Special Paper 480 . . THE GEOLOGICAL SOCIETY • OF AMERICA®
Mélanges: Processes of Formation and Societal Significance
edited by
John Wakabayashi Department of Earth and Environmental Sciences California State University, Fresno Fresno, California 93740 USA Yildirim Dilek Department of Geology Miami University Oxford, Ohio 45056 USA
Special Paper 480 3300 Penrose Place, P.O. Box 9140
Boulder, Colorado 80301-9140, USA
2011
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Contents Introduction: Characteristics and tectonic settings of mélanges, and their significance for societal and engineering problems . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v John Wakabayashi and Yildirim Dilek Part I. Mélange Generation in Oceanic Fracture Zones in Abyssal Settings 1. Serpentinite matrix mélange: Implications of mixed provenance for mélange formation . . . . . . . 1 John W. Shervais, Sung Hi Choi, Warren D. Sharp, Jeffrey Ross, Marchell Zoglman-Schuman, and Samuel B. Mukasa 2. Geochemical mapping of the Kings-Kaweah ophiolite belt, California—Evidence for progressive mélange formation in a large offset transform-subduction initiation environment . . . . . . . . . . . 31 J. Saleeby Part II. Mélange Formation Associated with Subduction Initiation 3. Constraints on the evolution of the Mesohellenic Ophiolite from subophiolitic metamorphic rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 75 R. Myhill 4. Role of plutonic and metamorphic block exhumation in a forearc ophiolite mélange belt: An example from the Mineoka belt, Japan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 95 Ryota Mori, Yujiro Ogawa, Naoto Hirano, Toshiaki Tsunogae, Masanori Kurosawa, and Tae Chiba Part III. Mélange Development in Subduction-Accretion Complexes and in Collisional Settings 5. Mélanges of the Franciscan Complex, California: Diverse structural settings, evidence for sedimentary mixing, and their connection to subduction processes . . . . . . . . . . . . . . . . . . . . . . . 117 John Wakabayashi 6. Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili, central Turkey . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 143 Anne Dangerfield, Ron Harris, Ender Sarıfakıoğlu, andYildirim Dilek 7. Petrology of a Franciscan olistostrome with a massive sandstone matrix: The King Ridge Road mélange at Cazadero, California . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 171 Rolfe Erickson 8. Sedimentary block-in-matrix fabric affected by tectonic shear, Miocene Nabae complex, Japan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 189 Soichi Osozawa, Terry Pavlis, and Martin F.J. Flower
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Contents 9. Numerical estimation of duplex thickening in a deep-level accretionary prism: A proposal for network duplex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 207 Hikaru Ueno, Ken-ichiro Hisada, and Yujiro Ogawa 10. Tectonic, sedimentary, and diapiric formation of the Messinian mélange: Tertiary Piedmont Basin (northwestern Italy) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 215 Andrea Festa 11. Recognition of a trench-fill type accretionary prism: Thrust-anticlines, duplexes, and chaotic deposits of the Pliocene-Pleistocene Chikura Group, Boso Peninsula, Japan . . . . . . . . 233 Satoru Muraoka and Yujiro Ogawa 12. Implication of dark bands in Miocene–Pliocene accretionary prism, Boso Peninsula, central Japan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 247 Yoko Michiguchi and Yujiro Ogawa Part IV. Significance of Mélanges for Engineering and Applied Geology 13. Geopractitioner approaches to working with antisocial mélanges . . . . . . . . . . . . . . . . . . . . . . . . 261 Edmund W. Medley and Dimitrios Zekkos
The Geological Society of America Special Paper 480 2011
Introduction: Characteristics and tectonic settings of mélanges, and their significance for societal and engineering problems John Wakabayashi Department of Earth and Environmental Sciences, California State University at Fresno, Fresno, California 93740, USA Yildirim Dilek Department of Geology, Miami University, Oxford, Ohio 45056, USA
Mélanges occur widely in collisional and accretionary orogenic belts around the world and represent mappable geological units consisting of blocks of different ages and origin, commonly embedded in an argillitic, sandy, or serpentinite matrix showing high stratal disruption and a chaotic internal structure. Understanding the mélange-forming processes and the significance of mélanges and related units in the geological record is of first-order significance in documenting the tectonic evolution of mountain belts; therefore, these chaotic rock units have attracted much attention in field-based structural studies since the nineteenth century. The term mélange has evolved to cover tectonic, sedimentary, and/or diapiric processes (Silver and Beutner, 1980) and tectonic settings of mélange formation, since its first use in 1919 by the British geologist Edward Greenly for the “Gwna Group” of the Mona Complex in Anglesey, north Wales (Greenly, 1919). In his classic work in the Franciscan Complex, Hsü (1968) proposed to use “mélange” only for tectonic mélanges and therein started a long-lived debate on the definition of the mélange term as well as on the processes involved in mélange formation. This controversy on the definition and formation of mélanges is livelier than ever in present time and requires a more systematic approach in mélange studies and better communication among the mélange researchers. To this end, we organized and convened a topical session on mélanges during the Geological Society of America (GSA) Annual Meeting in Denver in 2007. The session was well attended by many international scientists from North America, Europe, and the Pacific Rim countries. This session was sponsored by the GSA International and the Structural Geology and Tectonics Divisions, and it brought together earth scien-
tists in research communities from around the world, who do not ordinarily interact at the same meetings, in order to add an interdisciplinary dimension to our discussions on mélanges. It provided an excellent forum to discuss the new advances on the mélange concept as well as the diverse mélange types and mélange-forming processes based on some case studies. This Special Paper emanated from this successful GSA topical session. The benchmark GSA Special Paper 198 on mélanges, published more than 25 years ago (Raymond, 1984a), continues to influence research on mélanges, as testified by the frequent citation of the papers in it. The papers in the current volume build upon the solid foundation provided by the papers of Special Paper 198, as well as those published before and since, and include applications of new methodologies, exploration of new subjects, and a more international focus. The geographic spread of mélange localities around the world is also broader in parallel with the larger international authorship in the current volume. Given the three-dimensional (3-D) complexity of mélanges, it is of little surprise that field studies form the foundation of all of the research presented in this volume. Beyond this common linkage, the papers here span a broad spectrum of features and focus. We streamlined the chapters according to a relative conceptual chronology related to mélange development (or impact) and related processes that begin with formation on the abyssal ocean floor (Part I), then proceed to subduction initiation (Part II), and accretionary wedge development and orogenic belt formation (Part III). The final section concentrates on the impact of mélanges on societies by way of their engineering properties (Part IV). We provide below first a synoptic summary of the
Wakabayashi, J., and Dilek, Y., 2011, Introduction: Characteristics and tectonic settings of mélanges, and their significance for societal and engineering problems, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. v–x, doi:10.1130/2011.2480(00). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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chapters in this volume, and then we conclude with some general statements about mélanges and their significance. PART I. MÉLANGE GENERATION IN OCEANIC FRACTURE ZONES IN ABYSSAL SETTINGS The papers by Shervais et al. (Chapter 1) and Saleeby (Chapter 2) present integrated field, petrological, and geochemical data on serpentinite-matrix mélanges of different ages in California. Both studies provide field and geochemical evidence that blockin-matrix fabrics developed in serpentinite within an abyssal fracture zone. Shervais et al. present field evidence for subduction initiation along a fracture zone based on the presence of highpressure, high-temperature (HP-HT) garnet-amphibolite blocks from a serpentinite shear zone structurally beneath unmetamorphosed ophiolitic rock of the Jurassic Coast Range ophiolite. The serpentinite matrix mélange may represent the original subduction interface. The authors suggest that exhumation was aided not only by the lower density of the serpentinite matrix compared to the top of the lower plate (oceanic crust) and the upper plate (suboceanic mantle), but was also driven by the positive volume change associated with serpentinite in a confined zone that might have served to force the serpentinite upward. Saleeby presents evidence for subduction initiation along an abyssal fracture zone, based on his work in the Kings-Kaweah ophiolite belt of the southern Sierra Nevada. He suggests a long duration of time (~190 m.y.) between the early Ordovician abyssal ocean crust formation and the Permo-Carboniferous mélange development in an abyssal fracture zone environment. He argues that initiation of subduction along this fracture zone was followed by the development of supra-subduction zone igneous rocks. The evidence for subduction initiation comes from a ca. 255 Ma Sm/Nd age and (HP-HT) metamorphism of garnet-amphibolite blocks in a serpentinite matrix mélange. Saleeby argues that the emplacement of garnet-amphibolite blocks in the mélange was related to serpentinite diapirism through the upper plate of a subduction system, instead of back along the subduction interface through channel flow. PART II. MÉLANGE FORMATION ASSOCIATED WITH SUBDUCTION INITIATION The rock record of subduction initiation is evaluated in Myhill’s paper (Chapter 3) on the metamorphic sole of the Vourinos and Pindos ophiolites in the western Hellenides (Greece). This topic of subduction initiation is also covered in the papers by Shervais et al. and Saleeby (Part I), and to a lesser extent, in the papers by Mori et al. (Mineoka belt, Japan, Chapter 4), and Wakabayashi (Franciscan Complex, California, Chapter 5) in Part III. Myhill presents detailed metamorphic evidence and argues that metamorphic soles, which are the thin high-grade metamorphic sheets commonly found beneath Tethyan ophiolites, were formed at lower pressures than commonly thought, and that they
were therefore not necessarily associated with subduction initiation as has been widely assumed. He demonstrates that the hightemperature metamorphism of the metamorphic sole beneath the Mesohellenic ophiolites (Pindos and Vourinos) occurred during intra-oceanic thrusting (but not subduction) near a ridge crest and soon after subduction initiation, and that slices and blocks of the sole were incorporated into a subophiolitic mélange during further thrusting associated with ophiolite emplacement. Mori et al. (Chapter 4) present a model for the complex tectonic evolution of the Mineoka ophiolitic mélange belt in the Boso Peninsula of central Japan, based on field relations, geochronology, and petrology. The evolution of this mélange includes early HP-HT, possibly associated with a subduction initiation event, followed by considerable deformation and mixing involving triple junction interaction and evolution. An early stage of ductile deformation at deep crustal levels was associated with syn-subduction exhumation of metamorphic rocks, including HP-HT rocks, followed by a brittle phase of deformation developed at much shallower levels, as rocks were incorporated into the mélange zone. Their geochronological data suggest initiation of subduction at ca. 33–39 Ma, followed by development of the ophiolitic mélange between 15 and 18 Ma. Mori et al. propose that the present Izu arc may be an analog of the Mineoka mélange belt. PART III. MÉLANGE DEVELOPMENT IN SUBDUCTION-ACCRETION COMPLEXES AND IN COLLISIONAL SETTINGS In his process-oriented approach to delineating the tectonic settings of mélange formation, Wakabayashi (Chapter 5) divides and examines the classic Jurassic-Eocene Franciscan Complex mélanges of coastal California into distinct structural groups. The structurally highest mélanges in the Franciscan Complex may have formed at or shortly after subduction initiation, marking the initial subduction interface, whereas the mélanges separating coherent nappe sheets may represent later-developed paleo megathrust horizons within the accretionary prism. He argues that the large displacements associated with these nappes may have been largely accommodated along the borders of the mélanges rather than within them. He presents field and petrographic evidence supporting pre-tectonic sedimentary mixing of mélanges (development of block-in-matrix structure and introduction of exotic blocks), including those most likely to be classified as entirely tectonic mélanges (internappe mélanges). He shows the presence of “two cycle” high-P rocks, which were subducted to blueschist-facies depths, then exhumed and re-worked as sedimentary deposits, and then resubducted again to blueschist depths and exhumed. Dangerfield et al. (Chapter 6) present structural, geochronological, and geochemical data from the Eldivan ophiolite, which occurs as a coherent block in the Ankara Mélange in north-central Turkey. The Ankara Mélange is part of the İzmir-Ankara-Erzincan
Introduction: Characteristics of mélanges, and their significance for societal and engineering problems suture zone and represents a classic Tethyan colored mélange. Dangerfield and her co-authors show development of the Eldivan ophiolite in a suprasubduction zone setting, followed by its integration into the Ankara Mélange as an oceanic block. Detrital zircon U/Pb analyses from the mélange and the overlapping epiclastic sandstones show that mélange development occurred between ~143 Ma and 105 Ma, consistent with the regional geochronological data. The authors argue that although the development of the İzmir-Ankara-Erzincan suture zone involved continental collision tectonics, its overall evolution resembles the formation of mélange terrains in the southwest Pacific rather than that of a Himalayan-type continental collision. Erickson (Chapter 7) presents field, petrographic, and geochronological data from a Cretaceous sandstone-matrix olistostrome in the Franciscan Complex in northern California that collectively provide critical constrains on the exhumation age and patterns of various blocks and the depositional age of the olistostrome. The majority of these blocks themselves are Franciscan-derived. The evolution of the olistostrome includes initial subduction burial of the blocks; their subsequent exhumation and exposure as blueschist-, eclogite-, and amphibolite-facies blocks; their deposition some time after 83 Ma; and partial re-subduction to prehnite-pumpellyite facies conditions subsequently. As shown in the paper by Wakabayashi, Erickson’s work also demonstrates the sedimentary reworking of previously metamorphosed Franciscan rocks, including “high-grade” blocks formed during the earliest stages of Franciscan subduction at ca. 165 Ma. Osozawa et al. (Chapter 8) use map and outcrop relationships from excellent coastal exposures of the Miocene Nabae complex of Japan and petrofabric studies to show that the blockin-matrix fabric observed in this mélange was a result of early sedimentary sliding rather than tectonism. They demonstrate that the amount of shear strain associated with foliation development in the mélange matrix was minimal, and that this deformation was vastly inadequate to account for the introduction of exotic blocks of chert and basalt into the shale matrix. Their data also show evidence for reworking of clasts that include penetrative fabrics developed in an older subduction complex. Ueno et al. (Chapter 9) document complex duplex structures from accretionary complex rocks of the Jurassic-Cretaceous Chichibu Belt of Japan, and show how some of these structures have been previously (and mistakenly) interpreted as block-in-matrix features in the absence of good exposures. The superb coastal exposures allow detailed characterization of the structures and estimation of the amount of structural thickening associated with tectonic duplexing. They describe “network duplexes,” which are themselves composed of duplexes of smaller orders, and calculate thickening of a factor of ~6–13 at the greenschist facies level of this subduction-accretionary complex. Festa (Chapter 10) presents detailed field relationships from the Piedmont Basin in northwest Italy that developed as an episutural basin after the main stage Alpine collision, and he documents structures that formed at burial depths of 2–3 km. Utilizing sedimentary structures and sedimentary contact relationships (for
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establishing sedimentary origins), shear-sense indicators, and progressive strain and rotation of rocks toward faults, he distinguishes between mélanges that were formed by tectonic strain, sedimentary sliding, and diapiric emplacement. He then links the development of different types of mélanges to the regional tectonics, during which faulting (along with the formation of tectonic mélanges) may have triggered gas hydrate disassociation and rise of overpressured fluids (diapiric emplacement, preferentially following fault zones), triggering gravitational collapse and development of sedimentary mélanges. Muraoka and Ogawa (Chapter 11) present observations on mélanges, duplexes, and folds that they interpret to have formed in a trench-fill environment, the shallowest level of preservation of an accretionary prism. The evidence comes from fine coastal exposures of the Plio-Pleistocene Chikura Group on the Boso Peninsula of Japan. The Lower Chikura Group units are interpreted to have been deposited in the trench in advance of the thrust front and later incorporated into the accretionary wedge by seaward propagation of the thrust front; the Upper Chikura Group units, on the other hand, were originally deposited in a trench slope basin setting. The lower Chikura Group deposits include evidence for interaction of methane-rich fluids from associated chemosynthetic biocommunities that suggest a trenchfill environment similar to the modern Sagami Trough. Chaotic deposits or mélanges include those with diapiric (intrusive) field relationships as well as those that appear to represent submarine slides, whereas the duplexes and thrust anticlines record significant tectonic shortening in the coherent units. Michiguchi and Ogawa (Chapter 12) examine the internal structure of the Miocene-Pliocene accretionary prism complex exposed in the Boso Peninsula, Japan. They show that dark bands found in siltstones are the products of different deformation mechanisms in an accretionary prism toe and the frontal thrust region. The host rocks include both coherent stratal and chaotic units (as in mélanges). Their map, outcrop, and microscopic analysis suggests that some of these features formed as a result of high pore fluid pressure as shear fractures, whereas others formed as tensional fractures associated with different states of stress and deformation modes. One of their dark band types represents flexural-slip faults associated with folding, another type represents sliding planes formed during submarine landslides, whereas yet another type consists of thrust faults formed during accretion. PART IV. SIGNIFICANCE OF MÉLANGES FOR ENGINEERING AND APPLIED GEOLOGY The paper by Medley and Zekkos (Chapter 13) focuses on the geological engineering aspects of mélanges, bringing a societal relevance and significance to mélange studies and research. This topic has been largely overlooked in purely academic studies of mélanges, although it has been a subject of many detailed investigations in engineering geology, whose results have been published in the past 16 years. Up to now, many engineers and geologists in engineering and environmental geology force-fit
