Origin and Evolution of Deep Upper Mantle Rocks ...

GEOLOGICA ULTRAIECTINA Mededelingen van de Faculteit Geowetenschappen Universiteit Utrecht No. 266

Origin and Evolution of Deep Upper Mantle Rocks from Western Norway

Dirk Spengler

The PhD-project was conducted under the programme of the Netherlands Research Centre for Integrated Solid Earth Sciences (ISES) and the Vening Meinesz Research School of Geodynamics (VMSG) and carried out at: Department of Earth Sciences Utrecht University Budapestlaan 4 3584 CD Utrecht The Netherlands

Department of Petrology Vrije Universiteit Amsterdam De Boelelaan 1085 1081 HV Amsterdam The Netherlands

Lamont–Doherty Earth Observatory Columbia University Box 1000, 61 Route 9W Palisades, NY 10964 USA

Department of Earth Sciences Durham University South Road Durham DH1 3LE UK

ISBN-13: 978–90–5744–130–1 ISBN-10: 90–5744–130–6

Origin and Evolution of Deep Upper Mantle Rocks from Western Norway Entstehung und Entwicklung von Gesteinen des tieferen oberen Erdmantels in West–Norwegen (mit einer Zusammenfassung in deutscher Sprache)

Oorsprong en evolutie van diepe bovenmantelgesteenten van West Noorwegen (met een samenvatting in het Nederlands)

proefschrift ter verkrijging van de graad van doctor aan de Universiteit Utrecht op gezag van de rector magnificus, prof. dr. W.H. Gispen, ingevolge het besluit van het college voor promoties in het openbaar te verdedigen op maandag 30 oktober 2006 des middags te 12.45 uur door

Dirk Spengler geboren op 18 maart 1972, te Eberswalde, Duitsland

Promotor:

Prof. dr. S.H. White

Co–promotoren:

Dr. H.L.M. van Roermund Dr. M.R. Drury

The project was financiated by the Department of Earth Sciences, Utrecht University, and financially supported by the Department of Petrology, Vrije Universiteit Amsterdam, led by Prof. dr. G.R. Davies.

Examiners:

Prof. dr. R. Altherr (University of Heidelberg) Prof. dr. H.K. Brueckner (Columbia University and The City University of New York) Prof. dr. W.L. Griffin (Macquarie University, Sydney) Prof. dr. M.A. Menzies (Royal Holloway/University of London) Prof. dr. R.L.M. Vissers (Utrecht University)

“Ever tried. Ever failed. No matter. Try again. Fail again. Fail better.” Samuel Beckett

Summary/Zusammenfassung/ Samenvatting Summary The growth of the first continents on Earth marks a major change in the early planetary evolution. The mechanisms of continental lithosphere growth are controversial as most of the scarce accessible samples from deep cratonic lithosphere (the rigid mantle underneath the continents) hide their earliest records beneath secondary overprints during the last three billion years. This thesis investigates fragments of such old (≥3 Ga) subcontinental lithospheric mantle, which occur as isolated rock bodies (orogenic peridotites) in crustal gneiss exposed on Otrøy in the mountain chain of the Scandinavian Caledonides in western Norway. The Otrøy mantle fragments are compositionally banded and include garnet peridotite, spinel peridotite, garnetite and garnet pyroxenite. Garnet contains microscopic lamellae and comparably coarse grains of pyroxene in two mineral microstructures. The pyroxene lamellae and the host garnet have a strong crystallographic relationship, which characterise this microstructure as a break down product of a former majorite precursor mineral. Majorite is a typical component of the deep mantle (≥150 km). In situ analyses of rare Earth elements (REE) in the majorite break down products reveal that the break down reaction occurred at minimum temperatures of 1300 °C. Sm–Nd and Re–Os isotopes constrain the origin of most rock types between c. 3.3–2.9 Ga, in the Archaean era. Both temperature and age information demonstrate that the occurrence of the ultra-high pressure mineral indicator majorite in orogenic peridotite does not necessarily record conditions of the latest major geodynamic overprint, deep continental subduction, as recently proposed for other orogens. Coarse pyroxene associated with garnet in the other microstructure records similar initial minimum temperatures and have similar exceptionally low concentrations of light REE as pyroxene lamellae. This implies that the Otrøy mantle fragments originated at much greater depth (≥350 km) than the thickness of the oldest continents (c. 200 km). Mineral microstructures, mineral chemistry and melt modelling constrain that an up-welling, extremely hot (>1800 °C) mantle caused melting and emplacement of the peridotites at c. 150 km depth forming an ultra-depleted cratonic root in the early Earth. Subsequent metasomatism locally enriched the melt

viii

ZUSAMMENFASSUNG

depleted peridotites in melt-compatible elements (Fe, Ca and light REE), but also in Cr and heavy REE, elements which are usually regarded as melt-incompatible. Metasomatised peridotites preserve Archaean ages and have higher Cr content than primitive mantle. Cr and heavy REE are interpreted to be mobile during high degrees of peridotite melting (c. 50 %) at high pressures (c. 5 GPa). Sub-cratonic emplacement of the peridotites may have caused strong deformation, which explains the present compositional layering by mechanical mixing of garnet-rich and garnet-poor lithologies. Folding and recrystallization of this layering occurred during tectonic exhumation of the mantle fragments associated with the Caledonian orogeny more than 2 billion years later. Recrystallized assemblages lack pyroxene inclusions in garnet, preserve peak metamorphic conditions of 870 °C and 6.5 GPa (corresponding to c. 200 km depth) and have early Scandian ages of c. 430 Ma. This shows that continental subduction during orogeny in western Norway exceeded deep into the diamond stability field. The ultra-high pressure metamorphism appears to have lasted significantly longer than currently accepted in this region. Nevertheless evidence of lithosphere formation during early Earth evolution is preserved within the peridotite bodies, demonstrating that the Scandian tectonometamorphic overprint did not wipe out the earlier record as is generally thought by the scientific community. The chemical and microstructural relicts of a pre-orogenic history allows early Earth lithospheric studies to be performed in a totally new geological setting.

