geodynamics of sedimentary...
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Geodynamics ofsedimentary basins
Magdalena [email protected]
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who am I?
Magdalena Scheck-Wenderoth
Professor for Basin Analysis at RWTH Aachen in joint appointment with
German Research Centre for Geosciences GFZ Helmholtz Centre Potsdam
Director Department 4: GeosystemsHead Sect. 4.5: Basin Modelling
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TelegrafenbergPotsdam~1200 peopleconduct researchon Solid Earth
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Section 4.5 Basin Modelling
geodynamic processesin sedimentary basins:
quantify mass and energy fluxes
integrate observations into 3D subsurface models for key areas from regional to reservoir scale
4D process simulation of Thermal‐Hydraulic‐Mechanical processes across scales
Department 6 – Geotechnologies
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Structural analysisPotential field modellingTHM modelling
Subsidence analysisDeformation restoration
observations
Field
wellsStratigraphyLithologyTemperaturePaleo indicators
Seismic
Potential fields
Seismology
observation-based models
Interpretation &Structural modelling
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Literature –(text)books
Allen, P. and Allen, J., 2013. Basin Analysis. Principles and Applications. Wiley-Blackwell, 642 pp.
Fowler, C.M.R., 1996. The solid earth. Cambridge University Press, Cambridge, 472 pp.
Littke, R., Bayer, U., Gajewski, D. and Nelskamp, S., 2008. Dynamics of complex intracontinental basins : the Central European Basin System. Springer, Berlin.
Turcotte, D.L. and Schubert, G., 2002. Geodynamics. 2 ed. Cambridge University Press, Cambridge, United Kingdom, 456 pp.
K. Bjørlykke (ed.), Petroleum Geoscience: From Sedimentary Environments to Rock Physics, 1,DOI 10.1007/978-3-642-02332-3_1, © Springer-Verlag Berlin Heidelberg 2010
Hantschel & Kauerauf, 2009: Fundamentals of Basin and Petroleum Systems Modeling, Springer
http://www.geolsoc.org.uk/ks3/gsl/education/resources/rockcycle.html
Xie, X. and Heller, P. L. Plate tectonics and basin subsidence history. GSA Bulletin; January/February 2009; v. 121; no. 1/2; p. 55–64; doi: 10.1130/B26398.1
Literature –> +relevant articles theme-specific
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Overview
• Short Intro to Basin types• Mechanisms of basin formation• Heat transport in sedimentary basins• Geophysical characteristics of
– rift basins: – passive margins– intracontinental basins: – foreland basins: – strike-slip basins
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Sedimentary basins
• are the subsiding areas where sediments accumulate toform stratigraphic successions=> ARCHIVES
• form in response to different mechanisms in different plate-tectonic settings
• have characteristic geophysical fingerprint, subsidence andtemperature histories
• represent low pressure-high temperature reactors in whichaccumulated sediments experience changing conditions
• host georesources (water, petroleum, heat)
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Sedimentary basinswhy should we care?
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Sedimentary basinswhy should we care?
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global thickness distribution of sediments
Gabi Laske and Guy Mastershttps://igppweb.ucsd.edu/~gabi/sediment.html
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some sedimentary basins
valley Himalaya Lake Baikal
Atlantic Wannsee
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a few questions…
which sedimentary basins do you know?
which types of sedimentary basins do you know?
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a few questions…
which types of sedimentary basins do you know?
rift basin, intracontinental basin, passive continental margins, foreland basin, subduction trench, fore-arc- basin, back-arc- basin, pull-apart- basin,piggy-back - basin
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Basin Types
many classifications…
after tectonic setting:divergent=> extensional basinsconvergent => collisional basinsstrike-slip =>transtensional basins
after formation mechanismextensional basinsflexural basins
after lithosphere substrateoceaniccontinental
after petroleum prospectivity
intracontinental rift Intracontinental basin
passive margin
passive margin
ocean basin MOR
passive margin
subduction
trench foreland basin
continentalbreakup
post-rift deposits
syn-rift deposits
strike-slip basin(pull-apart basins; sheared margins)
Where do basins form?basins in their plate-tectonic setting
Wils
on
cyc
le
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Allen & Allen, 2013 17
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Basin Forming Mechanisms?
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in the end…its almost all aboutdensities & heat flow … (or their
changes)
Newton Archimedes Fourier
Some basics before looking at basins…
• Isostasy = f(load equilibria)= f(density) => controlsstresses induced by loads and therefore subsidence andrelated effects (changes in physical properties)
• Temperature => controls physical properties and rheologyof rocks
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compositional/rheological boundariesin the earth
21Allen & Allen, 2013, Fig. 1.4
Crust = sediments + upper silicic cryst. crust + lower mafic cryst. crust
upper mantle:peridotite
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The thermal lithosphere-
asthenosphereboundary
(LAB)
Allen & Allen, 2013
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how to make a basin?The ideal lithosphere in perfect equilibrium with theunderlying asthenosphere
almost nowhere on earth this is the casebecause the lithosphere is laterally and verticall
heterogenuous + under stress (extension/compression/buoyancy) and deforms.
