr

atio

nssu

bjec

tto

unce

rtai

nty

inph

otom

icro

grap

hic

estim

atio

nof

wei

ght

–m

axim

umer

ror±

20%

.d

Mode

lT

h/U

ratio

calc

ulat

edfrom

radi

ogen

ic208

Pb/

206

Pb

ratio

and

207

Pb/

235

Uag

e.e

Pb*

and

Pb c

repr

esen

tra

diog

enic

and

com

mon

Pb

resp

ective

ly.

fM

easu

red

ratio

corr

ecte

dfo

rfrac

tion

atio

nan

dsp

ike

only.

gCor

rect

edfo

rfrac

tion

atio

n,sp

ike,

and

com

mon

Pb

(see

ben

eath

).h

Cor

rect

edfo

rfrac

tion

atio

n,sp

ike,

blan

kPb

and

exce

ss206

Pb

resu

ltin

gfrom

230

Th

dise

quilb

ria

(Sch

arer

,19

84).

iErr

ors

are

2�%

,pr

opag

ated

usin

gth

eal

gorith

ms

of(S

chm

itz

&Sch

oen

e,20

07).

jRat

ioco

rrec

ted

for

spik

e,frac

tion

atio

nan

dbl

ank

Pb

only.

kCor

rela

tion

coe�

cien

tof

238

U/2

06

Pb

vs.

207

Pb/

206

Pb.

lCor

rela

tion

coe�

cien

tof

238

U/2

06

Pb

vs.

204

Pb/

206

Pb.

mCor

rela

tion

coe�

cien

tof

207

Pb/

206

Pb

vs.

204

Pb/

206

Pb.

Dat

are

duct

ion

and

erro

rpr

opog

atio

npa

ram

eter

s:

1.

Isot

ope

dilu

tion

was

per

form

edus

ing

the

205

Pb-

233

Th-

235

UEA

RT

HT

IME

trac

erso

lution

:206

Pb/

205

Pb

=0.

0382±

0.74

%;

207

/205

Pb

=0.

0005

01±

0.10

%;

204

/205

Pb

=0.

0000

323±

0.10

%;

208

Pb/

205

Pb

=0.

0011

4±

0.10

%;

207

Pb/

206

Pb

=0.

0131

±0.

10%

;204

Pb/

206

Pb

=0.

0008

46±

0.10

%;

238

U/2

35

U=

0.00

136

±0.

005%

;233

Th/

235

U=

1.00

00±

0.00

5%;

238

U/2

33

Th

=0.

0013

5±

0.00

5%;al

ler

rors

are

1�.

2.

Pb

frac

tion

atio

nfa

ctor

s:Far

aday

frac

tion

atio

n=

0.11

±0.

04%

.am

u�1

;D

aly

frac

tion

atio

n=

0.14

±0.

04%

.am

u�1

;er

rors

are

1�ab

s.

3.

Pro

cedu

ralbl

ank

assu

mm

edas

1.00

±50

%pg

Pb

and

0.10

±50

%pg

U(e

rror

sar

e1�

).Exc

ess

over

blan

kas

sign

edto

intial

com

mon

Pb.

Bla

nkco

mpos

itio

n:206

Pb/

204

Pb

=18

.50±

0.50

%;

207

Pb/

204

Pb

=15

.8±

0.32

%;

208

Pb/

204

Pb

=38

.02±

0.50

%(a

ller

rors

are

1�).

4.

Initia

lPb

com

pos

itio

nat

34M

aca

lcul

ated

from

Sta

cey

and

Kra

mer

s(1

975)

:206

Pb/

204

Pb

=18

.65±

0.8%

;207

Pb/

204

Pb

=15

.63±

0.32

%;

208

Pb/

204

Pb

=38

.57

±0.

74%

(err

ors

are

1�).

220

Fra

ctio

na

208P

b232T

h

b2�

i232T

h204P

b

b2�

i208P

b204P

b

b2�

i232T

h207P

b

b2�

i208P

b207P

b

b2�

i232T

h206P

b

b2�

i208P

b206P

b

b2�

i

X1

alla

nit

e2-1

1.55

0.63

60.5

90.

6039

.04

0.33

3.87

0.65

2.49

0.41

3.20

0.63

2.06

0.37

X1

alla

nit

e2-2

13.6

60.

9754

4.89

0.88

39.8

80.

3434

.75

0.92

2.54

0.42

28.4

10.

902.

080.

38X

1al

lanit

e2-4

20.5

00.

6682

4.12

0.62

40.2

00.

3752

.70

0.68

2.57

0.46

43.4

40.

652.

120.

42X

1al

lanit

e4-1

0.65

2.88

25.4

22.

6338

.97

0.37

1.63

2.64

2.49

0.45

1.35

2.63

2.07

0.42

X3

alla

nit

e3-1

20.5

50.

5681

8.03

0.53

39.8

10.

3552

.18

0.59

2.54

0.43

42.7

60.

562.

080.

39X

3al

lanit

e3-2

9.17

0.71

367.

240.

7040

.04

0.33

23.4

00.

742.

550.

4119

.00

0.71

2.07

0.37

X3

alla

nit

e3-3

11.2

90.

7344

8.05

0.69

39.6

70.

4028

.57

0.74

2.53

0.48

23.4

00.

722.

070.

44X

3al

lanit

e3-4

13.4

60.

7153

8.34

0.67

39.9

80.

3734

.26

0.72

2.54

0.45

28.0

10.

692.

080.

42

Tab

leA

.8:

aLab

elle

dfo

rth

inse

ctio

n(X

1),al

lani

tegr

ain

and

shar

dnu

mber

.Fra

ctio

nsco

rres

pon

dto

thos

eus

edfo

rU

–Pb

mea

sure

men

ts:

Tab

leA

.7.

bM

easu

red

ratio

corr

ecte

dfo

rfrac

tion

atio

nan

dsp

ike

only.

Dat

are

duct

ion

and

isot

ope

dilu

tion

calc

ulat

ions

wer

eper

form

edus

ing

anin

-hou

seEXCEL

spre

adsh

eet.

Err

orpr

opog

atio

npa

ram

eter

sar

eas

deta

iled

for

U–P

b

mea

sure

men

tsin

Tab

leA

.7w

ith

the

exce

ptio

nof

the

follo

win

g:

1.

Ava

lue

of0.

0012

8733

for

232

Th/

230

Th

ofth

eEA

RT

HT

IME

trac

erso

lution

was

used

for

Isot

ope

dilu

tion

calc

ulat

ions

.

2.

Mas

sfrac

tion

atio

nwas

cons

trai

ned

byan

alys

isof

a10

ppb

230

Th

spik

eso

lution

prep

ared

byR.

R.

Par

rish

atth

eG

eolo

gica

lSur

vey

ofCan

ada—

230

Th/

232

Th=

7.07

24±

0.00

7(1

�).

3.

Initia

lPb

com

pos

itio

nat

34M

aca

lcul

ated

from

Sta

cey

and

Kra

mer

s(1

975)

:208

Pb/

204

Pb

=38

.57±

0.74

%(e

rror

sar

e1�

).

221

Tab

leA

.9:

40

Ar/

39

Ar

sing

le-g

rain

fusion

data

.*Cor

rect

edfo

rat

mos

pher

icA

rw

hen

36

Ar

mea

sure

men

tex

ceed

stw

ice

the

36

Ar

blan

kva

lue.

40A

r1�

39A

r1�

38A

r1�

37A

r1�

36A

r1�

40A

r*40A

r⇤

39A

r1�

Age

1�C

orr?

*

J=

0.01173449

CW

T-17-1

0.1137

0.00045

0.04091

0.00010

0.00059

0.00003

0.00271

0.00573

0.00010

0.00003

75.3

2.0935

0.1949

43.78

4.03

Corr

CW

T-17-2

0.0579

0.00032

0.03429

0.00012

0.00043

0.00003

-0.00065

0.00574

0.00002

0.00003

92.3

1.6880

0.0109

35.39

0.57

Not

corr

CW

T-17-3

0.1099

0.00047

0.05596

0.00013

0.00071

0.00003

0.00159

0.00574

0.00001

0.00003

98.7

1.9640

0.0095

41.11

0.56

Not

corr

CW

T-17-4

0.6231

0.00340

0.30138

0.00230

0.00378

0.00004

-0.00178

0.00575

0.00030

0.00003

86.0

1.7784

0.0317

37.26

0.68

Corr

CW

T-17-5

0.5162

0.00097

0.27788

0.00039

0.00333

0.00002

-0.00684

0.00575

0.00025

0.00003

86.0

1.5973

0.0289

33.50

0.62

Corr

CW

T-17-6

1.0173

0.00392

0.47310

0.00227

0.00597

0.00006

-0.00741

0.00575

0.00055

0.00003

84.2

1.8100

0.0207

37.92

0.47

Corr

CW

T-17-7

0.6726

0.00053

0.35521

0.00046

0.00413

0.00004

-0.00178

0.00576

0.00029

0.00003

87.5

1.6563

0.0225

34.73

0.50

Corr

CW

T-17-8

0.1658

0.00037

0.08061

0.00010

0.00108

0.00003

-0.00009

0.00576

0.00013

0.00003

77.7

1.5993

0.0988

33.54

2.06

Corr

CW

T-17-9

0.6457

0.00112

0.33306

0.00041

0.00442

0.00007

-0.01073

0.00108

0.00023

0.00002

89.5

1.7346

0.0175

36.35

0.40

Corr

CW

T-17-10

0.8870

0.00203

0.46345

0.00034

0.00596

0.00004

-0.01073

0.00108

0.00033

0.00002

89.0

1.7035

0.0130

35.71

0.32

Corr

CW

T-17-11

0.4731

0.00047

0.25392

0.00030

0.00322

0.00002

-0.01016

0.00108

0.00015

0.00002

90.6

1.6887

0.0224

35.40

0.50

Corr

CW

T-17-12

0.4342

0.00074

0.19172

0.00042

0.00205

0.00003

-0.00096

0.00108

0.00025

0.00002

83.0

1.8797

0.0301

39.36

0.65

Corr

Av

blank

1-8

0.0023

0.00004

0.00003

0.00001

0.00002

0.00001

0.00103

0.00002

0.00008

0.00001

Av

blank

9-12

0.0021

0.00004

0.00008

0.00001

0.00004

0.00001

0.00109

0.00002

0.00008

0.00001

J=

0.01179633

ASA

-28-1

0.2196

0.00055

0.09498

0.00005

0.00121

0.00001

0.00036

0.00025

0.00007

0.00001

91.2

2.3122

0.0059

48.55

0.54

Not

corr

ASA

-28-2

0.1272

0.00055

0.05622

0.00003

0.00072

0.00001

0.00069

0.00025

0.00002

0.00001

96.4

2.2629

0.0098

47.53

0.62

Not

corr

ASA

-28-3

0.2213

0.00055

0.10768

0.00008

0.00135

0.00001

0.00036

0.00025

0.00006

0.00001

92.6

2.0555

0.0053

43.22

0.48

Not

corr

ASA

-28-4

0.0765

0.00054

0.03415

0.00004

0.00041

0.00001

0.00003

0.00025

0.00004

0.00001

86.3

2.2399

0.0162

47.05

0.82

Not

corr

ASA

-28-5

0.2142

0.00055

0.09871

0.00002

0.00120

0.00001

-0.00014

0.00025

0.00004

0.00001

95.1

2.1704

0.0056

45.61

0.51

Not

corr

ASA

-28-6

0.1415

0.00054

0.05196

0.00003

0.00066

0.00000

0.00069

0.00025

0.00002

0.00001

96.8

2.7231

0.0106

57.04

0.71

Not

corr

ASA

-28-7

0.1366

0.00055

0.05898

0.00004

0.00067

0.00001

0.00003

0.00025

0.00004

0.00001

92.3

2.3161

0.0094

48.63

0.62

Not

corr

ASA

-28-8

0.6656

0.00058

0.27185

0.00012

0.00332

0.00002

0.00301

0.00025

0.00012

0.00001

94.9

2.4485

0.0024

51.37

0.52

Not

corr

ASA

-28-9

0.0911

0.00054

0.03901

0.00006

0.00049

0.00000

0.00003

0.00025

0.00003

0.00001

91.7

2.3342

0.0144

49.00

0.77

Not

corr

ASA

-28-10

0.1211

0.00055

0.05179

0.00004

0.00070

0.00001

-0.00014

0.00026

0.00006

0.00001

86.5

2.3379

0.0107

49.08

0.66

Not

corr

ASA

-28-11

0.1519

0.00054

0.04480

0.00004

0.00060

0.00001

-0.00031

0.00026

0.00002

0.00001

97.0

3.3920

0.0125

70.78

0.86

Not

corr

ASA

-28-12

0.3891

0.00061

0.17130

0.00018

0.00220

0.00001

-0.00014

0.00026

0.00013

0.00001

90.5

2.2713

0.0042

47.70

0.50

Not

corr

ASA

-28-13

0.4072

0.00054

0.17980

0.00015

0.00210

0.00002

0.00003

0.00026

0.00008

0.00001

94.5

2.2649

0.0035

47.57

0.49

Not

corr

ASA

-28-14

0.4213

0.00055

0.17902

0.00011

0.00228

0.00001

0.00003

0.00026

0.00014

0.00001

90.5

2.3532

0.0034

49.40

0.51

Not

corr

Av

blank

1-9

0.0036

0.00001

0.00001

0.00000

0.00000

0.00000

0.00012

0.00000

0.00004

0.00000

Av

blank

10-14

0.0036

0.00002

0.00001

0.00000

0.00001

0.00000

0.00010

0.00000

0.00003

0.00000

222

Tab

leA

.9:

40

Ar/

39

Ar

sing

le-g

rain

fusion

data

.*Cor

rect

edfo

rat

mos

pher

icA

rw

hen

36

Ar

mea

sure

men

tex

ceed

stw

ice

the

36

Ar

blan

kva

lue.

40A

r1�

39A

r1�

38A

r1�

37A

r1�

36A

r1�

40A

r*40A

r⇤

39A

r1�

Age

1�C

orr?

*

J=

0.01181141

CW

T-8-1

0.3882

0.00049

0.20446

0.00021

0.00262

0.00003

-0.00229

0.00408

0.00018

0.00002

86.0

1.6321

0.0287

34.45

0.62

Corr

CW

T-8-2

0.3844

0.00022

0.22892

0.00011

0.00294

0.00003

0.00544

0.00408

0.00005

0.00002

95.8

1.6792

0.0012

35.43

0.35

Not

corr

CW

T-8-3

1.2137

0.00421

0.69864

0.00249

0.00884

0.00006

0.01125

0.00408

0.00022

0.00002

94.5

1.6424

0.0119

34.66

0.30

Corr

CW

T-8-4

0.2036

0.00025

0.10922

0.00010

0.00129

0.00003

-0.00874

0.00408

0.00010

0.00002

84.8

1.5819

0.0536

33.40

1.13

Corr

CW

T-8-5

0.4985

0.00037

0.28686

0.00032

0.00373

0.00003

-0.00617

0.00408

0.00013

0.00002

92.0

1.5991

0.0205

33.76

0.46

Corr

CW

T-8-6

0.1880

0.00035

0.10607

0.00021

0.00134

0.00003

-0.00100

0.00408

-0.00001

0.00002

100.9

1.7721

0.0048

37.37

0.42

Not

corr

CW

T-8-7

0.5910

0.00103

0.35609

0.00036

0.00447

0.00004

0.00545

0.00409

0.00007

0.00002

96.3

1.5979

0.0167

33.73

0.39

Corr

CW

T-8-8

0.3478

0.00062

0.19761

0.00037

0.00247

0.00003

-0.00101

0.00409

0.00009

0.00002

92.0

1.6187

0.0299

34.17

0.65

Corr

CW

T-8-9

0.2153

0.00051

0.11529

0.00019

0.00157

0.00003

0.00287

0.00409

0.00009

0.00002

87.0

1.6255

0.0510

34.31

1.08

Corr

CW

T-8-10

0.4528

0.00061

0.25645

0.00033

0.00299

0.00004

0.00029

0.00409

0.00010

0.00002

93.2

1.6454

0.0230

34.73

0.51

Corr

CW

T-8-11

0.1723

0.00037

0.09814

0.00012

0.00103

0.00003

0.00676

0.00409

0.00005

0.00002

90.7

1.7554

0.0043

37.02

0.41

Not

corr

CW

T-8-12

0.4092

0.00070

0.23361

0.00041

0.00293

0.00003

-0.01007

0.00410

0.00016

0.00002

88.1

1.5436

0.0254

32.60

0.55

Corr

CW

T-8-13

0.4376

0.00028

0.24839

0.00027

0.00344

0.00004

0.00353

0.00410

0.00015

0.00002

89.6

1.5782

0.0236

33.32

0.52

Corr

CW

T-8-14

0.3232

0.00028

0.18277

0.00029

0.00233

0.00003

0.00547

0.00410

0.00008

0.00002

92.3

1.6317

0.0321

34.44

0.69

Corr

CW

T-8-15

0.6800

0.00115

0.42336

0.00081

0.00530

0.00008

0.00029

0.00410

0.00007

0.00002

96.8

1.5543

0.0144

32.82

0.34

Corr

Av

Blank

0.0018

0.00004

0.00005

0.000011

0.00005

0.00001

0.00093

0.00002

0.00008

0.00001

J=

0.01171388

CW

T-12-1

0.1548

0.00202

0.07932

0.00068

0.00100

0.00001

0.00000

0.00068

-0.00002

0.00002

104.0

1.9516

0.0305

40.78

1.32

Not

corr

CW

T-12-2

0.0536

0.00129

0.02612

0.00015

0.00037

0.00001

0.00032

0.00068

-0.00002

0.00002

111.5

2.0514

0.0505

42.84

2.13

Not

corr

CW

T-12-3

0.2374

0.00188

0.12446

0.00055

0.00162

0.00001

-0.00081

0.00068

-0.00002

0.00002

102.6

1.9075

0.0173

39.87

0.82

Not

corr

CW

T-12-4

0.2374

0.00188

0.12446

0.00055

0.00162

0.00001

-0.00081

0.00068

-0.00002

0.00002

102.6

1.9569

0.0587

40.89

1.23

Corr

CW

T-12-5

1.5595

0.00177

0.76809

0.00057

0.00963

0.00001

0.00599

0.00068

0.00044

0.00003

91.7

1.8614

0.0102

38.91

0.29

Corr

CW

T-12-6

2.5144

0.00367

1.28759

0.00293

0.01564

0.00001

0.00437

0.00068

0.00053

0.00002

93.8

1.8313

0.0074

38.29

0.24

Corr

CW

T-12-7

2.1096

0.00131

0.97238

0.00044

0.01199

0.00001

0.00665

0.00068

0.00048

0.00002

93.3

2.0239

0.0074

42.27

0.26

Corr

CW

T-12-8

2.4667

0.00160

1.29342

0.00050

0.01598

0.00003

0.00032

0.00068

0.00035

0.00003

95.8

1.8273

0.0060

38.21

0.23

Corr

CW

T-12-9

0.9557

0.00133

0.47914

0.00025

0.00604

0.00003

-0.00016

0.00069

0.00015

0.00002

95.4

1.9027

0.0149

39.77

0.36

Corr

CW

T-12-10

1.4167

0.00128

0.62547

0.00020

0.00775

0.00001

0.00032

0.00069

0.00034

0.00002

92.9

2.1048

0.0114

43.94

0.32

Corr

CW

T-12-11

4.4933

0.00144

2.25201

0.00055

0.02774

0.00003

0.00065

0.00069

0.00051

0.00002

96.7

1.9284

0.0032

40.30

0.21

Corr

CW

T-12-12

1.2161

0.00133

0.59263

0.00026

0.00735

0.00002

-0.00032

0.00069

0.00025

0.00002

93.9

1.9277

0.0120

40.29

0.32

Corr

CW

T-12-13

1.8627

0.00136

0.95594

0.00037

0.01170

0.00003

0.00146

0.00069

0.00028

0.00002

95.6

1.8623

0.0075

38.93

0.25

Corr

CW

T-12-14

0.6321

0.00023

0.30750

0.00009

0.00387

0.00001

0.00088

0.00021

0.00020

0.00001

90.6

1.8634

0.0079

38.95

0.25

Corr

223

Tab

leA

.9:

40

Ar/

39

Ar

sing

le-g

rain

fusion

data

.*Cor

rect

edfo

rat

mos

pher

icA

rw

hen

36

Ar

mea

sure

men

tex

ceed

stw

ice

the

36

Ar

blan

kva

lue.

40A

r1�

39A

r1�

38A

r1�

37A

r1�

36A

r1�

40A

r*40A

r⇤

39A

r1�

Age

1�C

orr?

