Continental Strike-Slip Rifts and Their Stratigraphic Signature: Application to the Bangong/Nujiang Zone (Tibet) and the South Penninic Zone (Alps)

Geol Rundschau (1998) 87:206–224 © Springer-Verlag 1998

ORIGINAL PAPER

F. Mattern · W. Schneider · P. Wang · C. Li Continental strike-slip rifts and their stratigraphic signature: application to the Bangong/Nujiang zone (Tibet) and the South Penninic zone (Alps)

Received: 3 December 1997 / Accepted 21 February 1998

Abstract Our literature studies show that the thermal regime along continental strike-slip rifts is inconspicuous and Introduction that they are “low-volcanicity rifts” at best. Along with that, young continental strike-slip rifts exhibit no signs of Anomalous mantle structure, heat transfer into the litho- major thermally controlled doming. We suggest that the sphere, as well as domal uplifts are common features of larger the strike-slip component of a rift is, the less likely continental dip-slip rifts (Morgan and Baker 1983). An im- major thermal doming is causally associated with the rift portant process in the formation of such thermally uplifted zone. Since vertical lithosphere movements are reflected in rift areas is extension of the lithosphere which causes de- the stratigraphic record of a rifted area, different rift modes compression melt generation in the underlying asthenos- (strike-slip, dip-slip) may be distinguished by analyzing the phere and the diapiric ascent of the melts which can reach relevant sequences. Two ancient and especially suitable the surface (Ziegler 1990). Extensive uplifts associated strike-slip rift margins in Tethyan mountain belts, the Ban- with such rifts are typically on the order of 1–2 km with di- gong/Nujiang zone of Tibet and the South Penninic zone of ameters of a few hundred kilometers or more, and the dura- the Alps, were analyzed with regard to their uplift history. tion of the main phase is on the order of a few tens of mil- The results confirm recent regional rift models which indi- lion years or less (Morgan 1983). Examples of such rifts cate in both cases that rifting was dominated by strike-slip. with extensive uplifts are the Rhinegraben (Illies 1965, The stratigraphic approach may provide significant clues as 1975; Illies et al. 1979; Ziegler 1990) and the Central North to the mode of paleorifting when structural data are un- Sea rift (Ziegler 1990). The reviews by Illies (1965) and available. Pflug (1982) show for the Rhinegraben that the highest thermal anomalies occur all along the graben margins and Key words Continental strike-slip rifting · Continental that the 100°C isotherm appears to be depressed within the dip-slip rifting · Bangong/Nujiang · zone · Tibet · South graben with increasing thickness of the Tertiary basin fill. Penninic zone · Alps · Pantelleria rift · Romanche trans- Although the isotherm deepens towards the flanks over form · Suez rift · Dead Sea transform · San Andreas trans- short distance (Pflug 1982), the heat flow map of Germany form · Heat flow · Thermal regime · Thermal doming · Up- (Haenel 1971) demonstrates that the Upper Rhinegraben is lift · Rift stratigraphy · Volcanism the center of a regional positive anomaly with a diameter of approximately 300 km as indicated by the concentric pattern of isotherms. F. Mattern (✉) Doming of the crust is associated with continental dip- Institut für Geologie, Geophysik und Geoinformatik, Freie Universität Berlin, Malteserstrasse 74–100, slip rifting (Bott 1981), including continental back-arc set- D-12249 Berlin, Germany tings (Hsui and Toksöz 1981). Sengör and Burke (1978) put forward the concept of two basic types of rifting for dip-slip W. Schneider Institut für Geowissenschaften, Technische Universität rifts. The first and more important one (Khain 1992) is Braunschweig, Pockelsstrasse 4, D-38106 Braunschweig, Germany characterized by the active role of the mantle; convection plumes dome up and crack the lithosphere (active rifting). P. Wang In this case, doming predates rift formation. The second Department of Geoenergy, Changchun University of Science and Technology, 130026 Changchun, China one shows only a passive attitude of the mantle, and rifting is induced by horizontal plate movements which cause ex- C. Li tension of the lithosphere (passive rifting). In this case, Department of Applied Remote Sensing, Changchun, University of Earth Sciences, 79 Jianshe Street, doming possibly postdates rifting. According to Ziegler 130061 Changchun, China (1982, 1990) and Turcotte and Emerman (1983), passive 207

Rift Rift

Crust Crust

Lithosphere Lithosphere

Asthenosphere Asthenosphere

Active rifting / mantle plume model Passive rifting / tensional failure model

Fig. 1 Model for active and passive dip-slip rifting. (Compiled and drawn after Ziegler 1982, his Fig. 29 and Turcotte and Emerman 1983, their Fig. 1) rifting also causes uplift of rift flanks among dip-slip rifts. ence. They showed for very narrow basins that most of the Buck (1986) showed that passive rifting can induce convec- anomalous heat introduced by extension is also lost during tion beneath a dip-slip rift to produce uplift of the rift extension. These aspects point to differences between dip- flanks. slip and strike-slip rifts and encouraged us to examine dif- Ziegler (1982) and Turcotte and Emerman (1983) illus- ferent continental strike-slip rifts as documented in the lit- trated both rift mechanisms similarly through sections from erature in order to see to which degree such rifts are associ- the crust to the upper asthenosphere. They indicated the as- ated with thermal uplift. Although various effects of strike- sociation of rifting with uplift of the rift shoulders for both slip motion in continental rifting have been treated in the mechanisms (Fig. 1). According to Ziegler (1990), the ther- literature (e.g., Freund 1982; Scrutton 1982; Agnon et al. mal doming of a dip-slip rift zone can include uplift of the 1991), this aspect has remained neglected. We put the em- axial graben. The rate of thermal graben uplift may exceed phasis on the analysis of the regional stratigraphic evolu- coeval extension-induced rate of tectonic subsidence (Zie- tion, data pertaining to rift-related volcanism and heat flow. gler 1990). The uplift phenomenon is also known from dif- We found examples at various developmental stages. Our ferent continental back-arc regions such as the Mesozoic of intention is to draw attention to the possibility of distin- southernmost South America (Bruhn and Dalziel 1977) and guishing dip-slip and strike-slip rift modes via stratigraphic the Meso-Cenozoic of SE China (Mattern 1994). In these analysis which is of special significance and applicability cases, rifting was preceded or accompanied by uplift. The to the geodynamic reconstruction of ancient rift margins, Cenozoic Basin and Range province of North America, even if located in mountain belts. which includes the Great Basin, is also a continental back- arc region (Nelson 1981; Eaton 1984). The Great basin exhibits Cenozoic uplift, but the time of uplift seems to coin- Rifts not situated in mountainous settings cide with the cessation of subduction (compare uplift data as reviewed by Stewart 1978 with Irwin 1990, his Fig. Pantelleria rift system 3.12). Thus, uplift of the Great basin may have occurred while the Great basin was not entirely an active back-arc region. Nevertheless, the uplift seemed to be coeval with rift The active post-Miocene Pantelleria rift system (Fig. 2) of tectonics (Stewart 1978). the central Mediterranean Sea formed in the Strait of Sicily Apparently, thermal shoulder uplift is closely related to between Tunisia and Sicily on the Pelagian shelf (Reuther dip-slip rifting which progressively attenuates the litho- and Eisbacher 1985) as a consequence of Africa’s NW drift sphere and entails updoming of the asthenosphere (Favre during the Neogene (Boccaletti et al. 1990). Consistent and Stampfli 1992). Shoulder uplift provides a major clas- with that, Ben-Avraham et al. (1987) interpreted the Pantel- tic source for the rift basins whose stratigraphic record pro- leria rift system as a pull-apart region which represents a vides a sensitive tool for dating rift shoulder uplift (Favre segment of an approximately 1000 km long dextral strike- and Stampfli 1992). The sedimentary record of regions un- slip zone which terminates at areas of crustal extension, dergoing such rifting is characteristically dominated by i.e., the Tyrrhenian Sea to the WNW and the Gulf of Suez nonmarine deposits (Schneider 1972). to the ESE (Fig. 3). If spreading occurs only in isolated places along a The Pelagian shelf extends from the Maltese Islands to “leaky” strike-slip zone, thermal anomalies are expected to eastern Tunisia (Burollet et al. 1978; Bocaletti et al. 1990). be of small extent. Such pull-apart “leaks” may even be lo- According to Maldonado and Stanley (1977), 45% of the cated below sea level (Salton trough, Dead Sea). Modelling Strait of Sicily is shallower than 200 m and approximately by Pitman III and Andrews (1985) showed that in small ex- more than half of the sea floor is located in water depths be- tensional basins or pull-apart basins, lateral heat loss is very tween 200 and 700 m. Locally, the relief of the Strait’s sea important and accelerates lithospheric cooling and subsid- floor exceeds 1000 m (Maldonado and Stanley 1977). Ac- 208 Fig. 2 Tectonic map of the Pan- telleria rift. (Simplified after Reuther 1990, his Fig. 1)