Zusammenfassung Ein einschneidendes Ereignis in der fr¨ uhen Erdgeschichte ist die Entstehung der ersten Kontinente. Die Prozesse, die das Wachstum der Kontinente vor mehr als 3 Milliarden Jahren einleiteten, sind jedoch nicht eindeutig geklärt. Zu den Ursachen zählt, dass die klassischen bis zu wenigen Dezimeter großen Gesteinsproben (Xenolite) vom unteren Teil der Kontinente (des sub-kontinentalen lithosphärischen Mantels) durch eine j¨ ungere vulkanische Aktivität zur Erdoberfläche transportiert wurden, wodurch die alten Informationen, die Aufschluss u ¨ ber das Entstehen der Lithosphäre geben, häufig u ¨ berprägt wurden. In der vorliegenden Arbeit werden Kilometer große Fragmente des subkontinentalen Mantels untersucht, die auf der Insel Otrøy in West-Norwegen als tektonische Linsen in Gneiss an der Oberfläche eingeschlossen sind. Unter Druckentlastung gebildete Mineralmikrostrukturen zeigen, dass diese Mantelgesteine (Peridotite) aus einer enormen Tiefe stammen (≥350 km). Die Ursprungstiefe u ¨ berschreitet die Dicke der Lithossphäre unter den ältesten Kontinenten (ca. 200 km). Die chemische Zusammensetzungen der Minerale zeigen zudem, dass das aufsteigende Mantelgestein mit mindestens 1800 °C sehr heiß war und noch in großer Tiefe bei ca. 150 km teilweise schmolz (zu etwa 50 %). Der geschmolzene Teil des Gesteins spaltete sich schließlich als komatiitische Schmelze

SAMENVATTING

ix

vom Ausgangsgestein ab. Das u ¨ brige, nicht geschmolzene Mantelgestein bildete in der fr¨ uhen Erdzeit vor 3,3–2,9 Milliarden Jahren den Boden der lithosphärischen Platte unter Grönland. Durch ein temporäres Eintauchen (Subduktion) der kontinentalen Platte Baltica unter die Gronländische Platte mehr als 2 Milliarden Jahre später wurden Teile dieses schmelzverarmten lithosphärischen Bodens an die Erdoberfläche transportiert. Diese Plattenkollision verursachte die Enstehung des Gebirges in Skandinavien (Caledoniden). Die damit verbundene Gesteinsdeformation war intensiv, zerstörte jedoch nicht alle archaischen Mikrostrukturen in den Mineralen. Folglich können derartige Mikrostrukturen nicht als Beweis f¨ ur eine relativ junge und tiefe kontinentale Subduktion angef¨ uhrt werden, wie dies k¨ urzlich in anderen wissenschaftlichen Arbeiten formuliert wurde. Aufgrund dieser Erkenntnis ist es nun möglich, die Entstehung und Entwicklung der ältesten Kontinente an einem Gesteinstypus (orogene Peridotite) zu rekonstruieren, der bisher wenig auf diese Problemstellungen hin untersucht wurde.

Samenvatting De vorming van de eerste koude lithospherische platen (continenten) op aarde markeert een grote verandering in de evolutie van de aarde. Onze kennis hierover is controversieel omdat in de traditionele onderzoeks-monsters, sub-meter schalige mantel fragmenten (xenolieten) in vulkanische gesteenten, de oudste informatie over het ontstaan van de lithospheer vaak wordt overprint door jongere geologische gebeurtenissen (o.a. gerelateerd aan het magmatisme). Dit onderzoek toont aan dat kilometer schalige fragmenten van de (sub-)continentale lithospheer, nu gevonden in Noorwegen als tektonische lenzen ingesloten in continentale korst, afkomstig zijn van enorme diepte, ≥350 km. Deze diepte is groter dan de maximale dikte van de lithospheer onder oude continenten (c. 200 km). Decompressie microstructuren en chemische samenstellingen van mineralen laten tevens zien dat het opstijgende mantelgesteente erg heet was (≥1800 °C) en op grote diepte (c. 150 km) gedeeltelijk is gesmolten (tot c. 50 %). De gesmolten helft heeft zich uiteindelijk als een komati¨ıtische smelt fase van het moeder materiaal afgesplitst. Het overgebleven mantelgesteente vormde de bodem van een lithospherische plaat onder Groenland in de vroege aarde, 3.3–2.9 miljard jaar geleden. Twee miljard jaar later volgde de tektonische exhumatie van een deel van dit mantelgesteente naar het aardoppervlak als gevolg van de Caledonische gebergtevorming in Skandinavië. Daarbij schoof (=subduceerde) de Baltische- tijdelijk onder de Groenlandse plaat en haalde fragmenten van de continentale onderkant naar boven. De oude Archae¨ısche mineraal microstructuren bleven hierbij bewaard. Echter deze microstructuren kunnen niet als bewijs worden gebruikt voor diepe, continentale subductie, zoals onlangs is voorgesteld door andere onderzoekers. In plaats daarvan tonen ze aan dat de vroege aarde in een heel andere geologische omgeving gestudeerd kan worden.

Contents 1 Introduction 1.1

1.2

1.3

1

PhD project . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

3

1.1.1

Thesis introduction . . . . . . . . . . . . . . . . . . . . . . .

3

1.1.2

Problem statement . . . . . . . . . . . . . . . . . . . . . . .

4

1.1.3

Thesis outline . . . . . . . . . . . . . . . . . . . . . . . . . .

6

Upper mantle and upper mantle peridotites . . . . . . . . . . . . .

6

1.2.1

The Earth’s outer structure . . . . . . . . . . . . . . . . . .

6

1.2.2

Lithospheric mantle . . . . . . . . . . . . . . . . . . . . . . .

7

1.2.3

Classification of SCLM peridotite . . . . . . . . . . . . . . .

9

Geological setting . . . . . . . . . . . . . . . . . . . . . . . . . . . .

10

1.3.1

Scandinavian Caledonides . . . . . . . . . . . . . . . . . . .

