Volcanic margins
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Tectonically active rifts and passive-margins form a group of genetically related extensional basins which play an important role in the spectrum of basin types. A large number of major hydrocarbon provinces are associated with passive-margins.
Two major end members of passive margin types are present on Earth: volcanic and non-volcanic.Volcanic margins form part of large igneous provinces[1], which are characterised by massive emplacements of mafic extrusives and intrusive rocks over very short time periods. Recent reviews illustrate the importance and the wide distribution of such “atypical margins” that may represent 75-90 % of the global continental passive margins
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[edit] General Concepts
Volcanic margins are known to differ from classical passive margins in a number of ways, the main ones being:
- a huge volume of magma forms during the early stages of crustal accretion along the future spreading axis, typically as seaward-dipping reflector sequences (SDRs)
- the sharp ocean-continent boundary (e.g., Greenland, U.S. East Coast, Namibia)
- the presence of numerous sill/dike and vent complexes intruding the sedimentary basin,
- the lack of strong passive-margin subsidence[2] during and after the continental breakup is usually observed
- the presence of a thick lower crust with anomalously high seismic P-wave velocities (7.1-7.8 km/s), so-called lower crustal bodies (LCBs).
The high velocities (Vp > 7 kms) and large thicknesses of the LCBs are often used to support the theory of “hot” mantle plume involvement leading to the formation of a huge volume of magmatic rocks. LCBs are often located along the continent-ocean transition but can extend beneath the continental part of the crust. In the continental domain, there are few constraints on their nature and chronology. Better constraints of the timing of LCB emplacement, seismic velocities, size and geological context are pertinent to the plume/non-plume debate. Even if mafic intrusion if one of the common explanation to explain the LCB, recent publications suggest that part of the LCB could also represent old metamorphic basement terranes.
Figure 1: Main characteristics of volcanic margins versus non-volcanic passive margins. (a) Schematic crustal section of a wide non-volcanic “Galician type” margin characterised by the progressive exhumation of the underlying seprentinized mantle. (b) Structure and main characteristics of a narrow volcanic “Vøring type” margin. CLCB: continental lower-crustal body; OLCB: oceanic lower-crustal body; SDRs: Seaward Dipping Reflector sequences. S symbolizes the post-breakup subsidence of the non-volcanic margin, U represents the relative uplift recorded along the volcanic margin as an isostatic consequence of thick high velocity underplating observed along the continent-ocean transition (COT). Source: L. Gernigon (NGU)[[3]]
[edit] Rifting and magmatism
Geophysical, petrological and numerical studies have advanced the understanding of rift-related magmatic processes. Numerical models, based on geophysical and geological data, have contributed at lithospheric and crustal scales towards the understanding of dynamic processes which govern the evolution of rifted basins and volcanic margins. The intensity and timing of volcanic activity in rifts is highly variable. Many rifts are totally devoid of volcanic rocks or show only a very low level of volcanism. Other rifts display a high level of volcanic activity, sometimes shortly after the onset of crustal extension (e.g. Oslo Graben, Red Sea; East African rift system) Rifts can also become temporarily volcanic after an initial stage of non-volcanic subsidence (e.g. North Sea Central Graben). Major variations in the intensity and chronology of volcanic activity is also evident during the rifting stage of passive margins. For instance, the early rifting stage of the Labrador Sea was accompanied by a high level of volcanic activity, while the Arctic-North Atlantic rift remained non-volcanic during its Late Carboniferous to Cretaceous evolution until the latest Cretaceous- Palaeocene last rifting phase, giving rise to the North Atlantic volcanic margins initiation during the breakup between Eurasia and Greenalnd. Similarly the end Early Jurassic crustal separation in the Central Atlantic was preceded by a short pulse of flood-basalt extrusion whereas Mid-Cretaceous opening of the southern South Atlantic was preceded and accompanied by the extrusion of the Etendeka and Paraná flood basalts. However, the development of other passive margins was not accompanied by major syn-rift volcanism (e.g. Galician Margin, South Rockall-Hatton Margin).
Volcanic rocks associated with intra-continental rifts display a typically alkaline, mafic-felsic bi-modal composition. Mafic melts appear to be generally derived from an incompatible element- enriched mantle source, residing presumably in the mantle-lithosphere, the depleted asthenosphere and/or within contreversial "mantle plumes"[4]. Initial magma generation in intra-continental rifts generally occurs in the 100 - 200 km depth range, corresponding to the lower parts of the lithosphere and the upper asthenosphere. During the evolution of some rifts a decrease in alkalinity of the extruded mafic magmas and an increasing contribution of MORB-source melts (depleted mantle) can be recognized, both in time and generally towards the rift axis. This can be attributed to an increasing contribution of melts from the asthenosphere as the lithosphere is progressively thinned.