*

CW

T-12-15

0.9351

0.00048

0.51080

0.00025

0.00627

0.00001

0.00121

0.00021

0.00010

0.00001

96.8

1.7729

0.0049

37.08

0.21

Corr

CW

T-12-16

0.1373

0.00014

0.06659

0.00006

0.00085

0.00001

0.00005

0.00021

0.00003

0.00001

93.5

2.0626

0.0029

43.07

0.44

Not

corr

CW

T-12-17

4.2840

0.00072

2.13595

0.00035

0.02687

0.00004

0.00187

0.00021

0.00159

0.00001

89.0

1.7857

0.0018

37.35

0.19

Corr

Av

blank

1-14

0.0052

0.00002

0.00001

0.000000

0.00001

0.00000

0.00018

0.00000

0.00007

Av

blank

15-17

0.0040

0.00001

0.00002

0.000000

0.00000

0.00000

0.00012

0.00000

0.00004

J=

0.01169327

CW

T-13-1

0.2583

0.00068

0.13504

0.00056

0.00137

0.00003

0.00186

0.00391

0.00004

0.00002

95.8

1.9131

0.0094

39.91

0.55

Not

corr

CW

T-13-2

0.1845

0.00079

0.09186

0.00007

0.00119

0.00004

0.00465

0.00392

0.00004

0.00002

94.1

2.0088

0.0087

41.89

0.55

Not

corr

CW

T-13-3

0.1923

0.00052

0.06765

0.00012

0.00084

0.00002

-0.00372

0.00392

0.00011

0.00002

83.6

2.3763

0.0828

49.45

1.72

Corr

CW

T-13-4

0.1380

0.00045

0.06699

0.00025

0.00073

0.00004

0.00186

0.00392

0.00003

0.00002

94.3

2.0604

0.0102

42.95

0.60

Not

corr

CW

T-13-5

0.1420

0.00037

0.05978

0.00009

0.00076

0.00003

-0.00372

0.00392

0.00003

0.00002

94.5

2.3762

0.0071

49.44

0.57

Not

corr

CW

T-13-6

0.2306

0.00020

0.13338

0.00019

0.00153

0.00004

0.00186

0.00392

0.00002

0.00002

97.9

1.7288

0.0029

36.11

0.38

Not

corr

CW

T-13-7

0.4067

0.00049

0.16032

0.00038

0.00218

0.00003

-0.00163

0.00393

0.00011

0.00002

92.3

2.3403

0.0353

48.71

0.76

Corr

CW

T-13-8

0.6792

0.00067

0.35783

0.00051

0.00419

0.00003

0.00186

0.00393

0.00014

0.00002

94.1

1.7851

0.0159

37.27

0.38

Corr

CW

T-13-9

0.2606

0.00042

0.09174

0.00022

0.00118

0.00003

-0.01352

0.00393

0.00006

0.00002

93.6

2.8405

0.0082

58.95

0.67

Not

corr

CW

T-13-10

0.3990

0.00063

0.12574

0.00036

0.00149

0.00003

-0.00093

0.00393

0.00019

0.00002

86.2

2.7347

0.0453

56.79

0.97

Corr

Av

blank

0.0027

0.00004

0.00004

0.000010

0.00003

0.00001

0.00091

0.00002

0.00007

0.00001

J=

0.01177572

CW

T-15-1

0.0556

0.00008

0.01713

0.00002

0.00020

0.00001

0.00219

0.00022

0.00009

0.00001

54.5

1.7693

0.1257

37.20

2.62

Corr

CW

T-15-2

0.1195

0.00010

0.02704

0.00003

0.00031

0.00001

0.00404

0.00022

0.00023

0.00001

44.2

1.9513

0.0797

40.98

1.67

Corr

CW

T-15-3

0.0505

0.00008

0.01778

0.00003

0.00023

0.00001

0.00202

0.00022

0.00008

0.00001

55.7

1.5848

0.1212

33.36

2.53

Corr

CW

T-15-4

0.0410

0.00008

0.01422

0.00001

0.00017

0.00001

0.00202

0.00022

0.00005

0.00001

67.0

1.9305

0.1515

40.55

3.15

Corr

CW

T-15-5

0.0972

0.00011

0.02351

0.00002

0.00031

0.00001

0.00422

0.00022

0.00015

0.00001

55.7

2.3048

0.0917

48.31

1.91

Corr

CW

T-15-6

0.0591

0.00008

0.02546

0.00004

0.00032

0.00001

0.00405

0.00022

0.00007

0.00001

67.2

1.5604

0.0846

32.85

1.77

Corr

CW

T-15-7

0.0896

0.00008

0.02853

0.00003

0.00039

0.00001

0.00371

0.00022

0.00011

0.00001

65.1

2.0444

0.0755

42.92

1.58

Corr

CW

T-15-8

0.0578

0.00008

0.02154

0.00002

0.00029

0.00001

0.00355

0.00022

0.00009

0.00001

56.2

1.5085

0.1000

31.77

2.09

Corr

CW

T-15-9

0.0537

0.00009

0.02115

0.00004

0.00024

0.00001

0.00237

0.00022

0.00005

0.00001

74.9

1.9020

0.1019

39.96

2.13

Corr

CW

T-15-10

0.0711

0.00008

0.01435

0.00003

0.00019

0.00001

0.00237

0.00022

0.00016

0.00001

35.3

1.7487

0.1501

36.77

3.13

Corr

CW

T-15-11

0.0483

0.00008

0.01422

0.00002

0.00019

0.00001

0.00203

0.00022

0.00008

0.00001

53.7

1.8263

0.1515

38.39

3.16

Corr

CW

T-15-12

0.0909

0.00010

0.02401

0.00007

0.00038

0.00001

0.00304

0.00022

0.00018

0.00001

42.9

1.6250

0.0899

34.20

1.88

Corr

Av

blank

1-10

0.0045

0.00002

0.00001

0.000000

0.00001

0.00000

0.00010

0.00000

0.00003

0.00000

224

Tab

leA

.9:

40

Ar/

39

Ar

sing

le-g

rain

fusion

data

.*Cor

rect

edfo

rat

mos

pher

icA

rw

hen

36

Ar

mea

sure

men

tex

ceed

stw

ice

the

36

Ar

blan

kva

lue.

40A

r1�

39A

r1�

38A

r1�

37A

r1�

36A

r1�

40A

r*40A

r⇤

39A

r1�

Age

1�C

orr?

*

Av

blank

11-12

0.0045

0.00002

0.00002

0.000000

0.00001

0.00000

0.00011

0.00000

0.00004

0.00000

J=

0.01183755

ASA

-6a-1

0.0975

0.00027

0.05507

0.00010

0.00067

0.00002

-0.00619

0.01153

0.00004

0.00001

88.9

1.7696

0.0058

37.40

0.44

Not

corr

ASA

-6a-2

0.2514

0.00035

0.14132

0.00027

0.00183

0.00002

-0.00916

0.01153

0.00010

0.00001

88.6

1.5768

0.0290

33.36

0.63

Corr

ASA

-6a-3

0.1749

0.00033

0.07763

0.00042

0.00102

0.00003

-0.00099

0.01154

0.00003

0.00001

95.5

2.2534

0.0128

47.49

0.71

Not

corr

ASA

-6a-4

0.2604

0.00057

0.13161

0.00021

0.00166

0.00002

-0.00099

0.01154

0.00013

0.00001

85.6

1.6940

0.0313

35.82

0.68

Corr

ASA

-6a-5

0.1605

0.00067

0.08406

0.00026

0.00120

0.00003

-0.00695

0.01155

0.00012

0.00001

78.5

1.4989

0.0492

31.73

1.04

Corr

ASA

-6a-6

0.3025

0.00050

0.16837

0.00033

0.00213

0.00003

-0.00892

0.00440

0.00012

0.00002

88.8

1.5951

0.0378

33.75

0.81

Corr

ASA

-6a-7

0.2095

0.00034

0.12091

0.00021

0.00161

0.00003

-0.00289

0.00441

0.00006

0.00002

92.2

1.5986

0.0525

33.82

1.11

Corr

ASA

-6a-8

0.6128

0.00101

0.33768

0.00055

0.00410

0.00004

-0.00138

0.00441

0.00015

0.00002

93.0

1.6878

0.0192

35.69

0.44

Corr

ASA

-6a-9

0.0759

0.00023

0.03915

0.00016

0.00054

0.00003

-0.01119

0.00441

0.00005

0.00002

82.5

1.5985

0.1619

33.82

3.40

Corr

ASA

-6a-10

0.2007

0.00040

0.12069

0.00028

0.00156

0.00003

0.00390

0.00441

0.00004

0.00002

94.8

1.6632

0.0051

35.17

0.41

Not

corr

ASA

-6a-11

1.3152

0.00162

0.75440

0.00084

0.00954

0.00004

0.01523

0.00441

0.00031

0.00002

93.1

1.6239

0.0088

34.35

0.25

Corr

ASA

-6a-12

0.5945

0.00073

0.35552

0.00044

0.00443

0.00003

0.01071

0.00442

0.00017

0.00002

91.8

1.5350

0.0180

32.49

0.41

Corr

Av

blank

1-5

0.0020

0.00006

0.00011

0.000010

0.00005

0.00001

0.00096

0.00002

0.00005

0.00001

Av

blank

6-12

0.0017

0.00005

0.00005

0.000010

0.00002

0.00001

0.00101

0.00003

0.00005

0.00001

J=

0.01178615

ASA

-6b-1

0.3322

0.00033

0.15793

0.00027

0.00196

0.00003

-0.00144

0.00571

0.00017

0.00002

85.0

1.7875

0.0319

37.61

0.69

Corr

ASA

-6b-2

0.3275

0.00034

0.15067

0.00030

0.00171

0.00006

0.00317

0.00571

0.00023

0.00002

79.4

1.7249

0.0334

36.31

0.72

Corr

ASA

-6b-3

0.2125

0.00027

0.08396

0.00034

0.00115

0.00002

-0.00913

0.00571

0.00013

0.00002

82.1

2.0775

0.0602

43.64

1.27

Corr

ASA

-6b-4

0.5210

0.00083

0.20062

0.00079

0.00282

0.00002

-0.00144

0.00571

0.00029

0.00002

83.6

2.1718

0.0266

45.60

0.60

Corr

ASA

-6b-5

0.5272

0.00072

0.27040

0.00114

0.00345

0.00002

-0.00760

0.00572

0.00019

0.00002

89.4

1.7433

0.0201

36.69

0.46

Corr

ASA

-6b-6

0.4692

0.00200

0.22168

0.00107

0.00277

0.00003

0.01165

0.00572

0.00020

0.00002

87.5

1.8515

0.0259

38.95

0.57

Corr

ASA

-6b-7

0.7022

0.00094

0.29584

0.00203

0.00413

0.00004

-0.00530

0.00572

0.00033

0.00002

86.2

2.0453

0.0222

42.97

0.51

Corr

ASA

-6b-8

0.2900

0.00053

0.15289

0.00035

0.00200

0.00002

-0.00376

0.00572

0.00016

0.00002

83.8

1.5901

0.0331

33.50

0.71

Corr

ASA

-6b-9

0.2725

0.00029

0.14204

0.00042

0.00190

0.00002

-0.00916

0.00573

0.00017

0.00002

81.7

1.5671

0.0355

33.02

0.76

Corr

ASA

-6b-10

0.1940

0.00031

0.09307

0.00019

0.00106

0.00003

-0.00299

0.00573

0.00011

0.00002

83.4

1.7391

0.0539

36.61

1.14

Corr

ASA

-6b-11

0.6121

0.00277

0.21482

0.00069

0.00287

0.00002

-0.00994

0.00573

0.00036

0.00002

82.7

2.3558

0.0277

49.41

0.62

Corr

ASA

-6b-12

0.4221

0.00070

0.20799

0.00051

0.00243

0.00004

0.00241

0.00573

0.00016

0.00002

88.9

1.8040

0.0247

37.96

0.55

Corr

Av

blank

0.0019

0.00004

0.00003

0.000010

0.00001

0.00001

0.00086

0.00002

0.00006

0.00001

J=

0.01167266

225

Tab

leA

.9:

40

Ar/

39

Ar

sing

le-g

rain

fusion

data

.*Cor

rect

edfo

rat

mos

pher

icA

rw

hen

36

Ar

mea

sure

men

tex

ceed

stw

ice

the

36

Ar

blan

kva

lue.

40A

r1�

39A

r1�

38A

r1�

37A

r1�

36A

r1�

40A

r*40A

r⇤

39A

r1�

Age

1�C

orr?

*

CW

T-5-1

0.6913

0.00031

0.43869

0.00011

0.00533

0.00001

0.00070

0.00016

0.00014

0.00001

94.0

1.5758

0.0008

32.88

0.33

Not

corr

CW

T-5-2

0.7137

0.00047

0.44997

0.00048

0.00564

0.00002

0.00717

0.00016

0.00019

0.00001

92.1

1.4612

0.0053

30.51

0.19

Corr

CW

T-5-3

0.6158

0.00032

0.39768

0.00015

0.00491

0.00002

-0.00017

0.00016

0.00007

0.00001

96.6

1.5486

0.0010

32.32

0.32

Not

corr

CW

T-5-4

0.4326

0.00032

0.27047

0.00011

0.00326

0.00001

0.00035

0.00016

0.00009

0.00001

93.9

1.5995

0.0014

33.37

0.34

Not

corr

CW

T-5-5

0.5793

0.00023

0.37015

0.00013

0.00460

0.00001

0.00613

0.00016

0.00012

0.00001

93.9

1.4693

0.0061

30.68

0.20

Corr

CW

T-5-6

0.3214

0.00036

0.20059

0.00008

0.00250

0.00001

0.00018

0.00016

0.00009

0.00001

91.7

1.6023

0.0019

33.43

0.34

Not

corr

CW

T-5-7

0.2630

0.00027

0.14085

0.00007

0.00177

0.00001

0.00088

0.00016

0.00019

0.00001

78.7

1.8674

0.0022

38.90

0.40

Not

corr

CW

T-5-8

0.2550

0.00024

0.17097

0.00008

0.00207

0.00001

0.00018

0.00016

0.00001

0.00001

98.8

1.4915

0.0016

31.14

0.32

Not

corr

CW

T-5-9

0.3121

0.00031

0.20561

0.00004

0.00256

0.00001

0.00105

0.00016

0.00004

0.00001

96.2

1.5180

0.0015

31.69

0.32

Not

corr

CW

T-5-10

0.2989

0.00025

0.18954

0.00019

0.00228

0.00000

0.00018

0.00016

0.00007

0.00001

93.1

1.5769

0.0020

32.90

0.34

Not

corr

CW

T-5-11

0.2616

0.00030

0.16041

0.00012

0.00189

0.00000

0.00333

0.00016

0.00007

0.00001

92.1

1.5018

0.0141

31.35

0.33

Corr

CW

T-5-12

0.2230

0.00035

0.14694

0.00028

0.00172

0.00005

-0.00597

0.00722

0.00004

0.00001

94.4

1.5178

0.0038

31.68

0.35

Not

corr

CW

T-5-13

0.5424

0.00041

0.35148

0.00047

0.00430

0.00005

-0.01194

0.00723

0.00005

0.00001

97.1

1.5433

0.0024

32.21

0.33

Not

corr

CW

T-5-14

0.1200

0.00016

0.07732

0.00011

0.00091

0.00004

0.01016

0.00723

0.00002

0.00001

94.5

1.5525

0.0030

32.40

0.34

Not

corr

CW

T-5-15

0.3075

0.00040

0.16532

0.00024

0.00199

0.00005

0.00239

0.00723

0.00017

0.00001

83.4

1.5520

0.0198

32.39

0.44

Corr

CW

T-5-16

0.8423

0.00104

0.54444

0.00058

0.00675

0.00005

-0.00419

0.00724

0.00014

0.00001

95.0

1.4697

0.0064

30.69

0.20

Corr

CW

T-5-17

0.8234

0.00098

0.52772

0.00086

0.00683

0.00007

-0.00359

0.00724

0.00016

0.00001

94.2

1.4693

0.0068

30.68

0.21

Corr

Av.

Blank

12-17

0.0053

0.00002

0.00002

0.000000

0.00001

0.00000

0.00011

0.00000

0.00004

0.00000

Av.

Blank

1-12

0.0020

0.00004

0.00007

0.000010

0.00005

0.00001

0.00103

0.00002

0.00006

0.00001

J=

0.01175511

CW

T-7-1

0.4355

0.00063

0.19383

0.00032

0.00234

0.00002

0.00601

0.00737

0.00013

0.00001

91.2

2.0488

0.0202

42.93

0.47

Corr

CW

T-7-2

0.1379

0.00041

0.03782

0.00007

0.00046

0.00002

0.00320

0.00737

0.00004

0.00001

91.4

3.6460

0.0130

75.71

0.91

Not

corr

CW

T-7-3

0.2025

0.00025

0.12736

0.00028

0.00168

0.00003

-0.00324

0.00374

0.00002

0.00001

97.6

1.5898

0.0040

33.40

0.37

Not

corr

CW

T-7-4

0.3270

0.00095

0.21639

0.00061

0.00275

0.00004

0.00248

0.00375

0.00004

0.00001

96.7

1.5111

0.0061

31.76

0.41

Not

corr

CW

T-7-5

0.0594

0.00034

0.03458

0.00010

0.00032

0.00002

-0.00153

0.00375

0.00003

0.00001

86.7

1.7164

0.0109

36.04

0.58

Not

corr

CW

T-7-6

0.2622

0.00058

0.17202

0.00033

0.00192

0.00004

-0.00267

0.00375

0.00006

0.00001

93.6

1.5244

0.0045

32.04

0.37

Not

corr

CW

T-7-7

0.6008

0.00070

0.38016

0.00029

0.00492

0.00005

-0.00038

0.00375

0.00011

0.00001

94.8

1.4975

0.0118

31.48

0.29

Corr

CW

T-7-8

0.3398

0.00089

0.21388

0.00060

0.00262

0.00004

-0.00038

0.00375

0.00004

0.00001

96.8

1.5886

0.0061

33.38

0.42

Not

corr

CW

T-7-9

0.3435

0.00062

0.20505

0.00046

0.00234

0.00004

0.00535

0.00376

0.00015

0.00001

87.4

1.4641

0.0219

30.78

0.48

Corr

CW

T-7-10

0.1447

0.00023

0.08954

0.00012

0.00120

0.00002

0.01166

0.00376

0.00004

0.00001

92.5

1.6160

0.0033

33.95

0.36

Not

corr

CW

T-7-11

0.5408

0.00033

0.35857

0.00026

0.00460

0.00003

0.00880

0.00376

0.00003

0.00001

98.5

1.5081

0.0014

31.70

0.32

Not

corr

CW

T-7-12

0.5124

0.00055

0.30750

0.00057

0.00350

0.00005

-0.00268

0.00376

0.00015

0.00001

91.5

1.5254

0.0147

32.06

0.35

Corr

CW

T-7-13

0.0425

0.00036

0.01970

0.00020

0.00019

0.00003

-0.00432

0.00817

0.00005

0.00002

63.5

2.1552

0.0281

45.14

1.25

Not

corr

CW

T-7-14

0.6710

0.00144

0.44242

0.00103

0.00521

0.00007

-0.00015

0.00817

0.00003

0.00002

98.6

1.5166

0.0048

31.88

0.37

Not

corr

226

Tab

leA

.9:

40

Ar/

39

Ar

sing

le-g

rain

fusion

data

.*Cor

rect

edfo

rat

mos

pher

icA

rw

hen

36

Ar

mea

sure

men

tex

ceed

stw

ice

the

36

Ar

blan

kva

lue.

40A

r1�

39A

r1�

38A

r1�

37A

r1�

36A

r1�

40A

r*40A

r⇤

39A

r1�

Age

1�C

orr?

*

CW

T-7-15

0.1462

0.00038

0.09414

0.00029

0.00132

0.00003

-0.00194

0.00817

0.00000

0.00002

99.5

1.5530

0.0063

32.64

0.42

Not

corr

CW

T-7-16

0.4181

0.00051

0.26756

0.00050

0.00318

0.00004

-0.00551

0.00818

0.00014

0.00002

89.9

1.4051

0.0200

29.55

0.44

Corr

CW

T-7-17

0.5518

0.00039

0.35409

0.00054

0.00443

0.00003

0.00104

0.00818

0.00006

0.00002

96.7

1.5585

0.0026

32.75

0.34

Not

corr

CW

T-7-18

0.0460

0.00037

0.01879

0.00007

0.00025

0.00003

0.00641

0.00818

0.00006

0.00002

59.8

2.4478

0.0218

51.18

1.03

Not

corr

Av

blank

1-2

0.0021

0.00004

0.00002

0.000010

0.00003

0.00001

0.00100

0.00002

0.00008

0.00001

Av

blank

3-12

0.0021

0.00004

0.00005

0.000010

0.00005

0.00001

0.00096

0.00002

0.00006

0.00001

Av

blank

13-18

0.0020

0.00004

0.00004

0.000010

0.00006

0.00001

0.00103

0.00002

0.00007

0.00001

J=

0.01185351

ASA

-5-1

0.3581

0.00033

0.23577

0.00027

0.00306

0.00003

-0.00427

0.00377

0.00003

0.00002

97.6

1.5190

0.0022

32.19

0.33

Not

corr

ASA

-5-2

0.1598

0.00018

0.10459

0.00017

0.00140

0.00003

-0.00361

0.00377

0.00003

0.00002

94.7

1.5284

0.0030

32.39

0.34

Not

corr

ASA

-5-3

0.3541

0.00040

0.22729

0.00039

0.00276

0.00004

-0.00493

0.00378

0.00008

0.00002

93.4

1.4555

0.0240

30.86

0.53

Corr

ASA

-5-4

0.2761

0.00043

0.15944

0.00032

0.00195

0.00003

-0.01085

0.00378

0.00017

0.00002

81.9

1.4189

0.0342

30.09

0.73

Corr

ASA

-5-5

0.4234

0.00070

0.27641

0.00060

0.00352

0.00003

-0.00362

0.00378

0.00004

0.00002

97.3

1.5317

0.0042

32.46

0.37

Not

corr

ASA

-5-6

0.2200

0.00082

0.13977

0.00014

0.00171

0.00003

-0.00757

0.00378

0.00002

0.00002

97.5

1.5737

0.0061

33.34

0.42

Not

corr

ASA

-5-7

0.4633

0.00041

0.29311

0.00026

0.00357

0.00003

-0.00889

0.00378

0.00011

0.00002

93.1

1.4709

0.0186

31.18

0.42

Corr

ASA

-5-8

0.0818

0.00024

0.05127

0.00013

0.00061

0.00003

0.00033

0.00378

0.00005

0.00002

82.4

1.5965

0.0061

33.82

0.42

Not

corr

ASA

-5-9

0.0895

0.00026

0.05716

0.00018

0.00061

0.00003

0.00692

0.00379

0.00004

0.00002

87.2

1.5662

0.0066

33.18

0.43

Not

corr

ASA

-5-10

0.4953

0.00028

0.32092

0.00086

0.00386

0.00004

-0.00824

0.00379

0.00012

0.00002

92.9

1.4341

0.0173

30.41

0.39

Corr

ASA

-5-11

0.5840

0.00105

0.37388

0.00047

0.00447

0.00003

-0.01220

0.00379

0.00012

0.00002

94.0

1.4681

0.0149

31.12

0.35

Corr

ASA

-5-12

0.4731

0.00090

0.31443

0.00068

0.00403

0.00003

-0.00033

0.00379

0.00008

0.00002

95.1

1.4306

0.0177

30.34

0.40

Corr

ASA

-5-13

0.2431

0.00058

0.14068

0.00032

0.00172

0.00003

0.00239

0.00326

0.00005

0.00002

94.2

1.7279

0.0057

36.58

0.43

Not

corr

ASA

-5-14

0.3179

0.00043

0.19094

0.00024

0.00248

0.00003

-0.00376

0.00326

0.00007

0.00002

93.7

1.6647

0.0031

35.25

0.37

Not

corr

ASA

-5-15

0.1307

0.00049

0.07974

0.00016

0.00101

0.00003

0.00034

0.00326

-0.00002

0.00002

105.1

1.6394

0.0069

34.72

0.45

Not

corr

ASA

-5-16

0.1779

0.00031

0.08642

0.00022

0.00100

0.00003

-0.00308

0.00326

0.00017

0.00002

72.2

1.4857

0.0613

31.49

1.30

Corr

Av

blank

1-12

0.0017

0.00004

0.00005

0.000010

0.00004

0.00001

0.00098

0.00002

0.00006

0.00001

Av

blank

13-16

0.0018

0.00004

0.00008

0.000013

0.00005

0.00001

0.00090

0.00002

0.00008

0.00001

J=

0.01185816

CW

T-19-1

0.4544

0.00034

0.24047

0.00054

0.00289

0.00003

-0.00350

0.00716

0.00023

0.00001

85.3

1.6112

0.0173

34.14

0.40

Corr

CW

T-19-2

0.5640

0.00051

0.29198

0.00029

0.00388

0.00003

-0.00770

0.00716

0.00033

0.00001

82.9

1.6011

0.0141

33.93

0.34

Corr

CW

T-19-3

0.3345

0.00026

0.17600

0.00024

0.00217

0.00001

-0.00771

0.00717

0.00010

0.00001

91.5

1.7380

0.0232

36.80

0.52

Corr

CW

T-19-4

0.7096

0.00075

0.41989

0.00062

0.00506

0.00002

-0.00911

0.00717

0.00008

0.00001

96.8

1.6360

0.0101

34.66

0.27

Corr

CW

T-19-5

0.3392

0.00059

0.19167

0.00042

0.00213

0.00003

-0.00936

0.00503

0.00007

0.00002

94.2

1.6670

0.0251

35.31

0.55

Corr

227

Tab

leA

.9:

40

Ar/

39

Ar

sing

le-g

rain

fusion

data

.*Cor

rect

edfo

rat

mos

pher

icA

rw

hen

36

Ar

mea

sure

men

tex

ceed

stw

ice

the

36

Ar

blan

kva

lue.

40A

r1�

39A

r1�

38A

r1�

37A

r1�

36A

r1�

40A

r*40A

r⇤

39A

r1�

Age

1�C

orr?

