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Geodynamic processes and deformation in orogenic belts
Tectonophysics, 63 (1980) 261-273 @ Elsevier Scientific Publishing Company,
261 Amsterdam
- Printed
in The Netherlands
GEODYNAMIC PROCESSES AND DEFORMATION IN OROGENIC BELTS
JOHN G. DENNIS
and WOLFGANG
R. JACOBY
California St&e U~i~ersity, Long Beach, C&f. 90840 institut fiir ~eteoro~og~e und Geophysik der Johann Frankfurt/Main (Germany)
(U.S.A.) Wolfgang Goethe Uniuersit~t,
ABSTRACT Dennis, J.G. and Jacoby, W.R., 1980. Geodynamic processes and deformation in orogenie belts. In: M.R. Banks and D.H. Green (Editors), Orthodoxy and Creativity at the Frontiers of Earth Sciences (Carey S~posium). Tectonophysies, 63: 261-273. The development of geosynclines and orogenie belts is related to lithosphere convergence Initial sediment accumulation implying subsidence, and volcanic activity implying extension and rise of geotherms, are in most eases followed by folding and thrusting suggesting compression, and by uplift. In terms of recent analogs, sediment accumulation and crustal extension are characteristic of back-arc spreading; subsequent compression would indicate continent-continent collision; and rise of geotherms most likely requires localized thermal flow (convection) in the asthenosphere. These events are here shown to agree with Andrews and Sleep’s (1974) numerical model of asthenosphere flow at converging plate margins. Orthogeosynclinal subsidence appears to be a consequence of subcrustal ablation and lithosphere extension and thinning in active marginal basins. Arc and Andean type magmatism mark the reappearance of ablated and transported, relatively low-density subcrustal material. Collision slows and eventually stops the local convection cell, resulting in local heat accumulation, and hence high-T, low-P metamorphism and granitization while marginal basin (orthogeosynelinal) deposits are being compressed into Alpine style erogenic structures. Moreover, closing of the marginal basin leads to subsidiary subduction, which in turn may be responsible for some Alpine style structures. Oceanic trench deposits may become incorporated in erogenic zones, as high-P, low-2 metamorphic belts (thalassogeosynclines). Dynamic uplift is a fundamental characteristic of orogeny. Most rising and sinking in erogenic zones can be iinked to those asthenosphere processes which are a consequence of Andrews-Sleep convection.
INTRODUCTION
Ever since the plate tectonic model of crustal evolution became widely accepted, there has been some controversy as to the role of orogeny in the evolution of converging plate margins. We shall try to show that erogenic
262
belts could be a consequence of asthenosphere processes near subduction zones, and that these processes are compatible, within limits, with certain orthodox geological concepts. Let us first review some of these concepts. OROGENY
AND GEOSYNCLINES
What exactly is orogeny? The simplest answer to this question is that orogeny is the process which results in the formation of erogenic belts. So we should really ask: What constitutes an erogenic belt? Let us then agree on some criteria which help to define them. The most striking eh~cte~st~c of erogenic belts is that they are elongated belts of intensely deformed rocks. The dominant deformation style consists of folds, and, in many cases, overthrusts or slides, usually facing in the direction of a stable continental craton or foreland. Next, there is always strong relief, either still present or deduced from evidence of extensive denudation. Varying volumes of granitic rocks are emplaced in nearly all orogens, commonly within extensive areas of regionally metamorphosed rocks. Furthermore, most of the sedimentary rocks involved attest to deep subsidence immediately preceding orogeny, and the site of deposition is, of course, what is traditionally described as a geosyneline. We realize that geosync~nes have been out of fashion in some quarters lately (cf. Dietz, 1972) but we believe that the concept remains valid and useful. What do we mean by “geosyncline ?” First of all we must recognize that “geosyncline” is a stratigraphic, not a structural concept, for former geosynchnes are reconstructed from essentially statigraphic evidence. What Stille (1941) called orthogeosynclines are loci of long continued, thick sediment deposition on mobile crust. They are elongated, usually along continental margins, and of limited extent in space and time. They normally comprise two longitudinal zones: one, the eugeosynclinal zone, with intercalated mafic and ultramafic rocks, and a filling of commonly immature sediments whose sources are outside (oceanward) or within the geosyncline; and another, the miog~sync~~ zone, devoid of seat rn~a~m, and filled with mostly mature sediments whose source is the adjacent craton. Marshall Kay (e.g., 1951) has elaborated this scheme, and has shown that the sources of sediment supplied to the Paleozoic eugeosynclines of North America were probably chains of volcanic islands frin@ng the continent (Fig. 1). Field evidence shows that deposition is in large part on older continental basement. There is evidence for lateral extension during the geosynclinal phase (e.g., Triimpy, 1975) partly from paleogeographic and paleogeological reconstruction, but in any case, contemporaneous volcanism indicates crustal extension. Following the geosynclinal phase, the same belt undergoes folding, thrusting, rne~~~~ and ~1~~~; in &her words, orogenesis. At this stage there is no evidence for extension. On the contaary, the s&%@ur& style changes to indicate local or regional compression, but this does not neces-
263
Fig. 1. Paiaeogeography of early medial Ordovician of North America; synthesis by Marshall Kay (1951). Note that this map was drawn based on the then current gumption of permanence of ocean basins.
sarily involve the whole crust, and interpretations involving primary vertical movement must be considered. Certain types of orogeny seem to have occurred at a time of inferred cont~ent~ont~ent collision (e.g., Dewey and Bird, 1970), suggesting a mechanism for compression. These processes were first summarized by Stille (1941) in his geotectonicmagmatological cycle, which has four stages: (1) Geosynclinal subsidence with “initial” magmatism, that is emplacement of ophiolites and mafic volcanics. (2) Folding of the geosynclinal fill, accompanied by emplacement of predominantly granitic rocks (orogenesis and “synorogenic” m~at~m). (3) Consolidation of the hitherto mobile zone, accompanied by emplacement of predominantly granitic plutons (“subsequent” magmatism).