Fig. 3 Extent of the eastern Mediterranean strike-slip fault S. Tyr. fault Tur key system between the Gulf of Su- Sicily ez and the south Tyrrhenian fault system. The Pantelleria pull-apart region is a part of the Pantelleria rift eastern Mediterranean fault system. (Simplified after Ben- Avraham et al. 1987, their Fig. 8)

N AFRICA 250 km Gulf of Suez

Arc

Lampedusa/Lampione region Maltese Islands Carbonates, including patch reefs 60 m Tortonian Coralline limestone Carbonate mudstones & to Lower up to 162 m Tortonian wackestones 18 m Messinian Coralline limestone >15 m Greensand 0-12 m Serravallian Blue clay 0-65 m

Langhian Miocene Miocene Globigerina ? limestone Phosphorite thickness ? Burdigalian 23-207 m Aquitanian

Coralline limestone Chattian >140 m Carbonates 0-200 m Oligocene Oligocene

Dolomitized limestone Carbonates (Halk El Menzel Fm.) Lutetian to (Halk El Menzel equivalent) 36m (upto1.1km?) Priabonian Eo- cene thickness ? Eo- cene Fig. 4 Stratigraphy of the Pelagian shelf. (From Pedley et al. 1978; Bismuth and Bonnefous 1981; Grasso and Pedley 1985; Grasso et al. 1993) 209 cording to Reuther and Eisbacher (1985), the water depth is The continuous shallow marine Upper Oligocene to Up- generally less than 400 m except for NW-trending elongate per Miocene rock succession of the Maltese Islands is dom- basins (Fig. 2) which subsided rapidly since the beginning inated by limestone (e.g., coralline limestone; Fig. 4). On of the Pliocene where depths are in excess of 1000 m. Malta, Cretaceous to Eocene dolomitized limestone is Whereas the Upper Quaternary sedimentary basin fills con- found in the subsurface (Pedley et al. 1978). sist of gravity mass flow deposits (mainly turbidites), On Lampedusa, Upper Miocene reefs crop out above hemipelagites, and volcanic ash layers, coeval sediments of Oligocene Lepidocyclina beds (Pedley 1988). According to the shallow platform are characterized by coarse calcareous Grasso et al. (1993), however, there are no Oligocene strata sand layers interbedded with mud and sandy lutite (Mal- on Lampedusa. They also noticed a lack of Oligocene stra- donado and Stanley 1977). The Plio-Pleistocene basin suc- ta in most boreholes of the region. Nevertheless, they inter- cessions are 1–2 km thick and the contemporaneously de- preted a 200 m thick section of shallow marine carbonates posited platform sediments do not exceed 500 m in thick- recovered from a borehole 20 km SW of Lampedusa to be ness (Boccaletti et al. 1990). of Oligocene age. There is not much information on the The Moho depth is reduced beneath the basins (reviews Lower to Middle Miocene stratigraphic development of the by Reuther and Eisbacher 1985 and Boccaletti et al. 1990). Lampione/Lampedusa region. Grasso et al. (1993) found it The crust is thinned to less than 20 km below the Strait of difficult to say whether the Eocene to Late Miocene succes- Sicily, but of normal thickness in adjacent areas (Boccaletti sion is continuous or separated by major hiatuses. et al. 1990). High heat flow is confined to the northwestern According to Bismuth and Bonnefous (1981), the Halk part of the rift zone (Bocaletti et al. 1990; Della Vedova et el Menzel Formation is frequently enclosed by two sedi- al. 1995, and sources within these works). Quaternary vol- mentary gaps which are occasionally associated with volca- canic edifices are known from the volcanic islands of Pan- nic activity. These gaps, however, are no indication for ma- telleria and Linosa and from the sea floor (e.g., Burollet and jor regional prerift thermal uplifts. Concerning the lower Ellouz 1986). The map by Burollet and Ellouz (1986, their gap, it has to be noted that the parareefal Halk el Menzel Fig. 3) indicates that recent volcanic activity is associated Formation is characterized by the “total absence of argilla- with faults and that volcanic zones parallel the trend of the ceous layers and detrital siliceous material” (Bismuth and rift basins. Pantelleria mostly consists of peralkaline silicic Bonnefous 1981). Apparently, no large landmass could rocks and only minor amounts of trachyte and basalt (Cor- have been substantially uplifted prior to the deposition of nette et al. 1983). Pantelleria’s volcanites are known for the Halk el Menzel Formation. Similarly, no such landmass their peralkalinity (Cornette et al. 1983; Mahood and Hild- could have been exposed during the time lapse of the upper reth 1983; Civetta et al. 1984). The basalts, however, are of depositional gap since Oligocene bioherms were able to alkaline character (Mahood and Hildreth 1983). According form on the Maltese Islands (Fig. 4; Pedley et al. 1978) and to Barberi et al. (1969), Linosa is made up of alkalibasalts Miocene reefs developed during the Miocene on Lampe- and hawaiites with minor mugearitic and trachytic differen- dusa and the Maltese Islands (Pedley 1988). The incom- tiates. Grasso et al. (1991) distinguished basanitic, transi- plete regional Oligocene succession is probably due to a re- tional and alkalic rocks on Linosa. gional tectonic shortening event (Antonelli et al. 1988; The Pelagian shelf is a Meso-Cenozoic subsiding car- Grasso et al. 1993) or to major global regressive tendencies bonate shelf (Burollet et al. 1978; Boccaletti et al. 1990). In (compare Steckler 1985; Haq et al. 1988) or a combination the following text we focus on the Tertiary prerift sedimen- thereof. As indicated by the stratigraphic record of the Mal- tary development of the Pelagian platform. Seismic lines tese Islands and Lampedusa, there is no Miocene gap in the and wells indicate that the top of the Mesozoic strata in the sedimentary succession. Tunisian and Sicilian offshore area is at a depth of approxi- Rift-related uplift above sea level is known from the mately 2–2.5 km (Burollet et al. 1978). The Tertiary succes- Maltese Islands. However, synrift shoulder upwarp only sion has a relatively regular thickness and displays an open managed to expose Miocene strata 260 m above sea level marine facies (Burollet et al. 1978). The Meso-Cenozoic on the northeastern side of the rift providing for a gentle NE succession extends from the Tunisian shelf/Gulf of Gabès dip of the strata (Reuther 1990). Possibly the emergence of region eastward to the Ionian abyssal plain and beyond (Bu- the island of Lampedusa can also be explained by upwarp rollet et al. 1978, their Fig. 9). Well data indicate the contin- of the southwestern rift shoulders as pointed out by Grasso uous presence of Middle to Upper Eocene limestones and and Pedley (1985), although they noted that the strata are dolomites of the Halk el Menzel Formation between Tuni- dipping towards the SE. We have no information as to sia and the Pantelleria rift system at least as far as the area whether rift-related shoulder uplift is dynamically and/or around the island of Lampione where the upper member of thermally controlled. Since the Pantelleria rift system is the formation is exposed (Bonnefous and Bismuth 1982; presently active (Reuther 1990), we consider it as being in Grasso et al. 1993). Bismuth and Bonnefous (1981) noted its synrift stage. In any case, the uplift is of minor lateral that the formation is also present in the vicinity of Malta extent and minor vertical amount. and envisaged that time and facies equivalents exist in Italy. We conclude that evidence for upwarp or doming during The Halk el Menzel Formation represents deposits of a prerift and synrift stages is minimal and does not compare widely extending carbonate platform with parareefal char- to thermal uplifts associated with dip-slip rifts. Thus, the acter (Bismuth and Bonnefous 1981). general character of the Pantelleria rift was, and is, that of a partly deep shelf region of continuous marine deposition. 210 Great thickness of sediments and deep marine conditions 1976; Ponte and Asmus 1978; Asmus and Baisch 1983). occur in the pull-apart basins. Volcanic activity is restricted Bellieni et al. (1992a) reported tholeiitic and alkaline dikes to the immediate rift area. Boccaletti et al. (1990) have em- from northeastern Brazil. On the African side, rift-related phasized that the Pantelleria rift system is not a rift zone volcanites are also known to occur (review by Mascle et al. whose deformation is controlled by tensile stresses induced 1988). by a regional thermal instability, but rather by a mixed Presently available data indicate that no thermally con- mode of faulting within a strike-slip regime. In agreement trolled uplift was associated with this shear margin as long with that interpretation they related the crustal and thermal as shearing took place within continental lithosphere. Re- structure as well as the magmatic history to the evolution of ferring to modern offshore shelf basins of Brazil and Gha- composite pull-apart basins. na, Gorini and Bryan (1976) noticed that during early rifting no appreciable volcanism and no domal uplift affected these areas (see also Asmus and Baisch 1983). They recognized that in the basinal areas, early Paleozoic sediments Romanche transform zone are preserved and that the same sediments were shed from highs located only at or near the littoral zone. On the Afri- The 900 km long Romanche transform zone (Fig. 5) situat- can side, Cenomanian to Eocene uplift seems to have only ed between spreading centers of the Atlantic formed during occurred at a later (postrift) evolutionary stage of the Ro- the Lower Cretaceous between the Brazilian segment of manche transform zone after the change from an intraconti- South America’s northeastern continental margin and the nental transform to a continent/ocean transform (Fig. 5) Gold/Slave Coast continental margin of Africa when the when heat was transferred from oceanic to continental South Atlantic Ocean started to form (e.g., Emery et al. lithosphere (Basile et al. 1993; general models by Scrutton 1975; Zalan et al. 1985; Mascle and Blarez 1987; Bonatti et 1982 and Todd and Keen 1989). The stratigraphic architec- al. 1996). With progressive dextral transform motion the ture of coeval sediments within the Ivorian basin recorded nature of crustal contacts changed across the Romanche this uplift and older strata were tilted (Basile et al. 1993). zone from continent/continent to continent/ocean to Basile et al. (1993) estimated the amount of this vertical ocean/ocean contacts (Fig. 5; Mascle and Blarez 1987; Ba- motion to be 1 km 20 km north of the continent/ocean sile et al. 1993). boundary. In the course of thermal exchange the ridge- Both margins display morphostructural characteristics shaped transform may have been uplifted and its crest erod- uncommon in passive margins (Scrutton 1982; Zalan et al. ed (Basile et al. 1993). Other than this postrift thermal up- 1985; Mascle and Blarez 1987). Pull-apart basins and lift, only vertical tectonic motions were identified as post- transpressive belts are known to occur on either margin rift ridge uplift causes (Bonatti and Chermak 1981). In the (Zalan et al. 1985; Mascle and Blarez 1987). Piaui basin, offshore Brazil, a transpressive belt represents Although Meso-Cenozoic rift-related volcanites occur a structural high which was subjected to subaerial erosion on the mainland of northeastern Brazil (Sial 1976; Ponte from late Cretaceous to the Eocene (Zalan et al. 1985). On and Asmus 1978; Bellieni et al. 1992a, b), and although ba- the African side, contractional faulting related to changes in salt is known to occur in rift sediments offshore of Brazil the geometry of the equatorial Atlantic plate boundaries (Zalan et al. 1985), it seems that only minor amounts of vol- seemed to cause a postmid-Eocene ridge (Bonatti et al. canic rocks were produced during early rifting in the vicin- 1994, 1996) which, according to Bonatti et al. (1994), was ity of the Romanche transform zone (Gorini and Bryan above sea level at and before 5 Ma ago.