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block-in-matrix geology into a layer-cake stratigraphic interpretation, commonly with disastrous consequences, because this sector of applied geology has not kept pace with the most recent advances in recognition of mélanges. Medley and Zekkos fully describe the engineering issues of dealing with mélanges, including both theory and case studies. SUMMARY COMMENTS AND CONCLUSIONS Mélange Classification: Descriptive Rather than Genetic Schemes Recommended Most of the previous mélange studies, including those in GSA Special Paper 198, offered detailed classification or definitions of mélanges and their sub-types. It is clear that a uniform classification scheme has merit, given that the term mélange is used differently by many authors. However, caution is urged, especially when genetic significance is attached to a definition, given how difficult it may be to ascertain mélange origins from first-order field observations, particularly for mélanges that appear strongly deformed, such as those described by Osozawa et al. (Chapter 8) and Wakabayashi (Chapter 5). The classification schemes proposed by Cowan (1985) are largely descriptive and are therefore more useful than a genetic definition presented, for example, by Sengör (2003), wherein a purely tectonic origin is a requirement for the term mélange. Raymond (1984b) proposed a detailed classification scheme, but its ultimate application required some knowledge of the genesis of the mélange. Although this may seem a regressive definition, we recommend a broad definition of mélange as a bedrock unit with a matrix and variety of blocks included in it, similar to the recommendation of Silver and Beutner (1980). In fact Silver and Beutner (1980) noted that in addition to the more common block-in-matrix fabric, some mélanges have a block-on-block fabric, a structural style that appears to best fit the Mineoka ophiolite belt of Japan (Takahashi et al., 2003; also Mori et al., Chapter 4). Festa (Chapter 10) makes a similar recommendation for a descriptive, rather than genetic definition of the term mélange. Chapter 5 by Wakabayashi proposes mélange categories based on structural-tectonic settings that are derived from 3-D field relationships. This scheme has the primary goal of connecting the mélanges to large-scale processes during evolution of active plate margins, but it does not directly aid evaluation of strain and sedimentary processes in mélange formation in the way that a scheme such as Cowan’s (1985) does. Accordingly, we think that there is no single unifying classification or nomenclature scheme for mélanges, nor should there be, because different schemes serve different purposes. We recommend that authors writing about block-in-matrix units be as specific as possible about the descriptive aspects of these units, so that readers are not misled into applying their own definition of “mélange” that may differ markedly from that intended by the author. In many ways, the problem of mélange classification and nomenclature parallels that of the term ophiolite, for
which numerous definitions also exist (e.g., Dilek, 2003; Dilek and Furnes, 2009). Sedimentary versus Tectonic Mixing in Mélanges An increasing amount of field evidence has been presented in the past few decades, illustrating the significant contributions of sedimentary mixing to even some of the most (apparently) sheared mélanges (e.g., Aalto, 1989; Osozawa et al., 2009, Chapter 8; Wakabayashi, Chapter 5). These studies support the conclusions of earlier research (Cowan and Page, 1975; Cowan, 1978). Some of the most extreme examples include the sedimentary introduction of exotic blocks into nappe-bounding mélanges in the Franciscan Complex, which may have accommodated tens of km or more of displacement (Wakabayashi, Chapter 5). The studies of Osozawa et al. (2009, Chapter 8) also show that most or all exotic blocks in the Nabae Complex of the Shimanto Belt of Shikoku, Japan, and the Yuwan accretionary complex of the Ryukyu Islands, respectively, were integrated into the mélange by pre-tectonic phases of submarine sliding. Osozawa et al. (Chapter 8) argue that the deformation that produced the matrix foliation in the mélanges that they have examined records relatively minimal shear strain, which cannot account for introduction of exotic blocks or development of block-in-matrix fabrics. Aalto (1989) and Wakabayashi (Chapter 5) document a range of textures from undeformed sedimentary breccias to strongly foliated shale mélange matrix. Although evidence points to submarine sedimentary (gravity) sliding as a main contributor to the development of blockin-matrix fabrics in many of the most tectonized mélanges, sedimentary sliding was not a major process in the formation of block-in-matrix fabrics in all mélanges. Some mélanges clearly have a tectonic or diapiric origin. Festa (Chapter 10) summarizes effectively the criteria for distinguishing diapiric versus tectonic mélanges, and provides field examples of both. For diapiric mélanges, the diagnostic feature is opposing shear sense on opposite mélange contacts, a criterion that was first applied by Orange (1990) and subsequently used by Dela Pierre et al. (2007), as well as Muraoka and Ogawa (Chapter 11) and Festa (Chapter 10). For tectonic mélanges, an important field characteristic is an increasing degree of deformation and rotation of fabric elements as a fault or shear zone is approached (Festa, Chapter 10). Significance of Mélanges and Mélange Types in Orogenic Belt Development Mélanges are characteristic features of modern and ancient convergent plate boundaries, and rank with ophiolites and HP– low-temperature metamorphic rocks as critical recorders of convergent plate margin processes. Mélanges provide critical insights into sedimentary and structural evolution in the accretionary prism and forearc basin environments, including evidence for large-scale material movement (particularly in cross-sectional view) in accretionary wedges. Mélanges form as subduction of
Introduction: Characteristics of mélanges, and their significance for societal and engineering problems oceanic lithosphere is punctuated by a collisional process (see discussion in Dangerfield et al., Chapter 6, although they argue for a noncollisional origin for the particular mélange of their study), and/or terminated by the final stages of continental collision (Festa, Chapter 10). In addition to recording subduction- and collision-related sedimentary and tectonic processes, mélange formation may also include pre-subduction tectonics, including deformation along abyssal fracture zones (Shervais et al., Chapter 1; Saleeby, Chapter 2) and at oceanic core complexes (Saleeby, Chapter 2), as well as supra-subduction zone oceanic crust evolution (Dangerfield et al., Chapter 6; Shervais et al., Chapter 1). Mélanges also offer major insights into the most extreme vertical movements along convergent plate margins: the exhumation of high-pressure metamorphic rocks (Shervais et al., Chapter 1; Saleeby, Chapter 2; Mori et al., Chapter 4; Wakabayashi, Chapter 5; Erickson, Chapter 7). Societal Significance Mélanges, by the very nature of their chaotic block-inmatrix structure, pervasive and strong internal deformation and clay-rich soil contents, are prone to landsliding as well as creating problems because of the great contrast in ease of excavation of block and matrix (Medley and Zekkos, Chapter 13). Hence, they pose major challenges for engineering projects developed on them as well as for water supplies and infrastructure. Therefore, mélange terrains cause first-order societal problems for the people in California, Japan, Italy, Scotland, Greece, Cyprus, Turkey, the Philippines, and many other countries, where ophiolites and mélanges occur abundantly. Furthermore, most engineers and engineering geologists continue to treat mélangecontaining bedrock by using the basic principles of stratigraphy and by assuming a layered structure for their formation, and fail to account for the 3-D variation of many key parameters such as rock strength and ease of excavation. This ill-informed approach results in disastrous engineering problems, leading to significant property damage and casualties. It is thus highly important for the academic community and the practicing geological and civil engineers to convey their learned experience and knowledge on mélanges and mélange structures to each other through publishing in common literature and in conference proceedings in order to maximize the dissemination of their scientific and applied findings. The academic community in particular should continue to strive to remedy the knowledge gaps through interaction with the applied community as well as through implementing contemporary reforms in undergraduate education that would revive field instruction and field-based, observation-oriented earth science education. We hope that this GSA Special Paper presents an important step in this mission of closing the knowledge gap in the purely scientific and engineering aspects of mélanges and mélange-forming processes and their significance for engineering and societal issues. August 2010
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REFERENCES CITED Aalto, K.R., 1989, Franciscan Complex olistostrome at Crescent City, northern California: Sedimentology, v. 36, p. 471–495, doi:10 .1111/j.1365-3091.1989.tb00620.x. Cowan, D.S., 1978, Origin of blueschist-bearing chaotic rocks in the Franciscan Complex, San Simeon, California: Geological Society of America Bulletin, v. 89, p. 1415–1423. Cowan, D.S., 1985, Structural styles in Mesozoic and Cenozoic mélanges in the Western Cordillera of North America: Geological Society of America Bulletin, v. 96, p. 451–462, doi:10.1130/0016-7606(1985)96<451: SSIMAC>2.0.CO;2. Cowan, D.S., and Page, B.M., 1975, Recycled Franciscan material in Franciscan mélange west of Paso Robles, California: Geological Society of America Bulletin, v. 86, p. 1089–1095, doi:10.1130/0016 -7606(1975)86<1089:RFMIFM>2.0.CO;2. Dangerfield, A., Harris, R., Sarıfakıoğlu, E., and Dilek, Y., 2011, this volume, Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili, central Turkey, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(06). Dela Pierre, F., Festa, A., and Irace, A., 2007, Interaction of tectonic, sedimentary and diapiric processes in the origin of chaotic sediments: An example from the Messinian of Torino Hill (Tertiary Piedmont basin, northwestern Italy): Geological Society of America Bulletin, v. 119, p. 1107–1119, doi:10.1130/B26072.1. Dilek, Y., 2003, Ophiolite concept and its evolution, in Dilek, Y., and Newcomb, S., eds., Ophiolite Concept and the Evolution of Geologic Thought: Geological Society of America Special Paper 373, p. 1–16. Dilek, Y., and Furnes, H., 2009, Structure and geochemistry of Tethyan ophiolites and their petrogenesis in subduction rollback systems: Lithos, v. 113, p. 1–20, doi:10.1016/j.lithos.2009.04.022. Erickson, R., 2011, this volume, Petrology of a Franciscan olistostrome with a massive sandstone matrix: The King Ridge Road mélange at Cazadero, California, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(07). Festa, A., 2011, this volume, Tectonic, sedimentary, and diapiric formation of the Messinian mélange: Tertiary Piedmont Basin (northwestern Italy), in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(10). Greenly, E., 1919, The Geology of Anglesey: Great Britain Geological Survey Memoir, v. 1, 980 p. Hsü, K.J., 1968, The principles of mélanges and their bearing on the FranciscanKnoxville paradox: Geological Society of America Bulletin, v. 79, p. 1063– 1074, doi:10.1130/0016-7606(1968)79[1063:POMATB]2.0.CO;2. Medley, E.W., and Zekkos, D., 2011, this volume, Geopractitioner approaches to working with antisocial mélanges, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(13). Michiguchi, Y., and Ogawa, Y., 2011, this volume, Implication of dark bands in Miocene–Pliocene accretionary prism, Boso Peninsula, central Japan, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(12). Mori, R., Ogawa, Y., Hirano, N., Tsunogae, T., Kurosawa, M., and Chiba, T., 2011, this volume, Role of plutonic and metamorphic block exhumation in a forearc ophiolite mélange belt: An example from the Mineoka belt, Japan, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(04). Muraoka, S., and Ogawa, Y., 2011, this volume, Recognition of trench-fill type accretionary prism: Thrust anticlines, duplexes and chaotic deposits of Pliocene-Pleistocene Chikura Group, Boso Peninsula, Japan, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(11). Myhill, R., 2011, this volume, Constraints on the evolution of the Mesohellenic Ophiolite from subophiolitic metamorphic rocks, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal
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Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(03). Orange, D.L., 1990, Criteria helpful in recognizing shear-zone and diapiric mélanges: Examples from the Hoh accretionary complex, Olympic Peninsula, Washington: Geological Society of America Bulletin, v. 102, p. 935– 951, doi:10.1130/0016-7606(1990)102<0935:CHIRSZ>2.3.CO;2. Osozawa, S., Morimoto, J., and Flower, F.J., 2009, ‘Block-in-matrix’ fabrics that lack shearing but possess composite cleavage planes: A sedimentary mélange origin for the Yuwan accretionary complex in the Ryukyu island arc, Japan: Geological Society of America Bulletin, v. 121, p. 1190–1203, doi:10.1130/B26038.1. Osozawa, S., Pavlis, T., and Flowers, M.F.J., 2011, this volume, Sedimentary block-in-matrix fabric affected by tectonic shear, Miocene Nabae complex, Japan, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(08). Raymond, L.A., 1984a, editor, Mélanges: Their Nature, Origin and Significance: Geological Society of America Special Paper 198, 170 p. Raymond, L.A., 1984b, Classification of mélanges, in Raymond, L.A., ed., Mélanges: Their Nature, Origin and Significance: Geological Society of America Special Paper 198, p. 7–20. Saleeby, J., 2011, this volume, Geochemical mapping of the Kings-Kaweah ophiolite belt, California—Evidence for progressive mélange formation in a large offset transform-subduction initiation environment, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(02). Sengör, A.M.C., 2003, The repeated discovery of mélanges and its implications for the possibility and the role of objective evidence in the scientific enterprise, in Dilek, Y., and Newcomb, S., eds., Ophiolite Concept and the
Evolution of Geologic Thought: Geological Society of America Special Paper 373, p. 385–446. Shervais, J.W., Choi, S.H., Sharp, W.D., Ross, J., Zoglman-Schuman, M., and Mukasa, S.B., 2011, this volume, Serpentinite matrix mélange: Implications of mixed provenance for mélange formation, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(01). Silver, E.A., and Beutner, E.C., 1980, Melanges: Geology, v. 8, p. 32–34, doi:10 .1130/0091-7613(1980)8<32:M>2.0.CO;2. Takahashi, A., Ogawa, Y., Ohata, Y., and Hirano, N., 2003, The mode and nature of faulting and deformation and the emplacement history of the Mineoka Ophiolite, NW Pacific Rim, in Dilek, Y., and Robinson, P.T., eds., Ophiolites in Earth History: Geological Society of London Special Publication 218, p. 299–314. Ueno, H., Hisada, K.-I., and Ogawa, Y., 2011, this volume, Numerical estimation of duplex thickening in a deep-level accretionary prism: A proposal for network duplex, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(09). Wakabayashi, J., 2011, this volume, Mélanges of the Franciscan Complex, California: Diverse structural settings, evidence for sedimentary mixing, and their connection to subduction processes, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(05).
MANUSCRIPT ACCEPTED BY THE SOCIETY 21 DECEMBER 2010
Printed in the USA
The Geological Society of America Special Paper 480 2011
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation John W. Shervais Department of Geology, Utah State University, Logan, Utah 84322-4505, USA Sung Hi Choi Department of Geology and Earth Environmental Sciences, Chungnam National University, Daejeon 305-764, South Korea Warren D. Sharp Berkeley Geochronology Center, 2455 Ridge Road, Berkeley, California 94709, USA Jeffrey Ross* Department of Geosciences, Stony Brook University, Stony Brook, New York 11794-2100, USA Marchell Zoglman-Schuman† University of South Carolina, Columbia, South Carolina 29208, USA Samuel B. Mukasa Department of Geological Sciences, University of Michigan, Ann Arbor, Michigan 48109-1005, USA
ABSTRACT Serpentinite matrix mélange represents a significant, if less common, component of many accretionary complexes. There are two principal hypotheses for the origin of serpentinite mélange: (1) formation on the seafloor in a fracture zone–transform fault setting, and (2) formation within a subduction zone with mixing of rocks derived from both the upper and lower plates. The first hypothesis requires that the sheared serpentinite matrix be derived from hydrated abyssal peridotites and that the block assemblage consist exclusively of oceanic rocks (abyssal peridotites, oceanic basalts, and pelagic sediments). The second hypothesis implies that the sheared serpentinite matrix is derived from hydrated refractory peridotites with supra-subduction zone affinities, and that the block assemblage includes rocks derived from both the upper plate (forearc peridotites, arc volcanics, sediments) and the lower plate (abyssal peridotites, oceanic basalts, pelagic sediments). In either case, serpentinite mélange may include true mélange, with exotic blocks derived from other sources, and serpentinite broken formation, where the blocks are massive peridotite.
