14- obduction/kontinental kruste - dynamic earth · obduction (~94 ma) formation de la semelle...

14-Obduction/Kontinental Kruste

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Wilson Cycle

Subduction Collision

Rifting Ocean formation

Obduction?

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Oman, the best obducted ophiolite

Sultanatd’Oman

ÉmiratsArabesUnis

ArabieSaoudite

Champs pétrolifères

Golfe d’Oman

IranDétroit d’Hormuz

Muscat

JebelAkhdar

SaihHatat

Hawasina

Ibri

Fahut

Huqf

Haushi

Masirah

Sumeini

Dibba

Musandam

A

A’

23° N

22° N

21° N

20° N

24° N

59° E 60° E58° E57° E56° E55° E

Sédiments autochtones pré-Permien

Sédiments autochtones Permien-Crétacé

Sédiments allochtones (Hawasina-Sumeini)

Sédiments Maastrichien-Éocène

Ophiolites : roches mantelliquesOphiolites : roches crustales

Mar

ge a

rabe

Néo

téth

ys

Moho

Semelle métamorphique

Sédiments post-Éocène

Robert & Bousquet, 2010

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Oman, the best obducted ophiolite

Nappes d’Hawasina Ophiolite du Semail Fenêtre duSaih Hatat Muscat Golfe

d’Oman

0

10

200 50 100 150

km

km

Croûte continentale arabe

Roches sédimentaires

Protérozoïque supérieur-Paléozoïque

Roches sédimentaires

Permien-Crétacé supérieur

Sédiments océaniques

(nappes de l’Hawasina)

Croûte (Gabbros, Basaltes)

Paléo-Moho

Manteau (Hazburgites)Tertiaire discordant

Ophiolite de SemailPlateforme arabe

SSW NNE

As Sifah


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Sapphirine-Spinelle-Enstatite (orthopyroxene)

Garnet- brown HornblendeGreen Hornblende-Augite (clinopyroxene)

Green

Feldspath-Green Hornblende

Metacherts withChlorite-Micas

non metamorphicsediments

Metacherts with Epidote

Granulites

AmphibolitesGreenschist-Amphibolite

Transition(GAT)

Greenschists Decre

asing

Temp

ératu

re

at co

nstan

t pre

ssur

e (0,8

- 1 G

Pa)

late Granites(94 Ma)

late Gabbros(94 Ma)

Metamorphic sole

300 m

deformation associatedto the obduction and

the metamorphism

déformation associatedto the magmatic emplacement

Peridotites

Arabic plateformsediments

Pillows lavas

Sheeted dykes

gabbros isotropes

upper gabbros (foliated)

lower gabbros(layered)with intruded wehrlite

Transition zone MOHO

radiolaritesOceanic magmatism (97-95 Ma)


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Formation de la Néotéthys (->95 Ma)océanisationPlateforme arabe Eurasie

Obduction (~94 Ma)formation de la semelle métamorphique

métamorphisme de HTPlateforme arabe Eurasie

Subduction (~78 Ma)formation des éclogites et des schistes bleus

métamorphisme HP-BT

Plateforme arabe EurasieOphiolite du Sémail

Semelle métamorphique

Hawasina

Éclogites d’As Sifah

Schistes bleusd’As Sifah


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Wilson Cycle




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Wilson Cycle: Metamorphism

Blueschist Granite

Amphibolite

Eclogite

Eclogite

Grünschief

er

Niedrig

Grad

e

Greenschist

Greenschist

HP

Greenschist

Amphibolite

Low grade

Metamorphism

UHP

HP

Granulite

Heating of the crust

(Amphibolite, LP Granulite)

Diagenesis




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Wilson Cycle: Volcanism



Alcalin MORB=Tholeiite

Calco-alcalin

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Wilson Cycle: Sediment



Evaporites, conglomerat Radiolarite, carbonates

Flysch, turbidites Molasse

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Earth heat flow balance

Magnetic !eld

Cristallization heatRadioactivity

Radioactivity

Initial heat

Initial heat

Di"erential motion (shear)

Radioactivity

Radioactivity

Conduction

Convection

Tectonicprocesses

Cooling of the oceanic lithosphere

MagmatismMeteorites impact Magmatism

Lithosphere

Mantle

Core

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Oceanic crust ages

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Continental crust ages

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Oceanic lithosphere structure

Parsons & Sclater (1977) Stein & Stein (1992)

Age (Ma)

Depth (km)

50 100 15007

6

5

4

3 Half-space model

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Continental lithosphere structure