LAB:~1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
crust = 2800 kg/m3
mcrust =30000
lith = 3300 kg/m3
mlith =70000
asth =3200 kg/m3
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how to make a basin?
LAB:~1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
crust = 2800 kg/m3
mcrust =30000
lith = 3300 kg/m3
mlith =70000
asth =3200 kg/m3
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how to make a basin?
• stretch/extend• bend/flex
LAB:~1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
crust = 2800 kg/m3
mcrust =30000
lith = 3300 kg/m3
mlith =70000
asth =3200 kg/m3
mechanismsof basins formationLAB:1300°C
Lith
osp
here
asthenosphere
lithospheric mantle
crust
extension of lithospheric plates
flexure due to vertical orhorizontal loads
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mechanismsof basins formationLAB:1300°C
Lith
osp
here
asthenosphere
lithospheric mantle
crust
extension of lithospheric plates
• active vs passive rift models• pure shear vs simple shear models• uniform vs depth-dependent stretching
flexure due to vertical orhorizontal loads
• compressive lithosphere folding (horiz. compression)• orogenic loading=foreland basins• basal loading (phase transitions; subducted slab remains;
down-welling in the deeper mantle)
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Isostasy
is the concept that the elevation of the Earth's surface seeks a balance between the
weight of lithospheric rocks and the buoyancy of asthenospheric "fluid" (nearly-
molten rock).
Archimedes’ Principle
Simple example: iceberg
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Isostasy Lithosphere with lateral load variations
Local Isostasy
Lithosphere
Depth of isostatic compensation
lith* ylith * g = asth* yasth ylith * g
lith
ylith
asth
yasth
lith < asth
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Isostasy Lithosphere with lateral load variations
Local Isostasy
Lithosphere
Depth of isostatic compensation
lith* ylith = asth* yasth
lithmylithm
asth
yasth
crust* ycrust lithm* ylithm= asth* yasth
crust ycrust
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IsostasyLithosphere with lateral load variations
Local isostasy: 1855Pratt
No crustal root below orogens
mantle
crust
Isostasy is achieved by lateral density variations
Depth of isostatic compensation
G. H. Pratt hypothesized that the density of the crust varies, allowing the base of the crust to be the same everywhere. Sections of crust with high mountains, therefore, would be less dense than sections of crust where there are lowlands. This applies to instances where density varies, such as the difference between continental and oceanic crust.
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IsostasyLithosphere with lateral load variations
Local isostasy: 1855Airy
crustal root below orogens
PrattNo crustal root below orogens
mantle
crust
Isostasy is achieved by lateral density variations
Depth of isostatic compensation
Isostasy is achieved by lateral thicknessvariationsy1>y3>y2>y4
mantle
y1y2
m3 y4
George Bedell Airy proposed that the density of the crust is everywhere the same and the thickness of crustal material varies. Higher mountains are compensated by deeper roots. This explains the high elevations of most major mountain chains, such as the Himalayas.
G. H. Pratt hypothesized that the density of the crust varies, allowing the base of the crust to be the same everywhere. Sections of crust with high mountains, therefore, would be less dense than sections of crust where there are lowlands. This applies to instances where density varies, such as the difference between continental and oceanic crust.
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IsostasyLithosphere with lateral load variations
Local isostasyAiry
crustal root below orogens
PrattNo crustal root below orogens
Lithosphere
Isostasy is achieved by lateral density variations
Depth of isostatic compensation
Isostasy is achieved by lateral thicknessvariationsy1>y3>y2>y4
asthenosphere
asthenosphere
y1y2
y3y4
lith* ylith = asth* yasth
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Flexural isostasy
Airycrustal root below
orogens
PrattNo crustal root below orogens
asthenosphere
Lithosphere
Lithosphere has a finite strength and acts like an elastic beamBends down if loadedWavelenght and amplitude of deflection is a function of rigidity
load
Isostasy is achieved by lateral density variations
Depth of isostatic compensation
Isostasy is achieved by lateral thicknessvariations
asthenosphere
asthenospherelithosphere
Local IsostasyLithosphere with lateral load variations
load
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Extension of lithospheric plates
LAB:1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
linitial
Initial Lithosphere Structure
ßdinital
linital
dstretched
lstretched
dinital*linital=dstretched*lstretched
dinital lstretched
dstretched linital=ß=
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stretching factor ( ß )
• is the ratio of the initial lithosphere thickness to the final (stretched) thickness of the lithosphere.
• example: ß=2 means, that lithosphere has been thinned to
(a) ¼(b) ½(c) none of the two?
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stretching factor ( ß )
• is the ratio of the initial lithosphere thickness to the final (stretched) thickness of the lithosphere.
• example: ß=2 means, that lithosphere has been thinned to
(a) ¼(b) ½(c) none of the two?