*Current address: Bechtel Savannah River, Inc., Aiken, South Carolina 29801, USA. † Current address: 2623 Withington Heights Peak Drive NE, Rio Rancho, New Mexico 87144, USA. Shervais, J.W., Choi, S.H., Sharp, W.D., Ross, J., Zoglman-Schuman, M., and Mukasa, S.B., 2011, Serpentinite matrix mélange: Implications of mixed provenance for mélange formation, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 1–30, doi:10.1130/2011.2480(01). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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Shervais et al. The Tehama-Colusa serpentinite mélange underlies the Coast Range ophiolite in northern California and separates it from high-pressure/temperature (P/T) metamorphic rocks of the Franciscan complex. It has been interpreted both as an accreted fracture zone terrane and as a subduction-derived mélange belt. Our data show that the mélange matrix represents hydrated refractory peridotites with forearc affinities, and that blocks within the mélange consist largely of upper plate lithologies (refractory forearc harzburgite, arc volcanics, arc-derived sediments, and chert with Coast Range ophiolite biostratigraphy). Lower plate blocks within the mélange include oceanic basalts and chert with rare blueschist and amphibolite. Hornblendes from three amphibolite blocks that crop out in serpentinite mélange and sedimentary serpentinite yield 40Ar/39Ar plateau ages of 165.6–167.5 Ma, similar to published ages of highgrade blocks within the Franciscan complex and to crystallization ages in the Coast Range ophiolite. Other blocks have uncertain provenance. It has been shown that peridotite blocks within the mélange have low pyroxene equilibration temperatures that are consistent with formation in a fracture zone setting. However, the current mélange reflects largely upper-plate lithologies in both its matrix and its constituent blocks. We propose that the proto-Franciscan subduction zone nucleated on a large offset transform fault–fracture zone that evolved into a subduction zone mélange complex. Mélange matrix was formed by the hydration and volume expansion of refractory forearc peridotite, followed by subsequent shear deformation. Mélange blocks were formed largely by the breakup of upper plate crust and lithosphere, with minor offscraping and incorporation of lower plate crust. We propose that the methods discussed here can be applied to serpentinite matrix mélange worldwide in order to understand better the tectonic evolution of the orogens in which they occur.
INTRODUCTION Accretionary Complexes and Mélange Formation Accretionary complexes formed during plate convergence, subduction, and collision represent a fundamental component of the plate tectonics paradigm, along with ophiolites, island arcs, and paired metamorphic belts. Accretionary complexes typically include a large fraction of material derived from the upper plate as sediments and as tectonic slices eroded from the overlying crust and lithosphere, as well as material derived from the lower, subducting plate by faulting and subduction accretion (Bailey et al., 1964; Moore and Karig, 1980; Moore et al., 1980; Hamilton, 1988; von Huene and Scholl, 1991). Deformation is dominantly compressional, but thickening of the accretionary wedge typically results in extensional deformation of the thickest part of the accretionary wedge even as compressional deformation continues in the thin leading edge (Platt, 1986). Accretionary complexes may also be deformed by large-scale, trench-parallel shear that both erodes and accretes new material (Moore et al., 1980; Howell, 1989; Wright and Wyld, 2007). Understanding the origin and significance of mélange assemblages is critical to our understanding of convergent plate boundaries, and places important constraints on the tectonic history of ancient orogens. Accretionary mélange complexes are conceptually and formatively tied to plate tectonics, and their
presence in any orogen reveals its plate tectonic origin. This is especially important for the Archean and early Paleoproterozoic, where mélange formation may represent our best evidence for the onset of Phanerozoic-style plate tectonics (Shervais, 2006). In their classic configuration, accretionary complexes consist largely of clastic sediments deposited within the trench as turbidites or submarine fan deposits, which are subsequently deformed by thrust faulting to form coherent sheets of metasediment or disrupted to form shale-matrix mélange—the classic argille scagliose of the Ligurian Alps (Hsü, 1968; Page, 1978; Raymond, 1984). Formation of shale-matrix mélange may involve either or both tectonic disruption and olistostromes (debris flows), which commonly form broken formations that consist of formerly intercalated wackes and shales, with rare exotic blocks (Abbate et al., 1970; Cowan, 1978, 1982; Aalto, 1981; Raymond, 1984). Broken formations may also form by layer parallel extension of intercalated wackes and shales in response to extensional deformation of the accretionary wedge (Byrne, 1984). In contrast, true mélange requires some component of exotic material, whose provenance is unrelated to the matrix material. These exotic blocks may include igneous or sedimentary rocks of oceanic affinity (e.g., pillow lavas, chert) or metamorphic rocks formed from previously subducted material, mixed by either tectonic or debris flow processes (e.g., Bailey et al., 1964; Hsü, 1972; Cowan, 1978, 1982; Hall, 1980; Cloos, 1984). Regardless of their mode of formation, these complexes represent
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation subduction mélanges formed in response to the continuous and ongoing subduction of material deposited into the trench on an active plate margin. Other mélange formations represent passive margin sediments, seamounts, and fringing coral-reef limestones that formed during or after rifting along a passive continental margin, only to be dismembered during collision and emplacement of an overriding ophiolite (i.e., obduction) onto the passive margin (Yilmaz and Maxwell, 1984; Shervais, 2001; Metcalf and Shervais, 2008). Examples of these obduction mélanges (and their associated thrust complexes) include the Hawasina nappes and Haybi mélange, which underlie the Semail ophiolite of Oman (Robertson, 1986; Bechennec et al., 1988, 1990), and the Mamonia complex of Troodos (Lapierre et al., 2007). The primary distinction between an obduction mélange and a true subduction mélange is that the former is thrust over the passive margin from which it was derived (in the lower plate of a collision zone), whereas the latter consists largely of sediments derived from an active volcanic arc or its underlying plutons (in the upper plate of the subduction zone) and may be unassociated with a collisional event. Igneous elements of the passive margin obduction mélange typically date to the initiation of rifting and are unrelated to the later convergence.
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may also form as a result of volume expansion during the serpentinization process. Serpentinization involves ~50% volume expansion, depending on the primary mineralogy. Formation of block texture during serpentinization results in fractures where serpentinization is concentrated (O’Hanley, 1991, 1992). As the serpentinite expands, it may be forced out along these fractures, and blocks of partially serpentinized massive peridotite may be forced upward between adjacent fractures, forming the classic block-in-matrix texture of a broken formation without imposition of an external stress field; subsequent tectonic shear will exploit these preexisting “shear zones” to cause further deformation (e.g., Shervais et al., 2005a) (Fig. 1). Serpentinite mélange, like argille scagliose, may also form by sedimentary processes as debris flows or olistostromes. Sedimentary serpentinites are common in the Mesozoic of California (e.g., Moiseyev, 1970; Lockwood, 1971; Phipps, 1984) and have even been observed to have formed Quaternary debris flows (Cowan and Mansfield, 1970). In the Marianas forearc,
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Serpentinite Matrix Mélange Serpentinite matrix mélange represents a relatively minor fraction of most accretionary complexes, but their significance far outweighs their relative volume (e.g., Saleeby, 1979, 1984; Carpenter and Walker, 1992; Wright and Wyld, 1994; Malpas et al., 1994; Blake et al., 1995; Lennykh et al., 1995; Tankut et al., 1998; Coleman, 2000; Wakabayashi, 2004; Guillot et al., 2004; Choi et al., 2008a, 2008b). Serpentinite mélange exposes samples of the upper mantle and may be a prime carrier that brings high-grade metamorphic blocks (blueschist, eclogite, amphibolite, garnet amphibolite) back to the surface from depth (e.g., Bailey et al., 1964; Moore, 1984; Ross and Sharp, 1988; Oh and Liou, 1990; Baldwin and Harrison, 1992; Harlow et al., 2004; Beane and Liou, 2005; Tsujimori et al., 2006). The occurrence of chlorite-actinolite rinds on many blueschist and eclogite “knockers” implies transport within a serpentinite matrix, even when these blocks are found within a shale matrix mélange. A serpentinite mélange typically consists of a sheared serpentinite matrix with blocks of unsheared, partially serpentinized peridotite, volcanic rocks, chert, and high-grade metamorphic rocks. The sheared serpentinite matrix commonly comprises lizardite and chrysotile, with minor brucite, magnetite, and carbonate; antigorite is less common but may be the dominant serpentine mineral in some mélange belts. Serpentinite may also form a broken formation, in which blocks of massive, unsheared serpentinite float in a matrix of highly sheared or foliated serpentinite (e.g., Shervais et al., 2005a) (Fig. 1). Serpentinite broken formations may form tectonically in response to an external shear stress; alternatively, they
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Figure 1. Outcrop scale view of serpentinite broken formation, Black Diamond Ridge, Stonyford, California. (A) View to NE of massive lherzolite or harzburgite blocks in matrix of sheared serpentinite; crest of ridge is massive lherzolite. (B) Serpentinite broken formation with phacoidal lherzolite or harzburgite blocks (dark green, reddish brown) in sheared serpentinite matrix (pale blue green).
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serpentinite mud volcanoes are common and feed debris flows within the trench (e.g., Ishii et al., 1992; Fryer et al., 1995, 1999, 2000). These deposits may contain a variety of exotic clasts derived from their parent serpentinite body, such as blueschists, as well as intermingled clay, sand, and fossils. Their introduction into a trench setting may result in the tectonic recycling of this sedimentary mélange, forming many of the features we currently identify with tectonic mélange. Nonetheless, ultimately the source of these sedimentary serpentinites must itself have been mixed prior to their formation, especially where they carry high-pressure metamorphic rocks. A major question about the origin of serpentinite matrix mélange hinges on the provenance of the serpentinite itself: Does it represent oceanic mantle, hydrated on the seafloor long before accretion to the subduction complex, or does it represent the hanging wall of a subduction zone, hydrated during plate convergence? Is the mixed assemblage the result of tectonic processes, or does it form as a serpentinite diapir within forearc sediments (e.g., Ishii et al., 1992; Fryer et al., 2000)? These disparate origins require distinctly different histories, and each implies different assemblages of tectonic inclusions (e.g., Blake and Jayko, 1990; Blake et al., 1995; Wakabayashi, 2004; Choi et al., 2008a, 2008b). In this contribution we summarize evidence for the origin of the Tehama-Colusa serpentinite mélange in northern California, which is juxtaposed tectonically against both the Coast Range ophiolite and the Franciscan complex (Hopson and Pessagno, 2004, 2005; Shervais et al., 2004a, 2005a). This evidence includes mapping and petrologic studies by Shervais and co-workers (Shervais and Kimbrough, 1985, 1987; Shervais and Hanan, 1989; Shervais et al., 2004a, 2004b, 2005a, 2005b, 2005c), mineral chemistry and isotopic studies by Choi et al. (2008a, 2008b), and published reports by other workers (Jayko and Blake, 1986; Blake et al., 1987, 1992; Jayko et al., 1987; Huot and Maury, 2002; Hopson and Pessagno, 2004, 2005; Hopson et al., 2008). Previous studies going back to Saleeby (1983) have proposed a fracture zone origin for this mélange belt (Saleeby, 1983; Shervais and Kimbrough, 1987; Coleman, 2000; Hopson and Pessagno, 2004, 2005). Our data imply a more complex interpretation, in which a proto-Franciscan subduction zone nucleated on a largeoffset transform (oceanic fracture zone); subsequent extension of the upper plate, and magmatism associated with this extension, created the refractory forearc mantle and the overlying ophiolitic crust (Choi et al., 2008a, 2008b). Convergence during and after this process resulted in a variety of included blocks derived from both the upper and lower plates (Jayko and Blake, 1986; Shervais and Kimbrough, 1987; Blake et al., 1987, 1992; Huot and Maury, 2002; Hopson and Pessagno, 2004, 2005; McLaughlin et al., 1990; McLaughlin and Ohlin, 1984). METHODS The study of mélange formation typically involves two components: field observation and investigation of the prov-
enance of the mélange matrix and constituent blocks within the mélange. In this contribution we summarize the results of our field mapping projects in two locations, and the results of other field investigations in these locations. Field studies document the lithologies that occur within the mélange, their distribution, and how they relate to the surrounding rocks. Some lithologies are distinct and require an exotic source, e.g., amphibolite blocks and metasedimentary blocks. Sizes of the blocks range from several kilometers to less than a meter, so many blocks cannot be mapped even at larger scales. However, their location can be noted, along with their composition and distribution. We also review the provenance of knockers within the mélange, with particular focus on the petrology, mineral chemistry, and whole-rock geochemistry of metavolcanic and peridotite blocks within the mélange. For the mafic volcanic rocks, major and trace element concentrations can be used for comparison with volcanic rocks from known tectonic settings, using Harker diagrams and trace element tectonic discrimination diagrams. For the peridotite blocks and even the sheared serpentinite matrix, we can use the major element composition of relict mineral phases. Cr-spinel is especially useful because it is typically unaffected by the serpentinization process, with the exception of forming magnetite or ferrichromite rims. Abyssal peridotites are characterized by spinels with Cr#s [= 100Cr/(Cr+Al)] of ~10 to ~59; whereas supra-subduction zone–forearc peridotites are characterized by spinels with Cr#s ~35 to ~84 (Dick and Bullen, 1984; Ishii et al., 1992; Metcalf and Shervais, 2008). Although these groups overlap at Cr#s ~40–60, there are commonly samples that lie outside the overlap and provide diagnostic compositions (e.g., Choi et al., 2008a). The application of powerful new tools for trace element microanalysis (e.g., laser ablation ICP-MS [inductively coupled plasma–mass spectrometry]) will allow even more detailed studies in the future (e.g., Jean et al., 2010). Chert blocks of different provenance may also be distinguished if their radiolarian faunal assemblages are documented (e.g., Hopson and Pessagno, 2004, 2005). Cherts associated with the Coast Range ophiolite are characterized by a trend upsection from polytaxic Tethyan faunas to oligotaxic Boreal faunas (Pessagno and Blome, 1990; Pessagno et al., 2000; Murchey in Shervais et al., 2005c). Many of these cherts are also alumina rich, reflecting high volcanic ash contents (Hopson et al., 1981), but this is not true for all locales. In contrast, Franciscan abyssal cherts are characterized by higher silica contents, a wide range in ages, and classic ribbon chert intercalated with thin mudstone layers (Karl, 1984; Murchey, 1984). The mélange matrix is more difficult to characterize because of its variability and structural incoherence. However, X-raydiffraction studies can establish the phases present and distinguish among the serpentine phases, and microprobe studies of relict spinel compositions can establish the provenance of their protolith. The sheared matrix may preserve fabric elements indicative of shear sense and relative strain, but more commonly they simply reflect a poorly defined foliation.