10

1.3.2

Western Gneiss Region . . . . . . . . . . . . . . . . . . . . .

23

1.3.3

Otrøy island . . . . . . . . . . . . . . . . . . . . . . . . . . .

27

2 Characterisation of exsolution microstructures in Grt

33

2.1

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

35

2.2

Types of Pyx microstructures . . . . . . . . . . . . . . . . . . . . .

35

2.2.1

Pyroxene in garnet . . . . . . . . . . . . . . . . . . . . . . .

35

2.2.2

Pyroxene in pyroxene . . . . . . . . . . . . . . . . . . . . . .

41

Determining crystallographic relationships . . . . . . . . . . . . . .

43

2.3.1

Principle of EBSD . . . . . . . . . . . . . . . . . . . . . . .

43

2.3.2

Samples and results . . . . . . . . . . . . . . . . . . . . . . .

44

2.3.3

Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . .

46

Quantifying the intracrystalline microstructure . . . . . . . . . . . .

46

2.4.1

Method 1: Volume estimates using 2D image analysis . . . .

47

2.4.2

Method 2: Chemical integration using defocussed EMP mapping 52

2.4.3

Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . .

56

2.4.4

Application . . . . . . . . . . . . . . . . . . . . . . . . . . .

58

Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

60

2.3

2.4

2.5

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CONTENTS

3 Deep origin and hot melting of an Archaean orogenic peridotite massif in Norway 61 3.1

Abstract . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

63

3.2

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

63

3.3

Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

63

3.4

Petrography . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

65

3.5

Mineral chemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . .

66

3.5.1

Major elements . . . . . . . . . . . . . . . . . . . . . . . . .

66

3.5.2

REE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

67

3.5.3

Sm–Nd isotopes . . . . . . . . . . . . . . . . . . . . . . . . .

70

Discussion and conclusions . . . . . . . . . . . . . . . . . . . . . . .

71

3.6

4 Refractory history of Otrøy Grt-peridotite

75

4.1

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

77

4.2

Petrography and samples . . . . . . . . . . . . . . . . . . . . . . . .

78

4.2.1

Single and polycrystalline Grt – in harzburgite . . . . . . . .

80

4.2.2

Grt – in dunite . . . . . . . . . . . . . . . . . . . . . . . . .

82

4.2.3

Grt-free assemblages . . . . . . . . . . . . . . . . . . . . . .

82

4.2.4

Samples . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

83

4.3

Analytical methods . . . . . . . . . . . . . . . . . . . . . . . . . . .

84

4.4

Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

84

4.4.1

Major and minor elements . . . . . . . . . . . . . . . . . . .

84

4.4.2

REE . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

91

4.4.3

Melting estimates . . . . . . . . . . . . . . . . . . . . . . . .

94

4.4.4

PGE and Re . . . . . . . . . . . . . . . . . . . . . . . . . . .

97

4.4.5

Isotope chemistry . . . . . . . . . . . . . . . . . . . . . . . .

101

Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

108

4.5.1

Origin of Grt and the compositional layering . . . . . . . . .

108

4.5.2

Melting environment . . . . . . . . . . . . . . . . . . . . . .

117

4.5.3

Ni in Ol . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

121

Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

125

4.5

4.6

xiii

CONTENTS 5 Metasomatism of sub-continental lithospheric mantle

127

5.1

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

129

5.2

General petrography . . . . . . . . . . . . . . . . . . . . . . . . . .

129

5.3

Analytical techniques . . . . . . . . . . . . . . . . . . . . . . . . . .

131

5.4

Whole rock chemistry . . . . . . . . . . . . . . . . . . . . . . . . . .

132

5.4.1

Major elements . . . . . . . . . . . . . . . . . . . . . . . . .

132

5.4.2

Trace elements . . . . . . . . . . . . . . . . . . . . . . . . .

136


137

5.5.1

Major elements . . . . . . . . . . . . . . . . . . . . . . . . .

137

5.5.2

REE and Y in Grt . . . . . . . . . . . . . . . . . . . . . . .

145

Sm–Nd isotope chemistry . . . . . . . . . . . . . . . . . . . . . . . .

147

5.6.1

Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

147

5.6.2

Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

147

Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

148

5.7.1

Peridotite precursor chemistry . . . . . . . . . . . . . . . . .

148

5.7.2

Chemical modification during metasomatism . . . . . . . . .

150

5.7.3

Evolution of Grt . . . . . . . . . . . . . . . . . . . . . . . .

152

5.7.4

Timing . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

154

Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

155

5.5

5.6

5.7

5.8

6 Emplacement of mantle peridotite into the crust and subsequent evolution 157 6.1

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

159

6.2

Structural setting . . . . . . . . . . . . . . . . . . . . . . . . . . . .

160

6.2.1

Gneiss . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

160

6.2.2

Peridotite . . . . . . . . . . . . . . . . . . . . . . . . . . . .

162

Petrography of samples . . . . . . . . . . . . . . . . . . . . . . . . .

169

6.3.1

Peridotite . . . . . . . . . . . . . . . . . . . . . . . . . . . .

169

6.3.2

Pyroxenite . . . . . . . . . . . . . . . . . . . . . . . . . . . .

171


177

6.4.1

Peridotite . . . . . . . . . . . . . . . . . . . . . . . . . . . .

177

6.4.2

Pyroxenite . . . . . . . . . . . . . . . . . . . . . . . . . . . .

180

P–T estimates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

184

6.5.1

Geothermobarometers . . . . . . . . . . . . . . . . . . . . .

184

6.5.2

Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

187

Isotope geochemistry . . . . . . . . . . . . . . . . . . . . . . . . . .

189

6.3

6.4

6.5

6.6

xiv

CONTENTS

6.7

6.8

6.6.1

Samples and techniques . . . . . . . . . . . . . . . . . . . .

189

6.6.2

Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

191

Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

194

6.7.1

Meso- and microstructures . . . . . . . . . . . . . . . . . . .

194

6.7.2

P–T path . . . . . . . . . . . . . . . . . . . . . . . . . . . .

200

6.7.3

Geodynamic interpretation . . . . . . . . . . . . . . . . . . .