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Isostatic consequences of lithosphere extension
if plate boundary forces are absent, at the Lithosphere-Asthenosphere-Boundary, the load PLAB
in isostatic equilibrium isPLAB= c*yc*g + m*ym*g
Lithosphere
Asthenosphere
Lithospheric Mantle
Crust c yc
ym
m
level of isostatisc compensation
a
load of crust
load of lithospheric
mantle
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after stretching by ß the load PLABß is reduced
PLABß= c*ß-1yc*g + m*ß-1ym*g
Isostatic compensation isaccomplished by rising
asthenosphere ya with aLithosphere
c yc
ymm
ß-1yc
ß-1ym
ya
a
level of isostatisc compensationAsthenosphere
Lithospheric Mantle
Crust
Isostatic consequences of lithosphere extension
for isostatic equilibrium: c*yc*g + m*ym*g = c*ß-1yc*g + m*ß-1ym*g + a*ya*g
c*yc + m*ym = c*ß-1yc+ m*ß-1ym + a*ya
Lithosphere
c yc
ymm
ß-1yc
ß-1ym
ya
a
level of isostatisc compensationAsthenosphere
Lithospheric Mantle
Crust
Isostatic consequences of lithosphere extension
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c*yc + m*ym = c*ß-1yc+ m*ß-1ym + a*ya
Isostatic consequences of lithosphere extension
How much Asthenosphere has to flow in, if a lithosphere of initially 100 km thickness with a crust of 30 km is streched by ß=2 ?
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c*yc + m*ym = c*ß-1yc+ m*ß-1ym + a*ya
Isostatic consequences of lithosphere extension
How much Asthenosphere has to flow in, if a lithosphere of initially 100 km thickness with a crust of 30 km is streched by ß=2 ?
c*yc + m*ym - c*ß-1yc- m*ß-1ym)/ a = ya
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c*yc + m*ym = c*ß-1yc+ m*ß-1ym + a*ya
Isostatic consequences of lithosphere extension
How much Asthenosphere has to flow in, if a lithosphere of initially 100 km thickness with a crust of 30 km is streched by ß=2 ?
c*yc + m*ym - c*ß-1yc- m*ß-1ym)/ a = ya
dc=30kmdm= 70km
ß-1dc = 30/2=15km
ß-1dm = 70/2=35km
c=2800kg/m3
m=3300kg/m3
a=3200kg/m3
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(2800kg/m3*30km + 3300kg/m3*70km - 2800kg/m3*15km - 3300kg/m3*35km) 3300kg/m3 =ya
49218,75km = ya
c*yc + m*ym - c*ß-1yc- m*ß-1ym)/ a = ya
Isostatic consequences of lithosphere extension
How much Asthenosphere has to flow in, if a lithosphere of initially 100 km thickness with a crust of 30 km is streched by ß=2 ?
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what did we forget????
Lithosphäre
Asthenosphäre
lithosphärischer Mantel
Kruste c yc
dmm
ß-1yc
ß-1ym
ya
a
Isostatic consequences of lithosphere extension
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what did we forget????
The basin fill!!!!
For a water-filled basin:
c*yc + m*ym = c*ß-1yc+ m*ß-1ym + a*ya + w*yw
c*yc + m*ym = c*ß-1yc+ m*ß-1ym + a*ya + s*ys
c*yc + m*ym = c*ß-1yc+ m*ß-1ym + a*ya + s*ys + w*yw
Isostatic consequences of lithosphere extension
For a sediment-filled basin:
For a basin filled with water and sediment :
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For ß =2 => 30/2=15 km of the crust are filled with water
Isostatic consequences of lithosphere extension
(2800kg/m3*30km + 3300kg/m3*70km - 1000kg/m3*15km - 2800kg/m3*15km - 3300kg/m3*35km)/3300kg/m3 =ya
44531,25 km = ya
c*yc + m*ym = c*ß-1yc+ m*ß-1ym + a*ya + w*yw
c*yc + m*ym - w*yw - c*ß-1yc- m*ß-1ym)/ a = ya
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if the air-filled basin receives 5km Sediment
with sed=2400kg/m3
42343,75= ya
…….
Isostatic consequences of lithosphere extension
water-filled 44531,25 km = ya
air-filled 49218,75km = ya
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c=2800kg/m3
m=3300kg/m3
a=3200kg/m3
w=1000kg/m3
≈ 1/3 a
Isostatic consequences of someprocesses…How much isostatic rebound would a 200m sea level drop cause?
load that needs to be compensated byathenosphere:
W * yW = AyA
W * yW / A = yA
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c=2800kg/m3
m=3300kg/m3
a=3200kg/m3
w=1000kg/m3
≈ 1/3 a
Isostatic consequences of sea leveldrop
How much isostatic rebound would a 200m sea level drop cause?
load that needs to be compensated by athenosphere:
W * yW = AyA
W * yW / A = yA
1000kg/m3*200m/3200kg/m3= 62,5m
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c=2800kg/m3
m=3300kg/m3
a=3200kg/m3
w=1000kg/m3
≈ 1/3 a
Isostatic consequences of sea leveldrop
How much isostatic reboundwould a 200m sea level dropcause? 62,5 m
How much isostatic reboundwould the melting of a 2km thick ice sheet cause?
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c=2800kg/m3
m=3300kg/m3
a=3200kg/m3
w=1000kg/m3
≈ 1/3 a
Isostatic consequences of sea leveldrop
How much isostatic reboundwould a 200m sea level dropcause? 62,5 m
How much isostatic reboundwould the melting of a 2km thick ice sheet cause?