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation GEOLOGIC SETTING Coast Range Ophiolite The Coast Range ophiolite of California separates rocks of the Franciscan complex, a classic subduction zone accretionary complex in central and northern California, from forearc basin sedimentary rocks of the Great Valley Group (Bailey et al., 1970; McLaughlin et al., 1988; Shervais et al., 2004b; Hopson et al., 2008) (Fig. 2). Although the contact between the ophiolite and the Great Valley Group is commonly faulted, the original sedimentary contact is preserved locally (Blake et al., 1987). Igneous rocks of the ophiolite have been dated at 172 Ma to 161 Ma, although in the Diablo Range ages range down to 148 Ma (Hopson et al., 1981, 2008; Shervais et al., 2005c; Mattinson and Hopson, 2008). The ophiolite is commonly interpreted as a suprasubduction zone ophiolite that forms the basement of the Great Valley forearc basin (Evarts, 1977; Shervais and Kimbrough, 1985; Lagabrielle et al., 1986), but recent data show that more complex scenarios involving ridge collisions and late mid-oceanridge basalt (MORB) overprints are required (Giaramita et al., 1998; Evarts et al., 1999; Shervais et al., 2004b, 2005b, 2005c; and others). However, other investigators argue for more traditional models of formation at a mid-ocean-ridge spreading center (Hopson et al., 1981, 2008; Ingersoll, 2000; Dickinson, 2008). Serpentinized peridotite that underlies mafic igneous rocks of the Coast Range ophiolite has long been considered to represent the basement of mantle lithosphere upon which the ophiolite was constructed (e.g., Bailey et al., 1970; Hopson et al., 1981; Jayko et al., 1987; Shervais, 1990; Huot and Maury, 2002; Shervais et al., 2005a, 2005b). In contrast, Hopson and Pessagno (2004, 2005) proposed that the Tehama-Colusa serpentinite mélange in northern California represents an oceanic fracture zone unrelated to the ophiolite, and juxtaposed against it during subsequent accretion within the Franciscan subduction zone. This proposal echoes earlier suggestions by Saleeby (1983), Jayko and Blake (1986), and Shervais and Kimbrough (1987), but develops this hypothesis more fully based on their detailed synthesis (Hopson and Pessagno, 2004, 2005; Dickinson, 2008). Franciscan Complex The Franciscan complex of California is the classic example of an accretionary complex formed in response to the subduction of oceanic lithosphere (e.g., Bailey et al., 1964; Hsü, 1968; McLaughlin and Ohlin, 1984; McLaughlin et al., 1982, 1990; Blake et al., 1985, 1987, 1988; Ernst, 1993; Wakabayashi, 1999). The Franciscan complex in northern California comprises three NNW-trending belts that decrease in age and metamorphic grade from east to west: the Eastern Belt, Central Belt, and Coastal Belt (Bailey et al., 1964; McLaughlin et al., 1982). The Central Belt represents the classic shale-matrix mélange that is commonly associated with the Franciscan accretionary complex. It contains knockers of graywacke, greenstone, serpentinite,
5
chert, limestone, blueschist, eclogite, amphibolite, and garnet amphibolite (Bailey et al., 1964; Blake and Jones, 1974; Blake et al., 1988) as well as arkosic wackes interpreted as slope basin deposits (Becker and Cloos, 1985). The Coastal Belt, which lies outboard of the Central Belt, is a Schüppenzone that consists of east-tapered thrust wedges that generally young structurally downward toward the west and include rocks as young as Miocene in age, some of which are obducted over the older part of the Coastal Belt (McLaughlin et al., 1982; Aalto et al., 1995; Blake et al., 1985; McLaughlin et al., 2000). The Eastern Belt consists of coherent blueschist-facies metagraywacke and metabasalt, as well as exotic blocks in metamorphosed mélange with a metagraywacke matrix (Bailey et al., 1964; Blake and Jones, 1974; McLaughlin and Ohlin, 1984; Blake et al., 1988; Ernst, 1993). The metasediments have Early to mid-Cretaceous depositional ages and mid- to Late Cretaceous metamorphic ages (Blake et al., 1982). The easternmost unit in this belt is the Southfork Mountain schist, a penetratively deformed quartz-albite-lawsonite metagraywacke of Early Cretaceous age (120.5 Ma: Wakabayashi and Dumitru, 2007), and its blueschist-facies mafic member, the Chinquapin metabasalt. The Southfork Mountain schist crops out along the western margin of the Tehama mélange in the northern part of the study area, but it is replaced by other units of the Eastern Belt (Valentine Springs Formation, Yolla Bolly terrane) and in places, Central Belt rocks, farther south. TEHAMA-COLUSA SERPENTINITE MÉLANGE The Tehama-Colusa serpentinite mélange extends from Elder Creek in the north to Wilbur Springs in the south (Fig. 2). The northern portion of this belt was mapped by Blake et al. (1992), and it was described in some detail by Huot and Maury (2002) and by Hopson and Pessagno (2004, 2005); the reader is referred to their work for a complete survey of this belt. The southern end of this belt was mapped by McLaughlin et al. (1990). We present here an overview of the mélange between Elder Creek and Stonyford, California, where we have mapped parts of this mélange in detail. Jayko and Blake (1986) discuss the distribution and significance of foliated metasedimentary rocks in the Tehama mélange belt. The Tehama-Colusa mélange includes two main segments connected by a thin selvage of sheared serpentine that separates schists of the Franciscan complex on the west from the Coast Range ophiolite or wackes and mudstones of the Great Valley Group to the east: (1) the Tehama mélange segment, which structurally underlies the Elder Creek massif of the Coast Range ophiolite, and surrounds the Chrome peridotite block, a huge isolated knocker of harzburgite-dunite at the southern end of the Tehama mélange belt; and (2) the Colusa mélange, which underlies and partially surrounds the Stonyford volcanic complex and continues south to the area around Wilbur Springs (Fig. 2). The mélange is separated from adjacent rocks of the Great Valley Group in the south by the Stony Creek fault (Brown, 1964), and from the Elder
6
122° 44' W
40'
35'
Shervais et al. 122° 30' 25.5' W 40° N
Geologic map of Tehama-Colusa mélange Adapted from Jennings and Strand (1960) Hopson and Pessagno (2005) Shervais et al. (2005a, 2005b, 2005c)
Round Mtn
v
Crowfoot Point
Thomes Creek v
Tehama-Colusa mélange Elder Creek ophiolite Franciscan complex
45'
Hz
45'
Chrome
Stonyford volcanic complex
Grindstone Creek Alder Springs
Snow Mtn. volcanic complex Jurassic (?) Great Valley Series Cretaceous Great Valley Series Tertiary sediments
30'
30'
Quaternary alluvium Black Diamond Ridge
Places
ms v Lz
Hz = harzburgite blocks Stonyford
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Lz = lherzolite blocks
Dated amphibolite blocks v = volcanic blocks ms = metasediment blocks
Hz Hyphus Creek
0
v Little Stony Creek Hz
5 km
10 0
5 miles
10
15'
15'
Wilbur Springs
122° 40' W
122° 30' W
39° N
122° 20' W
Figure 2. Geologic map of the Tehama-Colusa mélange. Sheared serpentinite dominates the thinner outcrop belts of the mélange, whereas kilometer-scale and smaller blocks make up much of the outcrop belt where it is wider. Kilometerscale blocks of harzburgite (Hz) or lherzolite (Lz) form much of the mélange around Stonyford and Chrome, and volcanic blocks (v) are common in the north. After Jennings and Strand (1960), Blake et al. (1992), Hopson and Pessagno (2005), and Shervais et al. (2005b). Dashed-line boxes show outline of detailed maps in Figure 3.
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation Creek ophiolite in the north by the Beehive Flat fault (Jayko and Blake, 1986; Jayko et al., 1987). The Stony Creek fault trends N-S and forms a high-angle range front fault along this segment of the Coast Range. The Beehive Flat fault is a younger, NWtrending, low-angle detachment fault that cuts out section from the lower part of the Elder Creek ophiolite (Jayko et al., 1987). The serpentinite mélange belt is separated from the Franciscan complex to the west by the Coast Range fault, a high-angle struc-
ture that truncates an older tectonic contact between these units (Jayko and Blake, 1986) (Fig. 2). Tehama Mélange Exposures of sheared serpentinite north of Grindstone Creek (~39° 30′ N, Fig. 2 and Fig. 3A) constitute the Tehama mélange belt; the northern part of this belt has also been referred 122° 30' W 40° 00' N
122° 45'
Fault
Elder Creek Middle Fork
Tehama-Colusa mélange Massive peridotite blocks
Paskenta Fault Zone
Volcanic ± chert blocks
lt
Round Mountain
N
Elder Creek South Fork
Flat Fau
Coast Range
Elder Creek ophiolite Be eh ive Lz
Foliated metasediments
Digger Creek ssp
Elder Creek ophiolite
Paskenta Crowfoot Point
Franciscan complex Lo
Great Valley Group
Thomes Creek
0 St yC on lt
u Fa
lt
u Fa
he
rC
re
Chrome Harzburgite block
Hz
Creek
e Fa Coa
st R
ang
122° 45'
Chrome
Stony
ult
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ru
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2 3 4 miles 0 1 2 3 4 5 km
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ek
1
re
an
st
tR
hru
as
sT
Co
ng
pri
gS Su
7
Grindstone Creek
39° 40' 122° 30'
Figure 3 (Continued on following page). Geologic maps of the Tehama and Colusa mélange segments. (A) Tehama mélange segment, showing distribution of largest tectonic blocks (after Blake et al., 1992; Huot and Maury, 2002; and Shervais, unpublished). (B) Colusa mélange segment around Stonyford, showing distribution of massive peridotite blocks and the Stonyford volcanic complex (after Brown, 1964; Shervais et al., 2005a, 2005b, 2005c). Hz—harzburgite; Lz— lherzolite; ssp—sheared serpentinite; SFVC—Stonyford volcanic complex; Qal—Quaternary alluvium.
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to as the Round Mountain mélange (Jayko and Blake, 1986; Shervais and Kimbrough, 1987; Blake et al., 1992; Huot and Maury, 2002). The Tehama mélange lies southwest of the Elder Creek massif of the Coast Range ophiolite, separated from the ophiolite and its Crowfoot Point breccia unit by the Beehive Flat fault (Jayko and Blake, 1986; Blake et al., 1992; Huot and Maury, 2002). A high-angle fault along the western boundary of the mélange separates it from the Franciscan complex (South Fork Mountain schist: Blake et al., 1988, 1992); this fault is generally regarded as the Coast Range fault. Mélange blocks here consist largely of basalt and chert, commonly forming composite knockers of chert sitting depositionally on pillow lava. Small blocks of massive peridotite and diabase are less common (Huot and Maury, 2002). A second major exposure of the Tehama mélange occurs near the town of Chrome, ~20 km south of Elder Creek (Fig. 3A).
A large kilometer-scale block of harzburgite (~3 km across and ~8 km long) forms a bulge in the mélange belt, which is only a few tens or hundreds of meters wide on either side of this block. This block of partially serpentinized but massive harzburgitedunite underlies Red Mountain and contains several chromite mines, including the Grey Eagle mine, an open pit excavation ~500 m across. Several smaller pits are found upslope from the main pit. The podiform chromite deposits within dunite are now largely mined out, but remnants are still visible. Colusa Mélange The southern segment of the Tehama-Colusa mélange extends from Grindstone Creek in the north to Wilbur Springs in the south (Fig. 2). In the north it surrounds the Stonyford volcanic complex (Shervais and Hanan, 1989; Shervais et al.,
n Fra
39° 30'
nC
ca
cis
Great Valley Group
x ple
om
Tehama-Colusa mélange Massive peridotite blocks Stonyford volcanic complex
Black Diamond Ridge Volcanic sandstone blocks ssp
Franciscan complex
Lz Qal
Great Valley Group Gravelly Ridge Conglomerate
SFVC
Stonyford ssp
N
Hyphus Creek
Little Stony Creek Franciscan Complex
B 122° 40'
5 km
Basaltic sandstone Great Valley Group
Hz
Hz ssp
Hz 39° 15' N 122° 30' W Figure 3 (Continued).
Quaternary alluvium
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation 2005a, 2005b, 2005c); farther south the mélange is almost entirely peridotite, both massive and sheared, forming a serpentinite broken formation. The mélange near Stonyford appears to largely underlie the volcanic complex itself (as documented by deep exposures and fault offsets), but it also wraps around its eastern margin, at least in part (Fig. 3B). Blocks of cumulate gabbro, isotropic gabbro, quartz diorite, and pyroxenite underlie the volcanic complex and are also exposed to the northeast. Massive, unsheared harzburgite and lherzolite form large kilometer-scale blocks both north and south of the volcanic complex (Shervais et al., 2005a; Choi et al., 2008a, 2008b), whereas additional volcanic blocks and metamorphic blocks lie to the west. North of the Stonyford complex the crest of Black Diamond Ridge is underlain by massive lherzolite, and the lower elevations are underlain by sheared serpentinite with smaller (dekameter-scale to meter-scale blocks) of massive serpentinite (Fig. 1). South of the Stonyford complex, massive unsheared harzburgite blocks are exposed along Hyphus Creek and Little Stony Creek. Franciscan metasediments and metavolcanics— equivalent to the Valentine Springs Formation—crop out along the western boundary of the mélange, and Tithonian mudstones of the Great Valley Group are faulted against the eastern boundary of the mélange (Fig. 3B). Serpentinite mélange mapped in the Wilbur Springs area by McLaughlin et al. (1990)—their Grizzly Creek mélange unit— contains a similar assemblage of Coast Range ophiolite–derived blocks and is likely correlative with the Colusa mélange. South of Wilbur Springs (Fig. 2) sedimentary serpentinites are common, interbedded with Lower Cretaceous Great Valley Group mudstones (Moiseyev, 1970; Phipps, 1984; Carlson, 1981; McLaughlin et al., 1990; Campbell et al., 1993). These detrital serpentinites were derived from protrusions of sheared serpentinite mélange on the seafloor; they contain small amounts of clay and macrofossils that confirm their detrital origin and a Hauterivian age (McLaughlin et al., 1990). Similar deposits are found in the Mariana forearc (Fryer et al., 1995, 2000). SERPENTINITE MATRIX OF THE TEHAMACOLUSA MÉLANGE This mélange matrix consists of phacoidal lithons a few centimeters to tens of centimeters in size, encased by serpentinite microlithons and sheared scaly serpentinite that may be broken into continually smaller scaly fragments (Fig. 4A). The matrix is commonly homogeneous in appearance, but in places it displays distinct variations in color that reflect changes in serpentine mineralogy, possibly containing clasts of one serpentinite in a matrix of another (Figs. 4B, 4C). The scaly matrix defines a subvertical foliation that generally trends ~N-S (subparallel to the trend of the mélange outcrop belt), but in detail it wraps around the margins of mélange blocks, and in places it dips parallel to local fault contacts (Fig. 5). Huot and Maury (2002) also note zones of nonfoliated granular serpentinite that they interpret as late diapirs formed within the mélange.
9
S-C fabrics are observed near both Stonyford and Elder Creek (Dennis and Shervais, 1991; Huot and Maury, 2002). Dennis and Shervais (1991) described serpentinite mylonite in a fault block of mélange within the Franciscan complex near Stonyford (Hyphus Creek area) that has well-developed S-C fabrics, which indicate tops down-to-the-east–dextral shear along a low-angle
A
B
C
Figure 4. Outcrop close-ups of sheared serpentinite matrix, Tehama mélange. (A) Intercalated green lizardite and blue antigorite-bearing sheared serpentinite; S-C shear bands in blue serpentinite indicate dextral shear sense (foliation trends N-S). (B) Contact between green lizardite and blue antigorite-bearing serpentinite. (C) Centimeter-scale clast of green lizardite in sheared blue serpentinite; millimeter-scale clasts are common. All photos are north of Toomes Camp Road, Paskenta.