*

CW

T-19-6

0.6829

0.00054

0.42051

0.00065

0.00496

0.00005

-0.00345

0.00503

0.00002

0.00002

99.3

1.6241

0.0028

34.41

0.36

Not

corr

CW

T-19-7

0.7979

0.00146

0.46617

0.00058

0.00599

0.00004

0.00099

0.00503

0.00015

0.00002

94.6

1.6187

0.0108

34.30

0.28

Corr

CW

T-19-8

0.0424

0.00023

0.00589

0.00008

0.00007

0.00003

-0.00419

0.00503

0.00014

0.00002

4.7

0.3408

0.8031

7.28

17.11

Corr

CW

T-19-9

0.4609

0.00067

0.28636

0.00033

0.00345

0.00003

-0.01012

0.00504

0.00004

0.00002

97.6

1.6093

0.0030

34.10

0.36

Not

corr

CW

T-19-10

0.5270

0.00077

0.27463

0.00027

0.00338

0.00006

0.00173

0.00504

0.00018

0.00002

90.1

1.7289

0.0175

36.61

0.41

Corr

CW

T-19-11

2.0680

0.00140

1.19272

0.00120

0.01461

0.00004

0.02323

0.00504

0.00038

0.00002

94.6

1.6405

0.0044

34.76

0.20

Corr

CW

T-19-12

2.4333

0.00147

1.29805

0.00117

0.01580

0.00008

0.00915

0.00504

0.00080

0.00002

90.3

1.6932

0.0057

35.86

0.21

Corr

CW

T-19-13

1.1142

0.00104

0.61158

0.00055

0.00735

0.00004

0.01138

0.00505

0.00033

0.00002

91.3

1.6641

0.0081

35.25

0.24

Corr

CW

T-19-14

1.0718

0.00077

0.64197

0.00052

0.00824

0.00004

-0.00124

0.00505

0.00022

0.00002

94.0

1.5698

0.0076

33.27

0.23

Corr

Av

blank

1-4

0.0023

0.00004

0.00004

0.000013

0.0000

0.00001

0.00096

0.00002

0.00006

0.00001

Av

blank

5-14

0.0021

0.00010

0.00007

0.000010

0.0001

0.00001

0.00093

0.00001

0.00006

0.00001

J=

0.01180299

N45b-1

0.9502

0.00326

0.50500

0.00141

0.00673

0.00004

-0.00537

0.00500

0.00029

0.00002

90.93

1.7109

0.0151

36.07

0.36

Corr

N45b-2

0.8149

0.00054

0.48601

0.00066

0.00572

0.00004

0.00557

0.00501

0.00006

0.00002

97.76

1.6767

0.0025

35.35

0.37

Not

corr

N45b-3

0.3633

0.00058

0.20047

0.00089

0.00257

0.00003

-0.01389

0.00501

-0.00001

0.00002

100.68

1.8125

0.0086

38.19

0.52

Not

corr

N45b-4

0.2053

0.00029

0.09891

0.00025

0.00137

0.00003

-0.00416

0.00501

0.00005

0.00002

92.56

2.0760

0.0060

43.67

0.50

Not

corr

N45b-5

0.1449

0.00022

0.07719

0.00021

0.00088

0.00003

-0.00173

0.00502

0.00002

0.00002

95.58

1.8770

0.0058

39.53

0.46

Not

corr

N45b-6

0.6304

0.00100

0.32790

0.00041

0.00416

0.00005

-0.00234

0.00502

0.00004

0.00002

98.05

1.9226

0.0039

40.48

0.43

Not

corr

N45b-7

0.2802

0.00033

0.15905

0.00020

0.00205

0.00003

-0.00905

0.00502

0.00001

0.00002

98.77

1.7619

0.0030

37.13

0.39

Not

corr

N45b-8

0.5033

0.00065

0.29415

0.00063

0.00329

0.00004

-0.00966

0.00502

-0.00001

0.00002

100.49

1.7109

0.0043

36.07

0.40

Not

corr

N45b-9

0.8340

0.00059

0.47628

0.00037

0.00630

0.00005

-0.00234

0.00503

0.00005

0.00002

98.17

1.7511

0.0019

36.91

0.37

Not

corr

N45b-10

1.3991

0.00118

0.75939

0.00090

0.00959

0.00005

0.00132

0.00503

0.00018

0.00002

96.16

1.7717

0.0089

37.34

0.26

Corr

N45b-11

0.3761

0.00064

0.18942

0.00032

0.00227

0.00001

-0.00137

0.00256

0.00008

0.00001

93.98

1.8657

0.0179

39.30

0.42

Corr

N45b-12

0.1906

0.00018

0.10461

0.00017

0.00128

0.00001

0.01234

0.00256

0.00004

0.00001

94.32

1.7185

0.0314

36.23

0.68

Corr

N45b-13

0.9393

0.00022

0.54086

0.00068

0.00634

0.00005

0.00274

0.00257

0.00011

0.00001

96.64

1.7366

0.0022

36.61

0.37

Not

corr

N45b-14

0.5002

0.00072

0.28068

0.00058

0.00319

0.00008

-0.00480

0.00257

0.00007

0.00001

96.06

1.7120

0.0124

36.09

0.32

Corr

Av

blank

1-10

0.0024

0.00005

0.00005

0.000010

0.00004

0.00001

0.00103

0.00002

0.00008

0.00001

Av

blank

13-16

0.0014

0.00004

0.00002

0.000010

0.00005

0.00001

0.00089

0.00002

0.00005

0.00001

J=

0.01181694

ASA

-35b-1

0.3986

0.00035

0.23729

0.00024

0.00288

0.00005

0.00518

0.00058

0.00005

0.00001

96.29

1.6799

0.0022

35.46

0.36

Not

corr

ASA

-35b-2

0.0545

0.00016

0.01353

0.00006

0.00030

0.00002

-0.00288

0.00058

0.00006

0.00001

67.48

2.7190

0.2190

57.05

4.53

Corr

ASA

-35b-3

0.0357

0.00016

0.00820

0.00005

0.00008

0.00002

-0.00415

0.00217

0.00005

0.00002

54.88

4.3483

0.0343

90.40

1.65

Not

corr

228

Tab

leA

.9:

40

Ar/

39

Ar

sing

le-g

rain

fusion

data

.*Cor

rect

edfo

rat

mos

pher

icA

rw

hen

36

Ar

mea

sure

men

tex

ceed

stw

ice

the

36

Ar

blan

kva

lue.

40A

r1�

39A

r1�

38A

r1�

37A

r1�

36A

r1�

40A

r*40A

r⇤

39A

r1�

Age

1�C

orr?

*

ASA

-35b-4

0.0246

0.00018

0.00840

0.00005

0.00010

0.00002

0.00577

0.00217

0.00000

0.00002

94.66

2.9292

0.0271

61.39

1.27

Not

corr

ASA

-35b-5

0.1422

0.00022

0.03471

0.00007

0.00049

0.00003

0.00752

0.00217

0.00004

0.00002

90.77

4.0983

0.0104

85.32

0.94

Not

corr

ASA

-35b-6

0.0980

0.00018

0.02656

0.00010

0.00030

0.00004

0.00519

0.00218

0.00003

0.00002

89.61

3.6892

0.0153

76.98

0.98

Not

corr

ASA

-35b-7

0.1024

0.00034

0.03415

0.00022

0.00036

0.00003

0.00870

0.00218

0.00009

0.00002

72.76

2.1825

0.1671

45.94

3.48

Corr

ASA

-35b-8

0.0359

0.00017

0.01203

0.00011

0.00014

0.00003

0.00461

0.00218

-0.00002

0.00002

112.81

2.9842

0.0304

62.52

1.39

Not

corr

ASA

-35b-9

0.0916

0.00029

0.05604

0.00013

0.00071

0.00003

0.00169

0.00218

0.00002

0.00002

92.12

1.6349

0.0064

34.52

0.44

Not

corr

ASA

-35b-10

0.2315

0.00056

0.13637

0.00016

0.00165

0.00002

-0.00475

0.00218

0.00000

0.00002

99.43

1.6978

0.0046

35.84

0.40

Not

corr

ASA

-35b-11

0.2763

0.00050

0.17042

0.00023

0.00211

0.00003

-0.00182

0.00218

0.00003

0.00002

96.32

1.6216

0.0037

34.24

0.37

Not

corr

ASA

-35b-12

0.2219

0.00048

0.13549

0.00026

0.00144

0.00005

0.00931

0.00218

0.00004

0.00002

94.08

1.6381

0.0048

34.59

0.40

Not

corr

ASA

-35b-13

0.0514

0.00016

0.01535

0.00006

0.00022

0.00002

-0.00476

0.00218

0.00004

0.00002

74.47

3.3516

0.0172

70.07

0.98

Not

corr

ASA

-35b-14

0.1254

0.00052

0.03259

0.00014

0.00047

0.00002

0.01166

0.00218

0.00003

0.00002

91.89

3.8496

0.0229

80.26

1.22

Not

corr

ASA

-35b-15

0.0759

0.00019

0.02233

0.00007

0.00022

0.00002

0.00991

0.00219

0.00003

0.00002

86.59

3.3996

0.0137

71.06

0.90

Not

corr

ASA

-35b-16

0.4125

0.00130

0.24488

0.00024

0.00320

0.00004

0.00522

0.00219

0.00002

0.00002

98.25

1.6847

0.0056

35.56

0.42

Not

corr

ASA

-35b-17

0.0108

0.00016

0.00370

0.00005

0.00006

0.00002

0.00287

0.00219

-0.00002

0.00002

142.57

2.9195

0.0572

61.19

2.43

Not

corr

ASA

-35b-18

0.0765

0.00017

0.04833

0.00007

0.00061

0.00002

0.00052

0.00219

0.00001

0.00002

94.42

1.5820

0.0043

33.42

0.38

Not

corr

Av

blank

1-2

0.0021

0.00005

0.00004

0.000010

0.00003

0.00001

0.00105

0.00002

0.00006

Av

blank

3-18

0.0021

0.00003

0.00005

0.000010

0.00006

0.00001

0.00096

0.00002

0.00006

J=

0.01186193

ASA

-84b-1

0.0559

0.00035

0.02588

0.00009

0.00035

0.00002

0.02259

0.00410

0.00002

0.00001

90.75

2.1595

0.0157

45.63

0.80

Not

corr

ASA

-84b-2

0.0537

0.00032

0.02316

0.00006

0.00032

0.00002

0.01835

0.00410

0.00004

0.00001

79.35

2.3173

0.0152

48.92

0.80

Not

corr

ASA

-84b-3

0.1206

0.00053

0.04201

0.00005

0.00052

0.00003

0.01655

0.00410

0.00019

0.00001

54.06

1.5520

0.0923

32.91

1.95

Corr

ASA

-84b-4

0.0295

0.00033

0.01328

0.00002

0.00011

0.00002

-0.00046

0.00411

0.00005

0.00001

52.38

2.2197

0.0251

46.89

1.14

Not

corr

ASA

-84b-5

0.1464

0.00046

0.01374

0.00007

0.00024

0.00002

0.03714

0.00531

0.00043

0.00002

13.81

1.4715

0.3689

31.22

7.76

Corr

ASA

-84b-6

0.1111

0.00027

0.01836

0.00009

0.00042

0.00003

0.00088

0.00531

0.00031

0.00002

18.31

1.1085

0.2754

23.57

5.82

Corr

ASA

-84b-7

0.1018

0.00040

0.04630

0.00013

0.00050

0.00003

0.06542

0.00533

0.00016

0.00002

54.38

1.1952

0.1094

25.40

2.31

Corr

ASA

-84b-8

0.2809

0.00065

0.11053

0.00006

0.00142

0.00003

0.01118

0.00533

0.00042

0.00002

56.12

1.4265

0.0461

30.27

0.98

Corr

ASA

-84b-9

0.4015

0.00104

0.15395

0.00023

0.00198

0.00003

0.17820

0.00533

0.00052

0.00002

61.94

1.6155

0.0472

34.24

1.01

Corr

ASA

-84b-10

0.2780

0.00027

0.08751

0.00019

0.00112

0.00004

0.04557

0.00533

0.00046

0.00002

51.41

1.6333

0.0579

34.62

1.23

Corr

ASA

-84b-11

0.3068

0.00043

0.11435

0.00017

0.00146

0.00003

0.02428

0.00534

0.00020

0.00002

81.01

2.1731

0.0444

45.92

0.95

Corr

ASA

-84b-12

0.0997

0.00027

0.05136

0.00005

0.00064

0.00003

0.00846

0.00534

0.00007

0.00002

80.09

1.5544

0.0984

32.96

2.07

Corr

Av

blank

1-4

0.0026

0.00006

0.00008

0.000010

0.00005

0.00001

0.00087

0.00002

0.00008

0.00001

Av

blank

5-12

0.0018

0.00004

0.00003

0.000010

0.00003

0.00001

0.00091

0.00002

0.00005

0.00001

229

230

References

Abraham, K., Hormann, P. & Raith, M. (1974). Pro-gressive metamorphism of basic rocks from the south-ern Hohe Tauern area, Tyrol (Austria). N. Jb. Miner.Abh, 122, 1–35. 11, 38

Agard, P. (2006). Transient, syn-obduction exhumationof Zagros blueschists inferred from PTdt and kinematicconstraints: implications for wedge dynamics. In Geo-physical Research Abstracts, vol. 8, 02912. 189

Agard, P., Monie, P., Jolivet, L. & Goffe, B.(2002). Exhumation of the Schistes Lustres complex:in situ laser probe 40Ar/39Ar constraints and impli-cations for the Western Alps. Journal of MetamorphicGeology , 20, 599–618. 201

Agard, P., Omrani, J., Jolivet, L. & Mouthereau,F. (2005). Convergence history across Zagros (Iran):constraints from collisional and earlier deformation.International Journal of Earth Sciences, 94, 401–419.189

Agard, P., Yamato, P., Jolivet, L. & Burov, E.(2009). Exhumation of oceanic blueschists and eclog-ites in subduction zones: timing and mechanisms.Earth Science Reviews, 92, 53–79. 1, 60, 188, 189

Ague, J. (1995). Deep crustal growth of quartz, kyan-ite and garnet into large-aperture, fluid-filled fractures,north-eastern Connecticut, USA. Journal of Metamor-phic Geology , 13, 299–314. 107

Ague, J. (2011). Extreme channelization of fluid and theproblem of element mobility during Barrovian meta-morphism. American Mineralogist , 96, 333. 196

Ague, J. & Baxter, E. (2007). Brief thermal pulsesduring mountain building recorded by Sr di↵usion inapatite and multicomponent di↵usion in garnet. Earthand Planetary Science Letters, 261, 500–516. 164, 195

Albarello, D., Mantovani, E., Babbucci, D. &Tamburelli, C. (1995). Africa-Eurasia kinematics:main constraints and uncertainties. Tectonophysics,243, 25–36. 3

Allen, A. & Stubbs, D. (1982). An 40 Ar/39 Arstudy of a polymetamorphic complex in the AruntaBlock, central Australia. Contributions to Mineralogyand Petrology , 79, 319–332. 153

Amato, J., Johnson, C., Baumgartner, L. & Beard,B. (1999). Rapid exhumation of the Zermatt-Saasophiolite deduced from high-precision Sm-Nd and Rb-Sr geochronology. Earth and Planetary Science Let-ters, 171, 425–438. 189, 201

Anczkiewicz, R., Platt, J., Thirlwall, M. & Wak-abayashi, J. (2004). Franciscan subduction o↵ to aslow start: Evidence from high-precision Lu-Hf garnetages on high grade-blocks. Earth and Planetary Sci-ence Letters, 225, 147–161. 189

Arculus, R., Lapierre, H. & Jaillard, E. (1999).Geochemical window into subduction and accretionprocesses: Raspas metamorphic complex, Ecuador.Geology , 27, 547. 51

Argus, D., Gordon, R., DeMets, C. & Stein, S.(1989). Closure of the Africa-Eurasia-North Americaplate motion circuit and tectonics of the Gloria fault.Journal of Geophysical Research, 94, 5585–5602. 3

Arnaud, N. & Kelley, S. (1995). Evidence for excessargon during high pressure metamorphism in the DoraMaira massif (Western Alps, Italy), using an ultra-violet laser ablation microprobe 40 Ar-39 Ar technique.Contributions to Mineralogy and Petrology, 121, 1–11.133

Ashworth, J. & Birdi, J. (1990). Di↵usion modelling ofcoronas around olivine in an open system. Geochimicaet Cosmochimica Acta, 54, 2389–2401. 34, 69

Ayers, J., Dunkle, S., Gao, S. & Miller, C. (2002).Constraints on timing of peak and retrograde meta-morphism in the Dabie Shan ultrahigh-pressure meta-morphic belt, east-central China, using U-Th-Pb dat-ing of zircon and monazite. Chemical Geology , 186,315–331. 189

Bachmann, R., Glodny, J., Oncken, O. & Seifert,W. (2009). Abandonment of the South Penninic-Austroalpine palaeosubduction zone, Central Alps,and shift from subduction erosion to accretion: con-straints from Rb/Sr geochronology. Journal of Geolog-ical Society , 166, 217. 174, 183, 197

Baldwin, S., Monteleone, B., Webb, L., Fitzger-ald, P., Grove, M. & Hill, E. (2004). Pliocene

231

eclogite exhumation at plate tectonic rates in easternPapua New Guinea. Nature, 431, 263–267. 1, 133, 188,189

Baroux, E., Bethoux, N. & Bellier, O. (2001). Anal-yses of the stress field in southeastern France fromearthquake focal mechanisms. Geophysical Journal In-ternational , 145, 336–348. 3

Barrow, G. (1893). On an intrusion of muscovite-biotitegneiss in the south-eastern Highlands of Scotland, andits accompanying metamorphism. Quarterly Journal ofthe Geological Society , 49, 330. 195

Barth, S., Oberli, F. & Meier, M. (1989). U–Th–Pb systematics of morphologically characterized zir-con and allanite: a high-resolution isotopic study ofthe Alpine Rensen pluton (northern Italy). Earth andplanetary science letters, 95, 235–254. 131, 199, 201

Baxter, E. & DePaolo, D. (2000). Field measurementof slow metamorphic reaction rates at temperatures of500 to 600 C. Science, 288, 1411. 164

Baxter, E., Ague, J. & Depaolo, D. (2002). Pro-grade temperature-time evolution in the Barroviantype-locality constrained by Sm/Nd garnet ages fromGlen Clova, Scotland. Journal of the Geological Soci-ety , 159, 71. 164, 195

Beaumont, C., Jamieson, R., Nguyen, M. & Lee,B. (2001). Himalayan tectonics explained by extrusionof a low-viscosity crustal channel coupled to focusedsurface denudation. Nature, 414, 738–742. 99

Beaumont, C., Nguyen, M., Jamieson, R. & Ellis,S. (2006). Crustal flow modes in large hot orogens.Geological Society London Special Publications, 268,91. 99

Becker, H. (1993). Garnet peridotite and eclogite Sm-Nd mineral ages from the Lepontine dome (SwissAlps): New evidence for Eocene high-pressure meta-morphism in the central Alps. Geology , 21, 599. 131,201

Behrmann, J. & Frisch, W. (1990). Sinistral ductileshearing associated with metamorphic decompressionin the Tauern Window, Eastern Alps. Jahrbuch derGeologischen Bundesanstalt Wien, 133, 135–146. 10

Behrmann, J. & Ratschbacher, L. (1989).Archimedes revisited: a structural test of eclog-ite emplacement models in the Austrian Alps. TerraNova, 1, 242–252. 30, 38, 122, 190

Beitter, T., Wagner, T. & Markl, G. (2008). For-mation of kyanite–quartz veins of the Alpe Sponda,Central Alps, Switzerland: implications for Al trans-port during regional metamorphism. Contributions toMineralogy and Petrology , 156, 689–707. 107

Beltrando, M., Lister, G., Rosenbaum, G.,Richards, S. & Forster, M. (2010a). Recognizingepisodic lithospheric thinning along a convergent platemargin: The example of the Early Oligocene Alps.Earth-Science Reviews. 198, 199, 200

Beltrando, M., Rubatto, D. & Manatschal, G.(2010b). From passive margins to orogens: Thelink between ocean-continent transition zones and(ultra)high-pressure metamorphism. Geology , 38,559–562. 96

Berger, A. & Bousquet, R. (2008). Subduction-related metamorphism in the Alps: review of isotopicages based on petrology and their geodynamic con-sequences. Geological Society London Special Publica-tions, 298, 117. 3, 100, 201

Berger, A., Rosenberg, C. & Schaltegger, U.(2009). Stability and isotopic dating of monazite andallanite in partially molten rocks: examples from theCentral Alps. Swiss Journal of Geosciences, 102, 15–29. 131

Berger, A., Schmid, S., Engi, M., Bousquet, R. &Wiederkehr, M. (2011). Mechanisms of mass andheat transport during Barrovian metamorphism: Adiscussion based on field evidence from the CentralAlps (Switzerland/northern Italy). Tectonics, 30, 17.5, 195

Berman, R. (1988). Internally-consistent thermody-namic data for minerals in the system Na2O-K2O-CaO-MgO-FeO-Fe2O3-Al2O3-SiO2-TiO2-H2O-CO2.Journal of Petrology , 29, 445. 33

Bickle, M. (1973). Studies across the Matrei Zone.Ph.D. thesis, Oxford. 16, 83, 85, 86, 174, 183

Bickle, M. & Hawkesworth, C. (1978). Deformationphases and the tectonic history of the eastern Alps.Bulletin of the Geological Society of America, 89, 293.100, 168

Bickle, M. & McKenzie, D. (1987). The transport ofheat and matter by fluids during metamorphism. Con-tributions to Mineralogy and Petrology , 95, 384–392.164

Bickle, M. & Pearce, J. (1975). Oceanic mafic rocksin the Eastern Alps. Contributions to Mineralogy andPetrology , 49, 177–189. 106

Bickle, M., Hawkesworth, C., England, P. &Athey, D. (1975). A preliminary thermal model forregional metamorphism in the Eastern Alps. Earth andPlanetary Science Letters, 26, 13. 1, 6, 8, 96, 100, 130,164, 165, 166, 167, 168, 169, 170, 171, 172, 174, 175

232

Birch, F., Roy, R. & Decker, E. (1968). Heat flowand thermal history. Studies of Appalachian geology:northern and maritime, 437. 170

Bjornerud, M. & Austrheim, H. (2006). Hot fluidsor rock in eclogite facies metamorphism. Nature, 440.164

Blanckenburg, F., Villa, I., Baur, H., Morteani,G. & Steiger, R. (1989). Time calibration of a PT-path from the Western Tauern Window, Eastern Alps:the problem of closure temperatures. Contributions toMineralogy and Petrology , 101, 1–11. 10, 134, 147,194, 201

Bleibinhaus, F. & Gebrande, H. (2006). Crustal struc-ture of the Eastern Alps along the TRANSALP profilefrom wide-angle seismic tomography. Tectonophysics,414, 51–69. 4

Bocchio, R., De Capitani, L., Ottolini, L. & Cella,F. (2000). Trace element distribution in eclogites andtheir clinopyroxene/garnet pair: a case study fromSoazza (Switzerland). European Journal of Mineral-ogy , 12, 147. 106

Borsi, S., Del Moro, A., Sassi, F. & Zirpoli, G.(1978). On the age of the Vendrette di Ries (Rieser-ferner) massif and its geodynamic significance. Geolo-gische Rundschau, 68. 172

Bostock, M., Hyndman, R., Rondenay, S. & Pea-cock, S. (2002). An inverted continental Moho andserpentinization of the forearc mantle. Nature, 417,536–538. 194

Bousquet, R., Goffe, B., Vidal, O., Oberhansli,R. & Patriat, M. (2002). The tectono-metamorphichistory of the Valaisan domain from the Western tothe Central Alps: New constraints on the evolution ofthe Alps. Geological Society of America Bulletin, 114,207. 62, 158, 174, 185, 201

Brace, W. & Kohlstedt, D. (1980). Limits on litho-spheric stress imposed by laboratory experiments. J.geophys. Res, 85, 6248–6252. 172