Fig. 5 A, B Kinematic develop- AB ment of the Romanche transform zone. A Contours of conti- Spreading axis nents are shore lines at indicated 80 m/y AFRICA times. Heavy lines are mid- AFRICA ocean spreading ridges. The dia- gonally lined areas are transform/fracture belts. The large transform belt between the dis- SOUTH placement arrows includes the AMERICA SOUTH Romanche transform. (Simpli- AMERICA fied after Emery et al. 1975, Thinned continental crust their Fig. 34). B Depiction of AFRICA the Romanche transform’s transition from an intracontinental (above) to a continent/ocean transform (below). (Simplified SOUTH after Mascle and Blarez 1987, AMERICA their Fig. 3). The timing be- SOUTH tween the illustrations in A and AMERICA B does not necessarily correlate 110 m/y Oceanic crust 211

According to the review by Basile et al. (1993), early Eastern Mediterranean margin sedimentation at the then intracontinental Romanche transform zone was tectonically controlled as indicated by high subsidence rates due to faulting during the Aptian and Alb- ian. Sedimentation was terrestrial during the early Creta- ceous. During the Aptian/Albian on the Brazilian side and N Albian on the African side, tectonic activity ended and sed- Pelusium imentation generally changed from terrestrial to marine (Basile et al. 1993). Our literature studies revealed no indication for the presence of a thermal rift dome in the stratigraphic record. Thus, the region’s overall character remained that of sheared margins which implies shelf environments bathymetrically ranging from around sea level to deep shelf on little thinned continental crust (Scrutton 1982).

Suez rift Synrift The Suez rift (Fig. 6) is a 60 to 80 km wide depression, deposits whose central part is submerged below the Gulf of Suez Dead Sea transform (Lyberis 1988). Its length measures 480 km (Fig. 6). Whereas the submerged central and southern rift segments Anticline display the morphology of a rift valley and trend approxi- Prerift mately NW, the terrestrial northern segment’s strike is more structures Passive margin northerly and the rift valley character disappears. Joffe and hinge line Garfunkel (1987) attributed this to a decrease of extension Northern towards the north. They pointed out the en échelon arrange- Early Miocene Red Sea Late Miocene ment of basins located in the central depression of the Gulf Mean extension Pliocene/ and the small basins north of the Gulf (Bitter Lakes, Lake directions Quaternary Timsah; Fig. 6) and interpreted this geometry as an indication of strike-slip movement. Faulting cannot be traced 0 100 200 north of the Bitter Lakes (Steckler and Ten Brink 1986). km In the course of the Cenozoic breakaway of the Arabian plate from the African plate, the Red Sea and its two north- Fig. 6 Palinspastic structural map of the Gulf of Suez for the early ern branches (see Fig. 8), the Suez rift and the Dead Sea Miocene. (Simplified after Favre and Stampfli 1992, their Fig. 1) transform, formed (Garfunkel 1988). Since most of this divergent motion was taken up in the north by the Dead Sea the earliest synrift deposits (Fig. 7; Montenat et al. 1988; transform (Garfunkel 1988), the Suez rift remained in a rift- Patton et al. 1994), which was attributed by Steckler (1985) ing stage in contrast to the Red Sea where oceanic spread- to a general regression (compare Haq et al. 1988; see “Pan- ing occurs (Favre and Stampfli 1992). Therefore, the Suez telleria rift system”) as known from other continental mar- rift is entirely floored by continental crust as numerous drill gins rather than to rifting-related uplift. The only Oligocene holes show (Joffe and Garfunkel 1987). Extension esti- deposits known near the Gulf of Suez are a few locations mates indicate that extension was limited to a few tens of within the rift (Steckler 1985). He reasoned that “if an uplift kilometers and greater in the south than in the north (Patton had existed before any rifting, the crest of the dome is the et al. 1994, their Table 1). most unlikely place to contain preserved Oligocene beds”. Despite some regional tectonic pulses related to the dif- The first phase of Cenozoic rifting occurred near sea lev- ferent stages in the Tethys evolution, the Suez region’s el (Steckler 1985; Schütz 1994) and was accompanied by overall Mesozoic to early Tertiary character prior to rifting Oligocene to Miocene basaltic volcanism (Patton et al. was that of a shelf area, stable in the south and unstable in 1994) which, according to Moussa (quoted in Patton et al. the north (Patton et al. 1994). Prerift sequences of late Cre- 1994), is of subalkaline (tholeiitic) to alkaline character. Al- taceous to Eocene marine deposits are widespread (Monte- kaline olivine basalts were reported by Steinitz et al. (1978) nat et al. 1988). The Upper Paleocene (Thanetian) to Upper from the western part of the Sinai peninsula. This volcanism Eocene (Priabonian) stratigraphic record (Fig. 7) indicates is generally considered “minor” (Steckler 1985; Patton et continuous marine sedimentation, except for a non- al. 1994). The first rift deposits are the poorly age-con- depositional interval during the Bartonian (review by strained (Oligocene? to Miocene) terrestrial to shallow ma- Schütz 1994). rine Abu Zenima and Nukhul formations (Montenat et al. Apparently, there is a widespread Oligocene stratigraph- 1988; Patton et al. 1994), exhibiting “numerous and spec- ic gap between the Paleocene/Eocene prerift sequence and tacular changes in thickness and facies” (Montenat et al. 212 ing was more of the dip-slip type. The structural map (Fig. 6) indicates a systematic sinistral displacement of prerift Post-Zeit: clastics and marine structures of approximately 20 km (Fig. 6). Considering the limestone rift’s trend and the rotation of the extension direction, one could argue that the amount of apparent lateral offset was