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Thermal evolution of the continental lithosphere

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Oceanic vs continental crust

Moho

Dep

th (k

m)

4 6 80 20

5

10

15

ocean bottom

crust ~7 km thick

P-wave velocity (km s )– 1

Moho

Conrad

Dep

th (k

m)

4 6 80

10

20

30laminations

top of basement

P-wave velocity (km s )– 1

sea water

oceanic sediments

basalt

gabbro

ultramafics

Layer 1

Layer 2

Layer 3

upper mantle

ocean

upper crystalline basement

Cenozoic sediments

Mesozoic & Paleozoicsediments

granitic layer

Gneiss

amphibolites

granulites

ultramafics

near-surfacelow-velocity layer

sialic low-velocity layer

zone of positivevelocity gradient

middle crustal layer

high-velocity tooth

lower crustal layer

uppermost mantle

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Example of continental crust

Modeling results along the seismic refraction/wide-angle reflection profile at the northern margin of the Tibetan Plateau. Crustal and uppermost mantle seismic velocity (Vp) model. Both Vp and Vs can be estimated for the crust from these data. Pn velocity ranges from 7.8 to 8.2kms-1.

Crustal structure across the Altyn Tagh Rangeat the northern margin of the Tibetan Plateau

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Different types of continental crust

Mooney et al., 1998

Crust 5.1 model

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Chemical composition of the continental crust

Fe

Oxygen

Si

Mg

Ni

S

Ca

Al

Other

Na

K

Whole Earth Earth´s crust

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Chemical composition of the continental crust

364 CHAPTER 11

(Condie, 2000). On the basis of available data, Condie (2005b) concluded that 39% of the continental crust formed in the Archean, 31% in the Early Proterozoic, 12% in the Middle–Late Proterozoic, and 18% in the Phanerozoic.

Two of the most important mechanisms of Late Archean and Early Proterozoic continental growth and cratonic root evolution were magma addition (Section 9.8) and terrane accretion (Section 10.6). Several authors (e.g. Condie, 1998; Wyman & Kerrich, 2002) have sug-gested that the ascent of buoyant mafi c material in mantle plumes may have initiated crust formation and may have either initiated or modifi ed root formation during periods of rapid net growth (Section 11.3.3). Schmitz et al. (2004) linked the formation and stabiliza-tion of the Archean Kaapvaal craton in South Africa to subduction, arc magmatism, and terrane accretion at 2.9 Ga. In this and most of the other cratons, isotopic ages from mantle xenoliths and various crustal assem-blages indicate that chemical depletion in the mantle lithosphere was coupled to accretionary processes in the overlying crust (Pearson et al., 2002). This broad correspondence is strong evidence that the crust and the underlying lithospheric mantle formed more or less contemporaneously and have remained mechanically coupled since at least the Late Archean. A progressive decrease in the degree of depletion in the lithospheric mantle since the Archean (Fig. 11.14) indicates that the Archean–Proterozoic boundary represents a major shift in the nature of lithosphere-forming processes, with more gradual changes occurring during the Phanero-zoic (O’Reilly et al., 2001). The most obvious driving mechanism of this change is the secular cooling of the Earth (Section 11.2). In addition, processes related to subduction, collision, terrane accretion, and magma addition also helped to form and stabilize the con-tinental lithosphere.

Whereas these and many other investigations have identifi ed some of the processes that contributed to the formation and stabilization of Archean cratons, numer-ous questions still remain. Reconciling the composition of craton roots determined from petrologic studies with the results of seismic velocity studies is problem-atic (King, 2005). There are many uncertainties about how stability can be achieved for billions of years without suffering mechanical erosion and recycling in the presence of subduction and mantle convection. Another problem is that the strength of mantle materi-als required to stabilize craton roots in numerical exper-iments exceeds the strength estimates of these materials

derived from laboratory measurements (Lenardic et al., 2003). These issues, and the extent to which the cratonic mantle interacts with and infl uences the pattern of mantle convection, presently are unsolved. Improved resolution of the structure, age and geochemical evolu-tion of the continental crust and lithospheric mantle promise to help geoscientists resolve these problems in the future.

