20/11/13

13

25 Mas

ter1

Rés

ervo

irs

Géol

ogiq

ues Dyn

amique

des

Bas

sins

- M

iche

l Sér

anne

Stretched lithosphere & mantle denudation

Froitzheim & al

Fossil Non-volcanic Margins inverted and exposed in orogens

Ex: Albian Pyrenean basin

Lagabrielle & Bodinier 2008

corridor along the NFP. In the thirdcase, the polymictic clastic formationreworking the mantle debris would becompletely disconnected from its ori-ginal basement of unknown composi-tion.

Geodynamic implications andconclusions

Our new field observations in theregion of Etang de Lherz bear impor-tant constraints for the mode ofemplacement of the Pyrenean lherzo-lites. Following Choukroune (1973,1974, 1980), we confirm that thelherzolite bodies around Etang deLherz are enclosed within a sedimen-

tary clastic sequence resulting fromthe accumulation of debris reworkedfrom exposures of platform carbon-ates and ultramafic rocks. This inter-pretation is consistent with field datain the Bearn region where lherzoliticbodies also form restricted exposuresregarded as olistholiths emplacedwithin flysch formations of Creta-ceous age (Duee et al., 1984; Fortaneet al., 1986). In that sense, the sedi-mentary sequence of the Aulus basincan be compared to the Cretaceous-Eocene successions of the NorthernApennines where gravity depositsincluding ophiolitic debris flows(olistostromes) and olistoliths cropout extensively (Abbate et al., 1970;

Marroni and Pandolfi, 2001). Thelargest ophiolitic olistoliths of theApennines are 1–2 km long and 200–300 m thick, a size similar to that ofthe Lherz body. Gravity emplacedultramafic-rich bodies showing char-acteristics similar to the Lherz poly-mictic clastic sequences have been alsoreported from various orogenic beltssuch as the Californian Coast Ranges(e.g. Lockwood, 1971), or the internalAlps (Deville et al., 1992).From these observations, it is clear

that subcontinental mantle has beenexposed on the floor and ⁄or along theflanks of a deep, tectonically activebasin that now forms the Aulusregion. At present, exhumation ofmantle rocks is known to occur atthe foot of non-volcanic continentalmargins such as the Galicia-IberiaAtlantic margin (Boillot et al., 1980,1985; Abe, 2001; Whitmarsh et al.,2001; Manatschal, 2004), along theaxial reliefs of slow-spreading ridges(Lagabrielle et al., 1998; Karsonet al., 2006), within small oceanicbasins (Seyler and Bonatti, 1988), oralong the walls of oceanic fracturezones (Bonatti et al., 1971; Auzendeet al., 1989). In these different geody-namic settings, mantle exhumation isalways accompanied by sedimentsyielding large volumes of debris ofdominantly ultramafic composition,ranging from fine-grained turbiditesto debris-flows and rock slides, as firstreported on the flanks of the GorringeBank (Lagabrielle and Auzende,1982). In the Pyrenean case, mantleexhumation cannot be viewed as aprocess linked to the opening of awide ocean, as no relicts of oceaniclithosphere are present within themountain belt. In contrast to theTethyan ophiolites, the low degree ofpervasive serpentinization of most ofthe Pyrenean lherzolites suggests veryrapid exhumation followed by instan-taneous sedimentary reworking andburial within detrital sequences. Thisis consistent with transtensive condi-tions at the foot of a rapidly stretchedcontinental crust. This scenario isillustrated in Fig. 9. It may haveoccurred during the opening of aseries of pull-apart basins along theIberia ⁄Europe plate boundary due tooceanic spreading in the Bay of Bis-caye during the Albian (Le Pichonet al., 1970; Choukroune and Mat-tauer, 1978; Olivet, 1996).

Fig. 7 Antophyliite- and vermiculite-rich sandstones, polymictic and ultramaficbreccias exposed along the northern limit of the Lherz body at Site LHZ 7 (seeFig. 3 for location).

Fig. 8 Cartoon depicting possible geological setting and mode of emplacement of theultramafic bodies and associated clastic rocks.

Cretaceous exhumation of pyrenean mantle • Y. Lagabrielle and J.-L. Bodinier Terra Nova, Vol 20, No. 1, 11–21.............................................................................................................................................................