Pliocene 1-2 km ? ? ? ? somewhat reduced during the interval of dip-slip rifting. Zeit Fm., clastics & marine evaporites Messinian Rifting following the strike-slip interval was characterized aver. 500 m, max. 1 km by subsidence of large, elongate tilted blocks and uplift of ? ? ? ? the rift shoulders causing substantial erosion of the prerift Tortonian South Gharib Fm., marine evaporites and shale aver. 200 m, max. 2 km succession (Favre and Stampfli 1992). This Miocene tec- ? ? ?? Belayim Fm., marine evaporites, tonism lasted for 10–15 m.y. (Stampfli and Marthaler clastics, and limestone max. >500 m 1990). Serravallian The regional uplift history has received considerable at- Kareem Fm., anhydrite, shale, marl 200m tention in the literature (e.g., Kohn and Eyal 1981; Steckler Langhian Upper Rudeis Fm., fluvial & marine 1985; Garfunkel 1988; Montenat et al. 1988, Omar et al. Miocene clastics (marl, sst., cgl.), some 1989; Jarrige et al. 1990). There is consensus among nu- marine carbonates 1000 m merous workers that rifting was not preceded by precurso- Lower Rudeis Fm., mainly marine marl, Burdigalian some sandstone, limestone max. 1.2 km ry doming, and that uplift only occurred while rifting was already well underway (Steckler 1985; Montenat et al. Nukhul Fm., terr. clastics, marine limestone, sulfate 1988; Omar et al. 1989; Jarrige et al. 1990). Therefore, rift- Aquitanian ? ? ? max. 700 m ing was interpreted as passive (Steckler 1985; Omar et al. ? ? ? ?? Abu Zenima Fm. ??1989). Literature data pertaining to the timing of the struc- Chattian terr. clastics <110 m ???? tural rift development reveal the important observation that

Oligo- cene uplift of the rift flanks took place only after the change Priabonian Limestone & marl 200 m and more from strike-slip to dip-slip rifting when extension became Bartonian significant (e.g., Steckler 1985; Montenat et al. 1988; Jarr- Shale, marl & limestone, some gypsum ige et al. 1990; Stampfli and Marthaler 1990; Favre and Lutetian ~100 m

Eocene Mainly pelagic limestone, some Stampfli 1992). Presently, there is continuing regional up- Ypresian dolomite & sabkha anhydrite 140 m lift (Kohn and Eyal 1981; Steckler 1985) despite a waning or cessation of rifting (Steckler 1985; Patton et al. 1994). Thanetian Shale & limestone The amount of uplift (see Kohn and Eyal 1981; Garfunkel 15-40m 1988; Steckler et al. 1988) must be considered substantial. Paleocene The heat flow in the Gulf of Suez is elevated and in- Fig. 7 Stratigraphy of the Gulf of Suez region, compiled mainly af- creases towards the south, pointing to the southward inter Patton (1994) and Schütz (1994). Some information after James et crease in extension and lateral heat advection from the Red al. (1988) and Montenat et al. (1988) Sea as the likely causes (Feinstein et al. 1996). The present thermal conditions represent the maximum heat flow and temperatures for the basin’s sedimentary section (Feinstein 1988). Moreover, soft sediment deformation reflects abun- et al. 1996). dant seismic activity (sources in Montenat et al. 1988). Within the synrift deposits, basalt sheets and dikes and re- worked basalt pebbles occur (Montenat et al. 1988). One Dead Sea transform rift zone dike yielded a radiometric age of 26–25 Ma and a basalt flow was dated to be 22 Ma old. Evidence for major clastic The active >1000 km long Dead Sea transform rift zone input as of 17 Ma from the rift shoulders is manifested in (Fig. 8) extends from the southern tip of the Sinai peninsu- the Upper Rudeis Formation (Steckler and Ten Brink la to Turkey (Ben-Avraham 1987; Kovach et al. 1990). Its 1986). The Upper Rudeis Formation contains the first plate tectonic framework as outlined by Wilson (1965), widespread conglomerates which appear at the rift margins Kashai and Croker (1987) and Garfunkel and Ben- above a widespread unconformity (Steckler 1985). For fur- Avraham (1996) shows that it connects the divergent plate ther information pertaining to the stratigraphic synrift de- boundary of the Red Sea in the south with the contraction- velopment, see Scott and Govean (1985), Montenat et al. al province of east Anatolia of southern Turkey in the north (1988), and Patton et al. (1994). (Fig. 8), a part of the Taurus/Zagros collision zone (e.g., At the Suez rift, the rifting process which was initiated Ben-Avraham and von Herzen 1987). Sinistral motion during the late Oligocene can be subdivided into different along the Dead Sea transform started 26 m.y. ago deformational intervals due to rotation of the stress vectors (Wdowinski and Zilberman) or later than 20 Ma (e.g., (Fig. 6; Favre and Stampfli 1992). Whereas sinistral strike- Garfunkel and Ben-Avraham 1996). Since fault initiation, slip shearing was predominant from the late Oligocene to geologic offset amounted to 105 km (Freund et al. 1970). early Miocene spanning 5–10 m.y. and characterized by During the past 5 m.y., 30 km of displacement accumulated transtension and some local transpression, subsequent rift- (Joffe and Garfunkel 1987). 213 side only by transform faults and on the other by coeval parallel normal faults. However, the Dead Sea basin was interpreted more recently by Garfunkel und Ben-Avraham Tur key (1996) as “situated within a large pull-apart.” The area of the southern segment of the Dead Sea transform is a Phanerozoic continental platform on which mostly marine sediments of a few kilometers thickness accumulated (e.g., Freund et al. 1970; Garfunkel and Ben-Avraham N 1996). Before transform initiation, the region crossed by the transform was below sea level until approximately 40 Ma (Garfunkel and Ben-Avraham 1996). Between 15 and 12 Ma, the sea briefly extended 20–30 km inland from the present coast (Garfunkel and Ben-Avraham 1996) and dur- Cyprus ing the Pliocene, a marine transgression of the Mediterra- nean Sea took place through the Yisreel and Jordan valleys to reach the Dead Sea area in the south (Manspeizer 1985). Considering the thickness of the basin fill and not so much its lateral extent, the Dead Sea (Fig. 8) represents an impressive depocenter along the transform. Above more than 2000 m of terrestrial Miocene red beds, a marine to Mediterranean Sea brackish Pliocene to early Pleistocene sequence of clastic deposits and evaporites including diapir-forming rock salt Dead Sea measuring over 4000 m accumulated (Zak and Freund Bitter Lakes 1981; Manspeizer 1985). In the following time, younger area lacustrine evaporites and fluvio-deltaic sediments in excess of 3500 m were laid down (Zak and Freund 1981; Man- speizer 1985). The syntectonic deposition of the basin fill is characterized by abrupt lateral facies changes and un- Sinai conformities (Manspeizer 1985). Gulf of The Neogene depositional history of the Gulf of Elat Suez (Fig. 8) is described by Ben-Avraham et al. (1979). The fol- Gulf of lowing review is based on their data. The basin is 180 km Elat long, 15–28 km wide, and 1850 m deep. It generally lacks continental shelves and coastal plains. Sedimentation was syntectonic. The thickness of the basin fill, which consists of turbidites and pelagic sediments, exceeds 2–3 km. On AFRICA the steep basin slopes, alluvial fans seem to be the dominant deposits. During the early history of the gulf, marine strata, including evaporites of Miocene age, were laid down. They Red Sea occur mainly on the eastern side of the gulf. Extensive erosion affected the western side of the gulf and the region to 200 km the north. The timing of this event is poorly constrained but is probably of late Miocene age. Sediments of the gulf’s younger history, as seen on land, are generally thin alluvial and lacustrine deposits. Fig. 8 General tectonic setting of the Dead Sea transform and the The land around the gulf is generally uplifted. Uplift es- Gulf of Suez as well as regional fault pattern. (Simplified after Heimann et al. 1990, their Fig. 1) pecially affected the areas in the vicinity of the Red Sea. Uplift diminishes northward (Ben-Avraham et al. 1979). This coincides with the tendency of increasing heat flux in At the Dead Sea transform, the land surface is depressed the gulf from north to south and with the general trend of significantly below sea level [Jordan Valley, Arava Valley thinning of the continental lithosphere in the same direction (=Wadi Araba)]. Rhomb-shaped grabens (i.e., pull-apart (Ben-Avraham and von Herzen 1987). Although the gulf’s basins), such as the Gulf of Elat (=Gulf of Aqaba), Dead floor is generally continental, the southern part of the basin Sea (–409 m; Hall 1996), Sea of Galilee, and the Hula Val- floor seems to be more oceanic than in the north (Ben- ley, have been identified (e.g., Freund et al. 1968; Heimann Avraham and von Herzen 1987). According to Ben- and Ron 1987; Garfunkel and Ben-Avraham 1996). Alter- Avraham and von Herzen (1987), these observations reflect native to the pull-apart basin interpretation, Ben-Avraham propagation of rifting from the Red Sea northward. (1992) and Ben-Avraham and Zoback (1992) put forward The formation of the Dead Sea transform was accompa- the concept that the former two basins are bounded on one nied by widespread volcanism producing mainly alkali oli- 214 vine basalts, basanites and nephelinites (review by Garfun- der elevations should be lower by 500–800 m than they are kel and Ben-Avraham 1996). Although some volcanism oc- if topography was locally compensated. Ten Brink et al. curred along the northern two thirds of the transform as far (1988) pointed out the possibility that uplift is a conse- south as the Dead Sea basin, it often is not obviously relat- quence of mechanical stresses which accumulated in the ed to the transform since most of it affected the area east of course of Arabia’s collision in the north or due to loading of the transform up to a distance of a few hundred kilometers the Levant (passive) continental margin. Since heat flow (Garfunkel and Ben-Avraham 1996; for distribution of late within the rift and at its shoulders is normal, “any thermal Cenozoic volcanites see Fig. 13 in Reches 1987). cause for uplift must reside in the lower lithosphere” (Ten The areas bordering the Dead Sea transform were uplift- Brink et al. 1988). Ten Brink et al. (1990) summarized that ed by approximately 1 km and the area of the 15 to 12 Ma uplift can be explained by models implying dynamic sup- transgression now lies at an elevation of 500 m above sea port or regional compensation and by any thermal causes level (Garfunkel and Ben-Avraham 1996). After regression only if the heat source is deep-seated (i.e., in the upper man- of the Eocene sea, at least in southern Israel, a peneplain tle). developed (Garfunkel and Horowitz 1966). Observations According to Wdowinski and Zilberman (1996), uplift pertaining to the provenance and paleoslope of the clastic in the southern part of their study area can be explained as a Hazeva Formation indicate that during the early to late Mi- response to rift-normal extension and normal faulting ocene, the transform flanks were still not uplifted (Garfun- which induced a negative mass anomaly. For their northern kel and Ben-Avraham 1996). According to the review by study region, they attributed uplift to the combined effects Wdowinski and Zilberman (1996), the area around the of normal faulting, isostatic uplift, and loading by young southern segment of the Dead Sea transform was part of a deposits. single morphologic NW-drained unit of low surface relief From the foregoing paragraphs, we conclude that al- during the Oligocene–Miocene (38–5 Ma). This seems to though the uplift causes are under discussion, there is con- be in agreement with the conclusion that the transform sensus that a thermal uplift cause is not obvious. flanks were uplifted mainly after 10 Ma ago (Garfunkel and Ben-Avraham 1996). Judged by its lateral extent and by its vertical magnitude of several hundred meters (Ten Brink et San Andreas transform system al. 1990; Wdowinski and Zilberman 1996), the uplift cannot be considered of minor importance. Apparently, region- The active dextral San Andreas transform system started to al uplift appeared after the middle to late Miocene closure form during the Oligocene at North America’s western con- of the ocean basin once located between the Arabian plate tinental margin when the Pacific/Farallon mid-ocean ridge and Turkey (Sengör et al. 1985) and may be an effect of on- reached the continent causing the gradual cessation of sub- going plate convergence in eastern Anatolia. duction (Atwater 1970; Dickinson 1981; Irwin 1990). Although the general morphology of the Dead Sea trans- While the length of the regional subduction zone decreased, form region displays uplifted shoulders and a median valley the length of the transform zone increased due to migration somewhat similar to graben-type rift regions, Ten Brink et of the transform-bounding triple junctions in opposite di- al. (1990) pointed out some distinct differences of the Dead rections (Atwater 1970; Dickinson 1981; Irwin 1990). At Sea rift: the neglible amount of extension, major horizontal present, the length of the San Andreas transform system displacement, the much smaller width of the median valley measures approximately 2600 km from northern California (5–20 km compared with 40–60 km), and the surface heat to the mouth of the Gulf of California (Fig. 9). Whereas ear- flow along the Dead Sea transform and its surrounding area ly transform motion occurred between the oceanic Pacific which does not deviate from the global heat flow for the in- plate and the continental North American plate, it jumped terior of the continents. According to Feinstein (1987), tec- eastward one or more times (Irwin 1990). Close to the Mi- tonic models for the evolution of the Dead Sea basin requir- ocene/Pliocene time boundary, the peninsula of Baja Cali- ing a high thermal regime are inconsistent with thermal his- fornia began to separate from the North American plate tory data. At the Dead Sea, the heat flow is only locally el- (Dickinson 1981) and the Gulf of California opened. The evated (Garfunkel and Ben-Avraham 1996). It also must be dextral shear sense and the stepover geometry of en échelon noted that the crust on both sides of the Dead Sea is 30–35 arranged basins within the Gulf (Larson et al. 1972, their km thick (review by Garfunkel and Ben Avraham 1996) Fig. 1) as well as on the mainland of southern California and at a larger regional scale 20–37 km thick, locally (Salton trough, i.e., the Coachella Valley/Salton Sea/Impe- thinned by only 3–4 km beneath the rift (Wdowinski and rial Valley area; e.g., Fuis et al. 1982) indicates that region- Zilberman 1996). al extension occurs at pull-apart structures (Fig. 9; Irwin The gravity or gravity anomaly observed in profile 1990). across the Dead Sea transform was explained by the juxta- Unfortunately, in the case of the San Andreas system, in- position of two different crustal sections displaced left-lat- cluding the Gulf of California, it is difficult if not impossi- erally for 105 km at the transform (Ten Brink et al. 1988, ble to assess a possible relationship between strike-slip rift- 1990), suggesting an offset in Moho depth at the transform ing and uplift for the following reasons: The shear move- (Ten Brink et al. 1990). According to the gravity model by ments followed the Cenozoic active continental margin de- Ten Brink et al. (1988), the uplifted rift shoulders are not velopment and there are very special thermal implications compensated by local isostasy. They also noted that shoul- such as the shallow emplacement of hot asthenospheric 215