W * yW / A = yA
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c=2800kg/m3
m=3300kg/m3
a=3200kg/m3
w=1000kg/m3
≈ 1/3 a
Isostatic consequences of ice melting
How much isostatic reboundwould a 200m sea level dropcause?
How much isostatic reboundwould the melting of a 2km thick ice sheet cause?
W * yW / A = yA
1000kg/m3 x 2000m/3200kg/m3= 625 m
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c=2800kg/m3
m=3300kg/m3
a=3200kg/m3
w=1000kg/m3
≈ 1/3 a
Isostatic consequences of ice melting
How much isostatic reboundwould a 200m sea level dropcause?
How much isostatic reboundwould the melting of a 2km thick ice sheet cause?
Can 2000m oferosion/denudation beexplained by sea levelchanges or deglaciation?
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c=2800kg/m3
m=3300kg/m3
a=3200kg/m3
w=1000kg/m3
≈ 1/3 a
sed=2300kg/m3
≈ 2/3 a
Isostatic consequences of sea leveland glaciations
How much isostatic reboundwould a 200m sea level dropcause? ~60m
How much isostatic reboundwould the melting of a 2km thick ice sheet cause? ~600m
Can 2000m of denudation beexplained by sea levelchanges or deglaciation? NO!
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c=2800kg/m3
m=3300kg/m3
a=3200kg/m3
w=1000kg/m3
≈ 1/3 a
sed=2300kg/m3
≈ 2/3 a
How much isostatic reboundwould the erosion of 1 km ofsediments cause?
sed * ysed / A = yA
2300kg/m3 x 1000m/3200kg/m3= 718 m
Isostatic consequences of erosion
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But! isostasy may be just one part of a sequence of coupled processes.
c=2800kg/m3
m=3300kg/m3
a=3200kg/m3
w=1000kg/m3
≈ 1/3 a
sed=2300kg/m3
≈ 2/3 a
How much isostatic reboundwould the erosion of 1 km ofsediments cause?
sed * ysed / A = yA
2300kg/m3 x 1000m/3200kg/m3= 718 m
Isostatic consequences of erosion
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effects of initial water depth
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Basic concepts & definitionsRules to remember about isostasy:
• Ocean crust is thin and dense (5-10 km, density of ~ 3000 kg/m³).
• Continental crust is thick and less dense (30-50 km density ~ 2800 kg/m³).
• Less dense crust rises higher (continental crust, mountains) than does more dense crust (ocean crust). Oceans are low and continents are high.
• Loading the crust causes it to sink (sediments or ice sheet on top).
• Unloading the crust causes it to rise (as in eroding a mountain down).
Isostatic consequences of lithosphere extension
isostacy is a static concept.
geology of basins takes time.
-> how to integrate time dependence?
-> McKenzie: changes in temperature (and the related effects on density) over time are the key!
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Heat Transfer- Heat flow or transfer
is a processachieved in thelithospheredominantely byconduction
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The thermal lithosphere-asthenosphereboundary (LAB)
Allen & Allen, 2013
convective heat transport
conductive heat transport
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LAB
crust
lithospheric mantle
sediments
heat in basins…controlled by
heat sources
heat transport mechanisms
con
du
ction
convection
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LAB ~1300°C
crust
lith. mantle
sediments
grad
T
~10°C
heat sources
• heat input fromasthenosphere
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heat sources
• heat input fromasthenosphere
• radiogenic heatproduced in thelithospherecrust
lith. mantle
sediments
LAB ~1300°C
grad
T
~10°C
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radiogenic heat production
• contributes significantly to heat generation in the crust• most rocks contain radioactive minerals• concentrations are low, but due to large thickness may effect the
temperature distribution (source or sink in the heat equation)• uranium, potassium, thorium
• low carbonates & carbonates• medium sandstones• high granite, silt- and claystones
510)64.279.935.3( ThUK cccA (Schmucker, 1969)
A heat production densityc heat generation constants for known concentrations
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heat producing elements in sedimentary rocks
Allen & Allen, 2005 67
convectionconduction advection
by diffusion pressure driven bouyancy-driven
heat transport mechanism?
relevant at different scales
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Heat Transfer by Conduction
Diffuse and therefore SLOW process by which the kinetic energy is transferred by interatomic or intermolecular collision
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Conductive heat flow
Fourier’s law
Q: heat flowK: thermal conductivityT: temperaturey: coordinate in the direction of the
temperature variation (e.g. depth)
dy
dTKQ
gradTKQ
Heat flow is in units of mW m-2 or cal cm-2 s-1
HFU (heat flow units)=10-6 cal cm-2 s-1 or 41.84 mW m-2
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Conductive Heat Transfer
At
Tc
z
TK
zy
TK
yx
TK
x zyx
K – thermal conductivitiesc – specific heat- densityS – source or sink term (radiogenic heat, convection)
Any loss or gain in heat in a unit volume is balanced by a temperature change multiplied with volumetric heat capacity (Gretener, 1981).