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(4). Final cratonization commonly accompanied by mafic volcanism (“final” magmatism). AN ACTUALISTIC
PERSPECTIVE
Since it is our aim to connect deformation in erogenic belts with asthenosphere processes, we must now assume an actualistic perspective. What present-day tetonic processes can be related to the features and processes described so far? Naturally, we will turn to present zones of plate convergence: they are well-known zones of mobility, with active subduction and, in most cases, an active volcanic arc. In addition, we find small ocean basins behind some volcanic arcs, the marginal basins. Carey (1970) reeognized the extensional origin of these basins. There is evidence for their formation by creation of new oceanic crust. Rifting within them, with extension perpendicular to the adjacent island arcs is evidenced by fault scarps, thin sediment cover in the rift structures or fresh tholeiitic vofcanics exposed on the sea floor, high heat flow (Karig, 1971; Packham and Falvey, 1@71), as well as mostly irregular, but sometimes regular identifiable magnetic anomalies (e.g. Barker, 1972). The consequence of basin spreading is migration of island arcs away from the continent of which they may once have been a part (Matsuda and Uyeda, 1971). There is considerable evidence today that the upper mantle below marginal basins is anomalous. It is characterized by extremely low Q (Barazangi and Isacks, 1971), by low seismic velocities, particularly of S waves, and by low densities (Karig, 1971; Jacoby, 1975). These observations point to
-FHw A-
----_----_ =----;7 T
-I_-_-_-_-_
Hs i
(a)
fb)
Fig. 2. Density relations in marginal basin~subaidence (see text). a. Initial state. b. Final State. Horizontal dashed pattern = water; dotted pattern = sediments; vertical ruling = mantle; blank = crust.
265
extensive partial melting, high temperatures, and possible emplacement of low-density mantle material. We suggest that active marginalbasins of this kind have played an important role in the development of geosynclinesand erogenic belts. Large volumes of detrital sediment have been supplied from the “outer” or oceanic side, throughout the development of geosynclines (Kay, 1951). Atlantic-type continental margins cannot explain this easily except for the early stages of rifting and for the final stagesof continental collision. In contrast, marginal basins remain close to, and are surrounded by, sediment sources and are natural sinks for island arc products. Men~d (1967) has pointed out that the greatestthicknessesof sediment in the present oceans occur in present marginalbasins. A simple isostatic argument shows that sediment thicknessesin marginal basins may easily reach classical geosynclinal dimensions (see also Hsii, 1958). If we postulate essentialmass balance between initial, sedimentfree ocean basin and final, s~iment-~1~ basin (Fig. 2) we have: ori
where pw = density of water; ps = density of sediments; pm = density of mantle; h, = depth of water; h, = thicknessof sediments. The initial water depth can be magnified by a factor of 2 or 3 depending on the assumed sediment density. If we assume a basin depth of 5 km in a state of thermal equ~b~um after cooling from the originalspreadingphase, we can easily expect up to 15 km of sediments. Both eugeosynclinal and miogeosynclinal sequences appear to be largely deposited on sialic basement, while fully developed marginal basins are underlain by oceanic crust. This has been a great difficulty with plate tectonic models of orogeny. There is, however, evidence that in some cases the formation of marginalbasins involves necking and stretchingof continental crust. Refraction seismic data for some marginalbasinscan be interpretedas indicating portions of thinned continental crust, e.g. for the Sea of Japan (Beresnev et al., 1970; Kaseno, 1972). Summarizingsome of the principal features of geosynclines and marginalbasins we note certain analogies.Both features proved evidence for: (1) extension during fo~ation; (2) above normal heat flow and mafic volcanism; (3) subsidence leading to accumulation of great volumes of sediment; (4) sources of sedimentsoutside the continent itself. But we must mention some erogenic zones which differ in important respects from the prototype described so far: First, belts of high pressure, low temperature metamorphism (Miyashiro, 1961; Ernst, 1970) which commonly also include ophiolites and melanges. ~~cte~stic~y, they show no signs of contemporaneous graniteemplacement. These are the
266
thalassogeosynclines of Bogdanov (1969), and their actualistic counterparts are trench deposits. They are commonly associated with arc-trench gap type deposits (Dickinson, 1971). Then there are the Andean and arc style belts. Their backbone is a zone of vast volumes of mainly volcanic and plutonie rocks. Rock deformation in them can be adequately explained as a consequence of magmatic activity. Metamorphism is mostly low grade. Here magmatic activity can be documented for much of the lifespan of the associated subduction, though long gaps do exist. In contrast, plutonism associated with collision is of relatively short duration. ROLE OF THE ASTHENOSPHERE
Now we must examine the role of the asthenosphere. Essentially two types of models have been proposed to explain the high rate of heat flow and magmatic activity above dipping slabs: (1) frictional heating generated by the movement of the upper surface of the slab past virtually stagnant asthenosphere, with heat either conducted upward or convected by rising melts (e.g., Hasebe et al., 1970; Matsuda and Uyeda, 1971; Oxburgh and Turcotte, 1971); (2) secondary flow induced by the descending slab in the low Q wedge between the slab and the floating plate, with heat generated by viscous dissipation (Mackenzie, 1969) or advected by the induced flow which also thins the floating plate above (Sleep and Toksiiz 1971; Andrews and Sleep, 1974). The first model encounters serious difficulties, particularly if heat conduction is the dominant mechanism; in any case, it requires large stresses, hence high viscosities of mantle material to generate enough heat, and low viscosities to transport the heat to the surface fast enough, as pointed out by Andrews and Sleep (1974). The model, however, cannot be ruled out completely for all situations. The second model is more likely to be valid for marginal basins. Flow will occur in the wedge of asthenosphere, driven by some mechanical and/or thermal agent, if the viscosity is low enough. This is, in fact, implied by high temperatures below active marginal basins, inferred from heat flow data (Karig, 1971; Packham and Falvey, 1971). Viscosities are probably below 10zl poises. The subducting slab of lithosphere represents a moving boundary which then inevitably drives some flow by viscous drag. Furthermore, a strong lateral temperature variation is evident in the model which by itself must contribute to the driving mechanism. THE ANDREWS-SLEEP
MODEL
Our preferred model is that of Andrews and Sleep (1974), not only for its physical elegance and simplicity, but also because it provides some insight into a possible mechanism for the development of geosynclines and erogenic belts (Dennis, 1976). The model material has temperaturedependet viscos-
MSTANCE
267
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100
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2002
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TEMPERATURE 087
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MST= ”
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200
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.DEC MY
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Fig. 3. Stream function and temperature d~t~bution under active marginaI basins; model computed by Andrew8 and Sleep. From Audrews and Sleep (1974). a. 0.87 m.y. after initiation of subduction. b. 10 m.y. after initiation of subduction.