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detachment surface. Foliated serpentinite exposed east of Black Diamond Ridge (Stonyford) displays a crude S-C fabric that dips to the east and also indicates top down-to-the-east–dextral movement within the mélange. In the northern Tehama mélange, welldeveloped S-C fabrics indicate dextral shear along dominantly N-S–trending shear planes (Fig. 4A). An outcrop in Digger Creek displays both dextral shear S-C-C′ fabric and an asymmetric fold in foliated serpentinite adjacent to a N-S–trending fault contact with an ophiolite dike complex that indicates west-side-up offset (Fig. 6). The dextral shear S-C-C′ fabric is consistent with other indicators of dextral shear in the Colusa mélange. More commonly, sheared and foliated serpentinites do not display consistent shear sense indicators. X-ray-diffraction studies show that the sheared serpentinite matrix consists largely of lizardite, with minor chrysotile, brucite, Cr-spinel, and magnetite. Magnesite veins are reported locally. In parts of the Tehama mélange, the sheared matrix includes zones of indigo blue antigorite-bearing serpentine intercalated with more common green lizardite plus chrysotile. The foliated blue antigorite-bearing serpentinite commonly contains millimeter- to centimeter-scale clasts of pale-yellow-green lizardite-chrysotile (Fig. 4). The presence of antigorite as a relict phase (with dextral S-C fabrics) within the dominantly lizardite-chrysotile matrix implies original formation at temperatures of 300–640 °C, followed by retrogression to temperatures <300 °C (Evans et al., 1976; Evans, 2004). Relict Cr-spinel grains within the Tehama mélange near Round Mountain are largely Cr-rich (Cr#s > 55) consistent with derivation from highly refractory peridotite of supra-subductionzone affinity (Fig. 7A; Huot and Maury, 2002). The enrichment of these highly refractory spinels in TiO2 (Fig. 7B) implies reaction with a refractory magma similar in composition to boninite. The occurrence of high-Cr, high-Ti spinel is characteristic of a highly refractory harzburgite or dunite protolith, formed in a
supra-subduction setting by hydrous melt extraction (e.g., Choi et al., 2008a). LITHOLOGY OF BLOCKS IN TEHAMACOLUSA MÉLANGE In this section we review the variety of tectonic blocks (knockers) found within the mélange, grouped according to lithology. Blocks with inferred upper plate provenance include those derived from the Coast Range ophiolite or its underlying mantle lithosphere. Blocks with inferred lower plate provenance are those with clear oceanic affinities or those that have been subducted and metamorphosed along high-pressure/temperature (P/T) trajectories to blueschist or higher grade. A few blocks of low-grade metavolcanics and foliated metasediments cannot be firmly assigned to either of these groups (e.g., Jayko and Blake, 1986) and are considered separately.
A
East Dike complex
West Serpentinite
B
West C′
ssp C
Hz
Figure 5. Sheared serpentinite matrix (ssp) wraps around harzburgite block (Hz), Tehama-Colusa mélange. Edge of small, meter-scale harzburgite block (west) in sheared serpentinite matrix, which wraps around margin of boulder. Toomes Camp Road, Paskenta.
S
East Figure 6. Sheared serpentinite, Digger Creek, Tehama mélange segment. (A) Outcrop view of sheared serpentinite with dextral S-C-C′ fabric (west) in contact with dike complex (east); foliated serpentinite at contact deformed into asymmetric fold by down-to-east throw on fault; person is sitting on contact, which trends ~N-S. (B) Close-up of sheared serpentinite, with dextral S-C-C′ shear fabric. Scale aligned parallel to contact, which trends ~N-S.
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation Mantle Peridotites
11
the sheared serpentine matrix in preserving their primary mantle textures and mineral compositions, despite partial to extensive serpentinization. The massive blocks near Chrome and Hyphus– Little Stony Creeks are harzburgite, with less common dunite and chromite. Data presented by Shervais et al. (2005a) and Choi et al. (2008a, 2008b) show that the mineral compositions in these unsheared massive peridotites are highly refractory, with Cr-rich spinels (Cr#s > 35: Fig. 7A) and low minor-element contents in pyroxene. Spinels with Cr#s of 35–55 have low TiO2, but high Cr# spinels commonly have high TiO2, consistent with boninite melt reaction (Fig. 7B). Pyroxene compositions are refractory, with low
Mantle peridotites form the largest and most common blocks within the mélange (Fig. 1). These blocks range in shape from rectangular to phacoidal lozenges and in size from <1 m to kilometer scale, with the mélange foliation wrapping around block margins (Fig. 5). The largest blocks are up to several kilometers long and 1.5 km wide and occur primarily in the Colusa mélange segment and near Chrome (Shervais et al., 2005a; Choi et al., 2008a, 2008b), but smaller blocks are common throughout the mélange (e.g., Huot and Maury, 2002) (Fig. 1). They differ from 100
A
90
Abyssal peridotite spinel
80
Mg#
70 60 50 40
Hyphus-Little Stony Creek harzburgite blocks Black Diamond lherzolite block Chrome harzburgite block Forearc peridotite spinel Tehama peridotite blocks Tehama mélange matrix
30 20 10 0 0
10
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Hyphus-Little Stony Creek harzburgite blocks Black Diamond lherzolite block Chrome harzburgite block Tehama peridotite blocks Tehama mélange matrix
MO
TiO2
RB
me
1.0
Forearc spinel
Abyssal spinel
Boninite spinel
0.5
[Boninite melt reaction]
0.0 0
10
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Cr#
60
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Figure 7. Variations in Cr-spinel composition in mélange matrix and in massive peridotite blocks. (A) Spinel quadrilateral (Cr# = 100 × Cr/[Cr+Al] vs. Mg# = 100 × Mg/[Mg+Fe2+]) with fields for spinels from abyssal peridotites (PetDB database, 445 analyses) and forearc peridotites (GEOROC database, 45 analyses). Matrix spinels correlate with forearc peridotite spinels. Massive lherzolite block at Black Diamond Ridge and three peridotite blocks in the northern Tehama mélange correlate with abyssal peridotites, whereas two massive blocks in the northern Tehama mélange and other massive peridotite blocks at Chrome (southern Tehama mélange), Hyphus Creek, and Little Stony Creek (Colusa mélange) all correlate with forearc-supra-subduction-zone peridotite spinels. (B) Cr# in spinel vs. TiO2 in spinel plot. Most spinels have very low TiO2, reflecting moderate to extensive melt extraction. Spinels with intermediate Cr#s and high TiO2 reflect reaction with a MORB-like melt phase; spinels with high Cr#s and high TiO2 reflect reaction with a refractory melt phase similar to boninite. Data from Huot and Maury (2002) and Choi et al. (2008a).
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Al2O3 (0.5–3.3 wt%), TiO2 (<0.07 wt%), and Na2O (<0.24 wt%) in clinopyroxene (Choi et al., 2008a). The clinopyroxenes also have high Cr#s (>18) and Mg#s [= 100Mg/(Mg+Fe)] of 93–94.5. At least two mélange blocks in the northern Tehama mélange studied by Huot and Maury (2002) have similar mineral compositions. These compositions are consistent with pyroxenes and spinels from refractory supra-subduction-zone peridotites, indicating extraction of melt fractions exceeding 20% melting. Detailed studies of the Stonyford harzburgites have shown that their isotopic compositions reflect equilibration with subduction zone–derived fluids (Choi et al., 2008a, 2008b). These fluids have high concentrations of Sr and Pb and elevated 87Sr/86Sr, 207 Pb/204Pb and 208Pb/204Pb isotopic ratios that are higher than MORBs. These characteristics are consistent with arc peridotites (Choi et al., 2008a, 2008b).
A
The massive lherzolite block underlying Black Diamond Ridge (~3 km long and up to 1.5 km across), and three peridotite blocks within the Tehama mélange near Elder Creek, have mineral compositions similar to abyssal peridotites, with relatively Al-rich, low Cr# spinel in lherzolites and Cpx-bearing harzburgites, and intermediate Cr# spinel in Cpx-free harzburgites (Fig. 7A; Huot and Maury, 2002; Choi et al., 2008a, 2008b). Pyroxene compositions are relatively fertile as well, with high Al2O3 (3.7–6.2 wt%), TiO2 (~0.3 wt%) and Na2O (~0.9 wt%) in clinopyroxene (Huot and Maury, 2002; Choi et al., 2008a). The abyssal clinopyroxenes also have low Cr#s (<15) and Mg#s (91–93). Spinels have generally low TiO2 contents, but some spinels from one block with intermediate Cr#s are enriched in TiO2, indicating reaction with a MORB-like melt (Fig. 7B). These compositions are consistent with pyroxenes and spinels from
C
D
B
V ssp
Figure 8. Outcrop photos of volcanic blocks in mélange. (A) Basaltic volcanic block (Stonyford volcanic complex, V) in sheared serpentinite matrix, Gilmore Peak, Stonyford. (B) Meter-scale block of basalt in sheared serpentinite matrix (ssp) near Round Mountain (Toomes Camp Road) within the Tehama mélange segment. (C) Felsic volcanic block within the Tehama mélange segment NE of Round Mountain; sheared matrix is covered with Digger Pine and chamise. (D) Dekameter-scale block of partly sheared diabase, Tehama mélange near Round Mountain (Toomes Camp Road).
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation abyssal peridotites worldwide and show that their parent peridotite underwent relatively low fractions of melt extraction (<15% melting) based on the correlation between Cr# and the degree of melting from the model by Hellebrand et al. (2001). Detailed studies of the Black Diamond Ridge lherzolites by Choi and co-workers have shown that the isotopic compositions of these lherzolites reflect equilibration with subduction zone– derived fluids (Choi et al., 2008a, 2008b). These fluids have elevated 87Sr/86Sr, 207Pb/204Pb and 208Pb/204Pb isotopic compositions above the normal ranges observed in the MORBs toward values observed in the supra-subduction-zone peridotites. These data indicate that despite their original provenance as abyssal peridotites formed under normal oceanic crust, these peridotites resided for some time in the upper plate of a subduction zone. Volcanic Blocks Blocks of volcanic origin are the second most common block type within the Tehama-Colusa mélange (Fig. 8). Volcanic blocks comprise three distinct petrologic groups, based on their petrology, composition, metamorphic assemblages, and locations: (1) the Stonyford volcanic complex, a seamount complex with mixed oceanic-forearc affinities, and a zeolite facies metamorphic overprint in the Colusa mélange segment (Shervais and Hanan, 1989; Shervais et al., 2005a, 2005b); (2) volcanic blocks in the Tehama mélange near Round Mountain with low-grade greenschist-facies metamorphic assemblages (Shervais and Kimbrough, 1985, 1987; Huot and Maury, 2002); and (3) palegreen pumpellyite-facies metavolcanic blocks (Jayko and Blake, 1986; Shervais et al., 2005b). New analyses of rocks from the Stonyford volcanic complex that occur as blocks in mélange, and of pale-green metavolcanic rocks that occur as blocks in mélange, are presented in Table 1. Data for volcanic blocks in the Tehama mélange are from Shervais and Kimbrough (1987) and Huot and Maury (2002). These geochemical data are summarized in Figure 9. Stonyford Volcanic Complex The Stonyford volcanic complex (Shervais and Hanan, 1989; Shervais et al., 2005a, 2005b) comprises several large, kilometerscale blocks partially surrounded by sheared serpentinite mélange (Fig. 8A). The complex comprises pillow lava, sheet flows, and hyaloclastite breccias with minor intercalated chert that formed a seamount built on preexisting oceanic crust. Detailed mapping and petrologic studies by Shervais et al. (2005b) document three volcanic suites: an oceanic tholeiite suite similar to MORB, an alkali basalt suite, and a high-Al, low-Ti volcanic suite with refractory major and trace element compositions similar to arc tholeiites or boninites. These three suites are intercalated at all stratigraphic levels of the complex, showing that the oceanic suites (oceanic tholeiite, alkali basalt) were coeval with the arc-like high-Al, low-Ti suite, which is not found in Franciscan seamount associations (e.g., Marin Headlands, Shervais, 1989, 1990; Snow Mountain complex, MacPherson, 1983).
13
The preservation of unaltered volcanic glass in the hyaloclastite breccias shows that the Stonyford volcanic complex could not have been subducted during its emplacement, unlike the adjacent Snow Mountain complex, which displays incipient blueschist facies metamorphism (MacPherson, 1983). Metamorphic mineral assemblages in the volcanic rocks imply zeolite facies hydrothermal metamorphism (zeolites, smectites, chlorite), and a detailed petrologic study of pyroxene phase chemistry has shown that clinopyroxenes in the Stonyford volcanic complex contain no jadeite component (Shervais et al., 2005b). Secondary mineral assemblages imply that the glass-bearing levels of the complex were probably not heated to >100 °C (Shervais and Hanan, 1989). The occurrence of plutonic rocks derived from the Coast Range ophiolite structurally below the complex suggests that the Stonyford volcanic complex seamount was built on older crust of the Coast Range ophiolite. Radiolarian faunal assemblages in chert show that they formed over a prolonged time period of ~10 m. y. during the Middle Jurassic (Murchey in Shervais et al., 2005c). Finally, volcanic glass from the hyaloclastite breccias has been dated by 39Ar/40Ar at ca. 164 Ma—essentially the same age derived from zircon studies of the underlying ophiolite (Renne in Shervais et al., 2005c). All of this implies formation in the upper plate of the subduction complex as part of the Coast Range ophiolite. Volcanic Blocks, Tehama Mélange Volcanic blocks in the Tehama mélange range in size from <1 m (Fig. 8B) to >500 m, including the large knocker that underlies Round Mountain and gives its name to the Round Mountain mélange (Fig. 3A). The petrology and chemistry of these rocks were studied by Shervais and Kimbrough (1987) and Huot and Maury (2002). These blocks are MORB-like or ocean island basalt (OIB)-like low-Ti to high-Ti basalts that were interpreted by Huot and Maury (2002) as backarc basin basalts, and by Shervais and Kimbrough (1987) as oceanic basalts (Fig. 9). Because these blocks form slabs within the mélange, smaller blocks may be affected by rodingite mineralization, which results in higher than normal CaO, and lower than normal silica and alkalis (Fig. 9). They do not resemble arc basalts but share some affinities with oceanic tholeiites of the Stonyford volcanic complex (Shervais et al., 2005b). Most blocks are characterized by Ti/V ratios of 20–50—that is, within the range of normal midocean-ridge and backarc-basin basalts (Fig. 10). Blocks of felsic volcanics derived from the Coast Range ophiolite are found as rare blocks in the Tehama mélange near the Beehive Flat fault (Fig. 8C). These blocks are correlated with felsic volcanics and felsic dike complex rocks of the Elder Creek ophiolite (Shervais et al., 2004b). Volcanic and diabase blocks are typically undeformed but may also exhibit internal shear deformation (Fig. 8D) and folding (seen most clearly in composite blocks with chert). Pale-Green Metavolcanic Blocks Pale-green, pumpellyite-bearing metavolcanic rocks form blocks in the Colusa mélange south and west of the Stonyford
10.15
0.15
6.99
8.65
4.06
0.37
0.09
98.38
55.1
FeO*
MnO
MgO
CaO
Na2O
K2O
P2O5
SUM
Mg#
31
33 342 27 48
Cr V Sc Ba
57 313 34 86
38
6
71
300 350 48 154
131
11
28
47
140
10
12,170
3.76
1.30
54.2
99.64
0.01
0.36
4.18
6.43
8.09
0.24
12.17
17.20
1.98
49.00
SFVC basalt
Stony Creek
266 322 46 84
113
1
7
45
126
8
10,791
2.91
1.37
54.2
99.53
0.02
0.05
2.96
10.13
7.86
0.22
11.86
15.72
1.85
48.86
SFVC basalt
Stony Creek
82 356 32 16
32
6
41
29
46
7
8933
0.87
1.46
54.9
99.67
0.13
0.15
0.72
18.00
7.47
0.24
10.94
14.88
1.49
45.65
SFVC basalt
Auk-Auk
158 318 30 173
98
11
362
34
184
24
14,928
5.73
1.53
53.9
98.90
0.38
0.38
5.35
11.77
4.78
0.17
7.30
17.31
2.49
48.97
SFVC basalt
Auk-Auk
389 277 29 183
151
6
106
55
199
21
10,630
4.73
1.14
60.9
101.05
0.23
0.22
4.51
7.57
7.68
0.24
8.78
19.98
1.77
50.05
SFVC basalt
Black Diamond Creek
166 407 33 115
52
8
195
55
143
6
12,709
4.15
1.81
49.6
98.44
0.18
0.03
4.12
10.13
6.45
0.23
11.66
13.11
2.12
50.41
Palegreen metavolcanic
Black Diamond Creek
43 313 34 17
28
9
83
32
65
5
7134
5.10
1.59
52.9
98.25
0.10
0.14
4.96
6.23
6.92
0.23
11.00
14.91
1.19
52.57
Palegreen metavolcanic
Black Diamond Glade
296 200 26 45
68
4
76
30
64
5
7973
3.89
1.14
61.1
99.25
0.10
0.06
3.83
10.74
5.86
0.12
6.66
14.73
1.33
55.82
Palegreen metavolcanic
Black Diamond Glade
84 248 33 91
42
7
103
31
63
4
7074
4.73
1.48
54.7
100.31
0.07
0.48
4.25
8.88
5.64
0.06
8.32
14.97
1.18
56.46
Palegreen metavolcanic
Auk-Auk South
284 219 28 114
75
8
379
33
104
7
7973
4.59
1.19
60.0
99.28
0.10
0.39
4.20
9.05
6.53
0.13
7.75
18.44
1.33
51.36
Palegreen metavolcanic
Black Diamond Glade
358 252 32 133
99
16
276
29
119
15
7554
5.50
1.76
50.3
100.14
0.14
0.71
4.79
8.86
5.20
0.16
9.17
16.28
1.26
53.57
Palegreen metavolcanic
Stony Creek
23 183 20
7
6
80
27
45
3
3717
3.63
2.94
37.7
99.64
0.12
0.13
3.50
5.91
2.41
0.08
7.09
17.00
0.62
62.78
Palegreen metavolcanic
Elk Creek Rd.