Fig. 9 A Map of the San An- A B dreas transform system and its Coachella San Andreas plate tectonic setting. (Simpli- Valley Triple junction fault fied after Corona 1996, his Figs. N N 1 and 5, and Irwin 1990, his Fig. San Salton Sea 3.13). Box in center indicates San Francisco North Imper. Valley map area of B. B Geologic Oceanic Andreas River Bay American sketch map of the Salton trough spreading area with defined pull-apart axis Plate structures (dotted) and regional fault Salton Sea fault pattern. Map area is indi- Los Angeles cated in A. (Simplified after Ir- win 1990, his Fig. 3.8) Colorado

Pacific Plate Gulf of California

500 km Gulf Triple junction 50 km Oceanic spreading axis mantle into the vacated slab window after cessation of sub- transtensional tectonism affected the area is in agreement duction (Furlong 1993). Rifting in the Gulf area followed with findings by Herzig and Jacobs (1994) who investigat- extension of the uplifted Basin and Range province situated ed the geochemistry of Cenozoic volcanites. They interpret- in a Cenozoic back-arc environment (Nelson 1981; Eaton ed Oligocene–Miocene alkaline lavas to be related to Basin 1984) that is still heating at a regional scale (Zorin et al. and Range extension and post-4 Ma subalkaline subsurface 1984), and it has to be noted that the Gulf area is at least basalt and basalt xenoliths in Quaternary rhyolites to the contiguous with the modern Basin and Range province present geodynamic regime. They also suggested that the (Henry 1989; Zanchi 1994) and that the Salton trough’s ini- rhyolites and inclusions of granite xenoliths therein are dif- tial formation is perhaps related to the Basin and Range ex- ferentiates of mafic magma, probably represented by the tension (Crowell 1987). Another difficulty exists in the dis- basalt xenoliths. However, some workers previously con- tinction of recent possible thermal and tectonic uplift ef- sidered the rhyolites to be alkaline rhyolites (Robinson et fects in this seismically active region. al. 1976). These rhyolites are the only known late Quater- Heat flow determinations at segments of the San An- nary volcanites closely associated with the San Andreas dreas fault do not clearly show the existence of a heat flow transform zone north of the Gulf of California (Sharp anomaly (Henyey and Wasserburg 1971). On heat flow 1982). Comparing the recent regional volcanism with that maps of the western United States (Lachenbruch and Sass of the African-Arabian rift system, Herzig and Jacobs 1978; Lachenbruch et al. 1985; Sass et al. 1994), the San (1994) considered the Salton trough’s volcanism “insignifi- Andreas fault cannot be identified except for the Salton cant.” trough of southern Califonia which exhibits an elevated The Salton trough, located in a mountainous setting ad- heat flow. Nevertheless, the quoted maps indicate a similar jacent to the Colorado River delta, is an important depocen- thermal regime for other parts of the Basin and Range prov- ter along the San Andreas transform zone. The basin is at ince. least 100 km long and several tens of kilometers wide. Ac- Despite the aforementioned problems, the Salton trough cording to Fuis et al. (1982), the sedimentary basin fill, in- and the Gulf of California still provide important insights cluding metasediments in the Imperial Valley region, mea- into the understanding of continental strike-slip rifting. The sures 10–16 km in thickness. Sharp (1982) summarized that Salton trough is at the dawn of becoming a part of the open- marine and nonmarine sedimentation may have been essen- ing of the Gulf of California since it represents the structur- tially continuous since the late Miocene in the axial basin al and physiographic extension of the Gulf close to the part. Of interest to us is the thickness and facies of the stra- Gulf’s head (Fig. 9; Muffler and Doe 1968; Sharp 1982) ta which accumulated during the Pliocene–Quaternary and since the surface of the Salton Sea already lies 74 m be- transtensional basin cycle. A drill hole that probably bot- low sea level. Moreover, there are indications that the Sal- tomed in equivalents of the Pliocene–Pleistocene Palm ton trough is at least partly floored by oceanic/basic crust Spring Formation (Fig. 10) at a depth of more than 4 km in- (Elders et al. 1972; Fuis et al. 1982; Lachenbruch et al. dicates the great sediment thickness of this cycle (Muffler 1985; Herzig and Jacobs 1994). Modelling of geophysical and Doe 1968) and, of course, great subsidence. The data resulted in the interpretation that the Moho is as shal- transtensional history started during deposition of the Im- low as 22 km to 23.5 km beneath the Salton trough (Elders perial Formation (Fig. 10) when the transform boundary et al. 1972; Fuis et al. 1982; Lachenbruch et al. 1985). jumped eastward into the rift valley (Nilsen and Sylvester The idea that the Salton trough perhaps initially formed 1995). This formation is partly marine (e.g., Sharp 1982; during the Middle Miocene as a half graben in the course of McDougall 1996). Later, during the late Pliocene to early Basin and Range extension (Crowell 1987) before Pleistocene, coarse-grained and very thick alluvial fans 216 Pull-apart basins also occur in the northern gulf region, but SW side of Salton trough NE side of Salton trough there they lack such a morphology since they are filled up with sediment from the Colorado River (Einsele 1992). Alluvium, dune sand, lake deposits Shelf areas are best developed in the northern and eastern