Brewer, M. (1969). Excess radiogenic argon in meta-morphic micas from the Eastern Alps, Austria. Earthand Planetary Science Letters, 6, 321. 100

Broadhurst, C., Drake, M., Hagee, B. & Berna-towicz, T. (1992). Solubility and partitioning of Ne,Ar, Kr and Xe in minerals and synthetic basaltic melts.Geochimica et cosmochimica acta, 56, 709–723. 154

Brooker, R., Wartho, J., Carroll, M., Kelley, S.& Draper, D. (1998). Preliminary UVLAMP deter-minations of argon partition coe�cients for olivine andclinopyroxene grown from silicate melts. Chemical ge-ology , 147, 185–200. 154

Brouwer, F., Van de Zedde, D., Wortel, M. &Vissers, R. (2004). Late-orogenic heating during ex-humation: Alpine PTt trajectories and thermome-chanical models. Earth and Planetary Science Letters,220, 185–199. 195, 196, 197

Brouwer, F., Burri, T., Engi, M. & Berger,A. (2005). Eclogite relics in the Central Alps: PT-evolution, Lu-Hf ages, and implications for formationof tectonic melange zones. Schweizerische Mineralogis-che und Petrographische Mitteilungen, 85, 147–174. 6,196, 198, 201

Bruce Watson, E. & Brenan, J. (1987). Fluids inthe lithosphere, 1. Experimentally-determined wettingcharacteristics of CO2H2O fluids and their implica-tions for fluid transport, host-rock physical properties,and fluid inclusion formation. Earth and Planetary Sci-ence Letters, 85, 497–515. 162

Brunet, C., Monie, P., Jolivet, L. & Cadet, J.(2000). Migration of compression and extension in theTyrrhenian Sea, insights from 40Ar/39Ar ages on mi-cas along a transect from Corsica to Tuscany. Tectono-physics, 321, 127–155. 201

Bucher, S., Schmid, S., Bousquet, R. & Fugen-schuh, B. (2003). Late-stage deformation in a colli-sional orogen (Western Alps): nappe refolding, back-thrusting or normal faulting? Terra Nova, 15, 109–117. 201

Burg, J. & Gerya, T. (2005). The role of viscous heat-ing in Barrovian metamorphism of collisional orogens:thermomechanical models and application to the Lep-ontine Dome in the Central Alps. Journal of Metamor-phic Geology , 23, 75–95. 99, 164, 197

Burov, E., Jolivet, L., Le Pourhiet, L. & Poli-akov, A. (2001). A thermomechanical model of ex-humation of high pressure (HP) and ultra-high pres-sure (UHP) metamorphic rocks in Alpine-type collisionbelts. Tectonophysics, 342, 113–136. 99

Butler, R. (1992). Thrusting patterns in the NWFrench Subalpine chains. In Annales Tectonicae, vol. 6,150–172. 3

Butler, R., Matthews, S. & Parish, M. (1986). TheNW external Alpine thrust belt and its implications forthe geometry of the western Alpine orogen. GeologicalSociety, London, Special Publications, 19, 245. 3

Caddick, M., Bickle, M., Harris, N., Holland, T.,Horstwood, M., Parrish, R. & Ahmad, T. (2007).Burial and exhumation history of a Lesser Himalayanschist: Recording the formation of an inverted meta-morphic sequence in NW India. Earth and PlanetaryScience Letters, 264, 375–390. 184

233

Calais, E., Nocquet, J., Jouanne, F. & Tardy, M.(2002). Current strain regime in the Western Alps fromcontinuous Global Positioning System measurements,1996-2001. Geology , 30, 651. 2, 3

Camacho, A., Lee, J., Hensen, B. & Braun, J.(2005). Short-lived orogenic cycles and the eclogitiza-tion of cold crust by spasmodic hot fluids. Nature, 435,1191–1196. 153, 155, 164

Cannic, S., Mugnier, J. & Lardeaux, J. (1999). Neo-gene extension in the Western Alps. Mem. Sci. Geol.Padova, 51, 33–45. 201

Carlson, W. (2002). Scales of disequilibrium and ratesof equilibration during metamorphism. American Min-eralogist , 87, 185. 34

Carlson, W. (2010). Dependence of reaction kinetics onH2O activity as inferred from rates of intergranular dif-fusion of aluminium. Journal of Metamorphic Geology ,28, 735–752. 34

Carmichael, D. (1969). On the mechanism of progrademetamorphic reactions in quartz-bearing pelitic rocks.Contributions to Mineralogy and Petrology , 20, 244–267. 34

Carpenter, M. (1982). Omphacite microstructures astime-temperature indicators of blueschist-and eclogite-facies metamorphism. Contributions to Mineralogy andPetrology , 78, 441–451. 68

Carry, N., Gueydan, F., Brun, J. & Marquer, D.(2009). Mechanical decoupling of high-pressure crustalunits during continental subduction. Earth and Plan-etary Science Letters, 278, 13–25. 61, 199

Carslaw, H. & Jaeger, J. (1959). Conduction of heatin solids. Oxford: Clarendon Press. 169

Catlos, E., Sorensen, S. & Harrison, T. (2000). Th-Pb ion-microprobe dating of allanite. American Min-eralogist , 85, 633. 111

Ceriani, S., F”ugenschuh, B. & Schmid, S. (2001). Multi-stagethrusting at the” Penninic Front” in the Western Alpsbetween Mont Blanc and Pelvoux massifs. Interna-tional Journal of Earth Sciences, 90, 685–702. 5

Cesare, B., Poletti, E., Boiron, M. & Cathelin-eau, M. (2001). Alpine metamorphism and veining inthe Zentralgneis Complex of the SW Tauern Window:a model of fluid-rock interactions based on fluid inclu-sions. Tectonophysics, 336, 121–136. 10, 94

Challandes, N., Marquer, D. & Villa, I. (2003).Dating the evolution of CS microstructures: a com-bined 40Ar/39Ar step-heating and UV laserprobe

analysis of the Alpine Ro↵na shear zone. Chemical ge-ology , 197, 3–19. 201

Chamberlain, C. & Rumble, D. (1988). Thermalanomalies in a regional metamorphic terrane: An iso-topic study of the role of fluids. Journal of Petrology ,29, 1215. 164

Chauvet, A., Kienast, J., Pinardon, J. & Brunel,M. (1992). Petrological constraints and PT path of De-vonian collapse tectonics within the Scandian moun-tain belt (Western Gneiss Region, Norway). Journalof the Geological Society , 149, 383–400. 52

Cherniak, D. (1993). Lead di↵usion in titanite and pre-liminary results on the e↵ects of radiation damage onPb transport. Chemical Geology , 110, 177–194. 103

Chernoff, C. & Carlson, W. (1999). Trace elementzoning as a record of chemical disequilibrium duringgarnet growth. Geology , 27, 555. 123

Chopin, C., Henry, C. & Michard, A. (1991). Geol-ogy and petrology of the coesite-bearing terrain, DoraMaira massif, Western Alps. European Journal of Min-eralogy , 3, 263–291. 60

Choukroune, P., Ballevre, M., Cobbold, P., Gau-tier, Y., Merle, O. & Vuichard, J. (1986). Defor-mation and motion in the western Alpine arc. Tecton-ics, 5, 215–226. 3

Christensen, J., Selverstone, J., Rosenfield,J. & DePaolo, D. (1994). Correlation by Rb–Srgeochronology of garnet growth histories from di↵erentstructural levels within the Tauern Window, EasternAlps. Contributions to Mineralogy and Petrology , 118,1–12. 100

Clark, S. (1966). Handbook of physical constraints. Ge-ological Society of America Bulletin, 97. 180

Cliff, R. & Cohen, A. (1980). Uranium-lead isotopesystematics in a regionally metamorphosed tonalitefrom the Eastern Alps. Earth and Planetary ScienceLetters, 50, 211–218. 13

Cliff, R., Norris, R., Oxburgh, E. & Wright, R.(1971). Structural, metamorphic, and geochronologicalstudies in the Reisseck and southern Ankogel groups,the Eastern Alps. Jb. Geol. B.-A, 114, 121–272. 16

Cliff, R., Barnicoat, A. & Inger, S. (1998). EarlyTertiary eclogite facies metamorphism in the MonvisoOphiolite. Journal of Metamorphic Geology , 16, 447–456. 201

Cloos, M. (1982). Flow melanges: numerical modelingand geologic constraints on their origin in the Francis-can subduction complex, California. Geological Societyof America Bulletin, 93, 330. 183, 194

234

Cloos, M. (1983). Comparative study of melange matrixand metashales from the Franciscan subduction com-plex with the basal Great Valley sequence, California.The Journal of Geology , 91, 291–306. 195

Cloos, M. (1986). Blueschists in the Franciscan Com-plex of California: Petrotectonic constraints on upliftmechanisms. Geological Society of America Memoir ,164, 77–93. 189, 195

Coggon, R. & Holland, T. (2002). Mixing propertiesof phengitic micas and revised garnet-phengite ther-mobarometers. Journal of Metamorphic Geology , 20,683–696. 43, 44, 50, 52, 53, 56, 67

Connolly, J. & Thompson, A. (1989). Fluid andenthalpy production during regional metamorphism.Contributions to Mineralogy and Petrology, 102, 347–366. 160

Corfu, F., Hanchar, J., Hoskin, P. & Kinny, P.(2003). Atlas of zircon textures. Reviews in Mineralogyand Geochemistry , 53, 469. 106

Cornelius, H., Clar, E. & fur Bodenforschung,R. (1939). Geologie des Grossglocknergebietes. Zweig-stelle Wien der Reichstelle fur Bodenforschung fruherGeologische Bundesanstalt. 6, 11, 38, 164

Corrie, S. & Kohn, M. (2008). Trace-element distribu-tions in silicates during prograde metamorphic reac-tions: implications for monazite formation. Journal ofMetamorphic Geology , 26, 451–464. 123, 125, 126

Cortiana, G., Dal Piaz, G., Del Moro, A., Hun-ziker, J. & Martin, S. (1998). 40Ar-39Ar andRb-Sr dating of the Pillonet klippe and Sesia-Lanzobasal slice in the Ayas valley and evolution of theAustroalpine-Piedmont nappe stack. Mem Soc GeolItal , 50, 177–194. 201

Coward, M. & Dietrich, D. (1989). Alpine tectonicsan overview. Geological Society, London, Special Pub-lications, 45, 1. 2

Coyle, D. & Wagner, G. (1998). Positioning the ti-tanite fission-track partial annealing zone. ChemicalGeology , 149, 117–125. 103

Crank, J. (1975). The mathematics of di↵usion. 2nd.Ed., Clarendon Press, Oxford . 147

Crovetto, R., Fernandez-Prini, R. & Japas, M.(1982). Solubilities of inert gases and methane in HOand in DO in the temperature range of 300 to 600 K.The Journal of Chemical Physics, 76, 1077. 154

Cumbest, R., Johnson, E. & Onstott, T. (1994). Ar-gon composition of metamorphic fluids: Implicationsfor 40Ar/39Ar geochronology. Bulletin of the Geologi-cal Society of America, 106, 942. 157

Dachs, E. (1986). High-pressure mineral assemblagesand their breakdown-products in metasediments Southof the Grossvenediger, Tauern Window, Austria.Schweizerische Mineralogische und PetrographischeMitteilungen, 66, 145–161. 38

Dachs, E. (1990). Geothermobarometry in metasedi-ments of the southern Grossvenediger area (TauernWindow, Austria). Journal of metamorphic Geology ,8, 217–230. 10, 80, 81, 174

Dachs, E. & Proyer, A. (2001). Relics of high-pressuremetamorphism from the Grossglockner region, HoheTauern, Austria: Paragenetic evolution and PT-pathsof retrogressed eclogites. European Journal of Miner-alogy , 13, 67. 10, 62, 81, 82, 95, 97, 149, 174

Dachs, E. & Proyer, A. (2002). Constraints on the du-ration of high-pressure metamorphism in the TauernWindow from di↵usion modelling of discontinuousgrowth zones in eclogite garnet. Journal of metamor-phic Geology , 20, 769–780. 128

Dachs, E., Kurz, W. & Proyer, A. (2005). Alpineeclogites in the tauern window. Mitt. Osterr. Geol.Ges, 150, 189–216. 38

Daczko, N. & Halpin, J. (2009). Evidence for meltmigration enhancing recrystallization of metastableassemblages in mafic lower crust, Fiordland, NewZealand. Journal of Metamorphic Geology , 27, 167–185. 68

Dal Piaz, G., Cortiana, G., Del Moro, A., Mar-tin, S., Pennacchioni, G. & Tartarotti, P.(2001). Tertiary age and paleostructural inferences ofthe eclogitic imprint in the Austroalpine outliers andZermatt-Saas ophiolite, western Alps. InternationalJournal of Earth Sciences, 90, 668–684. 201

Dale, J. & Holland, T.J.B. (2003). Geothermobarom-etry, P–T paths and metamorphic field gradients ofhigh-pressure rocks from the Adula Nappe, CentralAlps. Journal of Metamorphic Geology , 21, 813–829.53

Dale, J., Powell, R., White, R., Elmer, F.& Holland, T. (2005). A thermodynamic modelfor Ca–Na clinoamphiboles in Na2O–CaO–FeO–MgO–Al2O3–SiO2–H2O–O for petrological calculations.Journal of Metamorphic Geology , 23, 771–791. 68

De Sigoyer, J., Guillot, S., Lardeaux, J. & Mas-cle, G. (1997). Glaucophane-bearing eclogites in theTso Morari dome (eastern Ladakh, NW Himalaya).European Journal of Mineralogy , 9, 1073. 1, 196

De Sigoyer, J., Chavagnac, V., Blichert-Toft, J.,Villa, I., Luais, B., Guillot, S., Cosca, M. &

235

Mascle, G. (2000). Dating the Indian continentalsubduction and collisional thickening in the northwestHimalaya: Multichronology of the Tso Morari eclog-ites. Geology , 28, 0091–7613. 1, 133, 162, 188, 189,195, 196

De Sigoyer, J., Guillot, S. & Dick, P. (2004). Ex-humation of the ultra high-pressure Tso Morari unitin eastern Ladakh (NW Himalaya): a case study. Tec-tonics, 23, 3. 1

Deer, W., Howie, R. & Zussman, J. (1992). The Rock-Forming Minerals. Prentice Hall, 2nd edn. 68, 106,109, 111

Deichmann, N. & Rybach, L. (1989). Earthquakes andtemperatures in the lower crust below the northernAlpine foreland of Switzerland. Properties and Pro-cesses of the Earth’s Lower Crust , 197–213. 3

Del Moro, A., Pardini, G., Quercioli, C., Villa, I.& Callegari, E. (1983). Rb/Sr and K/Ar chronol-ogy of Adamello granitoids, southern Alps. Mem. Soc.Geol. Ital , 26, 285–299. 199

Delacou, B., Sue, C., Champagnac, J. & Burkhard,M. (2004). Present-day geodynamics in the bend of thewestern and central Alps as constrained by earthquakeanalysis. Geophysical Journal International , 158, 753–774. 3

DeMets, C., Gordon, R., Argus, D. & Stein, S.(1990). Current plate motions. Geophysical journal in-ternational , 101, 425–478. 3

DeMets, C., Gordon, R., Angus, D. & Stein, S.(1994). E↵ect of recent revisions to the geomagneticreversal time scale on estimates of current plate mo-tions. Geophys. Res. Lett . 3

Deutsch, A. (1984a). Daiterung an Alkaliamphibolenund Stilpnomelan der sudlichen Platta-Decke. EclogaeGeol. Helv , 76. 201

Deutsch, A. (1984b). Young Alpine dykes south of theTauern Window (Austria): a K-Ar and Sr isotopestudy. Contributions to Mineralogy and Petrology , 85,45–57. 172

Dewey, J., Helman, M., Knott, S., Turco, E.& Hutton, D. (1989). Kinematics of the westernMediterranean. Geological Society London Special Pub-lications, 45, 265. 2, 184

Diener, J. & Powell, R. (In Press). Revised activity-composition models for clinopyroxene and amphibole.Journal of Metamorphic Geology . 68, 69

Diener, J., Powell, R., White, R. & Holland, T.(2007). A new thermodynamic model for clino-and or-thoamphiboles in the system Na2O–CaO–FeO–MgO–Al2O3–SiO2–H2O–O. Journal of Metamorphic Geol-ogy , 25, 631–656. 68, 69, 71, 75, 76, 77

Dingeldey, C., Dallmeyer, R., Koller, F. & Mas-sonne, H. (1997). P-T-t history of the Lower Aus-troalpine Nappe Complex in the Tarntaler Berge NWof the Tauern Window: implications for the geotec-tonic evolution of the central Eastern Alps. Contribu-tions to mineralogy and petrology , 129, 1–19. 201

Dodson, M. (1973). Closure temperature in coolinggeochronological and petrological systems. Contribu-tions to Mineralogy and Petrology , 40, 259–274. 102,133, 204

Dodson, M. (1976). Kinetic processes and thermal his-tory of slowly cooling solids. Nature, 40, 259–274. 102

Dodson, M. (1979). Theory of cooling ages. In E. Jager& J. Hunziker, eds., Lectures in Isotope Geology , 194–202, Springer-Verlag. 102

Droop, G. (1985). Alpine metamorphism in the south-east Tauern Window, Austria: 1 P-T variations inspace and time. Journal of Metamorphic Geology , 3,371–402. 174, 180, 184

Droop, G., Lombardo, B. & Pognante, U. (1990).Formation and distribution of eclogite facies rocks inthe Alps. Eclogite Facies Rocks, 225. 3

Duchene, S., Blichert-Toft, J., Luais, B., Telouk,P., Lardeaux, J. & Albarede, F. (1997). The Lu-Hf dating of garnets and the ages of the Alpine high-pressure metamorphism. Nature, 387, 586–589. 3, 131,197, 201

Dyar, M.D., Lowe, E.W., Guidotti, C.V. & De-laney, J.S. (2002). Fe3+ and Fe2+ partitioning amongsilicates in metapelites: A synchrotron micro-XANESstudy. American Mineralogist , 87, 514–522. 44

Ebbing, J., Braitenberg, C. & Gotze, H. (2001). For-ward and inverse modelling of gravity revealing insightinto crustal structures of the Eastern Alps. Tectono-physics, 337, 191–208. 173

Ebbing, J., Braitenberg, C. & Gotze, H. (2006).The lithospheric density structure of the Eastern Alps.Tectonophysics, 414, 145–155. 4, 173

Eichhorn, R., Loth, G., Holl, R., Finger, F.,Schermaier, A. & Kennedy, A. (2000). MultistageVariscan magmatism in the central Tauern Window(Austria) unveiled by U/Pb SHRIMP zircon data.Contributions to Mineralogy and Petrology, 139, 418–435. 13

236

Engi, M., Todd, C. & Schmatz, D. (1995). Ter-tiary metamorphic conditions in the eastern LepontineAlps. Schweizerische Mineralogische und Petrographis-che Mitteilungen, 75, 347–369. 196

Engi, M., Berger, A. & Roselle, G. (2001). Role ofthe tectonic accretion channel in collisional orogeny.Geology , 29, 1143. 164, 195, 197

England, P. (1978). Some thermal considerations ofthe Alpine metamorphism–past, present and future.Tectonophysics, 46, 21–40. 100, 164, 165, 168, 170,171, 174

England, P. (1981). Metamorphic pressure estimatesand sediment volumes for the Alpine orogeny: an in-dependent control on geobarometers? Earth and Plan-etary Science Letters, 56, 387–397. 180, 181, 184

England, P. & Holland, T. (1979). Archimedes andthe Tauern eclogites: the role of buoyancy in thepreservation of exotic eclogite blocks. Earth and Plan-etary Science Letters, 44, 287–294. 60, 173, 183, 190,191

England, P. & Richardson, S. (1977). The influenceof erosion upon the mineral facies of rocks from di↵er-ent metamorphic environments. Journal of GeologicalSociety , 134, 201. 1, 164, 165, 171

England, P. & Thompson, A. (1984). Pressure–Temperature–Time Paths of Regional MetamorphismI. Heat Transfer during the Evolution of Regions ofThickened Continental Crust. Journal of Petrology ,25, 894. 1, 6, 8, 102, 130, 164, 166, 167, 170, 171,196

Ercit, T. (2002). The mess that is” allanite”. CanadianMineralogist , 40, 1411. 42, 126

Eremin, K. (1994). Metamorphic Conditions within theEclogite Zone, Tauern Window, Austria. Ph.D. thesis,University of Cambridge. 38, 95, 97, 130, 190

Ernst, W. (1973a). Blueschist metamorphism and PTregimes in active subduction zones. Tectonophysics,17, 255–272. 68

Ernst, W. (1973b). Interpretative synthesis of metamor-phism in the Alps. Geological Society of America Bul-letin, 84, 2053. 3

Ernst, W. & Liou, J. (1995). Contrasting plate-tectonicstyles of the Qinling-Dabie-Sulu and Franciscan meta-morphic belts. Geology , 23, 353. 189

Ernst, W., Maruyama, S. & Wallis, S. (1997).Buoyancy-driven, rapid exhumation of ultrahigh-pressure metamorphosed continental crust. Proceed-ings of the National Academy of Sciences of the UnitedStates of America, 94, 9532. 188, 189

Eva, E. & Solarino, S. (1998). Variations of stress di-rections in the western Alpine arc. Geophysical JournalInternational , 135, 438–448. 3

Evans, B. (1990). Phase relations of epidote-blueschists.Lithos, 25, 3–23. 94

Exley, R. (1980). Microprobe studies of REE-rich acces-sory minerals: implications for Skye granite petrogen-esis and REE mobility in hydrothermal systems. Earthand Planetary Science Letters, 48, 97–110. 109

Exner, C. (1971). Geologie der peripheren Hafnergruppe(Hohe Tauern). Jahrbuch der Geologischen Bunde-sanstalt , 114, 1–119. 9, 14

Exner, C. (1980). Geologie der Hohen Tauern beiGmund in Karnten. Jahrbuch der Geologischen Bun-desanstalt , 123. 9, 14

Faccenda, M., Gerya, T. & Chakraborty, S. (2008).Styles of post-subduction collisional orogeny: Influ-ence of convergence velocity, crustal rheology and ra-diogenic heat production. Lithos, 103, 257–287. 99

Federico, L., Capponi, G., Crispini, L., Scambel-luri, M. & Villa, I. (2005). 39Ar/40Ar dating ofhigh-pressure rocks from the Ligurian Alps: Evidencefor a continuous subduction-exhumation cycle. Earthand Planetary Science Letters, 240, 668–680. 201

Feenstra, A. (1996). An EMP and TEM-AEM Studyof Margarite, Muscovite and Paragonite in Polymeta-morphic Metabauxites of Naxos Cyclades, Greece) andthe Implications of Fine-scale Mica Interlayering andMultiple Mica Generations. Journal of Petrology , 37,201–233. 44

Fei, Y. (1995). Thermal Expansion, 29–44. AmericanGeophysical Union, Washington, D.C. 180

Feininger, T. (1980). Eclogite and related high-pressure regional metamorphic rocks from the Andesof Ecuador. Journal of Petrology , 21, 107–140. 51

Feininger, T. & Silberman, M. (1982). K–Argeochronology of basement rocks on the NorthernFlank of the Huancabamba deflection, Ecuador. Open-File Rep.-US Geol. Surv , 82–206. 51

Ferry, J. & Spear, F. (1978). Experimental calibrationof the partitioning of Fe and Mg between biotite andgarnet. Contributions to mineralogy and petrology , 66,113–117. 34

Fick, A. (1855). Ueber di↵usion. Ann. Phys., 170, 59–89. 144

237

Finger, F., Frasl, G., Haunschmid, B., Lettner,H., Schermaier, A., Schindlmayr, A., Steyrer,H. & von Quadt, A. (1993). The Zentralgneise of theTauern window (Eastern Alps): insight into an intra-alpine Variscan batholith. The pre-Mesozoic Geologyin the Alps, 375–391. 9, 13