radiogenic heat enters as a source term A
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post-rift deposits
syn-rift deposits
crust
Lithospheric mantle
asthenosphere
LAB
If conduction has time enough to equilibrate on the scale of the lithosphere (dT/dt=0)
=> steady state conditions:
A – radiogenic heat production [mW/m³]T – temperature [°C]k- conductivity [W/mK]
At
Tc
z
TK
zy
TK
yx
TK
x zyx
Az
TK
zy
TK
yx
TK
x zyx
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Heat Transfer by Conduction
Thermal Conductivity of rocks= change in temperature per unit length for a spec. amount of heatHigh thermal conductivity => low thermal gradientLow thermal conductivity => high thermal gradient
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thermal conductivity of sedimentary rocks
Thermal conductivity λ W m∙K 1 heat W transported over a distance m with a drop in Temperature °C or K
Bulk thermal conductivity λb ϕλf 1 ϕ λS
oPorosity and fluid content water, brine, oil, gas oFunction of depth porosity loss with burial - compaction oMineralogy of the framework grains oType and amount of material in the matrix clay minerals
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thermal conductivity of sedimentary rocks
low reduced heat transfer high thermal gradient
high efficient heat transfer low thermal gradient
76Allen & Allen, 2013
thermal properties of sedimentary rocks
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77Allen & Allen, 2013
blanketing effects
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Allen &Allen, 2013, Fig.2.22
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relevant propertiessediments: mostly insulating, moderate radiogenic heatproduction, pore fluids heat transport, „thermal blanketing“
crust: conductive, high radiogenicheat production heat transport, heat budget
lithospheric mantle: conductive, low radiogenic heat production heat transport
depth to LAB: thermal gradient, heat budget
LAB ≈ 1300°C
cruste
lith. mantle
sediments
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typical HF values
Allen &Allen, 2013, Fig.10.31
Allen &Allen, 2013, Fig.3.8
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Observations
• Sediment fill can be a few 100m and up to >22km thick
• Structure of sediments often indicates 2 major phases:– Syn-rift: normal faults and grabens define sub-basins
=>initial, FAST subsidence– Post-rift:discordantly on synrift, onlaps on sub-basins, spatial
widening of subsiding area, continuous layering, decreasingrate of subsidence with time=>thermal subsidence
• Crust: thinned, seismic velocities and densities of the crustindicate changes of crustal configuration and composition in response to extension, strongly variable for different basins
• Lithospheric mantle: strongly variable for different basins
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Isostatic consequences of lithosphereextension
isostacy is a static concept.
geology of basins takes time.
-> how to integrate time dependence?
-> McKenzie: changes in temperature (and the related effects on density) over time are the key!
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LAB:1300°CLAB:1300°C
extensionLAB:1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
McKenzie´s uniform stretching model: Passive pure shear extensional model
considers isostatic and thermal consequences of lithosphere stretching
extensional rift models…
linitial
lstretchedUniform and symmetric lithosphericthinning by ß is accomodated bybrittle faulting in the upper crustand by ductile flow in the lower
crust and upper mantle
Stretching by ß = lstretched/linitial
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LAB:1300°CLAB:1300°C
extensionLAB:1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
Passive Rift Model
linitiallstretched- As the lithosphere is thinned, it is
partly replaced by hotter mantledue to PASSIVE asthenosphericrise to compensate the extensionisostatically
- subsidence/ uplift upon mechanical stretching can beobtained
- The amount of initial (instantaneous) subsidence will depend on the initial amount of stretching (ß) due to isostasy
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immediate subsidence and
elevated isotherm
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initial subsidenceMcKenzie´s uniform stretching model
smvm
mmv
L
cmv
L
ccmL
i T
TyyT
yy
y
S
)1(
)/11(22
1)(
*
***
yL = initial thickness of the lithosphereyc = initial thickness of the crustm
*= mantle density (at 0°C)c
*=crustal density (at 0°C)c=density of the basin infillv=thermal expansion coefficent for crust and mantleTm=temperature of the asthenosphere
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initial subsidenceMcKenzie´s uniform stretching model
smvm
mmv
L
cmv
L
ccmL
i T
TyyT
yy
y
S
)1(
)/11(22
1)(
*
***
uniform stretchingmodel predictsimmediate subsidenceif crustal thickness ismore than 1/8th of thelithosphere thickness
(For a 100km thicklithosphere this wouldmean a crust of ~12 km)
isostatic and thermal consequences oflithosphere extension
-> cooling/heating changesdensity:
increase in temperature=> increase in volume=>decrease in density
-> decrease in temperature=>decrease in volume =>increase in density
LAB:1300°CLAB:1300°C
extension
cooled LAB:1300°C
lstretched
Stretching by ß = lstretched/linitial
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=> Thermal subsidence
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Thermal subsidence & Plate Cooling Model
Oceanic plates form by lithospheric cooling
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North Pacific
North Atlantic
Depth and Heat Flow of ocean floor
𝑑 5.65 2.47𝑒
𝑑 2.6 0.365 𝑡
Good approx. till ~70 Ma
Slower for greater ages𝑞
510
𝑡
𝑞 48 96𝑒
90
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isostatic consequences…Seafloor Depth
𝜌 𝑑𝑧 𝜌 𝑧, 𝑇 𝑑𝑧 𝜌 𝑑𝑧 𝜌 𝑑𝑧 𝜌 𝑧, 𝑇 𝑑𝑧
Column of lithosphere Ridge
𝜌 𝑧, 𝑇 𝜌 · 1 𝛼 𝑇 𝑧, 𝑡 𝑇 𝑇 𝑧, 𝑡 𝑇 · erf𝑧
2 𝜅 · 𝑡
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Allen& Allen 2005: Basin Analysis
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LAB:1300°CLAB:1300°C
extension
cooled LAB:1300°C
Passive Rift ModelMcKenzie´s uniform stretching modelPassive pure shear extensional model
thermal subsidence
LAB:1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
linitiallstretched
Stretching by ß = lstretched/linitial
Gradual cooling (50 - 100 m.y.) & related density increaseleads to basin subsidenceabove and progressive infillwith sediment.