203

Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

207

7 Synthesis

209

7.1

Deep upper mantle origin of Otrøy peridotite . . . . . . . . . . . . .

211

7.2

Lithosphere evolution . . . . . . . . . . . . . . . . . . . . . . . . . .

213

7.3

Crustal peridotite emplacement and evolution . . . . . . . . . . . .

215

References

217

A Procedures for sample preparation and data acquisition

241

A.1 Whole rock powder preparation . . . . . . . . . . . . . . . . . . . .

243

A.1.1 At Universiteit Utrecht: . . . . . . . . . . . . . . . . . . . .

243

A.1.2 At Vrije Universiteit Amsterdam: . . . . . . . . . . . . . . .

243

A.2 XRF . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

244

A.3 Solution ICP-MS . . . . . . . . . . . . . . . . . . . . . . . . . . . .

244

A.3.1 Trace and minor elements (Universiteit Utrecht) . . . . . . .

244

A.3.2 PGE (Durham University) . . . . . . . . . . . . . . . . . . .

246

A.4 LA-ICP-MS . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

247

A.5 TIMS and N–TIMS . . . . . . . . . . . . . . . . . . . . . . . . . . .

248

A.5.1 Sm–Nd isotopes (Lamont-Doherty Earth Observatory) . . .

248

A.5.2 Sm–Nd isotopes (Vrije Universiteit Amsterdam) . . . . . . .

249

A.5.3 Os isotopes (Durham University) . . . . . . . . . . . . . . .

251

A.6 EMPA . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

251

B Detailed petrographic profile

253

Acknowledgements

265

Curriculum Vitae

267

Enclosure Enclosure Enclosure Enclosure

1: 2: 3: 4:

Geological map of the Raudhaugene peridotite body Geological map of the Ugelvik peridotite body Structural map of the Raudhaugene and Ugelvik peridotite bodies Geological map of the Midsundvatnet peridotite body

Abbreviations Mineral names Modified after Kretz (1983) and Spear (1993): Acm Adr Alm Amp Brc Bt Chl Chr Chu Coe Cpx Crn Di Dia En Fa Fo Gau

acmite andradite almandine amphibole brucite biotite chlorite chromite clinohumite coesite clinopyroxene corundum diopside diamond enstatite fayalite forsterite gaukinite (Grs+Adr+Uva)

Gr Grs Grt Hd Hbl Hem Hu Hyp Ilm Jd Kfs Ky Mag Mc Mgs Mj Mnz Ms Ol

Olg graphite Omp grossularite Opx garnet Prg hedenbergite Pyx hornblende Pl hematite Prp humite Qtz hyhersthene Rt ilmenite Sps jadeite Spl K-feldspar Spn kyanite Srp magnetite Tep microcline Ts magnesite Uva majorite Wd monazite Zrn muscovite olivine (α-olivine)

oligoclase omphacite orthopyroxene pargasite pyroxene plagioclase pyrope quartz rutile spessartine spinel sphene (titanite) serpentine tephroite tschermakite uvarovite wadsleyite (β-olivine) zircon

Chemical parameters and system abbreviations CHUR CMAS Cr# CrMAS IPGE LREE, MREE, HREE MAS Mg# PGE PM PPGE REE

chondritic uniform reservior CaO–MgO–Al2 O3 –SiO2 Cr/(Cr+Al) Cr2 O3 –MgO–Al2 O3 –SiO2 Ir-group PGE light, medium and heavy REE respectively MgO–Al2 O3 –SiO2 Mg/(Mg+Fe) Pt-group elements primitive mantle Pd-group PGE rare Earth elements

xvi

ABBREVIATIONS

Physical and chemical symbols C L CNi D Ol/L DNi F L LP, HP, UHP LT, HT P R t T TCHUR TMA TRD X γOs εNd εNd(i) λ

concentration concentration of Ni in L distribution coefficient distribution coefficient for Ni in Ol relative to L melt fraction melt (liquid) low, high and ultra-high pressure respectively low and high temperature respectively pressure melt residue time temperature Sm–Nd model age Re–Os model age Re depletion age mole fraction Os isotope record of a sample relative to CHUR Nd isotope record of a sample relative to CHUR εNd recalculated for the time i before present constant of decay

Technical acronyms BSE EBSD EDS EMP EMPA ICP-MS LA-ICP-MS N-TIMS PPL SEM SIMS TIMS WDS XPL XRF

backscatter electron electron backscatter diffraction energy dispersive spectrometry electron microprobe electron microprobe analysis inductively coupled plasma mass spectrometry laser-ablation ICP-MS negative TIMS plane polarized light scanning electron microscopy secondary ion mass spectrometry thermal ionization mass spectrometry wavelength dispersive spectrometry cross polarized light X-ray fluorescence spectrometry

