Plate convergence usually commences with intra-oceanic
Subduction, Andean margins commonly start after ophiolite
obduction and subduction flip.
CONVERGENT PLATE MARGINS
1)
2)
3)
Intra-oceanic (ensimatic) subduction
Andean margins
Continent - continent collision zones
1)
2)
3)
REMEMBER, IN 3-D A CONVERGENT MARGIN MAY
HAVE 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 = gmTt [1 + (m/(m-w)) 2T/] = 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 [ms],  - 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)
Fsp(z) =






z

T1
d+L
L
Re
8g m T1L2Re
4
2z
[exp(-
2ReL
2d
) - exp()] = ca 2x1013Nm-1
2ReL
– depth (d = z give Fsp = 0)
– coefficient of thermal expansion
– mantle temp,
– thickness of the upper mantle
– Lithosphere (plate) thickness
– Thermal Reynolds number
Re =(mcpvL)/2k
k - conductivity
Thermal Reynolds number
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 COMPLEX
Outer
non-volcanic
island
Compression
sea level
Fore-arc
basin
alternating
compression
amd tension
Active
volcanic
arc
Back-arc
basin/spreading
Remnant-arcs
from
arc-splitting
Tension
Trench
Accreationary
prism
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 regions
give 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
Blueshists normally
originate here!
Compression
sea level
alternating
compression
amd tension
Ophiolites normally
originate here!
Tension
Trench
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. Subduction
roll-back.
High heat-flow in the supra-subductions zone regime give rise to
relatively low shallow sealevel above the back-arc basins. Most
ophiolite complexes have their origin is a supra-subduction environment
3 - D MORPHOLOGY
SEISMICITY
NB! NOTICE INTRA-SLAB
EARTHQUAKES
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 between
approximately 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.