265 281 51 101
67
1
60
30
85
4
6891
0.17
1.26
58.5
100.46
0.10
0.01
0.16
23.80
6.72
0.14
8.48
14.04
1.15
45.86
Rodingite
Salt Creek
32 323 38 124
9
4
610
34
203
14
12,405
3.82
4.09
30.3
99.74
0.31
0.21
3.61
22.50
2.02
0.10
8.25
12.81
2.07
47.87
Rodingite
Hornet Nest Ridge
312 343 46 799
81
2
161
34
103
4
9009
0.26
1.81
49.7
100.52
0.12
0.01
0.25
21.55
5.51
0.15
9.94
15.06
1.50
46.42
Rodingite
Salt Creek
TABLE 1. MAJOR AND TRACE ELEMENT COMPOSITION OF VOLCANIC AND METAVOLCANIC ROCKS IN SERPENTINITE MÉLANGE SFV-2-1 SFV-2-2 SFV-38-1 SFV-39-1 SFZ-35-1 SFV-64-1 SFV-74-1 SFV-77-1 SFZ-55-1 SFZ64-1 SFZ85-1 SFZ93B-2 SFV-150-1 SFV-209-1 SFV-163-1
Notes: Major elements in wt% oxide; trace elements in parts per million. See Shervais et al. (2005b) for methods. SFVC—Stonyford volcanic complex.
12
28
Rb
Ni
27
167
Y
Sr
4
64
3
63
Nb
Zr
8162
6774
4.10
2.16
45.2
99.04
0.06
0.12
3.98
13.28
4.50
0.22
9.74
14.49
1.36
Ti
4.43
15.51
Al2O3
Alkalis
1.13
1.45
51.31
51.28
SiO2
TiO 2
FeO/MgO
SFVC basalt
SFVC basalt
Rock type:
Auk-Auk South
Auk-Auk South
Location:
SFV102-1
SFV100-1
Sample#
14 Shervais et al.
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation
15
20
Al2O3
18 16 14 12 14
FeO*
12 10 8 6 4 10
MgO
8 6 4 2 0 25
CaO
20 15 10 5
Alkalis
0 8 6 4 2 0 45
50 Stonyford basalt blocks Stonyford basalt rodingite Pale green metavolcanics
55
SiO2
60
65
Tehama basalt-diabase blocks Tehama diabase rodingite
Figure 9. Harker diagrams for metavolcanic knockers in Tehama-Colusa mélange. Five samples have been Ca-metasomatized during serpentinization to form rodingites; others are enriched in silica. The rodingites are depleted in alkalis, consistent with Ca-alkali exchange during metasomatism. Data for Tehama mélange blocks from Huot and Maury (2002) and Shervais and Kimbrough (1987); data for Stonyford basalt and pale-green metavolcanic rocks from Zoglman (1991).
16
Shervais et al.
600 550
Ti/V=10
500 450
Ti/V=20 MORB
Arc
V (ppm)
400 Ti/V=50
350
Figure 10. Ti-V diagram for metavolcanic rocks. Almost all metavolcanic rocks have Ti/V ratios of 20–50, suggesting that the protoliths were largely oceanic basalts. Data as in Figure 9. MORB— mid-ocean-ridge basalt, ppm—parts per million.
300 250
Alkali
200 Stonyford basalt blocks Stonyford basalt rodingite Pale-green metavolcanics Tehama basalt-diabase blocks Tehama diabase rodingite
150 100 50 0 0
5000
10,000
15,000
20,000
Ti (ppm)
volcanic complex, and farther north in the Tehama mélange; in both areas they may be associated with slivers of foliated metasediments. These rocks are distinct from volcanic rocks of the Stonyford volcanic complex because they have fine-grained metamorphic textures with prehnite-pumpellyite facies assemblages, and are typically pale green or gray green (Fig. 11). Relict structures include massive volcanic flows and volcanic breccias (Fig. 11). Net veins of quartz are common in many blocks and may account for the elevated silica content in some analyses. These textures are distinctly different from textures in Stonyford volcanic complex rocks, which retain primary textures and mineralogy. These rocks are basalts and andesites (spilites and keratophyres) that differ significantly in composition from volcanic rocks of the Stonyford volcanic complex but are similar to those found in the Tehama mélange (Fig. 9). Despite their wide range in silica contents (~50–63 wt% SiO2: Fig. 9), all of these volcanic blocks have Ti/V ratios of 20–45, consistent with their origin as oceanic basalts (Fig. 10). One of these blocks was dated by K-Ar at 148 Ma by McDowell et al. (1984)—considerably younger than the Middle Jurassic (166–163 Ma) ages obtained for plutonic blocks derived from the Coast Range ophiolite by Shervais et al. (2005c).
Jayko and Blake (1986) found pale-green metavolcanic rocks intercalated with the foliated metasediments, which they correlate with the Galice Formation (Smith River subterrane of the Klamath Mountains: Blake et al., 1992). Most of these metavolcanics lack any noticeable foliation or penetrative deformation fabrics, suggesting that the metavolcanic rocks are less easily deformed. The metavolcanic blocks are similar to lowgrade metavolcanic rocks that occur with the Franciscan complex near Stonyford. Plutonic Blocks Plutonic blocks are rare in the Tehama mélange, the most prominent being a large (~100 m scale) gabbro block at the southern end of the mélange belt near Grindstone Creek (Jayko and Blake, 1986; Blake et al., 1992). Smaller gabbro bodies found in the mélange may represent dismembered dikes rather than exotic blocks. Plutonic blocks with Coast Range ophiolite affinity are common within the Colusa mélange near the Stonyford volcanic complex, immersed in a matrix of sheared serpentinite that both underlies and overlies the volcanic complex structurally. Blocks that structurally underlie the volcanic complex include
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation cumulate clinopyroxenite and wehrlite, cumulate gabbro, isotropic gabbro, and quartz diorite. Blocks that structurally overlie the volcanic complex include isotropic gabbro and quartz diorite (both foliated and isotropic). These plutonic blocks are petrologically identical to equivalent lithologies in the Elder Creek ophiolite (Shervais et al., 2004b; Shervais, 2008). Zircons from (1) a quartz diorite dike in an isotropic gabbro block beneath the Stonyford volcanic complex and (2) a foliated quartz diorite block in the mélange NE of the Stonyford volcanic complex yield 206Pb*/238U ages (163 ± 2–166 ± 2 Ma) that are essentially the same as 206Pb*/238U ages (165 ± 2–172 ± 2 Ma) obtained for the Elder Creek ophiolite (Kimbrough in Shervais et al., 2005c). The position of most plutonic blocks structurally below the Stonyford volcanic complex suggests that these lithologies formed the original crustal substrate that underlay the volcanic complex prior to its disruption. Chert and Composite Chert–Volcanic Blocks Chert can occur either as discrete blocks or in composite blocks with basalt (Fig. 12). Despite their overall similarity physically, detailed biostratigraphic studies of fossil radiolarians may be used to discern differences in age and depositional history.
17
Chert occurs at two locations: intercalated within the Stonyford volcanic complex in the Colusa mélange segment, and as blocks within the Tehama mélange near Round Mountain. The Stonyford cherts are interbedded with volcanic rocks of the Stonyford volcanic complex (Shervais et al., 2005b, 2005c), whereas cherts within the Tehama mélange form both discrete blocks of red radiolarian ribbon chert and composite blocks of red ribbon chert with basalt (Louvion-Trellu, 1986; Huot and Maury, 2002). In both cases the cherts are dark reddish brown and banded, and they physically resemble ribbon cherts of the Franciscan complex (Louvion-Trellu, 1986; Huot and Maury, 2002; Hopson and Pessagno, 2004, 2005; Shervais et al., 2005b); they are distinct from the pale-tan or buff volcanic-ashrich cherts of the Coast Range ophiolite volcano-pelagic section (Hopson et al., 2008). A detailed biostratigraphic study by Murchey (in Shervais et al., 2005c) showed that the Stonyford cherts contain the same radiolarian assemblages and faunal successions as cherts from
A
A
B
B
basalt
Figure 11. Outcrop photos of pale-green metavolcanic block in mélange. (A) Road cut showing metavolcanic block (to right) in sheared serpentinite matrix (left); note person for scale. (B) Close-up of metavolcanic breccia from outcrop. Both photos at Bennett Creek Road.
chert
Figure 12. Outcrop photos of chert blocks in the Tehama mélange. (A) Folded ribbon chert; Warren Sharp for scale. (B) Dekameter-scale composite block of basalt (left) and chert (right); hammer for scale. Both photos at Toomes Camp Road, Paskenta.
18
Shervais et al.
the Coast Range ophiolite, with ages ranging from late Bajocian or early Bathonian through early Kimmeridgian, a span of some 10 m.y. A definitive characteristic of these (and all) Coast Range ophiolite cherts is the change going upsection from relatively small, polytaxic radiolarian assemblages with Tethyan affinities to more robust, oligotaxic assemblages with Boreal affinities, thought to represent either or both northward drift of the ophiolite or changes in global sea circulation (Pessagno and Blome, 1990; Pessagno et al., 2000; Murchey in Shervais et al., 2005c). Chert blocks in the Tehama mélange near Elder Creek have been assigned ages of Callovian to early Oxfordian, based primarily on European-based biostratigraphic calibrations by Louvion-Trellu (1986) and on determinations by Pessagno (in Hopson and Pessagno, 2004, 2005). Species lists published by Hopson et al. (1981) correspond to late Bajocian to Bathonian using the biostratigraphic zonation of Baumgartner et al. (1995). Thus the Tehama mélange chert blocks overlap the Stonyford cherts in age and may represent oceanic crust associated with the Coast Range ophiolite. Alternatively, the Tehama mélange cherts may also represent lower plate crust accreted during convergence in the subduction zone. The latter interpretation is supported by the observation that these cherts are commonly tightly folded, and their underlying basalts may also be sheared. This style of deformation is not observed within the Stonyford volcanic complex cherts, which crop out as homoclinal sequences within undeformed pillow lava. Volcanic Sandstone Blocks Mélange blocks of volcanogenic sandstone up to 500 m across occur within the Colusa mélange south of Stonyford. These blocks were studied in detail by Seymore (1999), who showed that they were derived from juvenile basaltic andesite volcanics similar to clasts in the Crowfoot Point breccia near Elder Creek. These sandstones are fine to medium grained, moderately well sorted, thin-bedded sandstone with volcanic (tuffaceous) rock fragments, minor mafic plutonic fragments, and chert grains. Pyroxene is the most common mafic mineral clast. When plotted on the pyroxene discrimination diagrams of Leterrier et al. (1982), the pyroxenes indicate derivation from tholeiitic arc basalts (Seymore, 1999). Whole-rock geochemical analyses of the volcanic sandstones resemble arc-derived tholeiitic basalts and basaltic andesites (Seymore, 1999). The volcanic sandstones are distinguished from the foliated metasediments described below (“Galice” affinity) by their lack of any foliation or deformational fabric, the lack of mudstone intercalations that are common in the Galice slivers, and by their lack of detrital chert. They superficially resemble the “basaltic sandstones” of the Jurassic lower Great Valley Group (Brown, 1964) but are distinguished by the chemistry of their pyroxene clasts and by their whole rock geochemistry, both of which indicate a calc-alkaline affinity for the lower Great Valley Group sandstones that is distinct from the tholeiitic affinity of the sandstone mélange blocks.
Foliated Metasediments Jayko and Blake (1986) described large (hundreds of meters scale) blocks of mildly schistose clastic metasediments within the Tehama mélange, and along the contact between the Tehama mélange and the Franciscan complex. Jayko and Blake correlate these blocks with the Galice Formation of the Smith River subterrane, Klamath Mountains, based on their clastic composition (sandstones rich in chert clasts) and mild penetrative deformation (Nevadan orogeny). The occurrence of metamorphosed sediments with a penetrative fabric is significant, because neither the Coast Range ophiolite nor the Great Valley Group was affected by Nevadan deformation. Large blocks of Galice-like lithologies occur at 10 localities in the Tehama mélange segment (Jayko and Blake, 1986). Metagraywacke is the most common lithology, with or without intercalations of slaty argillite, gritstone, and pale-green metavolcanic rocks, in monolithic or composite blocks. Metavolcanic rock is typically intercalated with the metasediments but may also be found alone. The metasedimentary rocks are dominated by black and whitish-gray chert clasts, with volcanic fragments and <10% quartz or feldspar. The mixture of black and white chert clasts gives the metasedimentary rock a “salt and pepper” appearance that is distinctive and unlike metagraywackes in the Franciscan complex. Metamorphic assemblages are chlorite-quartz-albite in the metasediments and chlorite-albite-prehnite-pumpellyite in the metavolcanics (Jayko and Blake, 1986). Similar rocks are found as composite or monolithic blocks in the Colusa mélange segment near Stonyford. Metasedimentary blocks near Stonyford include (1) slightly schistose pebble conglomerate or gritstone with clasts of quartz, chert, shale, and possibly quartzite; (2) medium-grained to silty wacke with minor schistosity, grading into semi-schist; (3) pale-green to dark-green massive metavolcanics with no schistosity; and (4) black slate with pencil cleavage (Zoglman, 1991; Shervais et al., 2005a, 2005b). The first three lithologies occur both as monolithic blocks and as composite blocks intercalated with black slate. All of these blocks are lower in grade and less intensely foliated than metasediments of the adjacent Franciscan complex. The pale-green metavolcanic rocks intercalated with the foliated metasediments lack the penetrative schistosity seen in the metasediments, even though they were metamorphosed and deformed under the same conditions; this implies a significant strength contrast between the two rock types. These metavolcanic intercalations may correlate with the isolated blocks of palegreen metavolcanics described earlier. Metamorphic Blocks: Blueschist, Amphibolite, and Garnet Amphibolite Blocks of blueschist and the “high-grade” metamorphic rocks amphibolite, garnet amphibolite, and hornblende gneiss occur sporadically in the serpentinite mélange. Metamorphic blocks are found around the Stonyford volcanic complex and
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation scattered along the Tehama mélange segment farther north. Blueschist blocks up to 10 m × 3 m are rare (Fig. 13A) and are typically overprinted by chlorite-actinolite rinds. In one locality near Stonyford they form an array of blocks that trend parallel to foliation in the sheared serpentinite matrix. Amphibolite blocks are more common; they typically occur as blocks of massive or layered and foliated amphibolite (Fig. 13B) with retrograde sheaths of chlorite or chlorite-actinolite schist. These blocks are relatively coarse grained and strongly deformed, with distinct
A
B
C
19
mafic- and felsic-rich banding. They are unlike amphibolite facies metagabbro blocks that may form in the deeper oceanic crust. A large amphibolite knocker is also found along Elder Creek in a serpentinite diapir that pierces mudstones of the Great Valley Group. Jayko and Blake (1986) propose that the serpentinite diapir is derived from serpentinite mélange that underlies the ophiolite. This amphibolite contains melt lenses of tonalite (Fig. 13C) that appear to have formed by in situ dehydration melting (Shervais, 2008). Two blocks consisting of garnet amphibolite and hornblende gneiss within foliated serpentinite of the Tehama-Colusa mélange, and one block that crops out in sedimentary serpentinite within the Great Valley Sequence (see Fig. 2) (Blake et al., 1992; Carlson, 1981; McLaughlin et al., 1990; Campbell et al., 1993), have been dated via the 40Ar-39Ar incremental heating technique on hornblende. Locations, field settings, and petrography of the dated blocks are given in Appendix 1, and analytical data are given in Table 2. The methods used were similar to those of Ross and Sharp (1988), and age spectra are shown in Figure 14. Foliated garnet-epidote amphibolite west of Paskenta that is partially replaced by actinolite, albite, and white mica yields a hornblende plateau age of 165.6 ± 2.5 Ma (Fig. 14A). Hornblende gneiss that crops out west of Chrome yields a hornblende plateau age of 167.5 ± 3.3 Ma (Fig. 14B). Hornblende gneiss that crops out near Wilbur Springs in Hauterivian sedimentary serpentinite within Great Valley Sequence clastic strata yields a plateau age of 166.3 ± 3.6 Ma (Fig. 14C). Estimated temperature and pressure conditions for similar amphibolites elsewhere in the Franciscan complex range from 550 to 700 °C and 1.0–1.4 GPa, respectively (e.g., Wakabayashi, 1990), and such amphibolite blocks are interpreted to represent remnants of a high-temperature metamorphic sole formed by thrusting of the basal ophiolite peridotites over oceanic crust during the initial stages of subduction (e.g., Wakabayashi, 1990; Blake et al., 1992) or later during a spreadingridge collision event (Shervais et al., 2004b). The occurrence of tonalitic melt lenses within the amphibolite (Fig. 13C) document dehydration melting temperatures >750–850 °C (Beard and Lofgren, 1991). Likely closure temperatures of metamorphic hornblende to Ar loss are ~450 °C; accordingly, the 40Ar39 Ar ages are interpreted to reflect cooling after peak metamorphic conditions. We note that the new ages for hornblendite and amphibolite blocks reported herein overlap with U-Pb zircon ages obtained for the Coast Range ophiolite at Elder Creek (206Pb*/238U, ~165– 172 ± 2 Ma, Kimbrough in Shervais et al., 2005c; 168 ± 0.3 Ma,
Figure 13. Outcrop photos of metamorphic blocks in mélange. (A) Highly deformed blueschist block; Colusa mélange near Stonyford. (B) Banded mafic-rich amphibolite; South Fork of Elder Creek. (C) Tonalite melt lenses indicate high metamorphic temperatures; South Fork of Elder Creek.