Foland, K. (1979). Limited mobility of argon in a Meta-morphic Terrain. Geochimica et Cosmochimica Acta,43, 793–801. 143, 153, 157

Foster, G. & Parrish, R. (2003). Metamorphic mon-azite and the generation of PTt paths. Geochronology:linking the isotopic record with petrology and textures,220, 25–47. 99

Foster, G., Kinny, P., Vance, D., Prince, C. &Harris, N. (2000). The significance of monazite U–Th–Pb age data in metamorphic assemblages; a com-bined study of monazite and garnet chronometry.Earth and Planetary Science Letters, 181, 327–340.99

Foster Jr, C. (1986). Thermodynamic models of re-actions involving garnet in a sillimanite/stauroliteschist= Modele thermodynamique des reactions impli-quant le grenat dans un schiste a sillimanite/staurolite.Mineralogical Magazine, 50, 427–439. 34

Frank, W. (1987). Evolution of the Austroalpine ele-ments in the Cretaceous. Geodynamics of the EasternAlps, 379–406. 5, 9

Frank, W., Hock, V. & Miller, C. (1987). Metamor-phic and tectonic history of the central Tauern Win-dow. Geodynamics of the Eastern Alps, 34–55. 38, 100

Franz, G. & Spear, F. (1983). High pressure metamor-phism of siliceous dolomites from the central TauernWindow, Austria. American Journal of Science, 283,396–413. 38

Franz, G., Mosbruggr, V. & Menge, R. (1991).Carbo-Permian pteridophyll leaf fragments from anamphibolite facies basement, Tauern Window, Aus-tria. Terra Nova, 3, 137–141. 10, 83

Franz, L., Romer, R., Klemd, R., Schmid, R.,OberhaEnsli, R., Wagner, T. & Shuwen, D.(2001). Eclogite-facies quartz veins within metaba-sites of the Dabie Shan (eastern China): pressure-temperature-time-deformation path, composition ofthe fluid phase and fluid flow during exhumation ofhigh-pressure rocks. Contributions to Mineralogy andPetrology , 141, 322–346. 107

Frechet, J. (1978). Sismicite du Sud-Est de la Franceet une nouvelle methode de zonage sismique. Ph.D.thesis, Universite des Sciences Technologiques etMedicales, Grenoble. 3

Freeman, S., Inger, S., Butler, R. & Cliff, R.(1997). Dating deformation using Rb-Sr in white mica:Greenschist facies deformation ages from the Entrelorshear zone, Italian Alps. Tectonics, 16, 57–76. 198,201

Frisch, W. (1980). Tectonics of the western Tauern Win-dow. Mitteilungen der Osterreichischen GeologischenGesellschaft , 71, 65–71. 8

Frisch, W. & Raab, D. (1987). Early Paleozoic back-arc and island-arc settings in Greenstone sequencesof the Central Tauern Window (Eastern Alps). Jahrb.Geol. Bundesanst.(Austria), 129, 545–566. 9, 11, 13,16, 27, 84

Frisch, W., Vavra, G. & Winkler, M. (1993). Evo-lution of the Penninic basement of the Eastern Alps.The Pre-Mesozoic geology in the Alps. Springer, BerlinHeidelberg New York , 349–360. 9, 13, 89

Froitzheim, N. & Manatschal, G. (1996). Kinematicsof Jurassic rifting, mantle exhumation, and passive-margin formation in the Austroalpine and Penninicnappes (eastern Switzerland). Geological Society ofAmerica Bulletin, 108, 1120. 5, 198

Froitzheim, N., SCHMID, S. & Conti, P. (1994).Repeated change from crustal shortening to orogen-parallel extension in the Austroalpine units of Graub”unden. Eclogae Geologicae Helvetiae, 87, 559–612. 5,199

Fry, N. (1973). Lawsonite pseudomorphed in Tauerngreenschist. Mineral. Mag , 39, 121–122. 24

Fry, N. (1989). Southwestward thrusting and tectonics ofthe western Alps. Geological Society, London, SpecialPublications, 45, 83. 3

Fugenschuh, B., Loprieno, A., Ceriani, S. &Schmid, S. (1999). Structural analysis of the Sub-brianconnais and Valais units in the area of Mutiers(Savoy, Western Alps). International Journal of EarthSciences, 88, 201–218. 201

Gabriele, P., Ballevre, M., Jaillard, E. &Hernandez, J. (2003). Garnet-chloritoid-kyanitemetapelites from the Raspas Complex (SW Ecuador):a key eclogite-facies assemblage. European Journal ofMineralogy , 15, 977–989. 50, 51, 57, 58, 60

Gabudianu Radulescu, I., Rubatto, D., Gregory,C. & Compagnoni, R. (2009). The age of HP meta-morphism in the Gran Paradiso Massif, Western Alps:A petrological and geochronological study of ”silverymicaschists”. Lithos, 110, 95–108. 198, 201

238

Ganguly, J., Tirone, M. & Hervig, R. (1998). Di↵u-sion Kinetics of Samarium and Neodymium in Garnet,and a Method for Determining Cooling Rates of Rocks.Science, 281, 805–807. 103

Gebauer, D. (1999). Alpine geochronology of the Cen-tral and Western Alps: new constraints for a complexgeodynamic evolution. Schweizerische Mineralogischeund Petrographische Mitteilungen, 79, 191–208. 198

Gebauer, D., Basu, A. & Hart, S. (1996). A PTtpath for an (ultra?) high pressure ultramafic/maficrock-association and its felsic country-rocks based onSHRIMP-dating of magmatic and metamorphic zircondomains. Example: Alpe Arami (Central Swiss Alps).Earth processes: reading the isotopic code, 309–328.131, 201

Gebauer, D., Schertl, H., Brix, M. & Schreyer, W.(1997). 35 Ma old ultrahigh-pressure metamorphismand evidence for very rapid exhumation in the DoraMaira Massif, Western Alps. Lithos, 41, 5–24. 131,197, 198, 201

Genser, J. & Neubauer, F. (1989). Low angle nor-mal faults at the eastern margin of the Tauern window(Eastern Alps). Mitt.”osterr. geol. Ges, 81, 233–243. 10

Gerya, T. & Stockhert, B. (2006). Two-dimensionalnumerical modeling of tectonic and metamorphic his-tories at active continental margins. InternationalJournal of Earth Sciences, 95, 250–274. 61, 62, 96,99, 187, 194, 195

Gerya, T., Stockhert, B. & Perchuk, A. (2002). Ex-humation of high-pressure metamorphic rocks in a sub-duction channel: a numerical simulation. Tectonics,21, 1056. 194

Gerya, T., Yuen, D. & Maresch, W. (2004). Thermo-mechanical modelling of slab detachment. Earth andPlanetary Science Letters, 226, 101–116. 99, 200

Getty, S. & Selverstone, J. (1994). Stable isotopicand trace element evidence for restricted fluid migra-tion in 2 GPa eclogites. Journal of metamorphic Geol-ogy , 12, 747–747. 163

Giere, R. & Sorensen, S. (2004). Allanite and otherREE-rich epidote-group minerals. Reviews in Mineral-ogy and Geochemistry , 56, 431. 109, 111

Giere, R., Virgo, D. & Popp, R. (1999). Oxidationstate of iron and incorporation of REE in igneous al-lanite. Journal of Conference Abstracts, 4, 721. 109

Gleißner, P., Glodny, J. & Franz, G. (2007). Rb-Sr isotopic dating of pseudomorphs after lawsonite in

metabasalts from the Glockner nappe, Tauern Win-dow, Eastern Alps. European Journal of Mineralogy ,19, 723–734. 10, 63, 75, 76, 79, 80, 81, 102, 128, 129,152, 153, 174

Glodny, J., Ring, U., Kuhn, A., Gleissner, P. &Franz, G. (2005). Crystallization and very rapid ex-humation of the youngest Alpine eclogites (TauernWindow, Eastern Alps) from Rb/Sr mineral as-semblage analysis. Contributions to Mineralogy andPetrology , 149, 699–712. 1, 6, 11, 96, 100, 101, 128,129, 135, 152, 153, 166, 174, 189, 190, 193, 201

Glodny, J., Ring, U. & Kuhn, A. (2008). Coeval high-pressure metamorphism, thrusting, strike-slip, and ex-tensional shearing in the Tauern Window, EasternAlps. Tectonics, 27. 1, 9, 10, 190, 199, 201

Gottschalk, M. (1997). Internally consistent ther-modynamic data for rock forming minerals inthe system SiO2-TiO2-Al2O3-Fe2O3-CaO-MgO-FeO-K2O-Na2O-H2O-CO2. European Journal ofMineralogy-Ohne Beihefte, 9, 175–224. 33

Gratier, J., Menard, G. & Arpin, R. (1989). Strain-displacement compatibility and restoration of theChaınes Subalpines of the western Alps. Geological So-ciety, London, Special Publications, 45, 65. 3

Green, E., Holland, T. & Powell, R. (2007). Anorder-disorder model for omphacitic pyroxenes in thesystem jadeite-diopside-hedenbergite-acmite, with ap-plications to eclogitic rocks. American Mineralogist ,92, 1181. 68, 69

Greenwood, L. (2009). Study of the Metamorphic Evo-lution of the Eclogite Zone, Tauern Window, Austria.Master’s thesis, University of Cambridge. 43

Gregory, C., Rubatto, D., Allen, C., Williams, I.,Hermann, J. & Ireland, T. (2007). Allanite micro-geochronology: A LA-ICP-MS and SHRIMP U-Th-Pbstudy. Chemical Geology , 245, 162–182. 111, 113, 114,115, 116, 119, 120, 121, 219

Grujic, D., Casey, M., Davidson, C., Hollister, L.,Kundig, R., Pavlis, T. & Schmid, S. (1996). Duc-tile extrusion of the Higher Himalayan Crystalline inBhutan: evidence from quartz microfabrics. Tectono-physics, 260, 21–43. 164

Guillot, S., De Sigoyer, J., Lardeaux, J. & Mas-cle, G. (1997). Eclogitic metasediments from theTso Morari area (Ladakh, Himalaya): Evidence forcontinental subduction during India-Asia convergence.Contributions to Mineralogy and Petrology, 128, 197–212. 196

239

Guillot, S., Hattori, K. & de Sigoyer, J. (2000).Mantle wedge serpentinization and exhumation ofeclogites: insights from eastern Ladakh, northwest Hi-malaya. Geology , 28, 199. 194

Guillot, S., Maheo, G., De Sigoyer, J., Hattori,K. & Kecher, A. (2008). Tethyan and Indian sub-duction viewed from the Himalayan high-to ultrahigh-pressure metamorphic rocks. Tectonophysics, 451,225–241. 196

Guillot, S., Hattori, K., Agard, P., Schwartz, S.& Vidal, O. (2009). Exhumation processes in oceanicand continental subduction contexts: a review. Sub-duction Zone Geodynamics, 175–205. 188

Guillot, S. and Hattori, K.H. and de Sigoyer, J.and Nagler, T. and Auzende, A.L. (2001). Evi-dence of hydration of the mantle wedge and its rolein the exhumation of eclogites. Earth and PlanetaryScience Letters, 193, 115–127. 194

Haas, J., Kovacs, S., Krystyn, L. & Lein, R. (1995).Significance of Late Permian-Triassic facies zones interrane reconstructions in the Alpine-North Pannoniandomain. Tectonophysics, 242, 19–40. 5

Hacker, B. & Wang, Q. (1995). Ar/Ar geochronologyof ultrahigh-pressure metamorphism in central China.Tectonics, 14, 994–1006. 133

Hacker, B., Andersen, T., Root, D., Mehl, L.,Mattinson, J. & Wooden, J. (2003). Exhumation ofhigh-pressure rocks beneath the Solund Basin, West-ern Gneiss Region of Norway. Journal of MetamorphicGeology , 21, 613–629. 52

Hacker, B., Ratschbacher, L. & Liou, J. (2004).Subduction, collision and exhumation in the ultrahigh-pressure Qinling-Dabie orogen. Aspects of the Tectonicevolution of China, 157–75. 189

Handy, M., Herwegh, M., Kamber, B., Tietz, R. &Villa, I. (1996). Geochronologic, petrologic and kine-matic constraints on the evolution of the Err-Plattaboundary, part of a fossil continent-ocean suture in theAlps (eastern Switzerland). Schweiz. Mineral. Petrogr.Mitt , 76, 453–474. 201

Hansmann, W. & Oberli, F. (1991). Zircon inheritancein an igneous rock suite from the southern Adamellobatholith (Italian Alps). Contributions to Mineralogyand Petrology , 107, 501–518. 131

Harley, S., Kelly, N. & Moller, A. (2007). Zir-con behaviour and the thermal histories of mountainchains. Elements, 3, 25. 99

Harrison, T. (1981). Di↵usion of 40Ar in hornblende.Contributions to Mineralogy and Petrology , 78, 324–331. 103

Harrison, T. & McDougall, I. (1981). Excess 40Arin metamorphic rocks from Broken Hill, New SouthWales: implications for40Ar/39Ar age spectra and thethermal history of the region. Earth and Planetary Sci-ence Letters, 55, 123–149. 156, 157

Harrison, T. & McDougall, I. (1982). The thermalsignificance of potassium feldspar K-Ar ages inferredfrom 40Ar/39Ar age spectrum results. Geochimica etCosmochimica Acta, 46, 1811–1820. 103

Harrison, T. & McDougall, I. (1985). Di↵usion of40Ar in biotite: Temperature, pressure and composi-tional e↵ects. Geochimica et Cosmochimica Acta, 49,2461–2468. 103

Harrison, T., Celerier, J., Aikman, A., Hermann,J. & Heizler, M. (2009). Di↵usion of 40Ar in mus-covite. Geochimica et Cosmochimica Acta, 73, 1039–1051. 146, 147, 151, 152, 163

Harte, B. (1975). Determination of a pelite petroge-netic grid for the Determination of a pelite petroge-netic grid for the eastern Scottish Dalradian. Yearbookof the Carnegie Institution of Washington, 74, 438–446. 50

Harte, B. & Hudson, N. (1979a). Pelite facies seriesand the temperatures and pressures of Dalradian meta-morphism in E Scotland. Geological Society LondonSpecial Publications, 8, 323. 35

Harte, B. & Hudson, N.F.C. (1979b). Pelite faciesseries and the temperatures and pressures of Dalra-dian metamorphism in eastern Scotland. In A.L. Har-ris, C.H. Holland & B.E. Leake, eds., The Caledonidesof the British Isles Reviewed , vol. 8, 323–337, GeologySociety Special Publication, London. 50

Harvey, J. & Baxter, E. (2009). An improved methodfor TIMS high precision neodymium isotope analysis ofvery small aliquots (1-10 ng). Chemical Geology , 258,251–257. 204

Hawkesworth, C. (1974a). Geochemical studies in andaround the southest corner of the Tauern Window .Ph.D. thesis, Oxford. 13, 170, 175, 177

Hawkesworth, C. (1974b). Vertical distribution of heatproduction in the basement of the Eastern Alps. Na-ture, 249, 435–436. 175, 177

Heaman, L. & Parrish, R. (1991). U-Pb geochronol-ogy of accessory minerals. In L. Heaman & J. Lud-den, eds., Applications of Radiogenic Isotope Systemsto Problems in Geology , vol. 19, 59–102, MineralogicalAsociation of Canada. 103, 111

240

Helgeson, H., Delany, J., Nesbitt, H. & Bird, D.(1978). Summary and critique of the thermodynamicproperties of rock forming minerals. American Journalof Science, 278A, 1–229. 33, 34

Hermann, J. (2002a). Allanite: thorium and light rareearth element carrier in subducted crust. Chemical Ge-ology , 192, 289–306. 99, 109, 122

Hermann, J. (2002b). Experimental constraints onphase relations in subducted continental crust. Con-tributions to Mineralogy and Petrology , 143, 219–235.50

Hermann, J. & Rubatto, D. (2003). Relating zirconand monazite domains to garnet growth zones: age andduration of granulite facies metamorphism in the ValMalenco lower crust. Journal of Metamorphic Geology ,21, 833–852. 99, 102, 111

Hermann, J., Muntener, O. & Scambelluri, M.(2000). The importance of serpentinite mylonites forsubduction and exhumation of oceanic crust. Tectono-physics, 327, 225–238. 194

Hermann, J., Rubatto, D., Korsakov, A. &Shatsky, V. (2001). Multiple zircon growth duringfast exhumation of diamondiferous, deeply subductedcontinental crust (Kokchetav Massif, Kazakhstan).Contributions to Mineralogy and Petrology, 141, 66–82. 189

Hermann, J., Rubatto, D. & Trommsdorff, V.(2006). Sub-solidus Oligocene zircon formation in gar-net peridotite during fast decompression and fluid infil-tration (Duria, Central Alps). Mineralogy and Petrol-ogy , 88, 181–206. 6, 196, 198

Hess, P. (1969). The metamorphic paragenesis ofcordierite in pelitic rocks. Contributions to Mineralogyand Petrology , 24, 191–207. 35

Hickmott, D., Shimizu, N., Spear, F. & Selver-stone, J. (1987). Trace-element zoning in a metamor-phic garnet. Geology , 15, 573. 123

Hock, V. (1974). Coexisting phengite, paragonite andmargarite in metasediments of the Mittlere HoheTauern, Austria. Contributions to Mineralogy andPetrology , 43, 261–273. 44

Hock, V. & Miller, C. (1980). Chemistry of mesozoicmetabasites in the middle and eastern part of the HoheTauern. Mitt. Osterr. Geol. Ges, 71–72, 81–88. 17, 62,106

Hoernes, H., S. Friedrichsen (1974). Oxygen iso-tope studies on metamorphic rocks of the westernhohe tauern. Schweizerische Mineralogische und Pet-rographische Mitteilungen, 54, 769. 38, 80, 82, 93, 94

Hoke, L. (1990). The Altkristallin of the KreuzeckMountains, SE Tauern Window, Eastern Alps - Base-ment Crust in a Convergent Plate Boundary Zone. Jb.Geol. B.-A, 133, 5–87. 28, 30

Holland, T. (1977). Structural and metamorphic studiesof eclogites and associated rocks in the Central TauernRegion of the Austrian Alps. Ph.D. thesis, Universityof Oxford. 10, 11, 19, 38, 39, 63, 83, 122

Holland, T. (1979). High water activities in the gener-ation of high pressure kyanite eclogites of the TauernWindow, Austria. The Journal of Geology , 87, 1–27.10, 38, 39, 47, 50, 100

Holland, T. & Blundy, J. (1994). Non-ideal inter-actions in calcic amphiboles and their bearing onamphibole-plagioclase thermometry. Contributions toMineralogy and Petrology , 116, 433–447. 65

Holland, T. & Norris, R. (1979). Deformed pillowlavas from the central Hohe Tauern, Austria, and theirbearing on the origin of epidote-banded greenstones.Earth and Planetary Science Letters, 43, 397–405. 23,26, 79, 140

Holland, T. & Powell, R. (1985). An internally con-sistent thermodynamic dataset with uncertainties andcorrelations: 2. Data and results. Journal of metamor-phic Geology , 3, 343–370. 33, 62

Holland, T. & Powell, R. (1990). An enlargedand updated internally consistent thermodynamicdataset with uncertainties and correlations: the sys-tem K2O-Na2O-CaO-MgO-MnO-FeO-Fe2O3-Al2O3-TiO2-SiO2-CH 2-O2. Journal of Metamorphic Geol-ogy , 8, 89–124. 34

Holland, T. & Powell, R. (1998). An internally con-sistent thermodynamic data set for phases of petro-logical interest. Journal of Metamorphic Geology , 16,309–344. 33, 34, 43, 45, 46, 66, 67, 68, 72, 87, 126, 174

Holland, T. & Powell, R. (2003). Activity-composition relations for phases in petrological calcu-lations: an asymmetric multicomponent formulation.Contributions to Mineralogy and Petrology, 145, 492–501. 47, 67

Holland, T. & Ray, N. (1985). Glaucophane and py-roxene breakdown reactions in the Pennine units ofthe Eastern Alps. Journal of Metamorphic Geology , 3,417–438. 10, 24, 62, 63, 64, 65, 66, 68, 69, 75, 78, 79,130

Holland, T. & Richardson, S. (1979). Amphibolezonation in metabasites as a guide to the evolution ofmetamorphic conditions. Contributions to Mineralogyand Petrology , 70, 143–148. 38, 63, 81, 95, 97, 176

241

Holland, T., Baker, J. & Powell, R. (1998). Mix-ing properties and activity-composition relationshipsof chlorites in the system MgO-FeO-Al

2

O3

-SiO2

-H2

O.European Journal of Mineralogy , 10, 395–406. 44, 67

Holland, T.J.B. (1980). The reaction albite = jadeite+ quartz determined experimentally in the range 600-1200 C. American Mineralogist , 129–134. 62, 90, 94

Hollister, L. (1966). Garnet zoning: an interpretationbased on the Rayleigh fractionation model. Science,154, 1647. 125, 127

Holness, M. (1993). Temperature and pressure depen-dence of quartz-aqueous fluid dihedral angles: the con-trol of adsorbed H2O on the permeability of quartzites.Earth and Planetary Science Letters, 117, 363–377.162

Holness, M. & Graham, C. (1995). P-T-X e↵ects onequilibrium carbonate H2O-CO2-NaCl dihedral an-gles: constraints on carbonate permeability and therole of deformation during fluid infiltration. Contribu-tions to Mineralogy and Petrology , 119, 303–313. 162

Horstwood, M., Foster, G., Parrish, R., Noble,S. & Nowell, G. (2003). Common-Pb corrected insitu U–Pb accessory mineral geochronology by LA-MC-ICP-MS. Journal of Analytical Atomic Spectrom-etry , 18, 837–846. 112, 113

Hoschek, G. (2001). Thermobarometry of metasedi-ments and metabasites from the Eclogite zone of theHohe Tauern, Eastern Alps, Austria. Lithos, 59, 127–150. 38, 39, 97, 174

Hoschek, G. (2004). Comparison of calculated PT pseu-dosections for a kyanite eclogite from the Tauern Win-dow, Eastern Alps, Austria. European Journal of Min-eralogy , 16, 59. 38

Hoschek, G. (2007). Metamorphic peak conditions ofeclogites in the Tauern Window, Eastern Alps, Aus-tria: Thermobarometry of the assemblage garnet+omphacite+ phengite+ kyanite+ quartz. Lithos, 93,1–16. 38, 50

Hoschek, G., Konzett, J. & Tessadri, R.(2010). Phase equilibria in quartzitic garnet–kyanite–chloritoid micaschist from the Eclogite Zone, TauernWindow, Eastern Alps. European Journal of Mineral-ogy , 22, 721–732. 49, 107, 109, 122

Hsu, K. (1969). Role of cohesive strength in the mechanicof over-thrust faulting and of landsliding. GeologicalSociety of America Bulletin, 80, 927. 171, 173

Hubbard, M. & Harrison, T. (1989). 40Ar/39Ar ageconstraints on deformation and metamorphism in theMain Central Thrust zone and Tibetan Slab, easternNepal Himalaya. Tectonics, 8, 865–880. 162

Huerta, A., Royden, L. & Hodges, K. (1999). Thee↵ects of accretion, erosion and radiogenic heat on themetamorphic evolution of collisional orogens. Journalof Metamorphic Geology , 17, 349–366. 164

Hurford, A., Hunziker, J. & Stockhert, B. (1991).Constraints on the late thermotectonic evolution of thewestern Alps: evidence for episodic rapid uplift. Tec-tonics, 10, 758–769. 199, 201