18 ! 2008 Blackwell Publishing Ltd

Ex: Tethysian margin

26 Mas

ter1

Rés

ervo

irs

Géol

ogiq

ues Dyn

amique

des

Bas

sins

- M

iche

l Sér

anne

Continent-Ocean Boundary

PERON-PINVIDIC, Gwenn (2006)

• Mantle exhumation & serpentinisation • End of rifting (no more faulting)

•Break-up: 1st accretion •Formation of 1st magnetic anomalies,

•Oc.crust accretion •Drifting

•Conjugate margins

Upper plate (Newfoundland)

lower plate (Iberia)

20/11/13

14

27 Mas

ter1

Rés

ervo

irs

Géol

ogiq

ues Dyn

amique

des

Bas

sins

- M

iche

l Sér

anne

Continent-Ocean Boundary : S. Atlantic

Unternher & al 2009

28 Mas

ter1

Rés

ervo

irs

Géol

ogiq

ues Dyn

amique

des

Bas

sins

- M

iche

l Sér

anne

Break-up

-  Lithospheric rupture -  End of extensional faults activity -  Onset of oceanic accretion -  First oceanic crust magnetic anomaly -  Two distinct continental margins

20/11/13

15

29 Mas

ter1

Rés

ervo

irs

Géol

ogiq

ues Dyn

amique

des

Bas

sins

- M

iche

l Sér

anne

Break-up unconformity

Break-up unconformity : signature of stress release

30 Mas

ter1

Rés

ervo

irs

Géol

ogiq

ues Dyn

amique

des

Bas

sins

- M

iche

l Sér

anne

Break-up = isostatic flexure of rift basins

lithosphere structure => depth of necking => isostatic flexural response => vertical movement => erosion

GSA TODAY, JANUARY 2006 5

GRAVITY ANOMALIES AND TeGravity anomalies, especially their

departures from local isostatic models (e.g., Airy, Pratt), have long played a key role in the debate concerning the strength of the lithosphere. Modern iso-static studies follow either a forward or inverse modeling approach. In forward modeling, the gravity anomaly due, for example, to a surface (i.e., topographic) load and its flexural compensation is calculated for different values of Te and compared to the observed gravity anom-aly. The best fit Te is then determined as the one that minimizes the difference between observed and calculated grav-ity anomalies. In inverse (e.g., spectral) models, gravity and topography data are used to estimate Te directly by com-puting the transfer function between them as a function of wavelength and comparing it to model predictions. The different approaches should yield the same results.

Figure 1. Schematic diagram illustrating different models for the long-term strength of continental lithosphere. In the crème-brûlée model, the strength is confined to the uppermost brittle layer of the crust, and compensation is achieved mainly by flow in the weak upper mantle. In the jelly sandwich model, the mantle is strong and the compensation for surface loads occurs mainly in the underlying asthenosphere. (A) Models of deformation. Arrows schematically show the velocity field of the flow. (B) Brace-Goetze failure envelopes for a thermotectonic age of 150 Ma, a weak, undried granulite lower crust, a uniform strain rate of 10−15 s−1, and either a dry (jelly sandwich) or wet (crème brûlée) olivine mantle. Hm is the short-term mechanical thickness of the lithosphere; Te is the long-term elastic thickness. Other parameters are as given in Tables 1 and 2. The two envelopes match those in Figures 5B and 5D of Jackson (2002). They yield a Te of 20 km (e.g., Burov and Diament, 1995), which is similar to the thickness of the most competent layer. This is because the competent layers are mechanically decoupled by weak ductile layers and so the inclusion of a weak lower crust or strong mantle contributes little to Te. (C) Brace-Goetze failure envelopes for a thermotectonic age of 500 Ma. Other parameters are as in (B) except that a strong, dry, Maryland diabase has been assumed for the lower crust. The two envelopes show other possible rheological models: in one, the upper and lower crusts are strong and the mantle is weak (upper panel); in the other, the upper and lower crusts and the mantle are strong (lower panel). The assumption of a strong lower crust in the weak mantle model again contributes little to Te because of decoupling, although Te would increase from 20 to 40 km if the upper crust was strong at its interface with the lower crust. In contrast, a strong lower crust contributes significantly to the Te of the strong mantle model. This is because the lower crust is strong at its interface with the mantle and so the crust and mantle are mechanically coupled.