93
94
McKenzie´s uniform stretching modelPassive pure shear extensional model
subsidence caused by thermal contraction
/
0 1sin)( teEtS
smmvmL TyE *2*0 /4
wm
mvmcscc
TyS
)2/(
)/11(*
L
cm
vcm
y
yT
2
**
Final subsidence
48
LAB:1300°CLAB:1300°C
extension
cooled LAB:1300°C
Passive rise of hot/low-density material from the asthenosphere
sinking cold/high-density material in the asthenosphere
Passive Rift ModelMcKenzie´s uniform stretching modelPassive pure shear extensional model
2-phase subsidence
Predicts 2-stage subsidence:
1. Si as a result of tectonicstretching – on a short time scale, <20 my, and
2. Sth as a result of thermal subsidence – long time scale, ca. 50 – 100+ my.
Subs
iden
ceTime
LAB:1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
linitial
95
LAB:1300°CLAB:1300°C
extension
cooled LAB:1300°C
Passive rise of hot/low-density material from the asthenosphere
sinking cold/high-density material in the asthenosphere
quantifying extension of lithospheric plates
Passive Rift ModelMcKenzie´s uniform stretching model
- Where to know ß from?
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49
LAB:1300°CLAB:1300°C
extension
cooled LAB:1300°C
Passive rise of hot/low-density material from the asthenosphere
sinking cold/high-density material in the asthenosphere
extension of lithospheric plates
Passive Rift ModelMcKenzie´s uniform stretching model
- Where to know ß from?
estimate ß from deep seismic information
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98
A deep reflection seismic line across the Northern Rhine Graben Wenzel etal., 1991
crustini
ß*crustini
crustini
ß*crustini
50
ß from preserved present-day crustalthickness
crustini
ß*crustini
lithospheric mantle re-equilibrated!
does not conserve ß…
99
subsidence curves
Where to know ß from?
estimate ß from deep seismic information estimate ß from tectonic subsidence
100
51
LAB:1300°CLAB:1300°C
extension
cooled LAB:1300°C
Passive rise of hot/low-density material from the asthenosphere
sinking cold/high-density material in the asthenosphere
Passive Rift ModelMcKenzie´s uniform stretching modelPassive pure shear extensional model
extension of lithospheric plates
Structurally the uniform stretching model predicts :
Syn-rift extensional faults in the upper crust with graben filloverlain by onlapping post-rift sediments of the thermal subsidence phase.
101
102
pure shear: temperature of the uprising asthenosphere exceeds the solidus of the mantle and allows melting.
Magmatic Implications of the Pure Shear Model
for ß> 3 formation ofdecompressional melts!
melts are trapped around thecrust-mantle boundary as theyare less dense than thelithospheric mantle but denserthan the crust. underplating
melts weaken the lithosphere continental breakup
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after breakup: magma rich passive margin
103
What are the observations?
seismic data: structure and velocity distribution
synrift grabensbreakup unconformity below deposits of thermal subsidence
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53
What are the observations in the sediments…
reflectionseismic data: structure
Argentine margin, Franke et al., 2010
breakupunconformi
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Comparison of profilesfrom northeastGreenland and Lofoten-Vesteralen Margin (Voss et al., 2009)
What are the observations in the crust…?
refractionseismic data: velocity distribution
high velocity bodiesin the lower crustnear COT
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54
seismic exercise
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Distance[km]100 200 300 400 500 600 700 800Depth [km]
0
10
20
30
40
108
55
cold (old) magma-rich passive margin
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Additional Information
110
56
111
other passive riftmodels…
the simple shearmodel
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Simple Shear Models
Structurally the simple shear model predicts :
Syn-rift extensional faults in the upper crust with graben fill detached along listric master faults and offset from overlying post-rift sediments of the thermal subsidence phase.
57
113
Implications of the Simple Shear ModelsSome thermal consequences:
With McKenzie’s pure shear model decompression leads to melting of upper mantle material.=>magma-rich passive margins
With Wernicke’s simple shear model it is very difficult to produce sufficient decompression to allow magma formation. =>magma-poor passive margins
114
pure shear: temperature of the uprising asthenosphere exceeds the solidus of the mantle and allows melting.
different thermal consequences:
simple shear: temperature of the uprising asthenosphere never reaches the solidus - so no melting occurs.