Chapter 1 Introduction

1.1 PhD project

1.1 1.1.1

3

PhD project Thesis introduction

The evolution of the Earth from its accretion >4.5 Ga ago to its present highly differentiated structure is recorded in the compositional diversity of rocks formed through time. All rock forming processes from the Archaean (4.0–2.5 Ga) to the present are influenced by the long term major change in the planets heat budget (cooling), because the Earth constantly ‘loses’ energy in the form of radiation and friction. A cooled Earth may mark an end to chemical differentiation processes and plate tectonics (Takahashi, 1990). Cooling may also have influenced the formation of the first crust and the beginning of plate tectonics in the early Earth’s history. To understand early processes it is reasonable to assume that present day rock forming processes cannot directly be interpolated to the past, when the Earth was hotter than today. Early processes can be studied on the oldest continents (cratons), which preserve the oldest sediments and volcanic rocks (continental crust) and corresponding melt residua (sub-continental lithospheric mantle, SCLM). Archaean SCLM is more buoyant and may be rheologically stronger than younger SCLM such that the oldest SCLM stabilised the overlying crust for billions of years. Contrasting models have been proposed to explain the formation of the oldest SCLM. These models include extensive LP melting (≤3 GPa) at plate margin settings (Bernstein et al., 1998; Stachel et al., 1998) followed by tectonic accretion (‘lithospheric stacking’; Helmstaedt and Schulze (1989); De Wit et al. (1992)), extensive HP melting (>4 GPa) within plate settings (Boyd, 1989; Griffin et al., 2003; Walter, 2003) and underplating of residua associated with mantle plume events (Griffin et al., 1999a). Most models for SCLM formation are based on suites of mantle xenocrysts and xenoliths, which occur in kimberlite and alkali-basalt magma derived from different SCLM depth levels. Mantle xenocrysts and xenoliths are relatively small (several mm to several dm) in size, are spatially separated from the host rock environment and are often strongly chemically overprinted by the transporting magma during the eruption. An alternative way to study the lithospheric mantle is provided by direct sampling of abyssal peridotites and mantle sections of ophiolites. Both types represent oceanic lithospheric mantle. Abyssal peridotites occur with formation ages, which do not exceed the early Mesozoic (8 km s−1 . The lithosphere, being the stronger outer layer of the Earth, reacts to stresses as a solid and varies in thickness depending on the geotherm between 50 and 150 to 200 km (Wilson, 1989). The lower boundary of the lithosphere is marked by an abrupt decrease in S-wave velocity. Underneath follows a seismic low velocity layer thought to be caused by the presence of partial melts. This layer is called the asthenosphere (Fig. 1.1) and due to its ductility it is regarded to provide solid state convective motions necessary for lithospheric plate movements. The thickness of this layer varies depending on P, T, melting point and

1.2 Upper mantle and upper mantle peridotites

7

Figure 1.1: Subdivisions of the upper Earth based on seismic data (modified after Wilson (1989)).

the availability of H2 O (Winter, 2001). Two further seismic discontinuities mark the Transition Zone between upper and lower mantle: at 410 km (the phase transition for (Mg,Fe)SiO4 from Ol to Wd) and at 660 km (the change of the crystallographic coordination of Si in silicates from tetrahedral to octahedral). The lower boundary of the ductile asthenosphere to the more rigid mesosphere is poorly constrained but among geophysists believed to occur at around 700 km (Kearey and Vine, 1990).

1.2.2

Lithospheric mantle

Most compositional information of the upper mantle is derived directly from mantle xenoliths and mantle fragments of orogenic, abyssal and ophiolitic peridotites. Geochemical differentiation during decompression melting of rising mantle within mantle convection cells is thought to be the process that forms the lithosphere: The crust is enriched in elements which have a high affinity for silicate melts (K, Na, Ca, Al, Si), while the underlying uppermost mantle is depleted in those elements and residually enriched in the refractories (Mg, Cr) – compared to fertile PM compositions (Fig. 1.2). The typical lithospheric mantle mineral assemblage is known to be a variable combination predominantly of Ol ((Mg, Fe)2 SiO4 ), Opx ((Mg, Fe)2 Si2 O6 ), Cpx ((Ca, Mg, Fe2+ , Nan , Aln )2 (Si2−n , Aln )O6 with n≤1) and an Al bearing phase, which changes with increasing P from Pl (NaAlSi3 O8 − CaAl2 Si2 O8 ) to Spl ((Mg, Fe)1 (Al, Fe, Cr)2 O4 ) to Grt ((Mg, Fe, Ca, Mn)3 (Al, Cr)2 Si3 O12 ). Each of these minerals forms solid solutions and variable amounts of them characterise most of the rock types known to occur in the lithospheric mantle (Fig. 1.2). Genetically,

8

1. INTRODUCTION

Figure 1.2: Classification of ultramafic rocks based on mineral modes in the ternary system Ol–Opx–Cpx. Thick arrow indicates the compositional change of the residue with increasing melt extraction starting from PM.

the lithospheric mantle can be subdivided into sub-oceanic lithospheric mantle (SOLM) and old and young sub-continental lithospheric mantle (SCLM). Old, Archaean SCLM occurs below the oldest continents (cratons) and reaches ages of up to 3.8 Ga (Hanghøy et al., 2001; Bennet et al., 2002). Archaean SCLM has been regarded to represent the refractory residue of ancient HT differentiation processes, including komatiite extraction, leading to highly depleted residues after >40 % melt extraction (Boyd, 1989; Herzberg, 1999; Walter, 1999). The density of the lithospheric mantle decreases with increasing melt depletion and stabilises old SCLM on the fertile underlying asthenosphere (Poudjom Djomani et al., 2001). Archaean SCLM can be described as relatively cold, light and dry, has a high viscosity and is characterised by high seismic velocities. The lithospere below cratons may reach thicknesses of 250 km and generally decrease in the degree of melt depletion with increasing depth (Griffin et al., 1999b; Pearson and Nowell, 2002). Young (post-Archaean) SCLM is thinner, hotter and less melt depleted compared to Archaean SCLM (Su et al., 1994; Griffin and Ryan, 1995). The degree of depletion decreases with decreasing age: early Proterozoic SCLM experienced 3.19 Si cations per formula unit at 12 O), indicating this Grt was stable at minimum (T dependent) confining P of ≥11.5 GPa (Fei and Bertka, 1999; Gasparik, 2003) corresponding to ≥350 km (Fig. 3).