268
ity and yield strength. At the start, the top is cold high-viscosity lithosphere. A slab descending 10 cm/y at 45” forms the right-hand boundary (Fig. 3). Flow into and out of the wedge is imposed on the left-hand boundary and on the bottom as computed analytically for a wedge extending to infinity. The initial layered temperature distribution is in conductive equilibrium above 80 km and adiabatic below. As flow is initiated in the wedge, stress builds up in the floating plate with compression near the trench between the converging plates and tension at some distance from the trench. At 1 m.y. the lithosphere starts to neck at the point of first yielding. First the rise of the isotherms causes some uplift of the crust, but as necking proceeds and the tensile stresses are concentrated, true rifting starts with depression of the surface and formation of an oceanic marginal basin. Material at the bottom of the heated floating plate is entrained with the more rapid asthenospheric flow and is subducted with the slab, thus causing the floating plate to move toward the trench. Ablation is, however, rapid only during the first few million years; after adjusting to a new. temperature equilibrium the thinned floating plate is no longer strongly ablated (Fig. 4). The time may be longer than computed by Andrews and Sleep (1974) if the viscosities and the yield
Fig. 4. Evolution of-an active marginal hasin. Letters indicate individual particles, as diaplaced. Dotted pattern = lithosphere; solid black = zone of enrichment in low-density material. Based on Andrews an& Sbep ( 19?4), from Dennis ( 1976).
269
strength are lower than assumed, which we would prefer. Andrews and Sleep’s (1974) model would be modified somewhat if in the calculation the asthenosphere had a bottom. The effect would be the development of a series of eddies in the asthenosphere below the continent and the formation of undulations at the surface (Sleep and Toksiiz, 1971). The amplitudes of the undulations would be magnified by erosion and sedimentation and by the corresponding isostatic response (Walcott, 1970). It is the apparent contradiction of extensional marginal basins in an environment of convergent plate margins which suggests that basin closing and orogeny are the natural fate of geosynclines. Eventually, continental crust riding on the subducting oceanic pfate is rafted against the trench, the island arc, and the marginal basin. For as long as the descending slab remains attached to the continent and driving forces continue to act, collision will continue and the whole island-arc marginalbasin belt will be compressed. The buoyant forces on the subducted continent will grow until subduction must choke (McKenzie, 1969). Marginal basins also may be subducted during some relatively short stage of the collision. Thus deep thrusts may develop before converging continents are welded together. Some of the basement thrusts in the Alps and in other orogens may be examples, suggesting that it may not always be possible to deduce the sense of the main subduction from the vergence of exposed thrust faults. SOME CONSEQUENCES OF ANDREWS-SLEW
CONVENTION
In the beginning, as convection proceeds, the lithosphere above the eddy has a tendency to stretch and thin. Material from its lower layers is dragged along, causing sub-lithosphere ablation, and is finally subducted. High isotherms may reach the sialic crust and raise its temperature above the solidus of dry granite. Thus, partially molten granitic material may be dragged into the corner of the wedge: it cannot be subducted, because ofits low specific gravity. Hence, it will form a convenient source for relatively silicic magma, for arc magmatic activity and for related mineral deposits. Its buoyancy will ensure continuing uplift, and a continuous source of sedimentation. Extension and ablation of the ~thosphere lowers the surface to oceanic or near-oceanic depths, thus providing a receptacle for geosynclinal ~~en~tion, much of which is continuously supplied from the rising volcanic arc, that is, from outside the continent. A weaker, coupled, continentward directed convection cell may cause a subdued version of some of these effects. Secondary convection could concentrate low-density material wherever such material is trapped by appropriate topography on the underside of the lithosphere, resulting in isostatic uplift. This mechanism may explain the Rocky Mountains uplifts of the United States and their mineral deposits, and the Harz Mountains of central Germany. Now, how is a geosyncline or marginal basin transformed into an orogen?
270
Current evidence seems to show that “classical” orogenies, those that involve orthogeosynclines as previously defined, have coincided with continent-continent collision. Such collisions terminate subduction. The end of subduction, in turn, chokes Andrews-Sleep convection. The marginal basin and its filling are compressed, removing the former heat sink. Heat and fluid escape are blocked. Hence, the former geosynclinal zone now heats up, leading to regional metamorphism and granitization. Volcanism has ceased with the cessation of the tensional regime. Subduction of marginal basins and compression of their filling, incidentally, transform these fairly wide areas in plan into the relatively narrow erogenic belts we know. This evolution explains why nearly all classical eugeosynclinal successions end up in erogenic belts: sooner or later all marginal basin deposits must be affected by collision. On the other hand, it can happen that Andrews-Sleep convection does not open a marginal basin: Andean style orogens may originate in this way, where, apparently, continental lithosphere is too thick for the initial rifting which elsewhere develops into a marginal basin. RELIEF
Dynamic relief is a prominent feature of erogenic belts. It provides persisting sources of immature sediments, and is evident in all young mountain ranges. To create relief, we need to add material of relatively low density for practical purposes, that means enriched in silica - at the base of the crust. There are four basic mechanisms for doing this: underthrusting of continental crust; compression-thickening (Molnar and Tapponnier, 1975); lateral transfer of low-density, low-viscosity material at the base of the lithosphere by special convection such as Andrews-Sleep convection; and, finally, mineral transformations, such as feldspathization, which evidently proceed after collision has throttled the Andrews-Sleep convection cell. After collision, thermal inertia may drive modified relict Andrews-Sleep convection and this could be responsible for late erogenic uplift coupled with foredeep subsidence. RELATED
PHENOMENA
No model can explain everything, especially in geology. But a useful model should at least help in gaining new perspectives on some old problems. For instance, the geologic record in some collision orogens leaves little evidence for a former vokanic arc: there is no clear sign of such an are in the Alps; but there appear to be vestiges of former oceanic crust, such as the Platta nappe. There is much evidence that the Tethys never opened very far in the area of the present-day Alps (T’riimpy, 1976). If arc volcanism were derived from the descending slab, it &or&l be ubiquitous in all cases of lithosphere subduction; but if it is a result of Andrews-Sleep convection, it could quite conceivably be impeded by some special circumstances attending the
271
