Geodynamics of sedimentary basinsirtg-strategy.de/.../ScheckWenderothStrategyLecture10.01.2019s/index.pdf · Isostasy is the concept that the elevation of the Earth's surface seeks

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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


• 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


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


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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Local isostasy: 1855Airy

crustal root below orogens

PrattNo crustal root below orogens

mantle

crust



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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Local isostasyAiry

crustal root below orogens


Lithosphere



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


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 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


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


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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

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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




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what did we forget????

Lithosphäre

Asthenosphäre

lithosphärischer Mantel

Kruste c yc

dmm

ß-1yc

ß-1ym

ya

a


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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


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


(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

…….


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


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


How much isostatic reboundwould a 200m sea level dropcause? 62,5 m


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?


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?


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).


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



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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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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

86

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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87

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


extension

cooled LAB:1300°C

lstretched


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=> Thermal subsidence

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Thermal subsidence & Plate Cooling Model

Oceanic plates form by lithospheric cooling

89

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

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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

91

Allen& Allen 2005: Basin Analysis

47


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


Gradual cooling (50 - 100 m.y.) & related density increaseleads to basin subsidenceabove and progressive infillwith sediment.

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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

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extension

cooled LAB:1300°C

Passive rise of hot/low-density material from the asthenosphere

sinking cold/high-density material in the asthenosphere


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


extension

cooled LAB:1300°C



quantifying extension of lithospheric plates

Passive Rift ModelMcKenzie´s uniform stretching model

- Where to know ß from?

96

49


extension

cooled LAB:1300°C




Passive Rift ModelMcKenzie´s uniform stretching model

- Where to know ß from?

estimate ß from deep seismic information

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A deep reflection seismic line across the Northern Rhine Graben Wenzel etal., 1991

crustini

ß*crustini

crustini

ß*crustini

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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


extension

cooled LAB:1300°C





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

52

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

104

53

What are the observations in the sediments…

reflectionseismic data: structure

Argentine margin, Franke et al., 2010

breakupunconformi

105

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

106

54

seismic exercise

107

Distance[km]100 200 300 400 500 600 700 800Depth [km]

0

10

20

30

40

108

55

cold (old) magma-rich passive margin

109

Additional Information

110

56

111

other passive riftmodels…

the simple shearmodel

112

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

116

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,

117

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

119

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

121

LAB:1300°C

Lithosphere

asthenosphere

lithospheric mantle

crust

Uplift & Erosion

1300°C1300°C



cooled


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

122

62

LAB:1300°C

Lithosphere

asthenosphere

lithospheric mantle

crust

Uplift & Erosion

1300°C1300°C



cooled


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)

123

LAB:1300°C

Lithosphere

asthenosphere

lithospheric mantle

crust

Uplift & Erosion

1300°C1300°C



cooled


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

124

63

East African Rift System

EARS=THE active rift model:

high topography, high heat flow, strong volcanism

hot anomaly in asthenospheric mantlebelow

125

East African Rift System – active rifting: high topography

www. geology.com/articles/east-africa-rift.shtml

126

64

EARS volcanism, + high heat flow


127

East African Rift System – active rifting


128

65

Crustal structure of the northern Main

Ethiopian Riftfrom EAGLE

Ethiopia Afar Geoscientific Lithospheric Experiment

129

Main Ethiopian Rift from EAGLE

rather uniform crustal velocities

130

66

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

131

Margaret H. Benoit et al. 2006132

67

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

133

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

138

70

LAB:1300°C

Lithosphere

Asthenosphere

Flexure

asthenosphere

LAB:1300°Chot

cold

Lithosphere folding

subduction

trenchforeland

basin

compression

vertical loading

≈ 1000 km

cold

139

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

72

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

144

73

the influence of the deep mantle..

145

dynamic topography (+/-)

Uplift & Erosion

1300°C1300°C


subsidence

1300°C1300°C

sinking cold/high-density material in theasthenosphere (why?)rising hot/low-density material in the

asthenosphere

cooled


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

76

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

77

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

153

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.

154

78

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

161

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