Huw Davies, J. & von Blanckenburg, F. (1995). Slabbreako↵: a model of lithosphere detachment and itstest in the magmatism and deformation of collisionalorogens. Earth and Planetary Science Letters, 129,85–102. 197, 200

Inger, S. & Cliff, R. (1997). Timing of metamorphismin the Tauern Window, Eastern Alps: Rb-Sr ages andfabric formation. Journal of Metamorphic Geology , 12,695–707. 10, 11, 96, 101, 102, 128, 129, 147, 152, 153,174, 184, 201

Jaffey, A., Flynn, K., Glendenin, L., Bentley, W.& Essling, A. (1971). Precision Measurement of Half-Lives and Specific Activities ofˆ{235} U andˆ{238} U.Physical Review C , 4, 1889–1906. 113

Jager, E., Niggli, E. & Wenk, E. (1967). Rb-SrAltersbestimmungen an Glimmern der Zenteralpen.Beitrage zur Geologischen Karte der Schweiz , 134. 103

Jaillard, E., Soler, P., Carlier, G. & Mourier,T. (1990). Geodynamic evolution of the northern andcentral Andes during early to middle Mesozoic times:a Tethyan model. Journal of the Geological Society ,147, 1009–1022. 51

Jamieson, R., Beaumont, C., Fullsack, P. & Lee, B.(1998). Barrovian metamorphism: Where’s the heat?Geological Society London Special Publications, 138,23–51. 164, 197

Janak, M., Cornell, D., Froitzheim, N., De Hoog,J., Broska, I., Vrabec, M. & Hurai, V. (2009).Eclogite-hosting metapelites from the Pohorje Moun-tains (Eastern Alps): PT evolution, zircon geochronol-ogy and tectonic implications. European Journal ofMineralogy , 21, 1191. 131

Janots, E., Brunet, F., Goffe, B., Poinssot, C.,Burchard, M. & Cemic, L. (2007). Thermochem-istry of monazite-(La) and dissakisite-(La): implica-tions for monazite and allanite stability in metapelites.Contributions to Mineralogy and Petrology, 154, 1–14.41, 99, 109, 111, 123

Janots, E., Engi, M., Berger, A., Allaz, J.,Schwarz, J. & Spandler, C. (2008). Prograde meta-morphic sequence of REE minerals in pelitic rocks of

242

the Central Alps: implications for allanite–monazite–xenotime phase relations from 250 to 610 C. Journal ofMetamorphic Geology , 26, 509–526. 41, 99, 107, 109,111, 123, 128, 129

Janots, E., Engi, M., Rubatto, D., Berger, A.,Gregory, C. & Rahn, M. (2009). Metamorphic ratesin collisional orogeny from in situ allanite and mon-azite dating. Geology , 37, 11. 5, 41, 123, 128, 196,198, 201

Janots, E., Berger, A. & Engi, M. (2010). Physico-chemical control on the REE-minerals in chloritoid-grade metasediments from a single outcrop (CentralAlps, Switzerland). Lithos. 128

Jenkin, G., Ellam, R. & Stuart, F. (2001). An in-vestigation of closure temperature of the biotite Rb-Srsystem: The importance of cation exchange. Geochim-ica et Cosmochimica Acta, 65, 1141–1160. 103

Jin, Z., Zhang, J., Green, H. & Jin, S. (2001). Eclog-ite rheology: Implications for subducted lithosphere.Geology , 29, 667. 173

Johannes, W. & Holtz, F. (1996). Petrogenesis andexperimental petrology of granitic rocks, vol. 335.Springer Berlin. 90, 91

Johnson, C. & Carlson, W. (1990). The origin ofolivine-plagioclase coronas in metagabbros from theAdirondack Mountains, New York. Journal of Meta-morphic Geology , 8, 697–717. 34, 69

Jolivet, L., Faccenna, C., Goffe, B., Burov, E.& Agard, P. (2003). Subduction tectonics and ex-humation of high-pressure metamorphic rocks in theMediterranean orogens. American Journal of Science,303, 353. 200

Jolivet, L., Raimbourg, H., Labrousse, L., Avigad,D., Leroy, Y., Austrheim, H. & Andersen, T.(2005). Softening trigerred by eclogitization, the firststep toward exhumation during continental subduc-tion. Earth and Planetary Science Letters, 237, 532–547. 60

Kastrup, U., Zoback, M., Deichmann, N., Evans,K., Giardini, D. & Michael, A. (2004). Stress fieldvariations in the Swiss Alps and the northern Alpineforeland derived from inversion of fault plane solutions.J. geophys. Res. 3

Kelley, S. (2002). Excess argon in K-Ar and Ar-Argeochronology. Chemical Geology , 188, 1–22. 133, 154,155, 156, 157

Kelley, S., Turner, G., Butterfield, A. & Shep-herd, T. (1986). The source and significance of argon

isotopes in fluid inclusions from areas of mineraliza-tion. Earth and Planetary Science Letters, 79, 303–318. 157

Khvostova, V. (1963). On the isomorphism of epidoteand orthite. Doklady Academy of Sciences U.S.S.R,Earth Sciences Section, 141, 1307–1309. 109

Kiessling, W. (1992). Palaeontological and facial fea-tures of the Upper Jurassic Hochstegen marble(Tauern window, Eastern Alps). Terra Nova, 4, 184–197. 14

Kirby, S. & Kronenberg, A. (1984). Deformationof clinopyroxenite: Evidence for a transition in flowmechanisms and semibrittle behavior. Journal of Geo-physical Research, 89, 3177–3192. 173

Konopasek, J. (2001). Eclogitic micaschists in the cen-tral part of the Krusne hory Mountains (BohemianMassif). European Journal of Mineralogy , 13, 87–100.51, 58

Konrad-Schmolke, M., Zack, T., O’Brien, P. & Ja-cob, D. (2008). Combined thermodynamic and rareearth element modelling of garnet growth during sub-duction: Examples from ultrahigh-pressure eclogite ofthe Western Gneiss Region, Norway. Earth and Plan-etary Science Letters, 272, 488–498. 99

Kooijman, E., Mezger, K. & Berndt, J. (2010). Con-straints on the U-Pb systematics of metamorphic rutilefrom in situ LA-ICP-MS analysis. Earth and PlanetaryScience Letters, 293, 321–330. 103, 107

Koons, P., Rubie, D. & FRUCH-GREEN, G. (1987).The e↵ects of disequilibrium and deformation on themineralogical evolution of quartz diorite during meta-morphism in the eclogite facies. Journal of Petrology ,28, 679. 92

Kosler, J., Tubrett, M. & Sylvester, P. (2001).Application of Laser Ablation ICP-MS to U-Th-PbDating of Monazite. Geostandards and GeoanalyticalResearch, 25, 375–386. 113

Kreemer, C. & Holt, W. (2001). A no-net-rotationmodel of present-day surface motions. Geophysical re-search letters, 28, 4407–4410. 3

Kruhl, J. (1993). The P-T-d development at thebasement-cover boundary in the north-eastern TauernWindow (Eastern Alps): Alpine continental collision.Journal of Metamorphic Geology , 11, 31–47. 83

Kuhlemann, J. (2000a). Post-collisional sediment bud-get of circum-Alpine basins (Central Europe). Mem.Sci. Geol. (Padova), 52, 1–91. 181

243

Kuhlemann, J. (2000b). Post-collisional sediment bud-get of circum-Alpine basins (Central Europe). Mem.Sci. Geol.(Padova), 52, 1–91. 184, 199

Kuhlemann, J., Frisch, W., Szekely, B., Dunkl, I.& Kazmer, M. (2002). Post-collisional sediment bud-get history of the Alps: tectonic versus climatic con-trol. International Journal of Earth Sciences, 91, 818–837. 184, 198

Kuhlemann, J., Dunkl, I., Brugel, A., Spiegel, C.& Frisch, W. (2006). From source terrains of theEastern Alps to the Molasse Basin: Detrital recordof non-steady-state exhumation. Tectonophysics, 413,301–316. 184, 198

Kummerow, J., Kind, R., Oncken, O., Giese,P., Ryberg, T., Wylegalla, K., Scherbaum,F. & Group, T.W. (2004). A natural and con-trolled source seismic profile through the Eastern Alps:TRANSALP. Earth and Planetary Science Letters,115–129. 4

Kurz, W. (2005). Constriction during exhumation: Ev-idence from eclogite microstructures. Geology , 33, 37.30

Kurz, W. & Froitzheim, N. (2002). The exhumation ofeclogite-facies metamorphic rocks—a review of modelsconfronted with examples from the Alps. InternationalGeology Review , 44, 702–743. 189, 190, 194

Kurz, W., Neubauer, F. & Genser, J. (1996).Kinematics of Penninic nappes (Glockner Nappe andbasement-cover nappes) in the Tauern Window (East-ern Alps, Austria) during subduction and Penninic-Austroalpine collision. Eclogae Geologicae Helvetiae,89, 573–606. 184, 190

Kurz, W., Neubauer, F. & Dachs, E. (1998a). Eclog-ite meso-and microfabrics: implications for the burialand exhumation history of eclogites in the Tauern Win-dow (Eastern Alps) from PTd paths. Tectonophysics,285, 183–209. 19

Kurz, W., Neubauer, F., Genser, J. & Dachs, E.(1998b). Alpine geodynamic evolution of passive andactive continental margin sequences in the TauernWindow (eastern Alps, Austria, Italy): a review. Ge-ologische rundschau, 87, 225–242. 4, 9, 13, 14, 16, 23,27, 37, 83, 84, 94, 95, 130, 158, 170, 185

Kurz, W., Handler, R. & Bertoldi, C. (2008). Trac-ing the exhumation of the Eclogite Zone (Tauern Win-dow, Eastern Alps) by 40 Ar/39 Ar dating of whitemica in eclogites. Swiss Journal of Geosciences, 101,191–206. 101, 134, 135, 140, 160, 162

Kylander-Clark, A., Hacker, B. & Mattinson, J.(2008). Slow exhumation of UHP terranes: Titaniteand rutile ages of the Western Gneiss Region, Norway.Earth and Planetary Science Letters, 272, 531–540.189

Lachenbruch, A. (1968). Preliminary geothermal modelof the Sierra Nevada. Journal of Geophysical Research,73, 6977–6989. 169, 170

Lachenbruch, A. (1970). Crustal temperature and heatproduction: implication of the linear heat flow relation.Journal of Geophysical Research, 75, 3291–3300. 169,170

Lammerer, B. & Weger, M. (1998). Footwall uplift inan orogenic wedge: the Tauern Window in the EasternAlps of Europe. Tectonophysics, 285, 213–230. 10

Lancaster, P., Baxter, E., Ague, J., Breeding, C.& Owens, T. (2008). Synchronous peak Barrovianmetamorphism driven by syn-orogenic magmatism andfluid flow in southern Connecticut, USA. Journal ofMetamorphic Geology , 26, 527–538. 195

Lanzirotti, A. (1995). Yttrium zoning in metamorphicgarnets. Geochimica et Cosmochimica Acta, 59, 4105–4110. 123

Lapen, T., Johnson, C., Baumgartner, L., Mahlen,N., Beard, B. & Amato, J. (2003). Burial rates dur-ing prograde metamorphism of an ultra-high-pressureterrane: an example from Lago di Cignana, westernAlps, Italy. Earth and Planetary Science Letters, 215,57–72. 129, 197, 201

Laubscher, H. (1991). The arc of the Western Alps to-day. Eclogae Geologicae Helvetiae, 84, 359–631. 2

Lee, C. (2003). Compositional variation of density andseismic velocities in natural peridotites at STP con-ditions: Implications for seismic imaging of composi-tional heterogeneities in the upper mantle. J. Geophys.Res, 108, 2441. 190

Lee, J., Williams, I. & Ellis, D. (1997). Pb, U andTh di↵usion in natural zircon. Nature, 390, 159–162.103

Leech, M., Singh, S., Jain, A., Klemperer, S. &Manickavasagam, R. (2005). The onset of India–Asia continental collision: Early, steep subduction re-quired by the timing of UHP metamorphism in thewestern Himalaya. Earth and Planetary Science Let-ters, 234, 83–97. 1

Li, S., Wang, S., Chen, Y., Liu, D., Qiu, J., Zhou, H.& Zhang, Z. (1994). Excess argon in phengite fromeclogite: Evidence from dating of eclogite minerals bySm–Nd, Rb–Sr and 40Ar/39Ar methods* 1,* 2. Chem-ical Geology , 112, 343–350. 162

244

Li, S., Jagoutz, E., Lo, C., Chen, Y., Li, Q. &Xiao, Y. (1999). Sm/Nd, Rb/Sr, and 40 Ar/39 Ar iso-topic systematics of the ultrahigh-pressure metamor-phic rocks in the Dabie-Sulu belt, Central China: aretrospective view. International Geology Review , 41,1114–1124. 162

Liati, A. & Froitzheim, N. (2006). Assessing theValais ocean, Western Alps: U-Pb SHRIMP zircongeochronology of eclogite in the Balma unit, on topof the Monte Rosa nappe. European journal of miner-alogy , 18, 299. 201

Liati, A. & Gebauer, D. (2003). Geochronologicalconstraints for the time of metamorphism in theGruf Complex (Central Alps) and implications for theAdula-Cima Lunga nappe system. Swiss Bulletin ofMineralogy and Petrology , 83, 159–172. 198

Liati, A., Froitzheim, N. & Fanning, C. (2005).Jurassic ophiolites within the Valais domain of theWestern and Central Alps: geochronological evidencefor re-rifting of oceanic crust. Contributions to Miner-alogy and Petrology , 149, 446–461. 201

Lide, D. & Frederikse, H. (1995). CRC Handbook ofChemistry and Physics: A Ready-reference Book ofChemical and Physical Data, 1995-1996 . CRC Press.104

Liou, J. & Maruyama, S. (1987). Paragenesesand compositions of amphiboles from Franciscanjadeite–glaucophane type facies series metabasites atCazadero, California. Journal of Metamorphic Geol-ogy , 5, 371–395. 66

Liu, Y., Genser, J., Handler, R., Friedl, G. &Neubauer, F. (2001). 40Ar/39Ar muscovite ages fromthe Penninic-Austroalpine plate boundary, EasternAlps. Tectonics, 20, 526–547. 174, 183

Luais, B., Duchene, S. & de Sigoyer, J. (2001). Sm-Nd disequilibrium in high-pressure, low temperatureHimalayan and Alpine rocks. Tectonophysics, 342, 1–22. 102

Ludwig, K. (2003). Isoplot/Ex Version 3.00: aGeochronological Toolkit for Microsoft Excel. Berke-ley Geochronological Center, Berkeley, CA. 113, 114

Lustrino, M., Morra, V., Fedele, L. & Franciosi,L. (2009). Beginning of the Apennine subduction sys-tem in central western Mediterranean: Constraintsfrom Cenozoic orogenic magmatic activity of Sardinia,Italy. Tectonics, 28, 23. 200

Lysak, S. (1992). Heat flow variations in continentalrifts. Tectonophysics, 208, 309–323. 170

Lyubetskaya, T. & Ague, J. (2010). Modeling meta-morphism in collisional orogens intruded by magmas:I. Thermal evolution. American Journal of Science,310, 427. 164

Mahar, E., Baker, J., Powell, R., Holland, T. &Howell, N. (1997a). The e↵ect of Mn on mineral sta-bility in metapelites. Journal of Metamorphic Geology ,15, 223–238. 35

Mahar, E., Baker, J., Powell, R., Holland, T. &Howell, N. (1997b). The e↵ect of Mn on mineral sta-bility in metapelites. Journal of Metamorphic Geology ,15, 223–238. 44, 54, 56, 57, 58

Marquer, D., Challandes, N. & Baudin, T. (1996).Shear zone patterns and strain distribution at the scaleof a Penninic nappe: the Suretta nappe (Eastern SwissAlps). Journal of Structural Geology , 18, 753–764. 199

Marty, B., O’Nions, R., Oxburgh, E., Martel, D.& Lombardi, S. (1992). Helium isotopes in Alpineregions. Tectonophysics, 206, 71–78. 171, 175, 200

Maruyama, S. & Liou, J. (1988). Petrology of Fran-ciscan metabasites along the jadeite-glaucophane typefacies series, Cazadero, California. Journal of Petrol-ogy , 29, 1. 66

Massonne, H. (2000). Experimental aspects of UHPmetamorphism in pelitic systems. Ultrahigh-pressureMetamorphism and Geodynamics in Collision-typeOrogenic Belts: Final Report of the Task Group III-6(1994-1998) of the International Lithosphere Project ,105. 50

Massonne, H. & Schreyer, W. (1987). Phengite geo-barometry based on the limiting assemblage withK-feldspar, phlogopite, and quartz. Contributions toMineralogy and Petrology , 96, 212–224. 92

Massonne, H. & Schreyer, W. (1989). Stability field ofthe high-pressure assemblage, talc+ phengite and twonew phengite barometers. European Journal of Miner-alogy , 1, 391–410. 52

McDonough, W. & Sun, S. (1995). The composition ofthe Earth. Chemical Geology , 120, 223–253. 68, 124

McDougal, I. & Harrison, T. (1988). Geochronologyand Thermochronology by the 40Ar/39Ar method. Ox-ford Monographs on Geology and Geophysics. 103, 140,143, 153

McKenzie, D. (1984). The Generation and Compactionof Partially Molten Rock. Journal of Petrology , 25,713–765. 162

245

Meffan-Main, S., Cliff, R., Barnicoat, A., Lom-bardo, B. & Compagnoni, R. (2004). A Tertiaryage for Alpine high-pressure metamorphism in theGran Paradiso massif, Western Alps: a Rb–Sr mi-crosampling study. Journal of Metamorphic Geology ,22, 267–281. 1, 131, 198, 201

Megrue, G. (1971). Distribution and Origin of Helium,Neon, and Argon Isotopes in Apollo 12 Samples Mea-sured by In Situ Analysis with a Laser-Probe MassSpectrometer. Journal of Geophysical Research, 76,4956–4968. 142

Menard, G. (1988). Structure et cinematique d’unechaıne de collision: les Alpes occidentales et centrales.Ph.D. thesis, Univ. Joseph Fourier, Grenoble. 3

Meyre, C., De Capitani, C., Zack, T. & Frey, M.(1999). Petrology of high-pressure metapelites fromthe Adula nappe (Central Alps, Switzerland). Journalof Petrology , 40, 199–213. 5, 53

Michard, A., Chopin, C. & Henry, C. (1993). Com-pression versus extension in the exhumation of theDora-Maira coesite-bearing unit, Western Alps, Italy.Tectonophysics, 221, 173–193. 198

Miller, C. (1974). On the metamorphism of the eclog-ites and high-grade blueschists from the Penninic ter-rane of the Tauern Window, Austria. Schweiz. Min.Petr. Mitt , 54, 371–384. 9, 38, 100

Miller, C. (1977). Chemismus und phasenpetrologischeUntersuchungen der Gesteine aus der Eklogitzone desTauernfensters, Osterreich. Mineralogy and Petrology ,24, 221–277. 38, 62

Miller, C., Mundil, R., Th”oni, M. & Konzett, J. (2005). Refining the timingof eclogite metamorphism: a geochemical, petrolog-ical, Sm-Nd and U-Pb case study from the PohorjeMountains, Slovenia (Eastern Alps). Contributions toMineralogy and Petrology , 150, 70–84. 131

Miller, C., Konzett, J., Tiepolo, M., Armstrong,R. & Thoni, M. (2007). Jadeite-gneiss from the Eclog-ite Zone, Tauern Window, Eastern Alps, Austria:Metamorphic, geochemical and zircon record of a sed-imentary protolith. Lithos, 93, 68–88. 101, 106

Muller, W., Mancktelow, N. & Meier, M. (2000).Rb-Sr microchrons of synkinematic mica in mylonites:an example from the DAV fault of the Eastern Alps.Earth and Planetary Science Letters, 180, 385–397.172

Muntener, O., Hermann, J. & Trommsdorff, V.(2000). Cooling history and exhumation of lower

crustal granulites and upper mantle (Malenco, East-ern Central Alps). Journal of Petrology , 41, 175–200.5

Nabelek, P., Whittington, A. & Hofmeister, A.(2010). Strain heating as a mechanism for partialmelting and ultrahigh temperature metamorphismin convergent orogens: Implications of temperature-dependent thermal di↵usivity and rheology. Journal ofGeophysical Research, 115, 1–17. 171, 180, 181, 182

Naeser, C. (1979). Fission track dating and geologicalannealing of fission tracks. In E. Jager & J. Hunziker,eds., Lectures in Isotope Geology , 154–169, Springer-Verlag. 103

Nagel, T., Capitani, C., Frey, M., Froitzheim, N.,Stcnitz, H. & Schmid, S. (2002). Structural andmetamorphic evolution during rapid exhumation in theLepontine dome (southern Simano and Adula nappes,Central Alps, Switzerland). Eclogae Geologicae Helve-tiae, 95, 301–322. 199

Negulescu, E., Sabau, G. & Massonne, H. (2009).Chloritoid-Bearing Mineral Assemblages in High-Pressure Metapelites from the Bughea Complex,Leaota Massif (South Carpathians). Journal of Petrol-ogy , 50, 103. 49, 50, 51, 57, 58

Neubauer, F. (1995). Geological evolution of the inter-nal Alps, Carpathians and of the Pannonian basin: anintroduction. Tectonophysics, 242, 1–4. 10

Neubauer, F., Genser, J., Kurz, W. & Wang, X.(1999). Exhumation of the Tauern window, EasternAlps. Physics and Chemistry of the Earth, Part A:Solid Earth and Geodesy , 24, 675–680. 10

Neubauer, F., Genser, J. & Handler, R. (2000). TheEastern Alps: result of a two-stage collision process.Mitt.”osterr. geol. Ges, 92, 117–134. 131

Neufeld, K., Ring, U., Heidelbach, F., Dietrich, S.& Neuser, R. (2008). Omphacite textures in eclogitesof the Tauern Window: Implications for the exhuma-tion of the Eclogite Zone, Eastern Alps. Journal ofStructural Geology , 30, 976–992. 190

Nicolas, A., Polino, R., Hirn, A. & Nicolich,R. (1990). ECORS-CROP Working Group. ECORS-CROP traverse and deep structure of the western Alps:a synthesis. Mem Soc. geol. France, NS , 156, 15–27.201

Nier, A. (1950). A redetermination of the relative abun-dances of the isotopes of carbon, nitrogen, oxygen, ar-gon, and potassium. Physical Review , 77, 789. 134

246

Nockolds, S. (1954). Average chemical compositions ofsome igneous rocks. Geological Society of America Bul-letin, 65, 1007. 90

Norton, D. & Knapp, R. (1977). Transport phenomenain hydrothermal systems: the nature of porosity. Am.J. Sci.;(United States), 277. 158, 159

Oberhauser, R. (1968). Beitrage zur Kenntnis derTektonik und der Palaogeographie wahrend derOberkreide und dem Palaogen im Ostalpenraum. Jb.geol. BA, 101, 115–145. 100, 166, 174

Oberli, F., Meier, M., Berger, A., Rosenberg, C.et al. (2004). U-Th-Pb and 230Th/238U disequilib-rium isotope systematics: Precise accessory mineralchronology and melt evolution tracing in the AlpineBergell intrusion. Geochimica et Cosmochimica Acta,68, 2543–2560. 109, 131, 199, 201

Okay, A. (1980). Mineralogy, petrology, and phase re-lations of glaucophane-lawsonite zone blueschists fromthe Tavsanli Region, Northwest Turkey. Contributionsto Mineralogy and Petrology , 72, 243–255. 68