closing of small oceanic basins. In fact, if asthenosphere eddies are mainly responsible for orogeny, then it is, of course, conceivable that some such eddies could have been initiated by processes other than classical lithosphere subduction. This might be the case for erogenic belts between cratons that show no relative displacement, such as the Damara orogen in southern Africa. When subduction gives way to strike-slip displacement, as in the case of the San Andreas Fault, thermal inertia could account for continuing, modified relict Andrews-Sleep convection. This may be a factor in the development of Basin and Range volcanism and tectonics in the western United States. SUMMARY
Let us now summarize how our model of orogeny works. First, the lithosphere ruptures near a continental margin, and the oceanic portion, already depressed by continental margin sediments, descends into the asthenosphere, forming a trench. Below the trench, the crust is under compression. The motion of the descending slab then initiates Andrews-Sleep convection. As a result, the floating margin is stretched under tension at right angles to the trench. Silica enriched partially molten material now begins to accumulate in the corner of the asthenosphere wedge in front of the descending slab, causing isostatic uplift above, and intrusion as well as extrusion of part of the accumulated low-density wedge material, thus forming a volcanic ridge. The evolution of this ridge will result in an Andean or arc-style orogen. Continued tension induced in the lithosphere by the convection may open a marginal basin, which becomes filled with sediments derived both from the continental interior, forming the miogeosynclinal zone, and from the volcanic arc, forming the eugeosynclinal zone. Volcanics derived from basinal rifting are interbedded with the sediments of the eugeosynclinal zone. At the same time, the trench fills with sediments transported oceanward from the arc by different mechanisms, including sliding and flowage. The trench fill becomes the thalassogeosyncline. These processes continue until a continent arrives at the trench on the back of the subducting plate. The continent cannot be subducted, so Andrews-Sleep convection must eventually stop. Heat which so far has been able to escape through the marginal basin floor and the slab is impounded and raises the temperature of the basement and fill of the marginal basin, while the basin itself is under compression and is gradually closing. Under these circumstances, instead of escaping, the trapped heat raises the temperature in surrounding rocks, causing regional metamorphism and granitization while the crustal segment affected is under compression. Where previously sediments had accumulated in a subsiding basin, there is now uplift due to crustal shortening and to volume increase both by raised temperatures and by mineral transformations. The extent of
212
sub-lithosphere mass transfers of low-density material at this stage is uncertain, but they probably play a role also. CONCLUSION
This model, therefore, is compatible with Stille’s classical geotectonic cycle, if limited to marginal basin-collision type orogens. We must, in fact, distinguish clearly between marginal basin and Andean or arc style erogenic zones, a distinction foreshadowed by Argand (Aubouin, 1972) many years ago, but which certainly needs further elaboration in the light of new geodynamic concepts. ACKNOWLEDGMENTS
We are indebted to W.H. Bucher, S.W. Carey, and H. Ramberg for kindling our interest in this topic, and to D.J. Andrews, G. Angenheister, J.F. Dewey, H. Neugebauer and H.G.F. Win&r for most enlightening discussions. This work owes much to the incentive of the Carey Symposium, and the manuscript was significantly improved by participation of one of us (J.G.D.), supported by grants from the Organizers and from California State University, Long Beach. REFERENCES Andrews, D.J. and Sleep, N.H., 1974. Numerical modeling of tectonic flow behind island arcs. Geophys. J., 38: 237-251. Aubouin, J., 1972. Chafnes liminaires (Andines).et chafnss geosynclinales (Alpines). Int. Geol. Congr., 24th. Montreal, Proc., 3: 438-461. Barazangi, M. and Isacks, B., 1971. Lateral variations of seismic wave attenuation in the upper mantle above the inclined earthquake zone of the Tonga island arc: deep anomaly in the upper mantle. J. Geophys. Res., 76: 8493-8516. Barker, P.E., 1972. A spreading center in the east Scotia Sea. Earth Planet. Sci. Lett., 15: 123-132. Beresnev, A.F., Gainanov, A.G., Kovylin, V.M. and Stroev, P.A., 1970. Structure of the earth’s crust and upper mantle under the Japan Sea and the Pacific Ocean adjoining Japan. In: Problems of Structure of the Earth’s Crust and Upper Mantle, Nauka, Moscow, pp. 50-61 (in Russian). Bogdanov, N.A., 1969. Thalassogeosynclines of the circumpacific ring. Geotectonics, 3: 141-147. Carey, S.W., 1970. Australia, New Guinea and Meianesia in the current revolution in concepts of the evolution of the earth. Search, l(5): 178-189. Z. Dtsch. Geol. Dennis, J.G., 1976. Geosynklinale, Orogenese und Piattentektonik. Ges., 127: 73-85. Dewey, J.F. and Bird, J.M., 1970. Mountain belt and the new global tectonics. J. Geophys. Res., 75: 2626-2647. Dietz, R.S., 1972. Geosynclines, mountains, and continent building. Sci. Am., 226: 3038. Dickinson, W.R., 1971. Clastic sedimentary sequences deposited in shelf, slope, and trough settings between magmatic ares and associated trenches. Pacific Geol., 3: 1530.
273 Ernst, W.G., 1970. Tectonic contact between the Franciscan melange and the Great Valley sequence, crustal expression of a late Mesozoic Benioff zone. J. Geophys. Res., 75: 886-902. Hasebe, K., Fujii, N. and Uyeda, S., 1970. Thermal processes under island arcs. Tectonophysics, 10: 335-355. Hsii, K.J., 1958. Isostasy and a theory for the origin of geosynclines. Am. J. Sci., 256: 305-327. Jacoby, W.R., 1975. Velocity-density systematics from seismic and gravity data. Zentral Inst. Phys. Erde, Verijffentl., 31: 323-333. Karig. D.E., 1971. Origin and development of marginal basins in the western Pacific. J. Geophys. Res., 76: 2542-2561. Kaseno, Yoshio, 1972. Geological features of the Japan Sea floor. Pacific Geol., 4: 91111. Kay, M., 1951. North American geosynclines. Geol. Sot. Am. Mem., 48: 143 pp. Matsuda, T. and Uyeda, S., 1971. On the Pacific-type orogeny and its model - extension of the paired belts concept and possible origin of marginal seas. Tectonophysics, 11: 5-27. McKenzie, D., 1969. Speculations on the consequences and causes of plate motions. Geophys., J., 18: l-32. Menard, H.W., 1967. Transitional types of crust under small ocean basins. J. Geophys. Res., 72: 3061-3073. Miyashiro, A., 1961. Evolution of metamorphic belts. Petrology, 2: 277-311. Molnar, P. and Tapponnicr, P., 1975. Cenozoic tectonics of Asia: effects of a continental collision. Science, 189 : 419-426. Oxburgh, E.R. and Turcotte, D.L., 1971. Origin of paired metamorphic belts and crustal dilation in island arc regions. J. Geophys. Res., 76: 1315-1327. Packham, G.H. and Falvey, D.A., 1971. An hypothesis for the formation of marginal seas in the western Pacific: Tectonophysics, 11: 79-110. Sleep, N.H. and ToksSz, N., 1971. Evolution of marginal basins. Nature, 233: 548-550. Triimpy, R., 1975. Penninic-Austroalpine boundary in the Swiss Alps: a presumed former continental margin and its problems. Am. J. Sci., 275-A: 209-238. Stille, H., 1941. Einfiihrung in den Bau Amerikas. Berlin, Borntraeger (1940), 717 pp. Walcott, RI., 1970. Flexural rigidity, thickness, and viscosity of the lithosphere. J. Geophys. Res., 75: 3941-3954.