Melt contributions from deep, more primitive mantle sources and/or the boundary layer between the upper and lower mantle , appear to be lacking in many rift-related volcanic suites trough traces of earlier hot-spot magmatism are sometimes observed. This raises serious doubts about the applicability of the deep mantle plume driven “active” rifting model to many examples of intra-plate rifting.
However, plume-related flood basalt provinces are characterizedby a wide range of geochemical and isotopic signatures reflecting mixing of partial melts derived from the upper mantle, the mantle-lithosphere and the original plume material.
The amount of melt generated during rifting depends to a large extent on the potential temperature of the asthenosphere with enhanced melt production reflecting higher potential temperatures[5]. Such domains correspond either to areas where deep "mantle plumes" have impinged on the lithosphere or which are underlain by small-scale convection cells. Surface manifestations of rift-related magmatism represent only a fraction of the total volume of melt generated during rifting. Rift-related magmatic processes affect also the mantle-lithosphere and the lower crust, leading to magmatic underplating. As such they have a profound effect on the evolution of rift zones during their tectonically active, as well as their post-rifting stage (low subsidence).
The volume and composition of melts generated along rift zone and/or volcanic margin is usually a function of the amount of lithospheric extension, the thermal state of the asthenosphere and lithosphere at the onset of extension, the presence of volatiles and the thickness of the lithosphere. Lithospheric stretching factors play an significant role by controlling the degree of adiabatic decompression of the lower lithosphere and the asthenosphere and the upwelling of the latter. Partial melting occurs when the upwelling material crosses the mantle solidus line (the position of which in P-T space is a function of its composition). In this respect it must be realized, that stress-induced extension of the lithosphere is likely to generate a greater degree of partial melting in areas underlain by anomalously hot or anomalous fertile mantle. Moreover, if very thick (~150 km), cold lithosphere is extended,little magmatism can be expected unless a very high degree of extension has occurred. Conversely, stretching of lithosphere having a thickness of some 100 km by a factor of 1.2-1.3 can already result in the onset of significant volumes of magma, rapidly dominated by an asthenospheric source. Strain rates also appear to play an important role in the volume of melts generated. At Low strain rates, conductive and convective heat diffusion probably plays an important role in suppressing partial melting. Then, in the presence of volatiles, the solidus is significantly lowered and partial melting can start at much smaller stretching factors than under anhydrous conditions.
[edit] Seaward Dipping Reflectors Sequences (SDRs)
Seaward-dipping reflectors Sequences(SDRs) represent flood basalts rapidly extruded during either rifting or initially subaerial sea-floor spreading. They form the main characteristics features of volcanic rifted margins.
Several different models have been proposed for the formation of the breakup-related volcanic extrusive complexes. These models can be divided into two main categories: those that relate the construction of the volcanic complex to infilling and capping of a rifted structure and those that are focused on the volcanics being formed during a phase of subaerial seafloor spreading.
Internal reflectors in the SDRs may represent ponded or thick lava units or very fragmented or hydroclastic units deposited in a shallow-marine environment. It has further been suggested that deep marine-emplaced flood-basalt constructions may be imaged as SDRs (outer SDRs) but this ,hypothesis cannot be addressed by the existing borehole data. Several other distinct seismic facies units are identified close to the SDRs. Outer High, and Landward Flows, Lava Delta, Inner Flows, and Volcanic Basin units are also identified along numerous volcanic rifted margins [6]
[edit] Lower Crustal Bodies (LCBs)
In the oceanic domain and along the continent-ocean transition, LCBs most likely represent mafic intrusion intruding or/and underplating the pre-existing crust. Fundemental questions and interpretation of the continental LCBs beneath rift zones of the outer is still controversial and interrogations remain:
- What do we really know about the geological meaning of the LCBs?
- Are continental LCBs really magmatic and mafic features?
- Are LCBs fully representative of breakup-related underplating?
- Do we necessarily need a mantle plume[7] to generate underplating?
- Do high P-wave velocities values really reflect hot mantle temperatures.
- What are the relationships between LCBs, basin deformation and continental breakup?
In the case of volcanic margins in general, the answers to these questions could lead to a better estimate of the total amount of melt produced during breakup. An improved volume estimate would have significant implications for quantification of mantle temperatures and dynamics, which are still poorly constrained.