Grt nodules in harzburgite and dunite also constrain the melting history. Nodule DS0213 (∼2 cm) has Grt with mm-scale Opx inclusions. Grt in nodule DS0289 (∼5 cm) contains relicts of intracrystalline Pyx needles, comparable to the microstructures in the larger garnetite samples. The two Grt’s are similar in colour and major element chemistry to Grt from the studied garnetites (Table 3.1) except for higher Cr, lower Al. The nodular Grt’s are characterised by extremely depleted MREE and exceptionally strongly fractionated HREE with (Dy/Yb)N 1300 ◦ C). We therefore disagree with previous suggestions that the Caledonian continental subduction was responsible for all decompression textures in the Grt-peridotites (Brueckner et al., 2002). Consequently we favour a multistage exhumation history (Van Roermund and Drury, 1998; Green et al., 2000). It follows that the occurrence of UHP minerals in orogenic peridotites does not necessarily record conditions of deep continental subduction as recently argued (Ye et al., 2000; Brueckner et al., 2002; Song et al., 2004). The initial Nd isotope ratio of the mineral isochron, 0.5476 (Table 3.3, Fig. 3.3) is extremely radiogenic compared to mantle compositions at 1.4 Ga (ǫNd(i) =+719). Nd

71

3.6 Discussion and conclusions Sample Mineral

DS0297 grt(core)†

Sm Nd 147 Sm/144 Nd 143

Nd/144 Nd Age

DS0298 grt(core)†

DS0298 cpx(inter.)†

DS0298 opx(inter.)‡

DS0298 grt(maj.)‡

0.1214 ± 0.0007 0.0844 ± 0.0014 0.2173 ± 0.0021

0.0208 ± 0.0002 0.0091 ± 0.00015 0.1079 ± 0.0010 3.685 5.84 1.269

0.57940 ± 7 TCHUR =2.90 Ga

0.60151 ± 37 0.55928 ± 4 TIsochron =1.405 ± 0.013 Ga

2.337

5.138

0.56934

0.59503 TCHUR =2.53 Ga

Table 3.3: Mineral Sm–Nd isotope data for Grt(measured), Cpx (measured), Opx (recalculated Grt/Opx

Grt/Opx

based on DSm =100 and DNd =40 for HT xenoliths from Simon (2004)) and Mj (recalculated from the modality of the exsolved phases in combination with partition coefficients for Sm and Nd). Concentrations in ppm, s.d. of 2 σ († TIMS, ‡ calculated). A two point isochron indicates a mid-Proterozoic equilibration of the intercrystalline microstructure, but all exsolved phases bear extremely high radiogenic Nd concentrations. The Nd model age (TCHUR ) for Ol-bearing garnetite (which has only one significant REE bearing phase, Grt) is bases on the isotopic composition of Grt and that for the Ol-absent garnetite on the reconstructed Mj composition. Both model ages indicate an Archaean formation of the Mj precursors. Assuming that the Mj precursors of the garnetites were in equilibrum with the depleted host peridotite then nearly all REE of the rock is concentrated within the garnetites implying an Archaean peridotite origin.

model ages (TCHUR ) are 2.9 Ga and 2.5 Ga for Ol-bearing and Ol-absent garnetite, respectively (Table 3.3). These model ages indicate that the highly depleted Grtperidotites were formed in the Archaean. Archaean melting of the Grt-peridotites is also supported by a whole rock Re–Os model age (3.3 Ga) from Otrøy (Beyer et al., 2004). The Otrøy peridotites therefore appear to be the first reported case of orogenic peridotites produced by extensive melt depletion entirely in the Grtperidotite stability field during the Archaean.

3.6

Discussion and conclusions

There are two geodynamic models for SCLM formation that potentially explain the occurrence of peridotites derived from the Transition Zone at Otrøy. The simplest option (model 1, path black–grey in Fig. 3.4) is that Archaean adiabatic upwelling of mantle from the Transition Zone led to extensive melting and Mj exsolution. Melting stopped when the peridotites were underplated and accreted beneath an Archaean craton at depths of ∼150 km. Over the next >1.0 Ga the new SCLM cooled slowly but remained above the Nd closure T for Cpx (∼1100 ◦ C (Van Orman et al., 2002)). Cooling below 1100 ◦C occurred at 1.4 Ga, perhaps related to the Svecofennian or Gothian Orogeny. The peridotites remained in the SCLM until exhumation 1.0 Ga later during the continent–continent collision of the Caledonian Orogeny. A more complex alternative applies if the Mj formation postdates the melting (model 2, path black–grey hatched–grey in Fig. 3.4). After hot upwelling, deep melting and accretion in the Archaean, part of the lower lithosphere was subducted, or delaminated, sinking to the Transition Zone accompanied by Mj formation. The

72

3. DEEP ORIGIN AND HOT MELTING

Figure 3.3: Sm–Nd isochron diagram for Grt and intercrystalline Cpx from Ol-absent garnetite (DS0298) yielding an age of 1.405 Ga for the equilibration of the intercrystalline microstructure. The isotopic composition of Opx is reconstructed on the basis of D Grt/Opx for HT xenoliths (Simon, 2004) at the time of mineral equilibration (1.405 Ga), when the isotopic exchange between Cpx and Grt effectively stopped (Table 3.3). The isotope composition of the Mj (whole rock) was recalculated based on the observed modes and isotope ratios of each exsolved phase. Sm–Nd diffusion effectively stops in a garnetite when Cpx becomes a closed system (Van Orman et al., 2002). Estimated closure T for Nd in a 2 mm Cpx at HP are between 1050–1250 °C for low and fast cooling rates (Van Orman et al., 2002), but the initial T of the exsolved garnetites was significantly higher as indicated by the Ce partitioning (Fig. 3.2(c)) and by the implied melting conditions. Consequently, the exsolution occurred above the Sm–Nd closure T and the 1.4 Ga isochron age records cooling to and probably of the geotherm after the emplacement of the peridotites into the SCLM.

low Fe peridotite residua subsequently underwent thermal equilibration and became buoyant (Ringwood, 1994), rising in a subsolidus upwelling to underplate existing SCLM in the Mid-Proterozoic accompanied by Mj exsolution. This decompression could not involve partial melting, which would disrupt the Archaean Nd–Os isotope systematics. Important implications of both models are that: a) UHP minerals in orogenic peridotites may be relicts that are not evidence for deep continental subduction and b) deep asthenospheric underplating could form highly depleted SCLM with substantial amounts of Grt in the residue. We prefer the simplicity of model 1, but the more complex history of model 2 can not be ruled out and indeed has been reproduced in numerical studies of early mantle convection (De Smet et al., 2000). Model 2 implies that cratonic lithosphere retains the ancient isotopic signature of lithosphere that has been recycled back into the convecting mantle. A fundamental characteristic of the Otrøy peridotites is that they preserve high Mg/Si, low Ca/Al (almost no Cpx but 5–20 % Grt) and relative enriched HREE. These are all characteristics inferred for the source of Al-depleted komatiites (Herzberg, 2004a).