261 Amsterdam
- Printed
in The Netherlands
GEODYNAMIC PROCESSES AND DEFORMATION IN OROGENIC BELTS
JOHN G. DENNIS
and WOLFGANG
R. JACOBY
California St&e U~i~ersity, Long Beach, C&f. 90840 institut fiir ~eteoro~og~e und Geophysik der Johann Frankfurt/Main (Germany)
(U.S.A.) Wolfgang Goethe Uniuersit~t,
ABSTRACT Dennis, J.G. and Jacoby, W.R., 1980. Geodynamic processes and deformation in orogenie belts. In: M.R. Banks and D.H. Green (Editors), Orthodoxy and Creativity at the Frontiers of Earth Sciences (Carey S~posium). Tectonophysies, 63: 261-273. The development of geosynclines and orogenie belts is related to lithosphere convergence Initial sediment accumulation implying subsidence, and volcanic activity implying extension and rise of geotherms, are in most eases followed by folding and thrusting suggesting compression, and by uplift. In terms of recent analogs, sediment accumulation and crustal extension are characteristic of back-arc spreading; subsequent compression would indicate continent-continent collision; and rise of geotherms most likely requires localized thermal flow (convection) in the asthenosphere. These events are here shown to agree with Andrews and Sleep’s (1974) numerical model of asthenosphere flow at converging plate margins. Orthogeosynclinal subsidence appears to be a consequence of subcrustal ablation and lithosphere extension and thinning in active marginal basins. Arc and Andean type magmatism mark the reappearance of ablated and transported, relatively low-density subcrustal material. Collision slows and eventually stops the local convection cell, resulting in local heat accumulation, and hence high-T, low-P metamorphism and granitization while marginal basin (orthogeosynelinal) deposits are being compressed into Alpine style erogenic structures. Moreover, closing of the marginal basin leads to subsidiary subduction, which in turn may be responsible for some Alpine style structures. Oceanic trench deposits may become incorporated in erogenic zones, as high-P, low-2 metamorphic belts (thalassogeosynclines). Dynamic uplift is a fundamental characteristic of orogeny. Most rising and sinking in erogenic zones can be iinked to those asthenosphere processes which are a consequence of Andrews-Sleep convection.
INTRODUCTION
Ever since the plate tectonic model of crustal evolution became widely accepted, there has been some controversy as to the role of orogeny in the evolution of converging plate margins. We shall try to show that erogenic
262
belts could be a consequence of asthenosphere processes near subduction zones, and that these processes are compatible, within limits, with certain orthodox geological concepts. Let us first review some of these concepts. OROGENY
AND GEOSYNCLINES
What exactly is orogeny? The simplest answer to this question is that orogeny is the process which results in the formation of erogenic belts. So we should really ask: What constitutes an erogenic belt? Let us then agree on some criteria which help to define them. The most striking eh~cte~st~c of erogenic belts is that they are elongated belts of intensely deformed rocks. The dominant deformation style consists of folds, and, in many cases, overthrusts or slides, usually facing in the direction of a stable continental craton or foreland. Next, there is always strong relief, either still present or deduced from evidence of extensive denudation. Varying volumes of granitic rocks are emplaced in nearly all orogens, commonly within extensive areas of regionally metamorphosed rocks. Furthermore, most of the sedimentary rocks involved attest to deep subsidence immediately preceding orogeny, and the site of deposition is, of course, what is traditionally described as a geosyneline. We realize that geosync~nes have been out of fashion in some quarters lately (cf. Dietz, 1972) but we believe that the concept remains valid and useful. What do we mean by “geosyncline ?” First of all we must recognize that “geosyncline” is a stratigraphic, not a structural concept, for former geosynchnes are reconstructed from essentially statigraphic evidence. What Stille (1941) called orthogeosynclines are loci of long continued, thick sediment deposition on mobile crust. They are elongated, usually along continental margins, and of limited extent in space and time. They normally comprise two longitudinal zones: one, the eugeosynclinal zone, with intercalated mafic and ultramafic rocks, and a filling of commonly immature sediments whose sources are outside (oceanward) or within the geosyncline; and another, the miog~sync~~ zone, devoid of seat rn~a~m, and filled with mostly mature sediments whose source is the adjacent craton. Marshall Kay (e.g., 1951) has elaborated this scheme, and has shown that the sources of sediment supplied to the Paleozoic eugeosynclines of North America were probably chains of volcanic islands frin@ng the continent (Fig. 1). Field evidence shows that deposition is in large part on older continental basement. There is evidence for lateral extension during the geosynclinal phase (e.g., Triimpy, 1975) partly from paleogeographic and paleogeological reconstruction, but in any case, contemporaneous volcanism indicates crustal extension. Following the geosynclinal phase, the same belt undergoes folding, thrusting, rne~~~~ and ~1~~~; in &her words, orogenesis. At this stage there is no evidence for extension. On the contaary, the s&%@ur& style changes to indicate local or regional compression, but this does not neces-
263
Fig. 1. Paiaeogeography of early medial Ordovician of North America; synthesis by Marshall Kay (1951). Note that this map was drawn based on the then current gumption of permanence of ocean basins.
sarily involve the whole crust, and interpretations involving primary vertical movement must be considered. Certain types of orogeny seem to have occurred at a time of inferred cont~ent~ont~ent collision (e.g., Dewey and Bird, 1970), suggesting a mechanism for compression. These processes were first summarized by Stille (1941) in his geotectonicmagmatological cycle, which has four stages: (1) Geosynclinal subsidence with “initial” magmatism, that is emplacement of ophiolites and mafic volcanics. (2) Folding of the geosynclinal fill, accompanied by emplacement of predominantly granitic rocks (orogenesis and “synorogenic” m~at~m). (3) Consolidation of the hitherto mobile zone, accompanied by emplacement of predominantly granitic plutons (“subsequent” magmatism).
264
(4). Final cratonization commonly accompanied by mafic volcanism (“final” magmatism). AN ACTUALISTIC
PERSPECTIVE
Since it is our aim to connect deformation in erogenic belts with asthenosphere processes, we must now assume an actualistic perspective. What present-day tetonic processes can be related to the features and processes described so far? Naturally, we will turn to present zones of plate convergence: they are well-known zones of mobility, with active subduction and, in most cases, an active volcanic arc. In addition, we find small ocean basins behind some volcanic arcs, the marginal basins. Carey (1970) reeognized the extensional origin of these basins. There is evidence for their formation by creation of new oceanic crust. Rifting within them, with extension perpendicular to the adjacent island arcs is evidenced by fault scarps, thin sediment cover in the rift structures or fresh tholeiitic vofcanics exposed on the sea floor, high heat flow (Karig, 1971; Packham and Falvey, 1@71), as well as mostly irregular, but sometimes regular identifiable magnetic anomalies (e.g. Barker, 1972). The consequence of basin spreading is migration of island arcs away from the continent of which they may once have been a part (Matsuda and Uyeda, 1971). There is considerable evidence today that the upper mantle below marginal basins is anomalous. It is characterized by extremely low Q (Barazangi and Isacks, 1971), by low seismic velocities, particularly of S waves, and by low densities (Karig, 1971; Jacoby, 1975). These observations point to
-FHw A-
----_----_ =----;7 T
-I_-_-_-_-_
Hs i
(a)
fb)
Fig. 2. Density relations in marginal basin~subaidence (see text). a. Initial state. b. Final State. Horizontal dashed pattern = water; dotted pattern = sediments; vertical ruling = mantle; blank = crust.