In view of the high-velocity character of the lower crust and its position close to the SDRs, a mafic/ultramafic interpretation is usually proposes to explain the high Vp values observed along the breakup axis. High-Mg underplated bodies characterised by high-velocity lower crust should be a consequence of a “ hot mantle plume” impingement on the base of the lithosphere. Anomalous temperature is likely to produce High-Mg magma so called Picritic magma, underplating or intruding the preexsiting crust. It is, however, not clear if picritic magmas are really related to high mantle temperatures. Picritic magmas do not necessarily characterise high-degree or high-temperature melts but could simply be explained by extensive decompression of an uprising mantle (active or passive) and later differentiation.
Some geodynamic models can explain moderate amounts of pre-breakup magmatism without involving any mantle plume effect. Moderate temperature, fertile mantle patches (e.g. eclogites) in the upper mantle, and small-scale convection may explain significant pre- and syn-breakup melt production (see: http://www.mantleplumes.org/)[8]
Another geological model that may account for both the non-magnetic and high-velocity characteristics of the NGR lower crust is that this layer consists of pre-breakup crystalline rocks. The non-magmatic lower continental crust below the outer Vøring Basin is generally interpreted as granodioritic with P-wave velocities ranging between 6.5 and 7 km/s . However, it has been observed that the lower crust has locally higher velocities.
LCB may represent serpentinised mantle[9]. However, the highly saturated water condition required for this process is questionable and needs further supporting evidences.
High-pressure granulite/eclogitic material is known to have both high P-wave velocity (7.2-8.5 km/s) and high density (2.8-3.6 g/cm3). A LCB may be partly (or fully?) explained as pre-existing high-velocity, non-mafic metamorphic rocks such as eclogites[10] or migmatites].
[edit] World volcanic margins
- U.S East coast - Norwegian volcanic margin[11]
- East Greenland volcanic margin[12]
- Red Sea Volcanic margin
- Argentine volcanic margin
- Namibia volcanic margin
- Hatton volcanic margin
- Western Australian margin
- Yemen volcanic margin
[edit] References
Bauer, K., S. Neben, B. Schreckenberger, R. Emmermann, K. Hinz, N. Fechner, K. Gohl, A. Schulze, R.B. Trumbull, and K. Weber, Deep structure of the Namibia continental margin as derived from integrated geophysical studies. 2000. Journal of Geophysical Research, 105(B11), 25829-25853.
Boutillier, R.R. & Keen, C.E. 1999. Small-scale convection and divergent plate boundaries. Journal of Geophysical Research, 104, 7389-7403.
Eldholm, O. & Grue, K. 1994. North Atlantic volcanic margins: dimensions and production rates. Journal of the Geophysical Research, 99, 2955-2968.
Gernigon, L., Ringenbach, J.C., Planke, S., Le Gall, B. 2004. Deep structures and breakup along volcanic rifted margins: Insights from integrated studies along the outer Vøring Basin (Norway). Marine and Petroleum Geology, 21, 3, 363-372.
Menzies, M.A., Klemperer, S., Ebinger, C., & Baker, J. 2002. Characteristics of volcanic rifted margins. In: Menzies, M.A., Klemperer, S., Ebinger C. & Baker, J. (eds), Volcanic Rifted Margins. Geological Society of America Special Paper, 362, 14.
Planke, S., and Eldholm, O., 1994. Seismic response and construction of seaward dipping wedges of flood basalts: Vøring volcanic margin. Journal of the Geophysical Research, 99:9263-9278.
Skogseid, J., Planke, S., Faleide, J., Pedersen, T., Eldholm, O., & Neverdal, F. (2000). NE Atlantic continental rifting and volcanic margin formation. In: Nøttvedt, A., Larsen, B.T., Olaussen, S., Tørudbakken, B., Skogseid, J., Gabrielsen, R.H., Brekke, H. & Birkeland, Ø. (eds.), Dynamics of the Norwegian Margin. Geological Society of London, Special Publication 167, 295-326.
Van Wijk, J.W., Huismans, R.S., ter Voorde, M., Cloetingh, S.A.P.L. 2001. Melt generation at Volcanic Continental Margins: no need for a Mantle Plume?. Geophysical Research Letters, 28, 20, 3995-3998.
White, R.S., & McKenzie, D. 1989. The generation of volcanic continental margins and flood basalts. Journal of Geophysical Research, 94, 7685-7729.
[edit] External links
www.largeigneousprovinces.org/[13]
www.mantleplumes.org[14] --Barnabi 18:09, 26 November 2006 (UTC)
