14- obduction/kontinental kruste - dynamic earth · obduction (~94 ma) formation de la semelle...
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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
Robert & Bousquet, 2010
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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)
Robert & Bousquet, 2010
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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
Robert & Bousquet, 2010
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Wilson Cycle
Subduction Collision
Rifting Ocean formation
Robert & Bousquet, 2010
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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
Subduction Collision
Rifting Ocean formation
Robert & Bousquet, 2010
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Wilson Cycle: Volcanism
Subduction Collision
Rifting Ocean formation
Alcalin MORB=Tholeiite
Calco-alcalin
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Wilson Cycle: Sediment
Subduction Collision
Rifting Ocean formation
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 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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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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Age of the continental crust
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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Age of the continental crust
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)