265
extensive partial melting, high temperatures, and possible emplacement of low-density mantle material. We suggest that active marginalbasins of this kind have played an important role in the development of geosynclinesand erogenic belts. Large volumes of detrital sediment have been supplied from the “outer” or oceanic side, throughout the development of geosynclines (Kay, 1951). Atlantic-type continental margins cannot explain this easily except for the early stages of rifting and for the final stagesof continental collision. In contrast, marginal basins remain close to, and are surrounded by, sediment sources and are natural sinks for island arc products. Men~d (1967) has pointed out that the greatestthicknessesof sediment in the present oceans occur in present marginalbasins. A simple isostatic argument shows that sediment thicknessesin marginal basins may easily reach classical geosynclinal dimensions (see also Hsii, 1958). If we postulate essentialmass balance between initial, sedimentfree ocean basin and final, s~iment-~1~ basin (Fig. 2) we have: ori
where pw = density of water; ps = density of sediments; pm = density of mantle; h, = depth of water; h, = thicknessof sediments. The initial water depth can be magnified by a factor of 2 or 3 depending on the assumed sediment density. If we assume a basin depth of 5 km in a state of thermal equ~b~um after cooling from the originalspreadingphase, we can easily expect up to 15 km of sediments. Both eugeosynclinal and miogeosynclinal sequences appear to be largely deposited on sialic basement, while fully developed marginal basins are underlain by oceanic crust. This has been a great difficulty with plate tectonic models of orogeny. There is, however, evidence that in some cases the formation of marginalbasins involves necking and stretchingof continental crust. Refraction seismic data for some marginalbasinscan be interpretedas indicating portions of thinned continental crust, e.g. for the Sea of Japan (Beresnev et al., 1970; Kaseno, 1972). Summarizingsome of the principal features of geosynclines and marginalbasins we note certain analogies.Both features proved evidence for: (1) extension during fo~ation; (2) above normal heat flow and mafic volcanism; (3) subsidence leading to accumulation of great volumes of sediment; (4) sources of sedimentsoutside the continent itself. But we must mention some erogenic zones which differ in important respects from the prototype described so far: First, belts of high pressure, low temperature metamorphism (Miyashiro, 1961; Ernst, 1970) which commonly also include ophiolites and melanges. ~~cte~stic~y, they show no signs of contemporaneous graniteemplacement. These are the
266
thalassogeosynclines of Bogdanov (1969), and their actualistic counterparts are trench deposits. They are commonly associated with arc-trench gap type deposits (Dickinson, 1971). Then there are the Andean and arc style belts. Their backbone is a zone of vast volumes of mainly volcanic and plutonie rocks. Rock deformation in them can be adequately explained as a consequence of magmatic activity. Metamorphism is mostly low grade. Here magmatic activity can be documented for much of the lifespan of the associated subduction, though long gaps do exist. In contrast, plutonism associated with collision is of relatively short duration. ROLE OF THE ASTHENOSPHERE
Now we must examine the role of the asthenosphere. Essentially two types of models have been proposed to explain the high rate of heat flow and magmatic activity above dipping slabs: (1) frictional heating generated by the movement of the upper surface of the slab past virtually stagnant asthenosphere, with heat either conducted upward or convected by rising melts (e.g., Hasebe et al., 1970; Matsuda and Uyeda, 1971; Oxburgh and Turcotte, 1971); (2) secondary flow induced by the descending slab in the low Q wedge between the slab and the floating plate, with heat generated by viscous dissipation (Mackenzie, 1969) or advected by the induced flow which also thins the floating plate above (Sleep and Toksiiz 1971; Andrews and Sleep, 1974). The first model encounters serious difficulties, particularly if heat conduction is the dominant mechanism; in any case, it requires large stresses, hence high viscosities of mantle material to generate enough heat, and low viscosities to transport the heat to the surface fast enough, as pointed out by Andrews and Sleep (1974). The model, however, cannot be ruled out completely for all situations. The second model is more likely to be valid for marginal basins. Flow will occur in the wedge of asthenosphere, driven by some mechanical and/or thermal agent, if the viscosity is low enough. This is, in fact, implied by high temperatures below active marginal basins, inferred from heat flow data (Karig, 1971; Packham and Falvey, 1971). Viscosities are probably below 10zl poises. The subducting slab of lithosphere represents a moving boundary which then inevitably drives some flow by viscous drag. Furthermore, a strong lateral temperature variation is evident in the model which by itself must contribute to the driving mechanism. THE ANDREWS-SLEEP
MODEL
Our preferred model is that of Andrews and Sleep (1974), not only for its physical elegance and simplicity, but also because it provides some insight into a possible mechanism for the development of geosynclines and erogenic belts (Dennis, 1976). The model material has temperaturedependet viscos-
MSTANCE
267
, KM
MSTANCE , KM 400
3qo
,
0
100
290 100 300 WOO
/
/
-II00 Oo~~
if_
/
/
/’ /
2002
a
TEMPERATURE 087
.O
,
3y
(
/’
,QEG MY
MST= ”
/
.
c
KM
200
fO0
,
0
WSTANCE , KM
b
TEMPERATURE IO
.DEC MY
C
Fig. 3. Stream function and temperature d~t~bution under active marginaI basins; model computed by Andrew8 and Sleep. From Audrews and Sleep (1974). a. 0.87 m.y. after initiation of subduction. b. 10 m.y. after initiation of subduction.