3.6 Discussion and conclusions

73

Figure 3.4: P–T diagram that schematically portrays the potential evolution of the Otrøy peridotites. Model 1 (black–grey path) requires near isothermal decompression from depths where the mantle contains >19 % Mj in Grt (A). Extensive polybaric melting (>30 %) in the Archaean continued until the Grt-bearing rocks reached the base of the overlying lithosphere (∼150 km, B). Residua were stored and cooled towards the continental geotherm in the Proterozoic. Further decompression associated with the Svecofennian or Gothian Orogeny cooled the peridotites preserving the 1.4 Ga Nd mineral cooling age (C). An alternative model 2 (black–grey hatched– grey path) is that following Archaean melting, residual peridotites are subducted to at least the mantle Transition Zone and stored until thermal equilibration resulted in buoyancy driven upwelling. Subsolidus upwelling took the peridotites to the base of the lithosphere (B–A–C) where cooling produced the 1.4 Ga Nd mineral cooling age (C). In both models further cooling until the Phanerozoic resulted in P–T conditions in equilibrium with the present continental geotherm (D). Final peridotite exhumation to the surface occured during the Caledonian continent–continent collision (E) (Terry et al., 2000b). Cross-hatched squares represent equilibrated P–T conditions for Grt-websterite lenses (bright) at Ugelvik (Otrøy) and Bardane (Fjørtoft) (Brueckner et al., 2002; Van Roermund et al., 2002) and garnetites (grey, this study). Grey lines illustrate molar Mj components in solid solution with Grt (Gasparik, 2003).

Chapter 4 Refractory history of Otrøy Grt-peridotite

4.1 Introduction

4.1

77

Introduction

Orogenic Mg–Cr type Grt-peridotite has been considered as one of the few potential source rocks for the study of fundamentally different mantle processes, which are related to mantle convection (asthenospheric upwelling), to the stabilization and evolution of sub-continental lithospheric mantle (SCLM; craton formation) and to subduction related metamorphism of SCLM during orogeny. Studies on such mantle fragments over the last decade led to contrasting interpretations concerning the derivation and equilibration of orogenic Grt-peridotite (Dobrzhinetskaya et al., 1996; Green et al., 1997; Van Roermund and Drury, 1998; Liou, 1999; Brueckner and Medaris, 2000; Nimis and Trommsdorff, 2001; Brueckner et al., 2002; Bozhilov et al., 2003; Song et al., 2004; Lapen et al., 2005; Spengler et al., 2006). Interpretations are not always unequivocal, because Mg–Cr peridotite has superimposed records of different processes, which culminated for orogenic peridotite always in a subduction setting. The relatively robust behavior of Grt against deformation (Karato et al., 1995) and chemical re-equilibration (Griffin et al., 1999c) forms a potential tool to unravel complex mantle peridotite histories, provided that the origin of Grt is known. The aim of this chapter is to study Grt-peridotite from Otrøy in order to determine, if the origin of Grt is primary or secondary. Global studies on orogenic Grt-peridotite has shown that the interpretation of the Grt origin is not straight forward due to its involvement in contrasting, spatially closely associated assemblages. Orogenic peridotite, that equilibrated either in the Grt-, Spl- or Pl-peridotite stability field (O’Hara, 1967), typically shows a compositional layering with steep chemical gradients between the most refractory and more fertile peridotite layers (Avé Lallemant, 1967; Dickey, 1970; Niida, 1974; Bodinier and Godard, 2003). The classical model to explain this layering involves variable degrees of melt extraction (from ∼0 % in fertile lherzolite to >25 % in the most refractory peridotite) during adiabatic partial melting of the asthenospheric mantle (Frey et al., 1985; Bodinier et al., 1988). This model implies a refractory origin of Grt in Grt-peridotite, but requires an additional process of deformation to explain steep small-scale chemical gradients between closely intercalated, differently melt depleted rock types. An alternative model proposes homogeneous melting followed by localised refertilization as a result of the solidification of infiltrated melt. In this scheme, anomalous HT at shallow depth (stability field of Spl- and Pl-peridotite) may explain the formation of most refractory (dunitic) compositions implying a secondary origin of intercalated Grt bearing assemblages. This more complex model has been questioned by inconsistencies derived from melt models and replacive relationships of dunite (Nicolas and Prinzhofer, 1983; Kelemen et al., 1995; Suhr, 1999). Another model suggests the formation of layered refractory peridotite by Ol and Opx forming melt-rock-reactions in ‘porous-flow’ channels (Kelemen, 1990; Kelemen et al., 1992, 1995; Suhr, 1999) and has been recognised as the dominant mechanism for the generation of mantle dunite and some harzburgites. However, the occurrence of extremely depleted but Grt-bearing harzburgite closely associated with Grt-free

78

4. REFRACTORY HISTORY

dunite seems hardly consistent with a single interaction process between magma and mantle as it would require the generation and reactive trapping of the melt in the Grt-peridotite stability field leaving primary Grt in the residue and simultaneously removing Grt to form Grt-free dunite layers. Instead, a polystage model, which involves the juxtaposition of relatively fertile rocks and most depleted peridotite (by a combination of the models above or according to the ‘marble-cake’ model, Allègre and Turcotte (1986)), appears more powerful to explain the close association of different assemblages in orogenic Grtperidotite with secondary Grt. Nevertheless, for lithological juxtaposition to have occurred in the SCLM significant age differences may be required between different end-members, especially Grt-bearing and Grt-lacking lithologies. In case of the WGR mantle peridotite, this is not supported by recent isotope studies showing an Archaean origin for Spl-peridotite, Grt-peridotite and garnetite (Beyer et al., 2004; Lapen et al., 2005,; Chapter 3). Alternatively, a juxtaposition of highly depleted and less depleted peridotite can be envisaged during subsolidus mantle upwelling. This requires the downwelling of highly melt depleted (light) sub-oceanic peridotite beforehand, because delamination and downwelling of sub-continental peridotite has been found to be unlikely for the lightest and oldest mantle peridotite on Earth (Poudjom Djomani et al., 2001), comparable to that on Otrøy. This chapter demonstrates the refractory (primary) origin of Grt in the Ortøy peridotite. Models involving a non-refractory origin for Grt will be discussed. It will be shown that an extreme melt extraction event in the Grt-peridotite stability field was followed by mechanical re-distribution of Grt during subsequent deformation. This new model suggests the present compositional peridotite banding of relative fertile and depleted layers to have virtually formed in response of melting, exsolution and deformation without the necessity of secondary chemical interaction.