Implications of the Simple Shear Models
58
115
if rifting proceeds and leads tobreakup how will the evolving passive
margin look like?
after breakup: magma poor passive margin
Franke, 2012
Exhumed mantle
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No high velocity/high density bodies in the lower crust, no SDRs
59
Reston and Perez-Gussinye, 2007
low-angle normal faults, brittle, assymmetric, nonvolcanic,, exhumed mantle, low surface heat flow
Conjugate Flemish Cap—Galicia Bank margin pair general crustal structure,
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What are the observations to discriminate between different models?
seismic
gravity
seismology
heat flow
deformationmode
Franke, 2012 118
60
extension of lithospheric plates• may lead to continental rifting followed by breakup• depending on the mode and the rate of extension different
types of passive margins may develop
• magma rich margins• have lots of syn-rift volcanism, high velocity-high-
density bodies in the lower crust, SDRs• are hot or cold depending on age
• magma-poor margins• have very little syn-rift volcanism, uniform lower crust,
No SDRs, but large low-angle detachment fault systems
• are rather cold throughout their history
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passive riftingis NOT the only
extensionalmechanism of basins
formation!
120
61
(+dynamic topography)LAB:1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
Uplift & Erosion
1300°C1300°C
positive dynamic topography
rising hot/low-density material in the asthenosphere
cooled
cooling & subsidence
1300°C
Active Rift Model
Active rise of asthenosphereprovokes initial thermal uplift
and erosion.
Subsequent cooling leads tosubsidence and results in
lithosphere thinning =dinitial / dstretched
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LAB:1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
Uplift & Erosion
1300°C1300°C
positive dynamic topography
rising hot/low-density material in the asthenosphere
cooled
cooling & subsidence
1300°C
Active Rift Modelpredicts
1. Uplift as a result of mantleupwelling - short time scale, ca. 10 – 20+ my. (- thermal anomaly)
2. thermal subsidence as a resultof cooling – long time scale, ca. 50 – 100+ my.
Subs
iden
ce
Time
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62
LAB:1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
Uplift & Erosion
1300°C1300°C
positive dynamic topography
rising hot/low-density material in the asthenosphere
cooled
cooling & subsidence
1300°C
Active Rift Modelwhat may be left only…
1. An erosional unconformity at the base (Why?)
2. Little preserved faulting andwide-spanned subsidence(sag basin) with onlaps at themargins (bull-head structure)
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LAB:1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
Uplift & Erosion
1300°C1300°C
positive dynamic topography
rising hot/low-density material in the asthenosphere
cooled
cooling & subsidence
1300°C
Active Rift Model
Isostatically, thinning of the lithospheric mantle results in net surface uplift as heavier lithospheric mantle is replaced by less dense (partially molten) asthenospheric material (active rift model).
Only interaction with surface processes (erosion) and shallow extension due to upward flexure of the plate gives net subsidence due to subsequent cooling of the asthenospheric material.
time
Heat flow
hot
cold
shallow
deep
subs
iden
ce
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East African Rift System
EARS=THE active rift model:
high topography, high heat flow, strong volcanism
hot anomaly in asthenospheric mantlebelow
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East African Rift System – active rifting: high topography
www. geology.com/articles/east-africa-rift.shtml
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EARS volcanism, + high heat flow
www. geology.com/articles/east-africa-rift.shtml
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East African Rift System – active rifting
www. geology.com/articles/east-africa-rift.shtml
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Crustal structure of the northern Main
Ethiopian Riftfrom EAGLE
Ethiopia Afar Geoscientific Lithospheric Experiment
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Main Ethiopian Rift from EAGLE
rather uniform crustal velocities
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Upper mantle P-wave speed variations beneathEthiopia and the origin of the Afar hotspot
Margaret H. Benoit et al. 2006
passive seismology tounravel mantle properties
slow anomaly in the mantlebelow the rift
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Margaret H. Benoit et al. 2006132
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Take home• 2 main types of extensional basin forming mechanisms: active vs passive
rifting, but several sub-types (pure shear, simple shear, depth-dependent, rheology-dependent),
• deformation (subsidence history) depends on rate of extension andrheological structure of the lithosphere
1) magma-poor (pre-breakup), Andersonian faulting, Wernicke-simple shear model; slow, cold, exhumed & serpentinized mantle at seafloor
=>plate-driven?
transitional cases….
2) magma-rich: intensive pre-breakup magmatism (LIPs, Plumes oder hot spots in der Nähe); SDR, HDB/HVB in der extremely thinned continental crust next toCOB; often enigmatic high rift shoulders; hot, fast? breakup
=> mantle-driven?