Okuyama-Kusunose, Y. (1985). Margarite-paragonite-muscovite assemblages from the low-grade metapelitesof the Tono metamorphic aureole, Kitakami Moun-tains, Northeast Japan. Ganseki Kobutsu KoshoGakkaishi , 80, 515–525. 44

Oliver, G., Chen, F., Buchwaldt, R. & Hegner,E. (2000). Fast tectonometamorphism and exhumationin the type area of the Barrovian and Buchan zones.Geology , 28, 459. 164, 195

Onstott, T., Phillips, D. & Pringle-Goodell, L.(1991). Laser microprobe measurement of chlorine andargon zonation in biotite. Chemical geology , 90, 145–168. 154, 156

Otamendi, J., De La Rosa, J., Douce, A. & Castro,A. (2002). Rayleigh fractionation of heavy rare earthsand yttrium during metamorphic garnet growth. Ge-ology , 30, 159. 123, 125, 126

Overstreet, W. (1967). The geologic occurrence ofMonazite. U.S. geological Survey Professional Paper .107

Oxburgh, E. (1968). An outline of the geology of theCentral Eastern Alps. Proceedings of the Geologists’Association, 79, 1–46. 3, 5, 8, 164, 168, 177, 180, 183,184

Oxburgh, E. & England, P. (1980). Heat flow and themetamorphic evolution of the Eastern Alps. EclogaeGeologicae Helvetiae, 73, 379–398. 164, 165

Oxburgh, E. & Turcotte, D. (1974). Thermal gradi-ents and regional metamorphism in overthrust terrainswith special reference to the Eastern Alps. Schweiz-erische Mineralogische und Petrographische Mitteilun-gen, 54, 642–662. 1, 6, 8, 100, 130, 164, 165, 166, 168,171, 174

Oxburgh, E., Lambert, S., Baadsgaard, H. & Si-mons, J. (1966). Potassium-argon studies across thesoutheast margin of the Tauern Window, the EasternAlps. Verhandlungen der Geologischen, Bundesanstalt ,1–2. 100

Parrish, R., Gough, S., Searle, M. & Waters,D. (2006). Plate velocity exhumation of ultrahigh-pressure eclogites in the Pakistan Himalaya. Geology ,34, 989. 1, 102, 128, 133, 189, 190

Paterson, M. & Luan, F. (1990). Quartzite rheologyunder geological conditions. Geological Society LondonSpecial Publications, 54, 299. 190

Pavoni, N. (1961). Faltung durch Horizontalver-schiebung. Eclogae Geol. Helv , 54, 515–534. 3

Pavoni, N. (1980). Crustal stresses inferred from fault-plane solutions of earthquakes and neotectonic defor-mation in Switzerland. Rock Mech., Suppl , 9, 63–68.3

Pavoni, N. (1986). Regularities in the pattern of majorfault zones of the earth and the origin of arcs. Originof Arcs, 63–78. 3

Peacock, S. (1992). Blueschist-facies metamorphism,shear heating, and PTt paths in subduction shearzones. Journal of Geophysical Research-Solid Earth,97, 17693–17707. 60

Peacock, S. (1996). Thermal and petrologic structureof subduction zones. Geophysical monograph, 96, 119–133. 60, 130, 161

Peacock, S. (2003). Thermal structure and meta-morphic evolution of subducting slabs. GEOPHYS-ICAL MONOGRAPH-AMERICAN GEOPHYSICALUNION , 138, 7–22. 130

Peccerillo, A., Poli, G., Sassi, F., Zipoli, G. &Mezzacasa, G. (1979). New data on the Upper Or-dovician acid plutonism in the Eastern Alps. N. Jb.Miner. Abh, 137, 162–183. 186

Penniston-Dorland, S., Sorensen, S., Ash, R. &Khadke, S. (2010). Lithium isotopes as a tracer offluids in a subduction zone melange: Franciscan Com-plex, CA. Earth and Planetary Science Letters, 292,181–190. 204

247

Peter Gromet, L. & Silver, L. (1983). Rare earthelement distributions among minerals in a granodior-ite and their petrogenetic implications. Geochimica etCosmochimica Acta, 47, 925–939. 109

Petra, V., Frank, S., Friedrich, F. & Axel, G.(2010). Magmato-sedimentary Carboniferous to Juras-sic evolution of the western Tauern window, EasternAlps (constraints from U-Pb zircon dating and geo-chemistry). International Journal of Earth Sciences,1–35. 8

Petrik, I., Broska, I., Lipka, J. & Siman, P. (1995).Granitoid Allanite-(Ce) Substituition Relations, Re-dox Conditions and REE Distributions (on an Exam-ple of I-Type Granitoids, Western Carpathians, Slo-vakia). Geologica Carpathica, 46, 79–94. 42

Philippot, P. & Rumble, D. (2000). Fluid-rock in-teractions during high-pressure and ultrahigh-pressuremetamorphism. International Geology Review , 42,312–327. 163

Phillips, G., Hand, M. & Offler, R. (2008). PTt de-formation framework of an accretionary prism, south-ern New England Orogen, eastern Australia: Impli-cations for blueschist exhumation and metamorphicswitching. Tectonics, 27, TC6017. 68

Pickles, C., Kelley, S., Reddy, S. & Wheeler,J. (1997). Determination of high spatial resolutionargon isotope variations in metamorphic biotites* 1.Geochimica et Cosmochimica Acta, 61, 3809–3833.133

Platt, J. (1993). Exhumation of high-pressure rocks: areview of concepts and processes. Terra Nova, 5, 119–133. 189

Ploshko, V. & Bogdanova, V. (1963). Isomorphoussubstitutions in minerals of the epidote group from thenorthern Caucasus. Geochemistry , 1, 61–71. 109

Pollington, A. & Baxter, E. (2010). High resolutionSm-Nd garnet geochronology reveals the uneven paceof tectonometamorphic processes. Earth and PlanetaryScience Letters, 293, 63–71. 204

Powell, R. (1985). Geothermometry and geobarometry:a discussion. Journal of the Geological Society , 142. 34

Powell, R. & Holland, T. (1985). An internally con-sistent thermodynamic dataset with uncertainties andcorrelations: 1. Methods and a worked example. Jour-nal of Metamorphic Geology , 3, 327–342. 33, 62

Powell, R. & Holland, T. (1988). An internally con-sistent thermodynamic dataset with uncertainties andcorrelations: 3. Applications to geobarometry, workedexamples and a computer program. Journal of Meta-morphic Geology , 6, 73–204. 34, 35, 66

Powell, R. & Holland, T. (1990a). Calculated mineralequilibria in the pelite system, KFMASH (K2O-FeO-MgO-Al2O3-SiO2-H2O). American Mineralogist , 75,367–380. 35

Powell, R. & Holland, T. (1994). Optimal geother-mometry and geobarometry. American Mineralogist ,79, 120. 34, 37

Powell, R. & Holland, T. (2008). On thermobarom-etry. Journal of Metamorphic Geology , 26, 155–179.36, 47

Powell, R. & Holland, T. (2010). Using EquilibriumThermodynamics to Understand Metamorphism andMetamorphic Rocks. Elements, 6, 309–314. 35

Powell, R. & Holland, T.J.B. (1990b). Calculatedmineral equilibria in the pelite system, KFMASH(K

2

O-FeO-MgO-Al2

O3

-SiO2

-H2

O). American Miner-alogist , 75, 367–380. 50

Powell, R., Holland, T. & Worley, B. (1998). Cal-culating phase diagrams involving solid solutions vianon-linear equations, with examples using THERMO-CALC. Journal of Metamorphic Geology , 16, 577–588.35

Proyer, A. (2003). The preservation of high-pressure rocks during exhumation: metagranitesand metapelites. Lithos, 70, 183–194. 89, 92

Purdy, J. & Jager, E. (1976). K-Ar ages on rock-forming minerals from the Central Alps. Memoirs-Institute of Geology and Mineralogy, University ofPadova, 30, 1–31. 103

Putlitz, B., Cosca, M. & Schumacher, J. (2005).Prograde mica 40Ar/39Ar growth ages recorded inhigh pressure rocks (Syros, Cyclades, Greece). Chem-ical geology , 214, 79–98. 142

Pyle, J. & Spear, F. (1999). Yttrium zoning in garnet:coupling of major and accessory phases during meta-morphic reactions. Geological Materials Research, 1,1–49. 99, 125, 126

Radulescu, I., Rubatto, D., Gregory, C. & Com-pagnoni, R. (2009). The age of HP metamorphism inthe Gran Paradiso Massif, Western Alps: A petrologi-cal and geochronological study of ”silvery micaschists”.Lithos, 110, 95–108. 115, 198

Raith, M., Hoermann, P. & Abraham, K. (1977).Petrology and metamorphic evolution of the Pen-ninic ophiolites in the Western Tauern Window (Aus-tria). Schweizerische Mineralogische und Petrographis-che Mitteilungen, 57, 187–232. 38

248

Raith, M., Mehrens, C. & Thole, W. (1980).Gliederung, tektonischer Bau und metamorphe En-twicklung der penninischen Serien im sudlichenVenediger-Gebiet, Osttirol. Jb. Geol. B.-A, 123, 1–37.19

Ratschbacher, L. (1986). Kinematics of Austro-Alpinecover nappes: changing translation path due to trans-pression. Tectonophysics, 125, 335–356. 10

Ratschbacher, L., Frisch, W., Linzer, H. & Merle,O. (1991a). Lateral extrusion in the Eastern Alps, part2: structural analysis. Tectonics, 10, 257–271. 10, 166

Ratschbacher, L., Merle, O., Davy, P. & Cobbold,P. (1991b). Lateral extrusion in the Eastern Alps, part1: boundary conditions and experiments scaled forgravity. Tectonics, 10, 245–256. 10, 166

Ratschbacher, L., Dingeldey, C., Miller, C.,Hacker, B. & McWilliams, M. (2004). Formation,subduction, and exhumation of Penninic oceanic crustin the Eastern Alps: time constraints from 40Ar/39Argeochronology. Tectonophysics, 394, 155–170. 101,134, 135, 140, 160, 162, 201

Ray, N. (1986). Epidote Group Mineralogy in the EasternAlps. Ph.D. thesis, University of Cambridge. 64, 65,66, 67, 69, 70, 72, 140, 208

Rebay, G., Powell, R. & Diener, J. (2010). Calcu-lated phase equilibria for a MORB composition in aP–T range, 450–650 C and 18–28 kbar: the stabilityof eclogite. Journal of Metamorphic Geology , 28, 635–645. 68, 69, 70, 72, 160

Reddy, S., Kelley, S. & Wheeler, J. (1996). A40Ar/39Ar laser probe study of micas from the SesiaZone, Italian Alps: implications for metamorphic anddeformation histories. Journal of Metamorphic Geol-ogy , 14, 493–508. 133, 201

Reddy, S., Wheeler, J., Butler, R., Cliff, R.,Freeman, S., Inger, S., Pickles, C. & Kelley, S.(2003). Kinematic reworking and exhumation withinthe convergent Alpine Orogen. Tectonophysics, 365,77–102. 201

Reed, S. & Buckley, A. (1998). Rare-earth element de-termination in minerals by electron-probe microanal-ysis; application of spectrum synthesis. MineralogicalMagazine, 62, 1. 112

Renne, P., Swisher, C., Deino, A., Karner, D.,Owens, T. & DePaolo, D. (1998). Intercalibra-tion of standards, absolute ages and uncertainties in40Ar/39Ar dating. Chemical Geology , 145, 117–152.142

Richardson, S. & England, P. (1979). Metamorphicconsequences of crustal eclogite production in over-thrust orogenic zones. Earth and Planetary ScienceLetters, 42, 183–190. 164

Ring, U. & Glodny, J. (2010). No need for lithosphericextension for exhuming (U)HP rocks by normal fault-ing. Journal of the Geological Society , 167. 185, 190

Ring, U., Ratschbacher, L. & Frisch, W. (1988).Plate-boundary kinematics in the Alps: Motion in theArosa suture zone. Geology , 16, 696. 10

Robin, P. & Cruden, A. (1994). Strain and vorticitypatterns in ideally ductile transpression zones. Journalof Structural Geology , 16, 447–466. 190

Roddick, J., Cliff, R. & Rex, D. (1980). The evo-lution of excess argon in alpine biotites–A40Ar-39Aranalysis. Earth and Planetary Science Letters, 48,185–208. 143, 146, 154

Roeder, D. & Bachmann, G. (1996). Evolution, struc-ture and petroleum geology of the German MolasseBasin. Memoires du Museum national d’histoire na-turelle, 170, 263–284. 5

Romer, R. & Siegesmund, S. (2003). Why allanite mayswindle about its true age. Contributions to Mineralogyand Petrology , 146, 297–307. 115, 131, 172, 199, 201

Roselieb, K., Blanc, P., B”uttner, H., Jambon, A., Rammensee, W.,Rosenhauer, M., Vielzeuf, D. & Walter, H.(1997). Experimental study of argon sorption inquartz: Evidence for argon incompatibility. Geochim-ica et cosmochimica acta, 61, 533–542. 154

Roselieb, K., Wartho, J., Buttner, H., Jambon,A. & Kelley, S. (1999). Solubility and di↵usivity ofnoble gases in synthetic phlogopite: a UV LAMP in-vestigation. Terra Abstr., 4. 155

Rosenbaum, G., Lister, G. & Duboz, C. (2002). Rel-ative motions of Africa, Iberia and Europe duringAlpine orogeny. Tectonophysics, 359, 117–129. 184,185, 198

Rosenberg, C., Brun, J. & Gapais, D. (2004). Inden-tation model of the Eastern Alps and the origin of theTauern Window. Geology , 32, 997. 10

Roy, R., Blackwell, D. & Birch, F. (1968). Heatgeneration of plutonic rocks and continental heat flowprovinces*. Earth and Planetary Science Letters, 5, 1–12. 170

Rubatto, D. (2002). Zircon trace element geochemistry:partitioning with garnet and the link between U-Pbages and metamorphism. Chemical Geology , 184, 123–138. 99, 111

249

Rubatto, D. & Gebauer, D. (1999). Eo/Oligocene (35Ma) high-pressure metamorphism in the Gornergratzone (Monte Rosa, western Alps): Implications for pa-leogeography. Schweizerische Mineralogische und Pet-rographische Mitteilungen, 79, 353–362. 131

Rubatto, D. & Hermann, J. (2001). Exhumation asfast as subduction? Geology , 29, 3. 1, 6, 102, 128,131, 133, 188, 189, 190, 197, 198, 201

Rubatto, D. & Hermann, J. (2003). Zircon formationduring fluid circulation in eclogites (Monviso, WesternAlps): implications for Zr and Hf budget in subductionzones. Geochimica et Cosmochimica Acta, 67, 2173–2187. 189, 201

Rubatto, D., Gebauer, D. & Fanning, M. (1998).Jurassic formation and Eocene subduction of theZermatt–Saas-Fee ophiolites: Implications for the geo-dynamic evolution of the Central and Western Alps.Contributions to Mineralogy and Petrology, 132, 269–287. 201

Rubatto, D., Gebauer, D. & Compagnoni, R.(1999). Dating of eclogite-facies zircons: the age ofAlpine metamorphism in the Sesia-Lanzo Zone (West-ern Alps). Earth and Planetary Science Letters, 167,141–158. 201

Rubatto, D., Hermann, J., Berger, A. & Engi, M.(2009). Protracted fluid-induced melting during Barro-vian metamorphism in the Central Alps. Contributionsto Mineralogy and Petrology , 158, 703–722. 198

Rubatto, D., Regis, D., Hermann, J., Boston, K.,Engi, M., Beltrando, M. & McAlpine, S. (2011).Yo-yo subduction recorded by accessory minerals inthe Italian Western Alps. Nature Geoscience, 4, 338–342. 198, 201

Rubie, D. (1986). The catalysis of mineral reactions bywater and restrictions on the presence of aqueous fluidduring metamorphism. Mineralogical Magazine, 50,399–415. 163

Ruffet, G., Feraud, G., Balevre, M. & Kienast, J.(1995). Plateau ages and excess argon in phengites: an40Ar—39Ar laser probe study of Alpine micas (SesiaZone, Western Alps, northern Italy). Chemical Geol-ogy , 121, 327–343. 133

Ruffet, G., Gruau, G., Ballevre, M., Feraud, G.& Philippot, P. (1997). Rb—Sr and 40Ar—39Arlaser probe dating of high-pressure phengites from theSesia zone (Western Alps): underscoring of excessargon and new age constraints on the high-pressuremetamorphism. Chemical Geology , 141, 1–18. 133

Rutherford, E. & Soddy, F. (1902a). The cause andnature of radoactivity, Pt.1. Philosophical Magazine,4, 370–396. 104

Rutherford, E. & Soddy, F. (1902b). The cause andnature of radoactivity, Pt.2. Philosophical Magazine,4, 569–585. 104

Rutter, E. & Brodie, K. (2004). Experimental grainsize-sensitive flow of hot-pressed Brazilian quartz ag-gregates. Journal of structural geology , 26, 2011–2023.173

Rybacki, E., Konrad, K., Renner, J., Wachmann,M., Stockhert, B. & Rummel, F. (2003). Exper-imental deformation of synthetic aragonite marble.Journal of geophysical research, 108, 2174. 173

Sander, B. (1911). Geologische Studien am Westendeder Hohen Tauern. Denkschr. Akad. Wiss. Wien., 82,257–319. 9

Sander, B. (1925). Note illustrative della carta geologicadelle Tre Venezie foglio Bressanone: scale:1:100,000.Padova, U�cio Idografico del Magistrato alle AcqueSezione Geolocia, 56. 172

Sawka, W., Chappell, B. & Norrish, K. (1984).Light-rare-earth-element zoning in sphene and allan-ite during granitoid fractionation. Geology , 12, 131.109

Scaillet, S. (1996). Excess 40Ar transport scale andmechanism in high-pressure phengites: A case studyfrom an eclogitized metabasite of the Dora-Mairanappe, western Alps. Geochimica et CosmochimicaActa, 60, 1075–1090. 133, 158, 163

Schaltegger, U., Brack, P., Ovtcharova, M.,Peytcheva, I., Schoene, B., Stracke, A., Maroc-chi, M. & Bargossi, G. (2009). Zircon and titaniterecording 1.5 million years of magma accretion, crys-tallization and initial cooling in a composite pluton(southern Adamello batholith, northern Italy). Earthand Planetary Science Letters, 286, 208–218. 199, 201

Scharbert, H. (1954). Die eklogitischen Gesteine dessudlichen Großvenedigergebietes (Osttirol). Jb. Geol.BA, 87, 39–63. 38

Scharer, U. (1984). The e↵ect of initial 230 Th dis-equilibrium on young U-Pb ages: The Makalu case,Himalaya. Earth and Planetary Science Letters, 67,191–204. 102, 114, 119, 120, 220

Schellart, W. & Lister, G. (2004). Tectonic modelsfor the formation of arc-shaped convergent zones andback-arc basins. Geological Society of America SpecialPublication, 383, 237–258. 200

250

Scherer, E., Cameron, K. & Blichert-Toft, J.(2000). Lu-Hf garnet geochronology: closure temper-ature relative to the Sm-Nd system and the e↵ects oftrace mineral inclusions. Geochimica et CosmochimicaActa, 64, 3413–3432. 103

Schmid, S. & Kissling, E. (2000). The arc of the west-ern Alps in the light of geophysical data on deep crustalstructure. Tectonics, 19, 62–85. 3, 4

Schmid, S., Aebli, H., Heller, F. & Zingg, A. (1989).The role of the Periadriatic Line in the tectonic evo-lution of the Alps. Geological Society, London, SpecialPublications, 45, 153. 10

Schmid, S., Pfiffner, O., Froitzheim, N.,Schonborn, G. & Kissling, E. (1996). Geophysical-geological transect and tectonic evolution of theSwiss-Italian Alps. Tectonics, 15, 1036–1064. 4

Schmid, S., Pfiffner, O. & Schreurs, G. (1997).Rifting and collision in the Penninic zone of easternSwitzerland. Deep structure of the Swiss Alps-Resultsfrom NFP/PNR, 20, 160–185. 5

Schmid, S., Fugenschuh, B., Kissling, E. & Schus-ter, R. (2004a). Tectonic map and overall architectureof the Alpine orogen. Eclogae Geologicae Helvetiae, 97,93–117. 3, 6, 8, 131

Schmid, S., Fugenschuh, B., Kissling, E. & Schus-ter, R. (2004b). TRANSMED Transects IV, V andVI: Three lithospheric transects across the Alps andtheir forelands. The TRANSMED Atlas: The Mediter-ranean Region from Crust to Mantle. Springer, BerlinHeidelberg, attached CD (version of the explanatorytext available from the first author as a pdf-file uponrequest). 2, 3, 4, 5, 10, 184, 200

Schmidegg, V. (1961). Geologische Ubersicht derVenediger-Gruppe nach dem derzeitigen Stand derAufnahmen von F. Karl und O. Schmidegg. Verhand-lungen der Geologischen, Bundesanstalt , 35–54. 11, 13,38

Schmidt, M. & Thompson, A. (1996). Epidote in calc-alkaline magmas: An experimental study of stability,phase relationships, and the role of epidote in mag-matic evolution. The American mineralogist , 81, 462–474. 109

Schmitz, M. & Schoene, B. (2007). Derivation of iso-tope ratios, errors and error correlations for U-Pbgeochronology using 205 Pb-235 U-(233 U)-spiked iso-tope dilution thermal ionization mass spectrometricdata. Geochemistry, Geophysics, Geosystems, 8. 114,220

Schreyer, W. (1988). Experimental studies on meta-morphism of crustal rocks under mantle pressures.Mineralogical Magazine, 52, 1–26. 50

Schulz, B. & Bombach, K. (2003). Single zircon Pb–Pbgeochronology of the Early-Paleozoic magmatic evolu-tion in the Austroalpine basement to the south of theTauern Window. Jb. Geol. B-A, 143, 303–321. 186

Schulz, B., Siegesmund, S., Steenken, A.,Schonhofer, R. & Heinrichs, T. (2001). Geolo-gie des ostalpinen Kristallins sudlich des Tauernfen-sters zwischen Virgental und Pustertal. Z. Dtsch. Geol.Ges., 152. 186

Sclater, J., Jaupart, C. & Galson, D. (1980). Theheat flow through oceanic and continental crust andthe heat loss of the Earth. Reviews of Geophysics, 18,269–311. 170, 171, 175

Sclater, J., Parsons, B. & Jaupart, C. (1981).Oceans and continents: similarities and di↵erences inthe mechanisms of heat loss. Journal of GeophysicalResearch, 86, 11535. 170, 171

Searle, M., Law, R., Godin, L., Larson, K.,Streule, M., Cottle, J. & Jessup, M. (2008).Defining the Himalayan main central thrust in Nepal.Journal of Geological Society , 165, 523. 164

Selverstone, J. (1985). Petrologic constraints on imbri-cation, metamorphism, and uplift in the SW TauernWindow, Eastern Alps. Tectonics, 4, 687–704. 10, 95

Selverstone, J. (1988). Evidence for east-west crustalextension in the eastern Alps: Implications for the un-roofing history of the Tauern window. Tectonics, 7,87–105. 10

Selverstone, J., Spear, F., Franz, G. & Morteani,G. (1984). High pressure metamorphism in the SWTauern Window, Austria: P-T paths from hornblende-kyanite-staurolite schists. Journal of Petrology , 25,501–531. 10, 83, 93, 94, 100

Shea, W. & Kronenberg, A. (1992). Rheology anddeformation mechanisms of an isotropic mica schist.Journal of Geophysical Research, 97, 15201–15. 172,173

Shen, A.H. & Keppler, H. (1997). Direct observation ofcompletemiscibility in the albite–H