11.4.3 Proterozoic plate tectonics

Early tectonic models of Proterozoic lithosphere envis-aged that the Archean cratons were subdivided by mobile belts in which deformation was wholly ensialic, with no rock associations that could be equated with

0

1

2

3

4

0 1 2 3% Al2O3

ARCHEAN

Xenolithaverages

ArcheanProterozoic

S.AustraliaKaapvaal

EuropeanMassifs

PhanerozoicSpinelPeridotite

% C

aO

PROTEROZOIC

PHANEROZOIC

Zabargad

PRIMITIVE MANTLE(=ASTHENOSPHERE)

E. ChinaVitim

YOUNGER

Phanerozoic

Gnt-SCLM

4 5

+

+++

+

+

+

+

Fig. 11.14 CaO-Al2O3 plot showing the range of subcontinental lithospheric mantle (SCLM) compositions for selected cratons that have been matched with ages of the youngest tectonothermal events in the overlying crust (after O’Reilly et al., 2001, with permission from the Geological Society of America). Compositions have been calculated from garnet xenocrysts (Gnt). Xenolith averages shown for comparison. Plot shows that newly formed subcontinental lithospheric mantle has become progressively less depleted in Al and Ca contents from Archean through Proterozoic and Phanerozoic time. Garnet peridotite xenoliths from young extensional areas (e.g. eastern China, Vitim in the Baikal region of Russia, and Zabargad Island in the Red Sea) are geochemically similar to primitive mantle, indicating very low degrees of melt depletion.

CaO-Al2O3 plot showing the range of subcontinental lithospheric mantle (SCLM) compositions for selected cratons that have been matched with ages of the youngest tectonothermal events in the overlying crust (after O’Reilly et al., 2001

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Age of the continental crust

PRECAMBRIAN TECTONICS AND THE SUPERCONTINENT CYCLE 363

dikes that fan out over a 100° arc and extend for more than 2300 km (Ernst et al., 2001). Some of these shield regions also contain huge sills and layered intrusions of mafi c and ultramafi c rock that occupy hundreds to thousands of square kilometers. These intrusions provide information on the deep plumbing systems of Precambrian magma chambers and on crust–mantle interactions. Three of the best known examples are the !1.27 Ga Muskox intrusion in northern Canada (Le Cheminant & Heaman, 1989; Stewart & DePaolo, 1996), the !2.0 Ga Bushveld complex in South Africa (Hall, 1932; Eales & Cawthorn, 1996), and the !2.7 Ga Stillwater complex in Montana, USA (Raedeke & McCallum, 1984; McCallum, 1996). Unlike the layered igneous suites of the Archean high-grade gneiss terrains (Section 11.3.2), these intrusions are virtually undeformed.

Anorthosite massifs (Section 11.3.2) emplaced during Proterozoic times also differ from the Archean exam-ples. Proterozoic anorthosites are associated with gran-ites and contain less plagioclase than the Archean anorthosites (Wiebe, 1992). These rocks form part of an association known as anorthosite-mangerite-charnockite-granite (AMCG) suites. Charnockites are high temperature, nearly anhydrous rocks that can be of either igneous or high-grade metamorphic origin (Winter, 2001). The source of magma and the setting of the anorthosites are controversial. Most studies inter-pret them as having crystallized either from mantle-derived melts that were contaminated by continental crust (Musacchio & Mooney, 2002) or as primary melts derived from the lower continental crust (Schiellerup et al., 2000). Current evidence favors the former model. Some authors also have suggested that these rocks were emplaced in rifts or backarc environments following periods of orogenesis, others have argued that they are closely related to the orogenic process (Rivers, 1997). Their emplacement represents an important mecha-nism of Proterozoic continental growth and crustal recycling.

11.4.2 Continental growth and craton stabilization

Many of the geologic features that comprise Protero-zoic belts (Section 11.4.1) indicate that the continental lithosphere achieved widespread tectonic stability during this Eon. Tectonic stability refers to the general

resistance of the cratons to large-scale lithospheric recy-cling processes. The results of seismic and petrologic studies (Sections 11.3.1, 11.3.3) and numerical modeling (Lenardic et al., 2000; King, 2005) all suggest that com-positional buoyancy and a highly viscous cratonic mantle explain why the cratons have been preserved for billions of years. These properties, and isolation from the deeper convecting mantle, have allowed the mantle lithosphere to maintain its mechanical integrity and to resist large-scale subduction, delamination and/or erosion from below. Phanerozoic tectonic processes have resulted in some recycling of continental litho-sphere (e.g. Sections 10.2.4, 10.4.5, 10.6.2), however the scale of this process relative to the size of the cratons is small.