Holoc. Gulf (Einsele and Niemitz 1982). Sedimentation in the cen- ? Ocotillo Brawley Ocotillo Cgl. tral and southern gulf’s basin slopes and subbasins is pre- Cgl. Fm. dominantly fine-grained, although regional relief is high Borrego Borrego and several marine deltas exist (Einsele 1992). A large pro- Formation Formation Canebrace portion of slope deposits is redeposited as mud turbidites in Cgl. the deepest areas (Einsele 1992). The isopach map by Ein- Palm Palm Pleistocene Canebrace Spring Spring sele and Niemitz (1982) indicates highly varying thickness- Cgl. Formation Formation es of wet, unconsolidated sediment and maximum thick- ? nesses in excess of 1500 m. Imperial Formation (marine) Imperial Formation Split Mt. Fm. (marine)

Pliocene ? Fish Creek gypsum Ancient rift margins in mountainous settings Split Mt. Fm. (terrestr.) Lava Mecca Formation Mioc. Anza Formation In some mountain belts, the preophiolitic stratigraphic

Pre- Crystalline rocks

Cenoz. record is very similar across an intervening ophiolite zone. In these cases, it is reasonable to assume that both terranes adjacent to the ophiolites once belonged to a continuous Fig. 10 Cenozoic stratigraphy of the Salton trough’s flanks. Wavy continental plate which was rifted and later welded togeth- lines indicate unconformable contacts. (After Sharp 1982, his Fig. 3) er again at their common rift margin. In order to reconstruct the geodynamic history of such regions, it is of interest to deduce what kind of rifting took place. Dip-slip and strike- formed grading axially into lacustrine mud, and in the case slip rifting are end members between which oblique-slip of the Borrego Formation, thin layers of evaporite (Babcock rifting represents a gradational option. 1974; Nilsen and Sylvester 1995). The material was shed Continental strike-slip rift margins are expected to be from the nearby mountains (Babcock 1974). Numerous un- difficult to identify in mountain belts because unambiguous conformities reflect the syntectonic deposition of the Salton evidence of lateral motion is seldom preserved (Mitchell trough’s basin fill (Fig. 10; e.g., Sharp 1972). The younger and Reading 1989). If structural evidence is missing to rule depositional history is dominated by great amounts of sedi- in/out a strike-slip rift process, strategies are needed to help ment from the Colorado River delta (Nilsen and Sylvester solve the dilemma. The evaluation of the rift margins’ 1995; Kidwell 1996) which separates the Salton trough stratigraphic records can provide a useful clue as to the type from the Gulf. of rifting. Pitman III and Andrews (1985) showed through their The above review of relatively young strike-slip rift modelling that lateral heat loss is very important in small zones not situated in mountainous settings indicates that extensional basins or pull-apart basins, and that this accel- there are generally no major thermal anomalies associated erates lithospheric cooling and subsidence. They demon- with these rifts and that there is no evidence for major re- strated for very narrow basins that most of the anomalous gional thermal uplift of such rifts as long as continental sep- heat introduced by extension is also lost during extension. aration is dominated by strike-slip tectonics. Therefore, we They further noted that their findings are particularly appli- feel encouraged to interpret the following two examples of cable to many of the small extensional basins associated ancient rift zones situated in mountainous settings as strike- with the San Andreas system. In our view, their findings slip-dominated rift margins based on the “uplift analysis” may indicate that thermal uplift of the Salton trough region of the stratigraphic record. due to strike-slip motion alone is expected to be limited. For The two examples bear several advantages: the basin itself, such thermal uplift is difficult to imagine, as 1. Their relevant sequences were deposited mainly in a much of the basin’s surface is below sea level, in general shelf environment which allows easy reconstruction of contrast to graben-type basins (see Introduction). Our liter- the uplift history (e.g., emergence above sea level). ature studies revealed no clear evidence for regionally ef- 2. Continental accretion took place at the terranes’ com- fective thermal uplift related to strike-slip tectonics. With- mon rift margins and both rift flank couples can be ana- out interpreting the causes of regional uplift (tectonic?, lyzed. thermal?), Sharp (1982) noted Quaternary uplift of the 3. A different degree of strike-slip rifting was involved western half of the trough and part of its eastern flank. among both examples. The topography of the narrow and deep Gulf of Califor- 4. Recent tectonic models have indicated a rift history dom- nia is complex. Pull-apart basins within the gulf form de- inated by strike-slip in both cases so that our stratigraph- pressions which are 2000–3000 m deep (Einsele 1992). ic analysis can be viewed as a test for these models. 217 out having to assume a range of 150 km for ophiolite thrust- Songpan - Ganzi Terrane N ing. The rift model implies that many of the ophiolite bod- ies are autochthonous or parautochthonous. Our map analy- Rutog Qiangtang Block sis revealed that it is possible to interpret several ophiolite BNZ Gerze thrust units as emplaced by local transpression. The rift concept also eliminates the aforementioned other problems Coward et al. (1988) and Dewey et al. (1988) addressed. BNZ Lhasa Therefore, the Bangong/Nujiang zone may indeed not rep- Lhasa Block resent a classic collisional suture but rather a basin that was 300 km partly floored by oceanic crust and that was partly preserved despite regional accretionary processes. Of concern is the segment of the Bangong/Nujiang zone Fig. 11 Ophiolites (black; after Kidd et al. 1988, their Fig. 9) and ex- north and northwest of Lhasa. The Lhasa block and the tent of the Bangong/Nujiang zone (BNZ) western and central segment of the Qiangtang block display significant similarities as to their Carboniferous to Triassic shelf facies development (Bally et al. 1980; Chang and Pan 1981; Leeder et al. 1988; Chang et al. 1989; Taner and Bangong/Nujiang zone Meyerhoff 1990a, b) which are summarized in Fig. 12. The Scythian (Lower Triassic) is the main interval of terrestrial The ophiolite-bearing Bangong/Nujiang zone runs at an ap- deposition. Since Scythian terrestrial deposits are restricted proximate trend of 100° through central Tibet. It separates to only the Qiangtang block, and since the Scythian is pos- the northerly located Qiangtang block from the Lhasa sibly as short as only 4 m.y. (Harland et al. 1990), we do not block. Judged by the distribution of ophiolites (Fig. 11), attribute the occurrence of terrestrial conditions to a major, map studies (Liu et al. 1988; enclosure in Kidd et al. 1988) rift-related, regional uplift. Minor terrestrial intervals are and the definition of the northern limit of the Lhasa block known from the Upper Permian and Upper Triassic of the by Burg et al. (1983), the Bangong/Nujiang zone is relative- Qiangtang block where alternations of marine and terrestri- ly narrow in the west and widens significantly in eastern Ti- al deposits occur. We attribute the younger alternations to bet (Fig. 11). Its length may exceed 1700 km as its termina- uplift tendencies related to the late Triassic Songpan-Ganzi tions are ill-defined (Fig. 11). terrane/Qiangtang block accretion which occurred north of Early interpretations considered the Bangong/Nujiang the Qiangtang block (Dewey et al. 1988). Only from the zone a consequence of plate convergence sensu the Wilson late Triassic on were major facies differences recognized to cycle. For example, Pan and Zheng (1983), Zheng et al. the north and south of the Bangong/Nujiang zone by Chang (1984), and Coulon et al. (1986) viewed it as a collisional et al. (1989). They concluded that the Lhasa and Qiangtang suture, and Girardeau et al. (1985) considered a late Juras- blocks formed one continental terrane until the late Triassic sic to early Cretaceous ophiolite obduction connected with when the Bangong/Nujiang basin opened between them. long-range thrusting on the order of 150 km. Xu et al. Rifting resulted in the formation of oceanic crust, but still (1985) assumed a mid-Jurassic collision. Later, it was real- during the Triassic (Sengör 1981; Wang et al. 1987; Taner ized that the Bangong/Nujiang zone displays some features and Meyerhoff 1990b). We found the Jurassic basin fill to which are difficult to reconcile with a conventional colli- be generally fine-grained, dominated by dark shale with in- sion and a preceding subduction interval since “precolli- tercalations of fine-grained turbidites (W. Schneider and F. sional” and “postcollisional” deformation are minor (Cow- Mattern, unpublished data). Thus, the fine-grained facies is ard et al. 1988; Dewey et al. 1988) and signs for related closely comparable to that of the Gulf of Elat and the cen- magmatism are absent (Dewey et al. 1988). Coward et al. tral and southern Gulf of California as shown above. The (1988) also discussed the mechanical problems of long- Jurassic is also characterized by a high frequency of local range ophiolite thrusting. Realizing some features of the unconformities and great thicknesses of up to several kilo- Bangong/Nujiang zone which are unusual for a classic op- meters (Yu et al. 1991). Based on our field observations, we hiolite suture zone, Taner and Meyerhoff (1990b) put for- interpret the Bangong/Nujiang basin as asymmetric, with a ward the idea that it resembles a rift. More recently, Yu et steep northern basin slope from which olistostromes and al. (1991) deduced that the Bangong/Nujiang zone is a large olistoliths of northern provenance originated and a strike-slip/pull-apart rift. gentle southern basin slope which lacks such deposits. We This relatively new interpretation of the Bangong/Nuji- believe that Yu et al. (1991) arrived at the same interpreta- ang zone may take some geoscientists by surprise. Howev- tion of sharply contrasting basin slope gradients prior to us. er, this interpretation eliminates some major difficulties in However, their basin cross sections lack compass direc- the understanding of the Bangong/Nujiang zone. In agree- tions. The asymmetry of the Bangong/Nujiang basin is in ment with the rift interpretation is the occurrence of similar support of the strike-slip basin interpretation (compare late Paleozoic to Triassic sedimentary successions to the Mitchell and Reading 1989; Ben-Avraham 1992; Ben- north and south of the Bangong/Nujiang zone (see below). Avraham and Zoback 1992). Moreover, the rift interpretation is able to explain the scat- There are indications for rift processes. Late Triassic ba- tered ophiolite distribution in a relatively large region with- sic–intermediate volcanites and pyroclastic rocks of the 218