4.2

Petrography and samples

The compositional layering of the Otrøy peridotite bodies at Ugelvik and Raudhaugene is defined by variable mineral modes of Ol, Grt, Opx, Cpx, Spl and ±Amp (Carswell, 1968; Van Roermund et al., 2000b; Drury et al., 2001). Different layers show sharp contact relationships. Individual layers range in thickness between several mm to a few m, but are dominantly several cm to tens of dm wide. An example is given in form of a detailed petrographic profile in Appendix B. The latter demonstrates a total compositional peridotite variation across 32 m layering of Ol76−97 Grt0−20 Spl0−11 Cpx0−5 Opx0−5 . Some other layers contain higher modes of Opx. Layers with significantly more than 5 % Cpx have not been observed. Minor mafic rock types (Grt-websterite, Grt-clinopyroxenite, Grt-orthopyroxenite) occur layer-parallel and range in size between a few mm to several cm. Mafic layers in dm size occur exceptionally rare. Disregarding the mineralogical layering, the peridotites exposed at both bodies can be subdevided by the weathering colour, which has the end-members yellow

4.2 Petrography and samples 79

Figure 4.1: Simplified lithological map of the Ugelvik peridotite body (central part) showing a compositional layering, sample numbers and locations.

80


and brown. Transitions between yellow and brown occur. The yellow group dominantes and forms the focus of this chapter. The brown group is subparallel to the mineralogical layering and records a metasomatic overprint (Chapter 5). Grt-bearing and Grt-free harzburgite and dunite are the dominating rock types in the yellow weathering peridotites on Otrøy. This characterises the Otrøy peridotites as more melt depleted than other orogenic Grt-peridotites world wide, which are dominantly lherzolitic in composition (Bodinier and Godard, 2003). The rock nomenclature below follows the typical petrographic classification criterions with ‘harzburgite’ defined as an ultramafic rock with less than 5 % Cpx and ‘Cpx-Grt-harzburgite’ as a Cpx-bearing Grt-harzburgite. An exception is formed by the compositional fields in Fig. 4.7(a) with references given in the caption.

4.2.1

Single and polycrystalline Grt – in harzburgite

The occurrence of Grt in substantial amounts is one of the most striking features in the Ugelvik and Raudhaugene peridotite bodies. Two petrographic subtypes of Grt can be distinguished in yellow weathering harzburgite: single crystal Grt and polycrystalline Grt. Single crystal Grt is the dominating subtype and has crystal sizes between a few mm and less than 2 cm (Fig. 4.2(d)). Polycrystalline Grt occurs subordinately and is composed of mm–cm scale Grt crystals with minor Opx, Cpx, ±Ol (crystal sizes of ≤5 mm) either as inclusions in Grt or as interstitials in between Grt grains (Fig. 4.2(a)–(c) and (f)–(g)). The polycrystalline Grt occurs as isolated and elongated lenses (Fig. 4.2(a)–(c)). The term nodular Grt is used for polycrystalline Grt of ≤5 cm. Larger polycrystalline Grt is named garnetite having sizes of ≤25 cm (Fig. 4.2(a)). Both single and polycrystalline Grt are porphyroclastic and aligned parallel to the compositional layering (Fig. 4.2(c)). These Grt’s in the yellow weathering peridotite are purple in colour. Some Grt cores preserve the intracrystalline Pyx microstructure in form of crystallographically oriented Pyx needles (Chapter 2). Garnetite contains Grt with well preserved Pyx needles in Grt cores. This Grt is bifringing and has minor undulose extinction in thin sections under XPL, but is not dynamically recrystallized −→

Figure 4.2: Optical light micrographs of mineral textures (a)–(d) and microstructures (e)– (h) in peridotite, showing evidence for mechanical re-distribution of Grt by deformation. Early mineral microstructures in porphyroclasts survived the deformation. (a) Elongated garnetite lens (dm-scale) in Grt-peridotite. (b) Elongated and recrystallized Grt porphyroclast. (c) Cm- and mm-scale Grt porphyroclasts aligned parallel to the compositional layering. (d) Thick section of porphyroclastic Grt with relict amoeboid grain shapes in peridotite (PPL). (e) Thin section of Grtdunite showing undeformed amoeboid shaped Grt (with mm-scale Opx-inclusions; PPL). (f) Thin section of non-recrystallized Grt in garnetite showing textural equilibrium between interstitial Opx and single Grt (with oriented Pyx needles, PPL, reproduced Fig. 2.1(a)). (g) Thin section showing a detail of a deformed Grt porphyroclast (cm-scale) preserving an Opx inclusion (mm-scale; PPL). (h) Thin section showing a detail of a deformed Grt porphyroclast (cm-scale), which contains relics of crystallographically oriented Pyx needles within a distorted crystal lattice (almost XPL, reproduced Fig. 2.2(a)). Coin diameter 21 mm.

81

4.2 Petrography and samples

(a)

(b)

(c)

(d) Sample DS0278

(e) Sample DS0258

(g) Sample DS0213

(f) Sample DS0297

(h) Sample DS0289

82


(Fig. 4.2(f)). Nodular and some single crystal cm-scale Grt may preserve relicts of Pyx needles in the cores, which may show an anisotropy with a strong undulose (wavy) extinction in thin sections under XPL (Fig. 4.2(h)). Most sub-cm-scale Grt lack Pyx needles in Grt cores. Grt grains