The cores of the fi rst continents appear to have reached a suffi cient size and thickness to resist being returned back into the mantle by subduction or delam-ination some 3 billion years ago. Collerson & Kamber (1999) used measurements of Nb/Th and Nb/U ratios to infer the net production rate of continental crust since 3.8 Ga. This method exploits differences in the behavior of these elements during the partial melting and chemical depletion of the mantle. The different ratios potentially provide information on the extent of the chemical depletion and the amount of continental crust that was present on Earth at different times. This work and the results of isotopic age determinations (Fig. 11.13) suggest that crust production was episodic with rapid net growth at 2.7, 1.9, and 1.2 Ga and slower growth afterward (Condie, 2000; Rino et al., 2004). Each of these pulses may have been short, lasting "100 Ma

Stage 3 Stage 2 Stage 1

1.2 Ga

1.9 Ga

2.7 Ga

Fre

quen

cy (

% )

15

10

5

00.2 0.6 1.0 1.4 1.8 2.2 2.6 3.0 3.4 3.8

Age ( Ga )

Fig. 11.13 Plot showing the distribution of U-Pb zircon ages in continental crust (after Condie, 1998, with permission from Elsevier).

Super continents

Plot showing the distribution of U-Pb zircon ages in continental crust (after Condie, 1998)

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Spatial distribution of the juvenile crust

About 75% of the continental crust was produced during the first two super event cycles.

Map of the continents showing the distribution of juvenile continental crust

Condie, 1998

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Age of the Earth

Age of Earth inferred to be 4.55 billion years old (Ga) based on radiometric age determination of meteorites such as Allende.

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Oldest dated mineral is zircon crystal 4.4 billion years oldfound in Western Australia

Red quartzite and conglomerate,Jack Hills, Western Australia(Peck et al., 2000, 2001)

Zircon crystal (Peck et al, 2000)

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Oldest dated rock is ~ 4 billion years old

Isua Greenstone Belt, West Greenland3.8 Ga (David Green, Denison Univ.)

Acasta Gneiss, 4.04 Ga

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Models of continental protolith

a) Hotsubduction

b) Oceanic PlateauAccretion

c) Loose-PlateLoading

melting ofeclogites?*

a) hot subduction or subduction of mid-ocean ridges and young oceanic lithosphere makes production of felsic and intermediate liquids by melting of the hydrothermally altered oceanic crust possible. Contribution from the mantle wedge, the subcontinental lithosphere, and the overlying crust may be geochemically identified. The protolith is felsic to intermediate.* Archean TTG could be produced by melting of hydrous eclogites (Rapp et al., 2003)

b) accretion of oceanic plateaus created during superplume events (plume head).

c) protracted loading by plume magma of a loose plate (not entrained by sinking lithosphere). In the last two cases, the protolith is basaltic. Crust is thickened and reaches critical buoyancy.

(after Albarède, 1998

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(a) Model 1: Segregation of residue from an upwelling mantle plume

Komatiite lavas

Residue ofhigh-degree melting

Residue oflow-degree melting

(b) Model 2: Segregation of recycled refractory residue

Oceanic plateau

Extraction andaccumulation ofrecycled refractoryresidue

Crust

Flotation and/orin-situ crystallizationof ol opx

Liquid interiorof magma ocean

(c) Model 3: Preservation of remnants of the crust of a magma ocean

Three possible mechanisms that could allow the segregation and accumulation of high-Mg olivine and orthopyroxene near the surface of the Earth in the early Earth

after Arndt et al., 2002

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(a) ModernCFB

Lithosphere

Cont. crustOIB MORB

(b) Archean

0

40

150

Dep

th (k

m)

0

40

150

Dep

th (k

m)

Archean tholeiites

Plumes

Asthenosphere

Depth of meltingCFB>OIB>MORB

Lithosphere

Cont. crust

Cont. crust

Cont. crust

MORB

Plumes

Asthenosphere

Model shows the influence of lithospheric thickness on depth of melting where CFB is continental flood basalt, OIB oceanic island basalt, and MORB mid-ocean ridge basalt

Model of komatiitic and tholeiitic basalt formation involving mantle plumes

(after Arndt et al., 1997)

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Melting temperature vs age

Mid-oceanridge

1200

1400Tem

pera

ture

(°C)

1600

1800

2000 EstimatedMantle MeltingConditions

Plumeorigin

Subductionorigin

Boninite

Maturearc

Subductionzone

Modernplume

Archeankomatiite

The range of mantle melt generation temperatures estimated for various modern tectonic settings compared to temperatures inferred for komatiite melt generation by a plume model and a subduction model

(after Grove & Parman, 2004)

14- obduction/kontinental kruste - dynamic earth · obduction (~94 ma) formation de la semelle...

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