Fig. 12 Sedimentary prerift and Lhasa Block Central Qiangtang Block synrift margin development of the Lhasa block and central Volcanites Limestone, Alternate terr. and Qiangtang block. x Terrestrial T3 300-1000 m shale marine clastics and intercalations. Note the similar >500 m limestone 2000-3000 m Carbonate shallow marine facies and the T2 Limestone, sst. 100 m scarcity of terrestrial deposits on Limestone, intercalated platform Limestone, shale 450 m both terranes. (From Yin et al. with ignimbrite 300 m T1 Terr. Coal-bearing sst. 300 m 1988; Leeder et al. 1988; Taner and Meyerhoff 1990a) Sandstone, siltstone P2 Coal, sst., shale, and andesite >100 m limestone 230 m Limestone, partly Limestone, intercalated bioclastic >640 m Carbonate P1 w. shale and volcanites Argillaceous limestone ramp 1900 m and shale 40 m Dropstone diamictites, Clastic Dropst. diamict., siltstone, C2 mudstone, sandstone shelf basin sst., sandy limestone 200 m >1000 m Sandstone, siltstone, Sandstone, intercalated bioclastic limestone, C1 with shale, limestone intercalted with brec- and tuff >900 m ciated limestone 1900 m

Lhasa block which occur along the line of the Qiling Co tween the Austroalpine and Middle Penninic regions. The Lake and Coqen (north central Lhasa block) reflect the ini- paleogeographic positions of the interpreted stratigraphic tial opening (Chang et al. 1989). Permian volcanism within sections of different thrust units (Calcareous Alps, Ortler, the Lhasa and Qiangtang blocks was attributed to exten- Ela, Err, Arosa, Schams and Tasna nappes) are shown in sional tectonics of back-arc or rift affinities (Leeder et al. Fig. 13. 1988). These authors also envisaged the fragmentation of Effects of Triassic rifting, such as volcanism and fault- the Triassic Lhasa carbonate platform through extensional ing, are well-known for the eastern and southern Alps (e.g., faulting which accounts for observed bathymetric differ- Bechstädt et al. 1976, 1978). Triassic alkaline volcanites ences. Moreover, Permian volcanites of the Kaixin ridge in occur in the rifted region (e.g., Bechstädt et al. 1976, 1978) the northern part of the Qiangtang terrane were geochemi- as well as more calcalkaline (and calcalkaline to sho- cally interpreted as originating “within attenuated conti- shonitic and shoshonitic) ones (Barbieri et al. 1982; Bosse- nental lithosphere during rifting” (Pearce and Mei 1988). lini et al. 1982; Lucchini et al. 1982; Crisci 1984; Oben- These basalts to andesites and trachyandesites display holzner 1986; Castellarin et al. 1988; Sloman 1989). The some features indicating a transitional composition be- petrogenesis of magmatites of the latter type(s) has often tween tholeiitic and alkaline (Pearce and Mei 1988). been attributed to subduction beneath the region (Lucchini During the Jurassic spreading stage, marine conditions et al. 1982; Obenholzner 1986; Sloman 1989; Bonadiman prevailed in the partly deep marine basin (Leeder et al. et al. 1994) in a strike-slip setting (Sloman 1989; Bonadi- 1988; Yin et al. 1988). Lower Jurassic postrift sequences are largely missing on the Lhasa and Qiangtang blocks (e.g., Smith and Xu 1988; Yin et al. 1988). This is attributed to dynamic uplift induced by the Songpan-Ganzi terrane/Qiangtang block accretion. As indicated by the stratigraphic record (Kidd et al. 1988, their map; Leeder et al. 1988) of fairly continuous marine shelfal sequences from the late Paleozoic to the Tri- assic, there is no indication for a major rift-related thermal uplift prior to, as well as during, rifting – neither regionally widespread nor vertically significant, nor timewise important. Thus, the stratigraphic record supports the strike-slip rift model by Yu et al. (1991).