268
ity and yield strength. At the start, the top is cold high-viscosity lithosphere. A slab descending 10 cm/y at 45” forms the right-hand boundary (Fig. 3). Flow into and out of the wedge is imposed on the left-hand boundary and on the bottom as computed analytically for a wedge extending to infinity. The initial layered temperature distribution is in conductive equilibrium above 80 km and adiabatic below. As flow is initiated in the wedge, stress builds up in the floating plate with compression near the trench between the converging plates and tension at some distance from the trench. At 1 m.y. the lithosphere starts to neck at the point of first yielding. First the rise of the isotherms causes some uplift of the crust, but as necking proceeds and the tensile stresses are concentrated, true rifting starts with depression of the surface and formation of an oceanic marginal basin. Material at the bottom of the heated floating plate is entrained with the more rapid asthenospheric flow and is subducted with the slab, thus causing the floating plate to move toward the trench. Ablation is, however, rapid only during the first few million years; after adjusting to a new. temperature equilibrium the thinned floating plate is no longer strongly ablated (Fig. 4). The time may be longer than computed by Andrews and Sleep (1974) if the viscosities and the yield
Fig. 4. Evolution of-an active marginal hasin. Letters indicate individual particles, as diaplaced. Dotted pattern = lithosphere; solid black = zone of enrichment in low-density material. Based on Andrews an& Sbep ( 19?4), from Dennis ( 1976).
269
strength are lower than assumed, which we would prefer. Andrews and Sleep’s (1974) model would be modified somewhat if in the calculation the asthenosphere had a bottom. The effect would be the development of a series of eddies in the asthenosphere below the continent and the formation of undulations at the surface (Sleep and Toksiiz, 1971). The amplitudes of the undulations would be magnified by erosion and sedimentation and by the corresponding isostatic response (Walcott, 1970). It is the apparent contradiction of extensional marginal basins in an environment of convergent plate margins which suggests that basin closing and orogeny are the natural fate of geosynclines. Eventually, continental crust riding on the subducting oceanic pfate is rafted against the trench, the island arc, and the marginal basin. For as long as the descending slab remains attached to the continent and driving forces continue to act, collision will continue and the whole island-arc marginalbasin belt will be compressed. The buoyant forces on the subducted continent will grow until subduction must choke (McKenzie, 1969). Marginal basins also may be subducted during some relatively short stage of the collision. Thus deep thrusts may develop before converging continents are welded together. Some of the basement thrusts in the Alps and in other orogens may be examples, suggesting that it may not always be possible to deduce the sense of the main subduction from the vergence of exposed thrust faults. SOME CONSEQUENCES OF ANDREWS-SLEW
CONVENTION
In the beginning, as convection proceeds, the lithosphere above the eddy has a tendency to stretch and thin. Material from its lower layers is dragged along, causing sub-lithosphere ablation, and is finally subducted. High isotherms may reach the sialic crust and raise its temperature above the solidus of dry granite. Thus, partially molten granitic material may be dragged into the corner of the wedge: it cannot be subducted, because ofits low specific gravity. Hence, it will form a convenient source for relatively silicic magma, for arc magmatic activity and for related mineral deposits. Its buoyancy will ensure continuing uplift, and a continuous source of sedimentation. Extension and ablation of the ~thosphere lowers the surface to oceanic or near-oceanic depths, thus providing a receptacle for geosynclinal ~~en~tion, much of which is continuously supplied from the rising volcanic arc, that is, from outside the continent. A weaker, coupled, continentward directed convection cell may cause a subdued version of some of these effects. Secondary convection could concentrate low-density material wherever such material is trapped by appropriate topography on the underside of the lithosphere, resulting in isostatic uplift. This mechanism may explain the Rocky Mountains uplifts of the United States and their mineral deposits, and the Harz Mountains of central Germany. Now, how is a geosyncline or marginal basin transformed into an orogen?
270
Current evidence seems to show that “classical” orogenies, those that involve orthogeosynclines as previously defined, have coincided with continent-continent collision. Such collisions terminate subduction. The end of subduction, in turn, chokes Andrews-Sleep convection. The marginal basin and its filling are compressed, removing the former heat sink. Heat and fluid escape are blocked. Hence, the former geosynclinal zone now heats up, leading to regional metamorphism and granitization. Volcanism has ceased with the cessation of the tensional regime. Subduction of marginal basins and compression of their filling, incidentally, transform these fairly wide areas in plan into the relatively narrow erogenic belts we know. This evolution explains why nearly all classical eugeosynclinal successions end up in erogenic belts: sooner or later all marginal basin deposits must be affected by collision. On the other hand, it can happen that Andrews-Sleep convection does not open a marginal basin: Andean style orogens may originate in this way, where, apparently, continental lithosphere is too thick for the initial rifting which elsewhere develops into a marginal basin. RELIEF
Dynamic relief is a prominent feature of erogenic belts. It provides persisting sources of immature sediments, and is evident in all young mountain ranges. To create relief, we need to add material of relatively low density for practical purposes, that means enriched in silica - at the base of the crust. There are four basic mechanisms for doing this: underthrusting of continental crust; compression-thickening (Molnar and Tapponnier, 1975); lateral transfer of low-density, low-viscosity material at the base of the lithosphere by special convection such as Andrews-Sleep convection; and, finally, mineral transformations, such as feldspathization, which evidently proceed after collision has throttled the Andrews-Sleep convection cell. After collision, thermal inertia may drive modified relict Andrews-Sleep convection and this could be responsible for late erogenic uplift coupled with foredeep subsidence. RELATED
PHENOMENA
No model can explain everything, especially in geology. But a useful model should at least help in gaining new perspectives on some old problems. For instance, the geologic record in some collision orogens leaves little evidence for a former vokanic arc: there is no clear sign of such an are in the Alps; but there appear to be vestiges of former oceanic crust, such as the Platta nappe. There is much evidence that the Tethys never opened very far in the area of the present-day Alps (T’riimpy, 1976). If arc volcanism were derived from the descending slab, it &or&l be ubiquitous in all cases of lithosphere subduction; but if it is a result of Andrews-Sleep convection, it could quite conceivably be impeded by some special circumstances attending the
271