TABLE 1. SUMMARY OF THERMAL AND MECHANICAL PARAMETERS USED IN MODEL CALCULATIONS

Thermal Surface temperature (0 km depth) 0 °C Temperature at base of thermal lithosphere 1330 °C Thermal conductivity of crust 2.5 Wm–1 °C–1

Thermal conductivity of mantle 3.5 Wm–1 °C–1

Thermal diffusivity of mantle 10–6 m2 s–1

Radiogenic heat production at surface 9.5 × 10–10 W kg–1

Radiogenic heat production decay depth constant 10 km Thermo-tectonic age of the lithosphere 150 Ma (Fig. 1B); 500 Ma (Fig. 1C) Mechanical Density of upper crust 2700 kg m–3

Density of lower crust 2900 kg m–3

Density of mantle 3330 kg m–3

Density of asthenosphere 3310 kg m–3

Lamé elastic constants !, G (here, ! = G) 30 GPa Byerlee’s law—Friction angle 30° Byerlee’s law—Cohesion 20 MPa

TABLE 2. SUMMARY OF DUCTILE PARAMETERS ASSUMED IN MODEL CALCULATIONS* Composition Pre-exponential

stress constant, A MPa–n s–1

Power law exponent,

n

Activation energy, QKJ mol–1

Figure 1

Uppercrust

Wet quartzite 1.1 × 10–4 4 223 B, C

Lowercrust

Dry Maryland diabase

8 ± 4 4.7 ± 0.6 485 ± 30 C

Undried Pikwitonei granulite

1.4 × 104 4.2 445 B

Mantle Dry olivine 4.85 × 104 3.5 535 B, C (jelly sandwich) Wet olivine 417 4.48 498 B, C (crème brûlée) *The failure envelopes in this paper match those in Jackson (2002); the Jackson (2002) envelopes are based on Figure 4 in Mackwell et al. (1998), who did not list all the parameters, referring instead to primary references. We therefore list here the parameters we have used.

Burov &Watts 2006 Cloetingh

Upper crust

Lower crust

Mantle

Upper crust

Lower crust

Mantle

« jelly sandwich »

« crème brulée »

Rifting stage End of rifting Lithosph.rheology

20/11/13

16

31 Mas

ter1

Rés

ervo

irs

Géol

ogiq

ues Dyn

amique

des

Bas

sins

- M

iche

l Sér

anne

post-rift

Moho

Syn-rift

Cont. Cr. thinned Ocean. Cr. Moho

from Meyers et al. 1996

50km 10km

Non-volcanic passive margin structure : Gabon

32 Mas

ter1

Rés

ervo

irs

Géol

ogiq

ues Dyn

amique

des

Bas

sins

- M

iche

l Sér

anne

Non-volcanic passive margin structure : Angola

Unternher & al 2009

20/11/13

17

33 Mas

ter1

Rés

ervo

irs

Géol

ogiq

ues Dyn

amique

des

Bas

sins

- M

iche

l Sér

anne

N S

Lower Cont.crust

Intermediate crust Cr.Oc.

Lithosphere mantle

ECORS

from Séranne, 1999

Non-volcanic Passive margin structure : Gulf of Lion

Low angle faults, extreme crust thinning, antithetic lower crust/mantle detachment = Lower crust and/or mantle denudation

34 Mas

ter1

Rés

ervo

irs

Géol

ogiq

ues Dyn

amique

des

Bas

sins

- M

iche

l Sér

anne

Synr

ift

Postrift

Prer

ift

time

Sedi

men

t th

ickn

ess fast subsidence

= synrift

postrift subsidence: cooling of the lithosphere

Previously thinned

syn-rift sediment

post-rift sediment

Continental rift-> passive margin : subsidence

20/11/13

18

35 Mas

ter1

Rés

ervo

irs

Géol

ogiq

ues Dyn

amique

des

Bas

sins

- M

iche

l Sér

anne

backstripping from Gabon Margin

Subsidence due to geodynamics

Total Subsidence = geodynamics + load

theoritical curve

Subsidence of passive margins

theoritical vs computed = geodynamic events to integrate & interpret

36 Mas

ter1

Rés

ervo

irs

Géol

ogiq

ues Dyn

amique

des

Bas

sins

- M

iche

l Sér

anne

Subsidence of passive margins

backstripping from East Coast of USA

Theoritical curve of thermal subsidence

Increase in subsidence=> phase of deformation ?

Backstripped curve = load effect removed (error bars : bathymetry, eustacy,…)

Time (Ma)

Dep

th(k

m)