20
Shervais et al. 40
39
TABLE 2. Ar- Ar INCREMENTAL HEATING DATA FOR HORNBLENDES FROM GARNET AMPHIBOLITE AND HORNBLENDE GNEISS BLOCKS Sample Garnet amphibolite from Thomes Creek
Temp. (°C)
39
40
Arcum
% Ar*
K/Ca
Age (Ma)*
±2σ (Ma)
800
0.173
77.4
0.662
166.1
2.7
925
0.308
93.0
0.216
158.9
3.8
975
0.365
83.3
0.071
158.5
5.1
1025
0.624
93.6
0.074
166.8
3.5
1050
0.800
94.5
0.074
165.3
3.5
1100
0.887
93.9
0.071
164.7
4.9
1200
0.974
88.7
0.071
165.4
4.0
1350
1.000
59.3
0.070
165.2 39
8.6 †
Plateau age, 165.6 ± 2.5 Ma, includes 63.5% of the Ar; MSWD = 0.17
Hornblende gneiss from Alder Springs Road
800
0.039
27.7
0.492
169.4
950
0.338
91.9
0.674
149.2
8.5 2.4
1000
0.390
87.5
0.086
153.8
3.2
1050
0.473
89.3
0.030
164.6
7.4
1075
0.639
91.5
0.026
167.3
5.1
1110
0.891
88.8
0.027
167.0
5.5
1150
0.913
88.4
0.024
168.1
12.3
1225
0.941
83.0
0.023
173.9
11.6
1360
1.000
82.5
0.024
168.6
6.4
39
Plateau age, 167.5 ± 3.3 Ma, includes 61% of the Ar; MSWD = 0.41
Hornblende gneiss from Thompson Creek near Wilbur Springs
800
0.030
19.2
0.394
130.4
925
0.097
76.3
0.638
135.7
7.0
1000
0.149
85.8
0.064
160.7
4.4
1050
0.381
92.8
0.028
165.9
5.9
1075
0.650
95.8
0.028
165.7
6.8
1125
0.821
94.5
0.028
167.5
6.3
1175
0.878
878.1
0.027
161.5
11.3
1.000
83.1
0.028
168.3
8.0
1350
16.7
39
Plateau age, 166.3 ± 3.6 Ma, includes 72.4% of the Ar; MSWD = 0.29 Ages were determined using methods and values of interfering neutron reactions given in Ross and Sharp (1988). *Ages were calculated using an age of 523.1 Ma for irradiation standard MMhb (Renne et al., 1998). † Mean square of weighted deviates.
Mattinson in Hopson et al., 2008) and are within the age range of the quartz diorite blocks near Stonyford (206Pb*/238U, ~163– 166 ± 2 Ma, Kimbrough in Shervais et al., 2005c). DISCUSSION The Tehama-Colusa mélange of Hopson and Pessagno (2004, 2005) forms an important tectonic component of the Cordilleran margin. It lies between the Franciscan accretionary complex, which represents the lower plate of the subduction zone (Bailey et al., 1964), and the Coast Range ophiolite, which represents the Middle Jurassic basement of the Great
Valley forearc basin (Bailey et al., 1970). Several authors have proposed that this mélange represents an oceanic fracture-zone assemblage formed on the seafloor far from its current location on the continental margin (Saleeby, 1983, 1984, 1992; Shervais, 1986; Hopson and Pessagno, 2004, 2005). Coleman (2000) proposed a similar origin for sheared serpentinite belts within the Franciscan complex. In contrast, other authors proposed a mixed origin, with crustal assemblages derived from the subducting plate, and mantle assemblages derived from the upper plate (Jayko et al., 1987; Shervais and Kimbrough, 1987; Blake and Jayko, 1990; Blake et al., 1995; Huot and Maury, 2002; Choi et al., 2008a, 2008b).
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation
21
Formation of the Mélange Matrix 180
A
165.6 ± 2.5 Ma
Age (Ma)
170
160
150
Thomes Creek 140 0.0
0.2
0.4
0.6
0.8
1.0
B
Age (Ma)
190
167.5 ± 3.3 Ma
170
150
Alder Springs 130 0.0
0.2
0.4
C
190
0.6
0.8
1.0
166.3 ± 3.6 Ma
Age (Ma)
170
150
130
110
Wilbur Springs 90 0.0
0.2
0.4
0.6
0.8
1.0
Cumulative 39Ar fraction Figure 14. 40Ar-39Ar age spectra for hornblendes from amphibolite and hornblende gneiss blocks. Box heights are 2σ errors; boxes with heavy lines define the plateau age. (A) Garnet amphibolite from Thomes Creek. (B) Hornblende gneiss from Alder Springs Road. (C) Hornblende gneiss from Thompson Creek near Wilbur Springs.
The sheared serpentinite matrix of the mélange was derived from peridotites in the upper plate of a subduction zone. The mélange matrix is characterized by high-Cr spinels that overlap spinel compositions in supra-subduction zone peridotites. The occurrence of the high-temperature serpentine polymorph antigorite also suggests an upper plate origin, with fluids from the subducting slab penetrating upward into the hot base of the upper plate during the early stages of subduction. The antigorite-bearing serpentinite may also represent upper plate lithosphere that was detached from the upper plate and coupled with the subducting lower plate, causing it to be partly subducted and metamorphosed to higher grades than the more common lizardite matrix. Mélange matrix is produced by the process of serpentinization, which is focused initially along joints and fractures that allow penetration of water. The 50% volume expansion caused by serpentinization will also be focused on these fractures and joints, forcing the extrusion of peridotite blocks from between fractures and shearing the surrounding serpentine as the blocks extrude. Once this process has begun, the enhanced access of fluids along these pathways will form a positive feedback loop, with increased serpentinization along the margins of the blocks and continued volume expansion of the intervening serpentine, forcing further extrusion and shearing. We note that the process described here will result in the upward protrusion of serpentinite, even in the absence of a density differential, because the process is driven by volume expansion, not density, and because upward expansion involves the least resistance from the surrounding rocks. The process described here initially forms only serpentinite broken formation. Subsequent introduction of crustal blocks from the overlying ophiolite (as the underlying basement undergoes progressive disruption and expansion) and from the upper levels of subducted oceanic crust (as it scrapes against the overlying upper plate peridotite) will gradually transform the broken formation into a true mélange. The introduction of high-pressure metamorphic rocks requires that the serpentinite matrix flow upward within the subduction channel, carrying blocks that metamorphosed at some depth, much like the channel flow model of Cloos (1982, 1984) for shale matrix mélange. We propose that this process, in which volume expansion drives uplift of serpentinite diapirs with a foliated matrix, may also provide the impetus for serpentinite mud volcanoes (e.g., Fryer and Fryer, 1987). Phipps (1984) argued that the density differential between serpentine (~2.52 for completely serpentinized peridotite) and the overlying sediments of the Great Valley Group (~2.55–2.59 for lithified Great Valley Group sediments) is too small to drive the protrusion of serpentinite diapirs, especially given that unlithified sediments are expected to have even lower bulk densities. We suggest that expansion-driven uplift is sufficient to begin this process, and that saturation with fluids at lower pressures will lead to further decreases in bulk density as the serpentinite-water mixture is fluidized within the mud volcano conduit. Volume
22
Shervais et al.
expansion may also drive the intrusion of serpentine into fault zones, where primary expansion-driven shear foliations would be overprinted by subsequent fault zone shear. Provenance of the Mélange Blocks One of the central questions in mélange formation is the provenance of the blocks found within it. Do these blocks represent a single coherent terrane dismembered by large-scale faulting? Or do they derive from two or more distinct rock assemblages that were originally separate? More specifically, what does the provenance of the blocks tell us about where and how the mélange formed? Potential source terranes can be divided into two broad settings: (1) blocks derived from the lower (subducting) plate (including sediments derived from the upper plate but processed through the subduction zone), and (2) blocks derived from the upper plate of the subduction zone. Blocks with lower plate provenance include those formed at a mid-ocean-ridge spreading center or intraplate seamount, sediments deposited upon this oceanic crust, the mantle lithosphere that underlies this oceanic crust, and rocks metamorphosed along high P-T gradients within the subduction zone. Because similar rocks may be found in upper plate settings, the blocks derived from oceanic lithosphere are restricted here to volcanic rocks with MORB or OIB compositions that cannot be associated with upper plate stratigraphy, abyssal peridotites that lack evidence for residence in a forearc setting, and high P-T metamorphic rocks. Blocks with upper plate affinities include any rocks derived from supra-subduction zone crust or lithosphere: supra-subduction zone (forearc) ophiolites, their underlying mantle lithosphere, and sediments deposited upon the forearc crust or within the forearc basin. These distinctions are clear for lithologies that are only associated with the ophiolite (e.g., cumulate plutonic rocks, quartz diorite) but may be less clear for lithologies that could derive from either oceanic crust or the supra-subduction ophiolite (e.g., volcanic rocks, peridotites, sediments). For each of the lithologies listed here, we discuss our basis for their assignment as being upper plate derived, and note other possible origins. A number of blocks cannot be placed into either an upper plate or lower plate provenance with confidence; in fact, these blocks may not be part of the mélange technically but rather may represent intercalated thrust slices (Jayko and Blake, 1986). Blocks with Upper Plate Provenance Blocks of upper plate provenance include (1) volcanic and plutonic rocks derived from the Coast Range ophiolite, which forms basement to the forearc basin Great Valley Group and has been shown to have formed in a supra-subduction zone setting (Shervais et al., 2004a, 2004b, 2005b, 2005c; Shervais, 2008); (2) highly refractory forearc peridotites, which represent the mantle lithosphere that underlies the Coast Range ophiolite (Choi et al., 2008a, 2008b; Jean et al., 2010); and (3) sedimentary rocks deposited on the ophiolite or derived from a juvenile volcanic arc.
Rock assemblages associated with the Coast Range ophiolite in northern California are exposed in the Elder Creek ophiolite remnant that borders the northern margin of the Tehama-Colusa mélange; these include cumulate dunite, pyroxenite, and gabbro, isotropic gabbro, quartz diorite, and volcanic rocks with arc geochemical signatures preserved in the Crowfoot Point breccia (Robertson, 1990; Beaman, 1991; Shervais et al., 2004a, 2004b). The tie between these assemblages as mélange blocks and intact Coast Range ophiolite at Elder Creek is strengthened by zircon U-Pb ages from quartz diorites that are identical within analytical uncertainty in both groups (Shervais et al., 2005b). The quartz diorites are especially diagnostic, because similar rocks are rare in oceanic domains but are common within supra-subduction zone ophiolites and are particularly abundant in the Elder Creek ophiolite (Shervais, 2008). The Stonyford volcanic complex is similar in many ways to Franciscan seamount complexes (lower plate provenance) but differs in some critical aspects that link it firmly to the Coast Range ophiolite. Like Franciscan seamounts, the Stonyford volcanic complex comprises a low-Ti MORB-like tholeiitic basalt suite and a high-Ti alkaline basalt suite (e.g., Marin Headlands: Shervais, 1989). It is also intercalated with dark-red radiolarian ribbon cherts similar to those seen in the Marin Headlands terrane. Unlike Franciscan seamounts, however, oceanic basalt suites in the Stonyford volcanic complex are intercalated with a very low-Ti, high-Al basalt suite that has refractory major and trace element compositions. These rocks could only form in a supra-subduction environment by extensive second stage melting of previously depleted MORBsource mantle (Shervais et al., 2005b). In addition, the ribbon cherts contain radiolarian assemblages with the same biostratigraphic assemblages seen in Coast Range ophiolite cherts, and which span identical age ranges (Bajocian through Kimmeridgian: Murchey in Shervais et al., 2005c). In particular, the change upsection from Tethyan to Boreal fauna is common to all Coast Range ophiolite chert exposures and marks a shared history common to all localities. Forearc lithosphere is represented by refractory harzburgites and dunites with high-Cr spinels and low-Al pyroxenes. These include the massive harzburgite blocks south of Stonyford (Hyphus–Little Stony Creek blocks), the Chrome block, and smaller blocks in the northern Tehama mélange segment (Huot and Maury, 2002; Choi et al., 2008a, 2008b). Lherzolite and Cpxrich harzburgite blocks with mineral compositions that are the same as abyssal peridotite (e.g., high-Al spinel and pyroxene) underlie Black Diamond Ridge and form smaller blocks in the Tehama mélange. It would seem that these blocks require a lower plate provenance, but the lherzolites of Black Diamond Ridge have high Sr, Rb, and Pb concentrations, and elevated 87Sr/86Sr isotopic compositions, which indicate that despite their affinities to abyssal peridotites these lherzolites must have resided within the upper plate of a subduction zone subsequent to their formation and equilibrated with a slab-derived enriched fluid phase (Choi et al., 2008a, 2008b).