2

O system. Nature,385, 710–712. 47

Sherlock, S. & Kelley, S. (2002). Excess argon evo-lution in HP-LT rocks: a UVLAMP study of phengiteand K-free minerals, NW Turkey. Chemical geology ,182, 619–636. 133, 134, 142, 162, 163

251

Sherlock, S., Kelley, S., Inger, S., Harris, N. &Okay, A. (1999). 40 Ar-39 Ar and Rb-Sr geochronol-ogy of high-pressure metamorphism and exhumationhistory of the Tavsanli Zone, NW Turkey. Contribu-tions to Mineralogy and Petrology , 137, 46–58. 133

Sherlock, S., Lucks, T., Kelley, S. & Barnicoat,A. (2005). A high resolution record of multiple diage-netic events: ultraviolet laser microprobe Ar/Ar anal-ysis of zoned K-feldspar overgrowths. Earth and Plan-etary Science Letters, 238, 329–341. 142

Sherlock, S., Jones, K. & Park, R. (2008). Grenville-age pseudotachylite in the Lewisian: laserprobe40Ar/39Ar ages from the Gairloch region of Scotland(UK). Journal of the Geological Society , 165, 73. 142

Siegesmund, S., Heinrichs, T., Romer, R. & Do-man, D. (2007). Age constraints on the evolution ofthe Austroalpine basement to the south of the TauernWindow. International Journal of Earth Sciences, 96,415–432. 28, 30, 186

Simonetti, A., Heaman, L., Chacko, T. & Baner-jee, N. (2006). In situ petrographic thin section U-Pb dating of zircon, monazite, and titanite using laserablation-MC-ICP-MS. International Journal of MassSpectrometry , 253, 87–97. 113

Sinclair, H. (1997). Flysch to molasse transition in pe-ripheral foreland basins: The role of the passive marginversus slab breako↵. Geology , 25, 1123. 31, 184, 198

Skelton, A. (2011). Metamorphic carbon fluxes: howmuch and how fast? Geology , 39, 95–96. 159

Skelton, A., Valley, J., Graham, C., Bickle, M. &Fallick, A. (2000). The correlation of reaction andisotope fronts and the mechanism of metamorphic fluidflow. Contributions to Mineralogy and Petrology, 138,364–375. 158, 159

Smith, H. & Barreiro, B. (1990). Monazite U-Pb dat-ing of staurolite grade metamorphism in pelitic schists.Contributions to Mineralogy and Petrology, 105, 602–615. 107, 109

Smith, S. & Kennedy, B. (1983). The solubility of no-ble gases in water and in NaCl brine. Geochimica etCosmochimica Acta, 47, 503–515. 154, 157

Smye, A., Greenwood, L. & Holland, T. (2010).Garnet-chloritoid-kyanite assemblages: eclogite faciesindicators of subduction constraints in orogenic belts.Journal of Metamorphic Geology , 28, 753–768. 33, 36,43, 44, 57, 58, 67, 107, 109

Smye, A., Bickle, M., Holland, T., Parrish, R. &Condon, D. (2011). Rapid formation and exhumation

of the youngest Alpine eclogites: A thermal conun-drum to Barrovian metamorphism. Earth and Plane-tary Science Letters, 306, 193–204. 10, 99, 164

Spear, F. (1986). PT evolution of metasediments fromthe Eclogite Zone, south-central Tauern Window, Aus-tria. Lithos, 19, 219–234. 38, 48, 50

Spear, F. & Cheney, J. (1989). A petrogenetic gridfor pelitic schists in the system SiO2-Al2O3-FeO-MgO-K2O-H2O. Computer , 101, 149–164. 35

Spear, F. & Parrish, R. (1996). P-T-t evolution of theValhalla Complex, British Columbia, Canada. Journalof Petrology , 37, 733–765. 103

Spiegel, C., Kuhlemann, J., Dunkl, I., Frisch, W.,Von Eynatten, H. & Balogh, K. (2000). The ero-sion history of the Central Alps: evidence from zir-con fission track data of the foreland basin sediments.Terra Nova, 12, 163–170. 199

Stacey, J. & Kramers, J. (1975). Approximation ofterrestrial lead isotope evolution by a two-stage model.Earth and Planetary Science Letters, 26, 207–221.105, 115, 119

Stampfli, G. (2000). Tethyan oceans. Geological society,london, special publications, 173, 1. 5

Stampfli, G. & Marchant, R. (1997). Geodynamicevolution of the Tethyan margins of the Western Alps.Deep structure of the Swiss AlpsResults from NRP , 20,223–239. 5

Stampfli, G., Mosar, J., Marquer, D., Marchant,R., Baudin, T. & Borel, G. (1998). Subduction andobduction processes in the Swiss Alps. Tectonophysics,296, 159–204. 2

Staub, R. (1924). Der Bau der Alpen. Dalp. 9

Steenken, A., Siegesmund, S. & Heinrichs, T.(2000). The emplacement of the Rieserferner Pluton(Eastern Alps, Tyrol): constraints from field observa-tions, magnetic fabrics and microstructures. Journalof Structural Geology , 22, 1855–1873. 172

Steiger, R. & Jager, E. (1977). Subcommission ongeochronology: convention on the use of decay con-stants in geo-and cosmochronology. Earth and plane-tary science letters, 36, 359–362. 104, 134

Steinmann, M. (1994). Ein Beckenmodell fur das Nord-penninikum der Ostschweiz. Jahrbuch der Geologis-chen Bundesanstalt , 137, 675–721. 5

Stella, A. (1894). Relazione sul rilevamento eseguitonell’anno 1893 nelle Alpi Occidentali (Valli dell’Orcoe della Soana). Bolletino del Reale Comitato Geologicad’Italia, 25, 343–371. 50

252

Stıpska, P. & Powell, R. (2005). Constraining the P-T path of a MORB-type eclogite using pseudosections,garnet zoning and garnet-clinopyroxene thermometry:an example from the Bohemian Massif. Journal ofMetamorphic Geology , 23, 725–743. 68

Stockhert, B. & Gerya, T. (2005). Pre-collisional highpressure metamorphism and nappe tectonics at activecontinental margins: a numerical simulation. TerraNova, 17, 102–110. 62

Stockhert, B., Massonne, H. & Nowlan, E. (1997).Low di↵erential stress during high-pressure metapelitefrom the eclogite Zone, Tauern Window, Eastern Alps.Lithos, 41, 103–118. 49, 60, 190

Stockli, D., Farley, K. & Dumitru, T. (2000). Cali-bration of the apatite (U-Th)/He thermochronometeron an exhumed fault block, White Mountains, Califor-nia. Geology , 28, 983–986. 103

Storey, C., Jeffries, T. & Smith, M. (2006). Com-mon lead-corrected laser ablation ICP-MS U-Pb sys-tematics and geochronology of titanite. Chemical geol-ogy , 227, 37–52. 113

Stuwe, K. & Sandiford, M. (1995). Mantle-lithospheric deformation and crustal metamorphismwith some speculations on the thermal and mechan-ical significance of the Tauern Event, Eastern Alps.Tectonophysics, 242, 115–132. 164, 166

Sue, C., Thouvenot, F., Frechet, J. & Tricart, P.(1999). Widespread extension in the core of the west-ern Alps revealed by earthquake analysis. Journal ofgeophysical research, 104, 611–25. 3

Tanner, S., Cousins, L. & Douglas, D. (1994). Re-duction of space charge e↵ects using a three-aperturegas dynamic vacuum interface for inductively coupledplasma-mass spectrometry. Applied Spectroscopy , 48,1367–1372. 113

Tapponnier, P. (1977). Evolution tectonique du systemealpin en Mediterranee: poinconnement et ecrasementrigide-plastique. Bulletin de la Societe geologique deFrance, Geodynamique de la Mediterranee occidentaleet de ses abords. 2

Teller, F. (1889). Jahresbericht 1888 , vol. 5. Vehr GeolR A Wien. 172

Tenthorey, E. & Hermann, J. (2004). Compositionof fluids during serpentinite breakdown in subductionzones: Evidence for limited boron mobility. Geology ,32, 865. 204

Tera, F. & Wasserburg, G. (1972). U-Th-Pb system-atics in three Apollo 14 basalts and the problem ofinitial Pb in lunar rocks. Earth and Planetary ScienceLetters, 14, 281–304. 115

Tera, F. & Wasserburg, G. (1973). A response to acomment on U-Pb systematics in lunar basalts. Earthand Planetary Science Letters, 19, 213–217. 115

Tera, F. & Wasserburg, G. (1974). U-Th-Pb system-atics on lunar rocks and inferences about lunar evolu-tion and the age of the Moon. In Proceedings of the 5thLunar Conference, Geochim. Cosmochim. Acta, vol. 2,1571–1599. 115

Thoni, M. & Jagoutz, E. (1992). Some new aspectsof dating eclogites in orogenic belts: Sm–Nd, Rb–Sr,and Pb–Pb results from the Austroalpine Saualpe andKoralpe type-locality (Carinthia/Styria, SE Austria).Geochim Cosmochim Acta, 56, 347–368. 101, 152

Thoni, M., Miller, C., Blichert-Toft, J., White-house, M., Konzett, J. & Zanetti, A. (2008).Timing of high-pressure metamorphism in the eclogitetype-locality (Kupplerbrunn-Prickler Halt, Saualpe,south-eastern Austria): constraints from correlationsof the Sm-Nd, Lu-Hf, U-Pb and Rb-Sr isotopic sys-tems. Journal of Metamorphic Geology , 26, 561–581.131

Thouvenot, F. (1996). Aspects geophysiques et struc-turaux des Alpes occidentales et de trois autresorogenes (Atlas, Pyrenees, Oural). These dEtat . 3

Thouvenot, F., Frechet, J., Guyoton, F.,Guiguet, R. & Jenatton, L. (1990). SISMALP:an automatic phone-interrogated seismic network forthe western Alps. Cahiers du Centre Europeen deGeodynamique et de Seismologie, 1, 1–10. 3

Tilton, G., Schreyer, W. & Schertl, H. (1991). Pb-Sr- Nd isotopic behavior of deeply subducted crustalrocks from the Dora Maira Massif, Western Alps, Italy-II: what is the age of the ultrahigh-pressure metamor-phism? Contributions to Mineralogy and Petrology,108, 22–33. 201

Tinkham, D., Zuluaga, C. & Stowell, H. (2001).Metapelite phase equilibria modeling in MnNCKF-MASH: the e↵ect of variable Al 2 O 3 andMgO/(MgO+ FeO) on mineral stability. Geol. Mater.Res, 3, 1–42. 35, 36

Tomkins, H. & Pattison, D. (2007). Accessory phasepetrogenesis in relation to major phase assemblagesin pelites from the Nelson contact aureole, southernBritish Columbia. Journal of Metamorphic Geology ,25, 401–421. 107

Torgersen, T., Kennedy, B., Hiyagon, H., Chiou,K., Reynolds, J. & Clarke, W. (1989). Argon accu-mulation and the crustal degassing flux of40Ar in theGreat Artesian Basin, Australia. Earth and planetaryscience letters, 92, 43–56. 156, 157

253

Tropper, P., Essene, E., Sharp, Z. & Hun-ziker, J. (1999). Application of K-feldspar-jadeite-quartz barometry to eclogite facies metagranites andmetapelites in the Sesia Lanzo Zone (Western Alps,Italy). Journal of Metamorphic Geology , 17, 195–210.50

Trumpy, R. (1955). Remarques sur la correlation desunites penniques externes antre la Savoie et le Valaiset sur l’origine des nappes prealpines. Bulletin de laSociete Geologique de France, 6, 217–231. 5, 200

Trumpy, R. (1980). An Outline of the Geology ofSwitzerland, Part A. Schweizerische Geologische Kom-mission, Basel. 198

Turner, G., Huneke, J., Podosek, F. & Wasser-burg, G. (1971). 40Ar-39Ar ages and cosmic ray expo-sure ages of Apollo 14 samples+. Earth and PlanetaryScience Letters, 12, 19–35. 143, 146

Twiss, R. & Moores, E. (1992). Structural Geology .WH Freeman. 21

Udovkina, N., Muravitskaya, G. & Laputina, I.(1977). Talc-garnet-kyanite rocks of the Kokchetavblock, Northern Kazakhstan. Doklady Akad.NaukSSSR, 237, 202–205. 52

Van Breemen, O. & Hawkesworth, C. (1980). Sm-Ndisotopic study of garnets and their metamorphic hostrocks. Trans. Roy. Soc. Edinburgh, 71, 97–102. 100,102

van den Beukel, J. (1992). Some thermomechanical as-pects of the subduction of continental lithosphere. Tec-tonics, 11, 316–329. 61

Vance, D., Muller, W. & Villa, I. (2003).Geochronology: linking the isotopic record withpetrology and textures–an introduction. Geological So-ciety London Special Publications, 220, 1. 99

Venturini, G. (1995). Geology, geochemistry andgeochronology of the inner central Sesia Zone (WesternAlps, Italy). Memoirs de geologie (Lausanne). 201

Villa, I., Hermann, J., Muntener, O. & Tromms-dorff, V. (2000). 39 Ar- 40 Ar dating of multiplyzoned amphibole generations (Malenco, Italian Alps).Contributions to Mineralogy and Petrology, 140, 363–381. 201

von Blanckenburg, F. (1992). Combined high preci-sion chronometry and geochemical tracing using acces-sory minerals: applied to the Central-Alpine Bergellintrusion (central Europe). Chemical Geology , 100,19–40. 114, 117

Von Blanckenburg, F. & Huw Davies, J. (1995).Slab breako↵: a model for syncollisional magmatismand tectonics in the Alps. Tectonics, 14, 120–120. 131,197, 200

Von Eynatten, H., Schlunegger, F., Gaupp, R. &Wijbrans, J. (1999). Exhumation of the Central Alps:evidence from 40Ar/39Ar laserprobe dating of detritalwhite micas from the Swiss Molasse Basin. Terra Nova,11, 284–289. 133, 199

Vuichard, J. & Ballevre, M. (1988). Garnet-chloritoid equilibria in eclogitic pelitic rocks from theSesia zone (Western Alps): their bearing on phase re-lations in high pressure metapelites. Journal of Meta-morphic Geology , 6, 135–157. 50

Wagner, G., Miller, D. & Jager, E. (1979). Fissiontrack ages on apatite of Bergell rocks from central Alpsand Bergell boulders in Oligocene sediments. Earthand Planetary Science Letters, 45, 355–360. 201

Wagner, R., Rosenberg, C., Handy, M., M”obus, C. & Albertz, M. (2006). Fracture-drivenintrusion and upwelling of a mid-crustal pluton fedfrom a transpressive shear zoneThe Rieserferner Plu-ton (Eastern Alps). Geological Society of America Bul-letin, 118, 219. 172

Wakabayashi, J. (1990). Counterclockwise PTt pathsfrom amphibolites, Franciscan Complex, California:relics from the early stages of subduction zone meta-morphism. The Journal of Geology , 657–680. 189

Wallis, S. (1988). The structural and kinematic devel-opment of the Austroalpine Pennine boundary in theS.E.Tauern, E.Alps.. Ph.D. thesis, University of Ox-ford. 168, 174, 183, 186

Wallis, S. & Behrmann, J. (1996). Crustal stackingand extension recorded by tectonic fabrics of the SEmargin of the Tauern Window, Austria. Journal ofStructural Geology , 18, 1455–1470. 183

Walther, J. & Orville, P. (1982). Volatile produc-tion and transport in regional metamorphism. Contri-butions to Mineralogy and Petrology , 79, 252–257. 155

Warren, C., Parrish, R., Searle, M. & Waters, D.(2003). Dating the subduction of the Arabian conti-nental margin beneath the Semail ophiolite, Oman.Geology , 31, 889. 106

Warren, C., Sherlock, S. & Kelley, S. (2010). Inter-preting high-pressure phengite 40 Ar/39 Ar laserprobeages: an example from Saih Hatat, NE Oman. Contri-butions to Mineralogy and Petrology , 1–19. 133, 142,147, 162, 163

254

Warren, C., Smye, A., Kelley, S. & Sherlock,S. (2011). Using white mica 40Ar/39Ar data as atracer for fluid flow and permeability under high-P conditions: Tauern Window, Eastern Alps. Jour-nal of Metamorphic Geology , doi:10.1111/j.1525-1314.2011.00956.x. 133

Wartho, J., Kelley, S., Brooker, R., Carroll, M.,Villa, I. & Lee, M. (1999). Direct measurement ofAr di↵usion profiles in a gem-quality Madagascar K-feldspar using the ultra-violet laser ablation micro-probe (UVLAMP). Earth and planetary science let-ters, 170, 141–153. 154

Waters, D. (1976). Structural, metamorphic andgeochronological studies in the southeast Tauern. Ph.D.thesis, University of Oxford. 28, 30, 186

Wei, C. & Clarke, G. (2011). Calculated phase equi-libria for MORB compositions: a reappraisal of themetamorphic evolution of lawsonite eclogite. Journalof Metamorphic Geology , In Press. 68

Wei, C. & Powell, R. (2003). Phase relations in high-pressure metapelites in the system KFMASH(K

2

O–FeO–MgO–Al

2

O3

–SiO2

–H2

O) with application to nat-ural rocks. Contributions to Mineralogy and Petrology,145, 301–315. 50

Wei, C. & Powell, R. (2006). Calculated phaserelations in the system NCKFMASH (Na

2

O–CaO–K

2

O–FeO–MgO–Al2

O3

–SiO2

–H2

O) for high-pressuremetapelites. Journal of Petrology , 47, 385–408. 50, 59

Wei, C., Yang, Y., Su, X., Song, S. & Zhang, L.(2009). Metamorphic evolution of low-T eclogite fromthe North Qilian orogen, NW China: evidence frompetrology and calculated phase equilibria in the systemNCKFMASHO. Journal of Metamorphic Geology , 27,55–70. 68

Weiss, R. (1970). The solubility of nitrogen, oxygen andargon in water and seawater. In Deep Sea Research andOceanographic Abstracts, vol. 17, 721–735, Elsevier.154

Wetherill, G. (1966). Radioactive decay constants andenergies. Handbook of Physical Constants. GeologicalSociety America, Memoir , 97, 513–519. 104

Wheeler, J. (1996). DIFFARG: a program for simulat-ing argon di↵usion profiles in minerals. Computers &Geosciences, 22, 919–929. 146

White, R., Powell, R., Holland, T. & Worley, B.(2000). The e↵ect of TiO2 and Fe2O3 on metapeliticassemblages at greenschist and amphibolite facies con-ditions: mineral equilibria calculations in the systemK2O±FeO±MgO±Al2O3±SiO2±H2O±TiO2±Fe2O3.

Journal of Metamorphic Geology , 18, 497–511. 35,43, 56

White, R., Pomroy, N. & Powell, R. (2005). An insitu metatexite–diatexite transition in upper amphibo-lite facies rocks from Broken Hill, Australia. Journalof Metamorphic Geology , 23, 579–602. 43, 45, 47

White, R., Powell, R. & Holland, T. (2007).Progress relating to calculation of partial melting equi-libria for metapelites. Journal of Metamorphic Geol-ogy , 25, 511–527. 36, 67

Whitney, P. & McLelland, J. (1973). Origin of coro-nas in metagabbros of the Adirondack Mts., NY. Con-tributions to Mineralogy and Petrology , 39, 81–98. 34,69

Whittington, A., Hofmeister, A. & Nabelek, P.(2009). Temperature-dependent thermal di↵usivity ofthe Earth’s crust and implications for magmatism. Na-ture, 458, 319–321. 164, 180, 181, 182, 183

Wiedenbeck, M., Alle, P., Corfu, F., Griffin, W.,Meier, M., Oberli, F., Quadt, A., Roddick, J. &Spiegel, W. (1995). Three natural zircon standardsfor U-Th-Pb, Lu-Hf, trace element and REE analyses.Geostandards and Geoanalytical Research, 19, 1–23.113, 114, 219

Wiedenbeck, M., Hanchar, J., Peck, W.,Sylvester, P., Valley, J., Whitehouse, M.,Kronz, A., Morishita, Y. & Nasdala, L. (2004).Further Characterisation of the 91500 Zircon Crystal.Geostandards and Geoanalytical Research, 28, 9–39.115

Wiederkehr, M., Sudo, M., Bousquet, R., Berger,A. & Schmid, S. (2009). Alpine orogenic evolutionfrom subduction to collisional thermal overprint: The40Ar/39Ar age constraints from the Valaisan Ocean,central Alps. Tectonics, 28, TC6009. 133, 142, 162,198, 201

Will, T., Okrusch, M., Schmadicke, E. & Chen, G.(1998). Phase relations in the greenschist-blueschist-amphibolite-eclogite facies in the system Na

2

O-CaO-FeO-MgO-Al

2

O3

-SiO2

-H2

O (NCFMASH), with appli-cation to metamorphic rocks from Samos, Greece.Contributions to Mineralogy and Petrology , 132, 85–102. 72

Wilson, S. (1997). The collection, preparation, andtesting of USGS reference material BCR-2, ColumbiaRiver, Basalt. U.S. Geological Survey Open-File Re-port , 98. 218

Wing, B., Ferry, J. & Harrison, T. (2003). Progradedestruction and formation of monazite and allanite

255

during contact and regional metamorphism of pelites:petrology and geochronology. Contributions to Miner-alogy and Petrology , 145, 228–250. 99, 107, 109, 111

Winkler, M. (1996). Genese und geodynamische Stel-lung der Zentralgneise im Tauernfenster . Inst. undMuseum f”ur Geologie und Pal”aontologie. 89, 90, 91, 92

Yamada, R., Tagami, T., Nishimura, S. & Ito, H.(1995). Annealing kinetics of fission tracks in zircon:an experimental study. Chemical Geology , 122, 249–258. 103

Yamato, P., Agard, P., Burov, E., Le Pourhiet,L., Jolivet, L. & Tiberi, C. (2007). Burial and ex-humation in a subduction wedge: Mutual constraintsfrom thermomechanical modeling and natural PTtdata (Schistes Lustres, western Alps). Journal of Geo-physical Research, 112, B07410. 99

YANG, J. et al. (2008). Ultrahigh-pressure garnet peri-dotites from the devolatilization of sea-floor hydratedultramafic rocks. Journal of Metamorphic Geology , 26,695–716. 68

Zack, T. & Luvizottow, G. (2006). Application of ru-tile thermometry to eclogites. Mineralogy and Petrol-ogy , 88, 69–85. 106

Zanolla, C., Braitenberg, C., Ebbing, J., Bern-abini, M., Bram, K., Gabriel, G., Gotze, H.,Giammetti, S., Meurers, B., Nicolich, R. et al.(2006). New gravity maps of the Eastern Alps andsignificance for the crustal structures. Tectonophysics,414, 127–143. 173

Zaun, P. & Wagner, G. (1985). Fission-track stabilityin zircons under geological conditions. Nuclear Tracks,10, 303–307. 103

Zimmermann, R., Hammerschmidt, K. & Franz, G.(1994). Eocene high pressure metamorphism in thePenninic units of the Tauern Window (Eastern Alps):evidence from 40 Ar- 39 Ar dating and petrological in-vestigations. Contributions to Mineralogy and Petrol-ogy , 117, 175–186. 10, 39, 84, 93, 94, 100, 101, 102,134, 135, 140, 160, 162, 180, 181, 185, 201

Zucali, M., Spalla, M. & Gosso, G. (2002). Strainpartitioning and fabric evolution as a correlation tool:the example of the Eclogitic Micaschists Complex inthe Sesia-Lanzo Zone (Monte Mucrone-Monte Mars,Western Alps, Italy). Schweizerische Mineralogischeund Petrographische Mitteilungen, 82, 429–454. 50,58, 59

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