closing of small oceanic basins. In fact, if asthenosphere eddies are mainly responsible for orogeny, then it is, of course, conceivable that some such eddies could have been initiated by processes other than classical lithosphere subduction. This might be the case for erogenic belts between cratons that show no relative displacement, such as the Damara orogen in southern Africa. When subduction gives way to strike-slip displacement, as in the case of the San Andreas Fault, thermal inertia could account for continuing, modified relict Andrews-Sleep convection. This may be a factor in the development of Basin and Range volcanism and tectonics in the western United States. SUMMARY
Let us now summarize how our model of orogeny works. First, the lithosphere ruptures near a continental margin, and the oceanic portion, already depressed by continental margin sediments, descends into the asthenosphere, forming a trench. Below the trench, the crust is under compression. The motion of the descending slab then initiates Andrews-Sleep convection. As a result, the floating margin is stretched under tension at right angles to the trench. Silica enriched partially molten material now begins to accumulate in the corner of the asthenosphere wedge in front of the descending slab, causing isostatic uplift above, and intrusion as well as extrusion of part of the accumulated low-density wedge material, thus forming a volcanic ridge. The evolution of this ridge will result in an Andean or arc-style orogen. Continued tension induced in the lithosphere by the convection may open a marginal basin, which becomes filled with sediments derived both from the continental interior, forming the miogeosynclinal zone, and from the volcanic arc, forming the eugeosynclinal zone. Volcanics derived from basinal rifting are interbedded with the sediments of the eugeosynclinal zone. At the same time, the trench fills with sediments transported oceanward from the arc by different mechanisms, including sliding and flowage. The trench fill becomes the thalassogeosyncline. These processes continue until a continent arrives at the trench on the back of the subducting plate. The continent cannot be subducted, so Andrews-Sleep convection must eventually stop. Heat which so far has been able to escape through the marginal basin floor and the slab is impounded and raises the temperature of the basement and fill of the marginal basin, while the basin itself is under compression and is gradually closing. Under these circumstances, instead of escaping, the trapped heat raises the temperature in surrounding rocks, causing regional metamorphism and granitization while the crustal segment affected is under compression. Where previously sediments had accumulated in a subsiding basin, there is now uplift due to crustal shortening and to volume increase both by raised temperatures and by mineral transformations. The extent of
212
sub-lithosphere mass transfers of low-density material at this stage is uncertain, but they probably play a role also. CONCLUSION
This model, therefore, is compatible with Stille’s classical geotectonic cycle, if limited to marginal basin-collision type orogens. We must, in fact, distinguish clearly between marginal basin and Andean or arc style erogenic zones, a distinction foreshadowed by Argand (Aubouin, 1972) many years ago, but which certainly needs further elaboration in the light of new geodynamic concepts. ACKNOWLEDGMENTS
We are indebted to W.H. Bucher, S.W. Carey, and H. Ramberg for kindling our interest in this topic, and to D.J. Andrews, G. Angenheister, J.F. Dewey, H. Neugebauer and H.G.F. Win&r for most enlightening discussions. This work owes much to the incentive of the Carey Symposium, and the manuscript was significantly improved by participation of one of us (J.G.D.), supported by grants from the Organizers and from California State University, Long Beach. REFERENCES Andrews, D.J. and Sleep, N.H., 1974. Numerical modeling of tectonic flow behind island arcs. Geophys. J., 38: 237-251. Aubouin, J., 1972. Chafnes liminaires (Andines).et chafnss geosynclinales (Alpines). Int. Geol. Congr., 24th. Montreal, Proc., 3: 438-461. Barazangi, M. and Isacks, B., 1971. Lateral variations of seismic wave attenuation in the upper mantle above the inclined earthquake zone of the Tonga island arc: deep anomaly in the upper mantle. J. Geophys. Res., 76: 8493-8516. Barker, P.E., 1972. A spreading center in the east Scotia Sea. Earth Planet. Sci. Lett., 15: 123-132. Beresnev, A.F., Gainanov, A.G., Kovylin, V.M. and Stroev, P.A., 1970. Structure of the earth’s crust and upper mantle under the Japan Sea and the Pacific Ocean adjoining Japan. In: Problems of Structure of the Earth’s Crust and Upper Mantle, Nauka, Moscow, pp. 50-61 (in Russian). Bogdanov, N.A., 1969. Thalassogeosynclines of the circumpacific ring. Geotectonics, 3: 141-147. Carey, S.W., 1970. Australia, New Guinea and Meianesia in the current revolution in concepts of the evolution of the earth. Search, l(5): 178-189. Z. Dtsch. Geol. Dennis, J.G., 1976. Geosynklinale, Orogenese und Piattentektonik. Ges., 127: 73-85. Dewey, J.F. and Bird, J.M., 1970. Mountain belt and the new global tectonics. J. Geophys. Res., 75: 2626-2647. Dietz, R.S., 1972. Geosynclines, mountains, and continent building. Sci. Am., 226: 3038. Dickinson, W.R., 1971. Clastic sedimentary sequences deposited in shelf, slope, and trough settings between magmatic ares and associated trenches. Pacific Geol., 3: 1530.
273 Ernst, W.G., 1970. Tectonic contact between the Franciscan melange and the Great Valley sequence, crustal expression of a late Mesozoic Benioff zone. J. Geophys. Res., 75: 886-902. Hasebe, K., Fujii, N. and Uyeda, S., 1970. Thermal processes under island arcs. Tectonophysics, 10: 335-355. Hsii, K.J., 1958. Isostasy and a theory for the origin of geosynclines. Am. J. Sci., 256: 305-327. Jacoby, W.R., 1975. Velocity-density systematics from seismic and gravity data. Zentral Inst. Phys. Erde, Verijffentl., 31: 323-333. Karig. D.E., 1971. Origin and development of marginal basins in the western Pacific. J. Geophys. Res., 76: 2542-2561. Kaseno, Yoshio, 1972. Geological features of the Japan Sea floor. Pacific Geol., 4: 91111. Kay, M., 1951. North American geosynclines. Geol. Sot. Am. Mem., 48: 143 pp. Matsuda, T. and Uyeda, S., 1971. On the Pacific-type orogeny and its model - extension of the paired belts concept and possible origin of marginal seas. Tectonophysics, 11: 5-27. McKenzie, D., 1969. Speculations on the consequences and causes of plate motions. Geophys., J., 18: l-32. Menard, H.W., 1967. Transitional types of crust under small ocean basins. J. Geophys. Res., 72: 3061-3073. Miyashiro, A., 1961. Evolution of metamorphic belts. Petrology, 2: 277-311. Molnar, P. and Tapponnicr, P., 1975. Cenozoic tectonics of Asia: effects of a continental collision. Science, 189 : 419-426. Oxburgh, E.R. and Turcotte, D.L., 1971. Origin of paired metamorphic belts and crustal dilation in island arc regions. J. Geophys. Res., 76: 1315-1327. Packham, G.H. and Falvey, D.A., 1971. An hypothesis for the formation of marginal seas in the western Pacific: Tectonophysics, 11: 79-110. Sleep, N.H. and ToksSz, N., 1971. Evolution of marginal basins. Nature, 233: 548-550. Triimpy, R., 1975. Penninic-Austroalpine boundary in the Swiss Alps: a presumed former continental margin and its problems. Am. J. Sci., 275-A: 209-238. Stille, H., 1941. Einfiihrung in den Bau Amerikas. Berlin, Borntraeger (1940), 717 pp. Walcott, RI., 1970. Flexural rigidity, thickness, and viscosity of the lithosphere. J. Geophys. Res., 75: 3941-3954.