South Penninic zone

The South Penninic zone separates the continental Austro- alpine basement and cover nappes from those of the continental Middle Penninic domain at the western/eastern Alps’ transition. The ophiolites of the Arosa nappe repre- Fig. 13 Uppermost Jurassic paleogeography of the Penninic region of the western/eastern Alps transition (From Weissert and Bernoulli sents the rift and spreading zone of the narrow South Pen- 1985; Schmid et al. 1990). SPZ South Penninic zone. Distribution of ninic ocean basin (Frisch et al. 1994) which formed be- oceanic crust after Schmid et al. (1990) 219 man et al. 1994). Sloman (1989) suggested that the vol- The major rifting stage occurred between 200 and 165 canites may have inherited their “arc signature” from Ma (Lemoine and Trümpy 1987) when significant vertical Permo-Carboniferous subduction of oceanic lithosphere. and lateral facies changes as well as thickness variations Bonadiman et al. (1994) considered mantle sources for the over short distance developed on the platform accompanied volcanites which inherited subduction-related geochemical by intense faulting. Fault-controlled breccias, sediment- components from the preceding Variscan orogenic cycle. filled tectonic fissures, and high-angle unconformities be- Crisci et al. (1984) had previously assumed a source by par- tween Triassic and Jurassic strata are characteristic. This tial melting during the early stage of rifting from an upper stage during which the carbonate platform was fragmented mantle strongly modified by the Variscan orogeny and con- is especially well studied in the Austroalpine units and the taminated by crustal material. southern Alps (e.g., Winterer and Bosselini 1981; Froitz- During the Triassic, the area in which the future South heim and Eberli 1990). Uplift above sea level may have oc- Penninic ocean basin formed was a shallow marine conti- curred locally, but the distinction between subaerial and nental platform. This platform was gently but noticeably submarine erosive features is problematic (Schlager and subdivided into basins and structural highs by tectonic Schöllnberger 1974). Nevertheless, marine conditions were movements (e.g., Bechstädt et al. 1976, 1978; Winterer and still predominant during the Jurassic. Bosselini 1981). The structural highs exhibit stratigraphic In the course of divergent plate motion, subcontinental successions of reduced thickness (Winterer and Bosselini mantle was locally exhumed (Trommsdorff et al. 1993). 1981). They also served as the foundation of organic car- Oceanic spreading began during the mid–Jurassic (Chan- bonate buildups (e.g., Bechstädt et al. 1976, 1978). Lower nell and Kozur 1997) and the age of the oldest beds above Mesozoic shallow marine deposits are continuous through- the regional breakup unconformities is usually Dogger out the region (Naef 1987; Trümpy 1988). However, in (Lemoine and Trümpy 1987). Deep marine conditions are proximity to the future South Penninic basin, the thickness indicated by Upper Jurassic radiolarites in the Austroalpine of Triassic formations is markedly reduced (Fig. 14) which and South Penninic realms, although it has to be considered is attributed to rift-related doming prior to a rift stage of in- that the calcite compensation depth was relatively shallow creased faulting. Doming is laterally widespread but verti- then (Bosselini and Winterer 1975). Among the ophiolites cally limited. Earlier workers viewed the Triassic rift dome of the Arosa nappe, basalts displaying the characteristics of as a “geanticline” (Cadisch 1932) or a “persistent high” within-plate tholeiite and mid-ocean ridge basalts are (Naef 1987). Doming did not create a significant general known (Frisch et al. 1994). Alkaline dikes intruded the change in the depositional environment. Shallow marine strata of the Calcareous Alps during the mid-Cretaceous conditions were still predominant despite the fact that Carn- when no subduction took place beneath them (Trommsdorff ian deposits were at times subaerially exposed and the et al. 1990). The setting could have been a system of horsts Norian dolomite may show signs of supratidal conditions and grabens possibly related to a transpressive tectonic set- (Naef 1987). The overall character of the rift region re- ting (Trommsdorff et al. 1990). mained that of a carbonate platform. The shallow marine Triassic and Jurassic stratigraphic record indicates no major vertical uplift. Again, strike-slip- dominated rifting is assumed. However, a “pure” strike-slip mode appears unlikely. Since there is some evidence for Mid - Penn. SPZ Austroalpine doming, an oblique type of rifting with a dominant strike- TASNA SCHAMS AROSA ERR ELA ORTLER NCA 0 km slip component is considered. This view is in agreement with recent tectonic models by numerous workers (Weissert and Bernoulli 1985; Trümpy 1988; Schmid et al. 1990 and sources within these works) who interpreted the entire Pen- ninic region as a Mesozoic transform zone. The South Pen- 1 km ninic basin is portrayed as an oblique rift with a dominant component of strike-slip motion (Weissert and Bernoulli Rhaetian Kössen Fm. 1985; Schmid et al. 1990). According to Baud and Sept- Rhaetian Bunter fontaine (1980), Jurassic erosion removed Upper Triassic Keuper Fm. and Lower Jurassic strata of the Préalpes Médianes which Norian dolomite 2 km belong to the southeastern margin of the Middle Penninic region. Stampfli and Marthaler (1990) interpreted this ero- Carnian Raibl Fm. sional truncation as an indication for shoulder uplift during Mid-Triassic carbonate the late Lias. This view is not in disagreement with our in- NCA Northern Calcareous Alps terpretation since rifting of the South Penninic zone was in- (west part of Lechtal nappe) 3 km creasingly more orthogonal towards the west–southwest (Fig. 14). Fig. 14 Thickness variations of Triassic deposits of Austroalpine, South, and Middle Penninic nappes of the eastern/western Alps transition. The Arosa nappe is a mélange. Volcanites occur in different formations (From Schmid 1965; Gwinner 1978; Schnabel 1980; Naef 1987) 220 rift basin as well as in the opposite direction. In such a case, Conclusion pronounced stratigraphic gaps caused by erosion are expected to be present especially on the rift shoulders. Along strike-slip margins, pull-apart depocenters may Our examples show that strike-slip rifting is an important occur exhibiting great thicknesses of sediments which ac- mechanism to split continental crust. Strike-slip rift zones cumulated in a relatively short time span. Syntectonic dep- can measure several hundred kilometers (Suez rift, South osition is typical and expressed by local angular un- Penninic zone) to 1000 km (east Mediterranean strike-slip conformities, abrupt lateral facies, and thickness changes. fault including the Pantelleria rift as shown in Fig. 3, Ro- This also holds true for strike-slip-dominated oblique rifts. manche transform, Dead Sea transform) in length or even These, however, are expected to display reduced thickness- significantly in excess of 1000 km (e.g., Bangong/Nujiang es along the central rift axis and/or erosional truncations zone, San Andreas transform). (South Penninic zone). Strike-slip rifts seem to occur not far from continental Continental strike-slip rifts represent “low-volcanicity margins which in some cases are of overall active character rifts” at best (compare Barberi et al. 1982). Thus, the pres- (those located between Africa and Eurasia; San Andreas). ence of minor amounts of volcanites is a part of the rift re- The Romanche fracture zone, however, formed in the midst gion’s stratigraphy. The investigated examples showed that of a large continental plate because it transformed the mo- such volcanic rocks are generally alkaline and at times per- tion between two segments of a huge dip-slip rift system. alkaline and subalkaline. To which degree the volcanic Strike-slip failure of continental crust without a signifi- stratigraphic record can directly serve as a diagnostic crite- cant dip-slip component neither requires precursory nor in- rion for strike-slip rifts is premature to say. On one hand, duces synrift thermal doming. The investigated examples we are aware of the fact that alkaline volcanites also occur show that there is no evidence of persistent major vertical in dip-slip rifts among which “low-volcanicity rifts” are and lateral thermal uplift in direct genetic relation with con- known (Barberi et al. 1982); on the other, there may be par- tinental strike-slip rifts. Only in the case of the oblique-slip ticular circumstances related to strike-slip magmatism South Penninic rift zone can moderate lateral and minor which are distinctly different from graben-type vertical uplift be detected prior to and during rifting. This magmatism, such as an overall normal heat flow (at least at uplift can be interpreted as being thermally supported. the surface), low regional attenuation of continental crust, Close inspection of strike-slip rifts indicates that uplift and only minor decompression melting. These circum- may be attributed to regional effective plate convergence stances may (or may not?) have an impact on magma evo- (Dead Sea transform: Taurus/Zagros collision zone), lateral lution (e.g., degree of contamination from continental heat transfer from relatively young oceanic lithosphere in crust). Scrutton (1982) pointed to the low igneous activity the course of transform evolution (Romanche transform), or at sheared margins in their early, active phase and suggest- to a change from strike-slip to dip-slip rifting (Suez rift). ed that lateral shear zones are “not conduits for magma, These uplift causes are, however, not genetically related to probably simply because there is no magma beneath them” the strike-slip rift process. (see also De Caprona 1992, p. 22; for localization of mag- At a local scale, some short-lived uplift among strike- matic activity along strike-slip faults see Aydin et al. 1990). slip rifts may be related to complex structural fault-block We regard the aspect of possible differences between strike- interaction as is the style of strike-slip environments (e.g., slip and graben volcanites as a future research opportunity. Harding 1974, 1985; Sylvester 1988). We pointed out the Our analysis of the stratigraphic record of the South presence of transpressive effects in some examples. Penninic zone and the Bangong/Nujiang zone lends support We conclude that the larger the strike-slip component of to recent strike-slip rift models for these two areas which a rift is, the less likely major doming is associated with the represent ancient rifted margin in mountainous settings. rift zone. Thus, analyzing the prerift and synrift stratigraphic record of rifted margins with regard to their uplift or sub- Acknowledgements This manuscript was reviewed by W. R. Buck sidence history represents a significant potential to decipher and B. C. Burchfiel. Their comments are greatly appreciated. We thank K.-J. Reutter for reviewing a previous version of the manuscript the mode of paleorifting. Strike-slip rift margin successions and Z. Ben-Avraham for supplying difficult-to-get literature. We are can be identified by the lack of major prerift and synrift indebted to the “Deutsche Forschungsgemeinschaft” for financial thermal doming in their stratigraphic record. The lack of support. We hope we did justice to all quoted authors. such doming can be most readily determined in submerged platform settings. In these cases, the overall character of the rifted regions remains that of a platform and the marine References strata are generally continuous. 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