• McKenzie‘s uniform streching model describes postrift subsidence amazinglywell, but dynamic models considering rheological heterogeneities needed toexplain the large variety of different extensional basins and magmaticexpressions
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more specific models of basinformation…
134
• consider stratified and heterogenuous rheology ofthe lithosphere
• variations in stretching rate
• predict many different rifting/breakup scenarios
68
Strength profiles:
135Allen & Allen, 2013, Fig. 2.38
jelly sandwich
versus
crème brulée
more sophisticated thermo-mechanicalmodels of basin formation…
136
jelly sandwich
versus
crème brulée
Allen & Allen, 2013
69
meanwhile models go 4D!
137
Koptev et al., 2015
plume-lithosphere interaction slow and longlastingintracontinental basins
Cacace & Scheck-Wenderoth 2016
Mechanisms of basins formation
LAB:1300°C
Lithosphere
asthenosphere
lithospheric mantle
crust
1.extension of lithospheric plates
2.flexure due to vertical or horizontal loads
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70
LAB:1300°C
Lithosphere
Asthenosphere
Flexure
asthenosphere
LAB:1300°Chot
cold
Lithosphere folding
subduction
trenchforeland
basin
compression
vertical loading
≈ 1000 km
cold
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Basins due to flexure
Flexure = long-wavelength bending of the elastic lithosphere in response to vertical or horizontal loads
wavelength of flexure depends on flexural rigidity of lithosphere and magnitude of load
flexural rigidity of lithosphere is age dependent (old cold strong versus young hot weak)
flexural rigidity of lithosphere NOT time-dependent=> viscous relaxation = relatively fast process (<105-106 years)
Cloetingh & Burov, 2010140
71
Basins due to flexure
D= flexural rigidity
ω= deflection
α=flexural parameter
=density
x0=half wavelength
x=0:maximum ωx0: ω=0
Allen and Allen, 2013141
142
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143Allen and Allen, 2013
crème brulée
jelly sandwich
Folded lithosphere of Iberia
Cloetingh & Burov, 2010
analogue modelling experiments,displaying pop-up structures underlying topographic highs, accompanied by folding of the Moho (Fernandez-Lozano et al.,2010)
Topography displaying cylindrical patterns of alternating parallel trending highs and lows and corresponding gravity anomalies
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the influence of the deep mantle..
145
dynamic topography (+/-)
Uplift & Erosion
1300°C1300°C
positive dynamic topography
subsidence
1300°C1300°C
sinking cold/high-density material in theasthenosphere (why?)rising hot/low-density material in the
asthenosphere
cooled
cooling & subsidence
1300°C
negative dynamic topography
time
Heat flo
w
hot
cold
shallow
deep
sub
sid
ence
cold
Subsidence history similiar tolithosphere folding
sub-type: active rift model
146
74
“Transtensional”, pull-apart basins, e.g., Salton Sea, Dead Sea etc.
147
148
75
orientation of principal stress axes in strike-slip regime?
Sigma 1 horizontal
Sigma 2 vertical
Sigma 3 horizontal
DIFFERENT from extensional basins
Sigma 1 vertical
Sigma 2 horizontal
Sigma 3 horizontal
DIFFERENT from compressive basins
Sigma 1 horizontal
Sigma 2 horizontal
Sigma 3 vertical
149
strike-slip basins
Sigma1
Sigma 3
150
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Transform --- plate (lithosphere-) scale e.g. San Andreas FTranscurrent --- crustal scale e.g. Garlock Fault CA
(USGS)
Transform?Transcurrent?
151
narrow (km-10-nerkm), elongated (up to1000km)
asymmety, abrupt facies changes
negative/positive flower structures
Sediments above steep basement fault
fast, short-lived subsidence
high sedimentation rates (1m/1000a)
lateral migration of depo centres
strike-slip basin types
152
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how to find out that its not a rift?Strike-slip PDZ in Cross-section
• Flower-structures
• Variable offsets along the same fault because of movement in and out of the plane of observation
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CONTINENT-OCEAN TRANSFORM MARGINS Sheared Margins
Figure 1. Generic three-stage model for shear margin formation (after Lorenzo, 1997): (1) rift: continent-continent shearing; (2) drift: continent-ocean transform boundary (active margin); and (3) passive margin: continent- ocean fracture zone boundary.
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Sheared Margins Gulf of Guinea
Antobreh et al., 2009
155
Antobreh et al., 2009
crustal structure @ Sheared Margins
156
N-S reflection seismic line across the Côte d’Ivoire-Ghana transform margin, Basile et al., 1993
79
To sum up…
157
Typical subsidence histories
Plate tectonics and basin subsidence historyXiangyang Xie and Paul L. HellerGSA Bulletin; January 2009; v. 121; no. 1-2; p. 55-64; DOI: 10.1130/B26398.1
158
80
159
typical tectonic subsidence rates
Allen& Allen, 2013, Fig. 9.17
Different thermal histories
depending on mechanismraterheology
very characteristicthermal and subsidence historiesdeformation pattern (brittle vs ductile)
160
81
factors influencing the thermal field
• basal heat flow (thinning mechanism, extensional/ flexural?)• age (time since basin initiation, depth of thermal LAB)• sedimentation rate• lithology of basin fill and basement ( densities, thermal
conductivities & radiogenic heat production)• compaction history (highly compacted ->dense, highly
conductive)• crystalline basement structure
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