Plate convergence usually commences with intra-oceanic Subduction, Andean margins commonly start after ophioliteobduction and subduction flip.

CONVERGENT PLATE MARGINSCONVERGENT PLATE MARGINS1) Intra-oceanic (ensimatic) subduction2) Andean margins3) Continent - continent collision zones

1) 2)

3)

REMEMBER, IN 3-D A CONVERGENT MARGIN MAYHAVE DIFFERENT MATURITY ALONG STRIKE!

legend and estimates of plate-tectonic forces • Fsp - Slab-pull(+)

• Frp - Ridge-push(+)• Fsu - Suctional force (+)• For - Orogenic spreading(+)• Fdf - Mantle drag-force (+ or -)• Fsr - Subduction ressistance (-)• Fcd - Extra continental-drag(-)• Ftr - Transform resistance (-)

Frp = g e (m – w) (L/3 +e/2) ≈ 2*1012 Nm-1

Can also be expressed as a function of age:

Frp = gmTt [1 + (m/(m-w)) 2T/] = 1.19x10-3 t (Unit MPa)

g - gravity ≈ 9,8 ms-2

e – elevation of spreading ridge above cold plate ≈ 3,3 km (e- is a function of the age [t])

m – mantle density, ≈ 3,2 g cm-3 w – water density

L – lithosphere thickness ≈ 85km

T - temperature(~1200C), -thermal diffusivity [ms], - coefficient of thermal expansion [ = 3*10-5 K-1]

Estimate of slab-pull force Fsp pr. unit length subduction zone (see Fowler: Solid Earth, Chap 7, for details)

z – depth (d = z give Fsp = 0) – coefficient of thermal expansion T1 – mantle temp, d+L – thickness of the upper mantle L – Lithosphere (plate) thickness Re – Thermal Reynolds number

Fsp(z) = 8g m T1L2Re4

[exp(-2ReL

2z) - exp(- 2d

2ReL)] = ca 2x1013Nm-1

Re =(mcpvL)/2k

Thermal Reynolds number

k - conductivity

cp - spesific heat

k - kinematic viscosity

v - subd. velocity

Plate velocity as a function of % subduction margin

Reactions and phase transitions affecting the forces in subduction

zones

• In addition to the thermal contraction and density change will the forces of the subducting litosphere be affected by

• Gabbro to eclogite transition (+)• Olivin-spinel transiton (+)• Spinel to oxides (perovskitt and periklas) (-)

Modeled density structure of subducted MORB, Hacker et al. 2003

Temperature variation across a subduction zone

• Notice the localization of the olivin-spinel and spinel-oxide transitions.

• Use the next fig to explain the phenomena

Phase diagrams for the transititions for olivin to spinel and spinel to

post-spinel (oxides)

THE ANATOMY OF A SUBDUCTION COMPLEXTHE ANATOMY OF A SUBDUCTION COMPLEX

Trench

Accreationaryprism

Fore-arcbasin

Active volcanic

arc

Back-arcbasin/spreading

Tension

alternatingcompressionamd tension

sea level

Remnant-arcsfrom

arc-splitting

Outernon-volcanic

island

Compression

Low geothermLow geotherm

High geothermHigh geothermHigh geothermHigh geotherm

Please notice that Benioff zones frequently have an irregular shape in 3-D (ex. Banda Arc). 80% of all seismic energy is released in Benioff zones.The low geotherm in subductions zones makes them the prime example of high P - low T regional metamorphic complexes. The high geotherm in the arc-region gives contemporaneous high-T low P regional metamorphism, together these two regionsgive rise to a feature known as”Paired Metamorphic Belts”

Accreationary Prism, Example from Scotland.Age:Late Ordovician to Late Silurian ca 450-420 Ma

PAIRED METAMORPHIC BELTS

Trench

Tension

alternatingcompressionamd tension

sea level

Compression

Low geothermLow geotherm

High geothermHigh geothermHigh geothermHigh geotherm

Seismic quality factor (Q): The ability to transmitt seismic energy without loosing the energy. Low Q in high-T regions. Seismic quiet zones---NB potential build-up to very large quakes!

Arc-splitting - tensional regime above subductions zones. Subductionroll-back.High heat-flow in the supra-subductions zone regime give rise torelatively low shallow sealevel above the back-arc basins. Mostophiolite complexes have their origin is a supra-subduction environment

Ophiolites normallyoriginate here!

Blueshists normallyoriginate here!

SEISMICITY

3 - D MORPHOLOGY

NB! NOTICE INTRA-SLABEARTHQUAKES

26/12-2004, Mag 9 earthquake of Sunda arc - Andaman sea

Link: fault plane solution

Link: displacement magnitude

Link: earthquake information in general

The amount of energy radiated by an earthquake is a measure of the potential for damage to man-made structures. Theoretically, its computation requires summing the energy flux over a broad suite of frequencies generated by an earthquake as it ruptures a fault. Because of instrumental limitations, most estimates of energy have historically relied on the empirical relationship developed by Beno Gutenberg and Charles Richter: log10E = 11.8 + 1.5MS where energy, E, is expressed in ergs. The drawback of this method is that MS is computed from an bandwidth betweenapproximately 18 to 22s. It is now known that the energy radiated by an earthquake is concentrated over a different bandwidth and at higher frequencies. With the worldwide deployment of modern digitally recording seismograph with broad bandwidth response, computerized methods are now able to make accurate and explicit estimates of energy on a routine basis for all major earthquakes. A magnitude based on energy radiated by an earthquake, Me, can now be defined, Me = 2/3 log10E - 2.9. For every increase in magnitude by 1 unit, the associated seismic energy increases by about 32 times. Although Mw and Me are both magnitudes, they describe different physical properites of the earthquake. Mw, computed from low-frequency seismic data, is a measure of the area ruptured by an earthquake. Me, computed from high frequency seismic data, is a measure of seismic potential for damage. Consequently, Mw [MW = 2/3 log10(MO) - 10.7] Mw= µ(area)(displacement)] and Me often do not have the same numerical value.

Frictional heating on faults may result in melting of any rock-coposition

Stress-measurements from grain-size and/or dislocation density (4 to 5 x1013m-2) in olivine associated with pseudotachylytes in peridotite indicate that peridotites (mantle rocks) may sustain extreme differential stress: 1-3≈ 3-600 MPa. Assuming a fault with a modest displacement of d ≈ 1m, and a differential stress of 300 MPa the release of energy according to equation

(1): Wf = Q + E where Q = heat and E = seismic energy is Wf = d n = d (1-3)/2 =1m(300MPa)/2 ≈ 1.5 x 108 J m-2 or 47 kWhm-2.

The seismic energy (E) is commonly estimated to be < 5% of Wf on a strong fault, ie.less than 2.3 kWh m-2 is radiated as seismic waves, the remaining energy (Q) turns to heat and surface energy (difficult to measure) along the fault. The process is adiabatic since the fault movement occurs in seconds and no heat is lost by conduction (thermal diffusivity ~1.5 mm2s-1). Taking the heat capacity of lherzolite, Cp = 1150 J kg-1 oC-1 and a heat of fusion (Fo) H = 8.6 x106 Jkg-1 the thermal energy (equation 4) required to melt one kg of peridotite:

(4) Q = Cp(T) + H = 1150Jkg-1oC-1 (1200oC) + 8.6 x106 Jkg-1 = 2.7 x 105 Jkg-1.

On a fault with D = 1m, ~60 kg lherzolite may melt pr m2 of the fault plane, corresponding to an approximately 2 cm thick layer of ultramafic pseudotachylyte.