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation Sedimentary rocks deposited on the Coast Range ophiolite or derived from proximal volcanic arc terranes include cherts intercalated with the Stonyford volcanic complex and the volcanogenic sandstone blocks of the Colusa mélange. The cherts have already been discussed in connection with the Stonyford volcanic complex. The volcanogenic sandstones are inferred to have an upper plate provenance, because they represent proximal detritus from a juvenile intra-oceanic arc terrane. However, their source has not been identified, and their introduction into the mélange is enigmatic. These sandstones may correlate with the Crowfoot Point breccia of the Elder Creek ophiolite (Robertson, 1990), but their whole rock and pyroxene geochemistry imply a tholeiitic arc source, in contrast to the calc-alkaline affinity of clasts in the Crowfoot Point breccia. A similar problem arises in correlating these sandstones with the “basaltic sandstone” zone in the lowermost Great Valley Group, which has also been shown to have a calc-alkaline source and is geochemically distinct from the volcanogenic sandstone blocks (Seymore, 1999). Blocks with Lower Plate Provenance Blocks derived from the subducting lower plate include basalts with MORB-like or backarc-like compositions, cherts associated with these basalts, fertile peridotites with abyssal compositions, and rocks metamorphosed within the subduction zone. Blocks of basaltic greenstone in the northern Tehama mélange studied by Shervais and Kimbrough (1987) and Huot and Maury (2002) have major element and trace element compositions consistent with their origin as MORBs (both normal and enriched) and backarc basin basalts. Huot and Maury (2002) note minor negative Nb anomalies and slight Th enrichments compared with MORB, which they suggest implies formation within a backarc basin spreading system rather than a true mid-oceanridge spreading center. Nonetheless, they place these basalts on the subducting lower plate prior to their incorporation into the mélange (Huot and Maury, 2002). Ribbon cherts deposited on these blocks have radiolarian assemblages consistent with Middle to Late Jurassic ages (Callovian to Oxfordian) that fall within the age range of cherts intercalated with the Stonyford volcanic complex but do not appear to document the transition from Tethyan to Boreal faunas characteristic of Coast Range ophiolite cherts (Louvion-Trellu, 1986; Hopson and Pessagno, 2004). Their ages and silica-rich compositions are consistent with deposition on either a mid-ocean ridge or backarc basin basement, as inferred from their underlying basaltic substrate. It is worth noting that these cherts are commonly folded complexly, something that is not observed within cherts of the Stonyford volcanic complex but is common in many Franciscan cherts (e.g., Marin Headlands: Karl, 1984; Murchey, 1984; Murchey and Jones, 1984). Fertile peridotites with whole-rock and mineral compositions similar to abyssal peridotites are found within the Colusa mélange at Black Diamond Ridge and within the Tehama
23
mélange near Elder Creek. In both instances, refractory forearc peridotites are found nearby. It is possible that these peridotites were incorporated into the mélange from the lower plate during active subduction. However, it is not uncommon to find abyssaltype peridotites in forearc settings, associated with more common refractory peridotites (e.g., Parkinson and Pearce, 1998), and there is reason to believe that the fertile peridotites in the Tehama-Colusa mélange were incorporated into the lithosphere prior to mélange formation. Choi et al. (2008a, 2008b) found that all peridotites in their study (including fertile and refractory peridotites) had elevated Pb concentrations in pyroxene, and all but two samples were characterized by elevated 87Sr/86Sr isotopic compositions. As these characteristics are found in primary high-temperature pyroxenes, they must have predated serpentinization and mélange formation. Thus, although the fertile peridotites may have originally formed in an oceanic setting, they were incorporated into the upper plate prior to development of the mélange and were not scraped off a descending lower plate into an already active mélange terrane. Metamorphic rocks that formed in response to subduction zone metamorphism are clear indicators of lower plate provenance, even if the protoliths were torn from the upper plate prior to subduction. The rare occurrence of blueschist in the mélange clearly requires subduction zone metamorphism of the block prior to incorporation into the mélange. Amphibolites are only slightly more common than blueschist blocks. These are strongly deformed (folded), and some contain melt lenses of tonalite or trondhjemite composition—characteristics that are not consistent with simple thermal metamorphism of lower oceanic crust. These rocks must have been metamorphosed against the base of the hot lithosphere during subduction initiation (Wakabayashi, 1999) or during a later ridge collision event (Shervais et al., 2004a, 2004b). In either case a lower plate provenance is indicated. 40Ar-39Ar cooling ages for hornblendes reported here (i.e., 165.6 ± 2.5 Ma and 167.5 ± 3.3 Ma) from two such blocks of the Tehama-Colusa mélange overlap with U-Pb zircon ages of 163– 172 Ma for upper plate igneous rocks of the Coast Range ophiolite and near Stonyford (Kimbrough in Shervais et al., 2005c, and Mattinson in Hopson et al., 2008), emphasizing that large-scale tectonic transport and removal of intervening lithosphere must have taken place in order to closely juxtapose lower and upper plate rocks formed at essentially the same time but at very different levels in the lithosphere. Hornblende gneiss that occurs as clasts in sedimentary serpentinite near Wilbur Springs is similar in rock type and age (40Ar-39Ar hornblende plateau age of 166.3 ± 3.6 Ma) to highgrade blocks of lower plate origin in the Tehama-Colusa mélange and high-grade metamorphic rocks that occur elsewhere at high structural levels of the Franciscan complex (e.g., Wakabayashi and Dumitru, 2007, and references therein). The Wilbur Springs gneiss therefore provides a stratigraphic link between the Franciscan complex and the Great Valley Sequence in Hauterivian time, as noted by Carlson (1981, 1984; see also McLaughlin et al., 1990; Campbell et al., 1993).
24
Shervais et al.
Blocks of Uncertain Provenance The mildly foliated metasediments and intercalated palegreen metavolcanic rocks correlated by Jayko and Blake (1986) with the Galice Formation are difficult to characterize in terms of provenance. If the foliation represents deformation associated with the Nevadan orogeny, they must have an upper plate provenance. This conclusion is supported by their chert-rich detrital modes, which echo the modal composition of Galice and Mariposa Formation sediments. However, their incorporation into the mélange may postdate mélange formation proper and may be related to Cretaceous or younger thrusting (Jayko and Blake, 1986). If so, these rocks would represent not mélange blocks but imbricate thrust slices tectonically interleaved with the mélange. A similar uncertainty applies to the blocks of pale-green metavolcanic rock as well. These rocks resemble similar metavolcanic rocks intercalated with the metasediments, but they occur in discrete monolithic blocks. Their compositions are consistent with altered oceanic basalts that have been enriched in silica. In the Colusa mélange these blocks are more common than the metasedimentary blocks and are commonly found in different parts of the mélange. They may represent lower-plate oceanic crust if distinct from the metasediments, or upper-plate rocks if they correlate with the metasediments.
with a fracture zone origin, such as the extremely deformed fabric of the serpentinite matrix, which penetrates along fractures in the more brittle knockers (e.g., Hopson and Pessagno, 2004, 2005). This dilemma may be resolved by the recent work of Choi et al. (2008b), who show that Black Diamond Ridge peridotites with Al-rich spinels have low equilibration temperatures— consistent with large offset transform peridotites but in contrast to the higher equilibration temperatures typical of abyssal peridotites of small offset or non-transform settings (Fig. 15). Choi et al. (2008b) propose that the Colusa mélange near Stonyford represents a preexisting large offset fault zone that ruptured to form the proto-Franciscan subduction zone in the Middle Jurassic. Following this, mixing of upper plate peridotite (serpentine) with lower plate crustal rocks occurred during the initial stages of subduction. Peridotites with compositions similar to abyssal peridotite represent the preservation of fracture zone peridotite in the upper plate of the subduction zone, whose leading edge was originally part of the fracture zone (Choi et al., 2008b). This interpretation is supported by 87Sr/86Sr compositions of clinopyroxene grains of the peridotites, which show that even the
Spinel peridotites Chrome Black Diamond Ridge (Stonyford) Hyphus-Little Stony Creeks (Stonyford)
Fracture Zone Model for Subduction Initiation 90
Large offset transform peridotites OFZ = Owen Fracture Zone (~300 km offset) VFZ = Vema Fracture Zone (~310 km offset) RFZ = Romanche Fracture Zone (~950 km offset)
80 70
100Cr/(Cr+Al) spinel
Hopson and Pessagno (2004, 2005) proposed that the Tehama-Colusa mélange formed on the seafloor in a fracture zone–transform fault setting, building on previous suggestions by Saleeby (1983), Shervais and Kimbrough (1987), and others. Pervasive faulting along the active transform segment of the fracture zone allowed the penetration of seawater to fuel serpentinization, and dismembered the oceanic crust and lithosphere to form knockers of basalt, chert, and peridotite. The rarity of gabbro is consistent with observations of modern fracture zones, which typically have a thin basaltic crust resting directly on peridotite, with little or no cumulate lower crust. In this model, amphibolites represent thermal metamorphism of lower oceanic crust, whereas the other lithologies represent upper oceanic crust (basalt, diabase), sediments deposited on oceanic crust (chert), and fertile mantle lithosphere (abyssal peridotites). The mélange matrix forms by the pervasive shear of serpentized peridotite in the active transform. A strictly oceanic origin for the Tehama-Colusa mélange requires that all of the assembled lithologies formed on the seafloor and that rocks of upper plate affinity cannot make up an integral part of the mélange. That is clearly not the case. In addition to the mixed provenance of lithologies found as knockers within the Tehama-Colusa mélange, the occurrence of quartz diorite blocks with ages identical to the adjacent Coast Range ophiolite is at odds with an origin on the seafloor long before its accretion to the subduction zone margin. Nonetheless, the Tehama-Colusa mélange displays characteristics consistent
Abyssal peridotites (Small offset or non-tranform settings)
60 50 40 30
VFZ OFZ
20 10
RFZ
0 800
900
1000
1100
1200
1300
T (°C) Figure 15. Cr# in spinel versus equilibration temperatures estimated by two-pyroxene thermometry of Wells (1977). Data sources for large offset transform-fault peridotites (Hamlyn and Bonatti, 1980; Bonatti et al., 1993; Brunelli et al., 2006) and small offset or non-transform setting peridotites (Hamlyn and Bonatti, 1980; Shibata and Thompson, 1986; Dick, 1989; Johnson et al., 1990; Edwards et al., 1996; Hellebrand et al., 2002).
Serpentinite matrix mélange: Implications of mixed provenance for mélange formation
25
abyssal-like peridotites have been affected by fluid enrichments (Choi et al., 2008b). Advanced proto-Franciscan subduction has modified the fracture zone peridotite in the upper plate of the subduction zone, producing the supra-subduction zone peridotites with high Cr-spinels such as Chrome and Hyphus–Little Stony Creeks. Relatively low equilibration temperatures for the suprasubduction zone peridotites support slow subsolidus cooling in the mantle wedge (Fig. 15). Black Diamond Ridge peridotites represent a snapshot of the mantle wedge composition prior to hydrous melting. The nucleation of subduction zones along former transform boundaries has long been proposed for both modern arc systems (e.g., Casey and Dewey, 1984; Bloomer et al., 1995) and for the Franciscan–Coast Range ophiolite system (Stern and Bloomer, 1992). Choi et al. (2008b) show that this process is likely to have occurred in western California during initiation of Franciscan subduction, as proposed by Stern and Bloomer (1992). We show here that during the early stages of this subduction, lithologies from both the upper and lower plate may be mixed together in a matrix of serpentine derived by hydration and abrasion of the lower, upper mantle portion of the upper plate.
analyses of relict primary minerals; trace element and isotopic studies are also required to fully understand the history of the ultramafic rocks. The development of new tools for analyzing small samples of relict primary minerals (secondary ion mass spectrometry) and laser ablation (ICP-MS) has now reached the stage where significant advances can be made in understanding serpentinite mélange.
CONCLUSIONS
APPENDIX 1. DATED FRANCISCAN METAMORPHIC BLOCKS: LOCATIONS, FIELD SETTINGS, LITHOLOGY, 40Ar/39Ar METHODS, AND AGE SPECTRA INTERPRETATIONS
Serpentinite matrix mélanges are common throughout the world’s accretionary complexes, yet they remain poorly understood, and their origin is controversial. The most popular proposals for their origin invoke either entrapment of an oceanic fracture zone or mixing of upper plate and lower plate assemblages along their interface in a subduction zone. Each model makes specific predictions about the composition of both the mélange matrix and its component blocks that may be investigated systematically. The fracture zone–subduction initiation model discussed here provides an additional wrinkle in interpreting these assemblages. Subduction initiation along a preexisting fracture zone may involve abyssal peridotite in the upper plate of the subduction zone or allow lower plate abyssal peridotite to be scraped off into the accumulating mélange. Our work shows that it is possible to unravel the complex tectonic history of a serpentinite mélange belt by careful study of its component lithologies and its matrix. Previous studies have focused largely (or exclusively) on the provenance of volcanic blocks within the mélange. We show that a central focus for any study of serpentinite matrix mélange must be the provenance of peridotite blocks within the mélange, and the mélange matrix itself (e.g., Huot and Maury, 2002; Choi et al., 2008a, 2008b). This is critical for addressing questions about oceanic versus subduction zone origin: The fracture zone model predicts that all of the ultramafic components will be abyssal in origin, whereas subduction zone models predict that the ultramafic components will be largely supra-subduction zone in origin. These studies should not be confined to major element
ACKNOWLEDGMENTS This research was supported by U.S. National Science Foundation grants EAR8816398, EAR9018721, and EAR0440255 (Shervais), and EAR0440238 (Mukasa). Insightful reviews by Trevor Dumitru and Terry Kato helped us organize our thoughts and improved the manuscript greatly, as did the editorial handling and review by John Wakabayashi; we are grateful to all three. We would also like to acknowledge years of discussions with our colleagues and mentors, who challenged our assumptions and helped us to look more deeply into the rocks. These include Cliff Hopson, M. Clark Blake, Angela Jayko, Robert McLaughlin, Robert Coleman, Steve Phipps, and others too numerous to mention.
Garnet amphibolite from Thomes Creek (39° 51′ 21″ N, 122° 39′ 53″ W) occurs as a block in foliated serpentinite of the Tehama-Colusa mélange just north of “The Gorge” on Thomes Creek (Blake et al., 1992). Foliated, early-formed minerals, including hornblende, plagioclase, epidote, garnet, apatite, sphene, and opaque minerals, are ~50% replaced by clinozoisite, actinolite, albite, white mica, chlorite, calcite, pumpellyite, and anhedral sphene. Hornblende gneiss from Alder Springs Road (39° 38′ 30″ N, 122° 36′ 10″ W) occurs as a block in foliated serpentinite of the TehamaColusa mélange (Blake et al., 1992). It consists primarily of foliated hornblende, rutile, and sphene, partly replaced by actinolite, white mica, chlorite, and anhedral sphene. Hornblende gneiss from Thompson Creek near Wilbur Springs crops out as a block within Hauterivian sedimentary serpentinite breccia that lies conformably within clastic strata of the Great Valley Sequence (Carlson, 1981; McLaughlin et al., 1990; Campbell et al., 1993). The gneiss consists primarily of hornblende and rutile, partially replaced (~5%) by blueschist facies minerals including blue amphibole, white mica, and sphene. 40 Ar/39Ar ages were determined using methods and values of interfering neutron reactions given in Ross and Sharp (1988). Criteria for plateau ages are those of Sharp and Renne (2005). Ages herein were calculated using an age of 523.1 Ma for irradiation standard MMhb (Renne et al., 1998). Hornblendes of the three dated samples yield well-defined, high-temperature plateaus characterized by K/Ca ratios appropriate for amphibole, whereas lower temperature increments have considerably higher K/Ca ratios that we interpret to be the result of degassing of trace amounts of younger, high-K white mica intergrowths (see Table 2). This interpretation is similar to that documented more fully by Ross and Sharp (1988) for other Franciscan amphibolite ages.
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Printed in the USA
The Geological Society of America Special Paper 480 2011
Geochemical mapping of the Kings-Kaweah ophiolite belt, California—Evidence for progressive mélange formation in a large offset transform-subduction initiation environment J. Saleeby* Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, California 91125, USA
ABSTRACT The Kings-Kaweah ophiolite belt of the southwestern Sierra Nevada Foothills was generated in two pulses of mid-oceanic-ridge basalt (MORB) magmatism. The first was in the Early Ordovician, which resulted in the generation of a complete abyssal crust and upper mantle section. The crustal section was rendered from convecting mantle whose Nd, Sr, and Pb isotopic systematics lie at the extreme end of the subPacific mantle regime in terms of time integrated depletions of large ion lithophile (LIL) elements. Semi-intact fragments of this Early Ordovician oceanic lithosphere sequence constitute the Kings River ophiolite. Following ~190 m.y. of residence in the Panthalassa abyssal realm, a second pulse of MORB magmatism invaded the Early Ordovician lithosphere sequence in conjunction with intensive ductile shearing and the development of ocean floor mélange. This Permo-Carboniferous magmatic and deformational regime produced many of the essential features observed along spreading ridge–large-offset transform fracture zones of the modern ocean basins. During this regime, Early Ordovician upper mantle–lower crustal rocks were deformed in the ductile regime along what appears to have been an oceanic metamorphic core complex, as well as along steeply dipping strike-slip ductile shear zones that broke the ophiolite into semi-intact slabs. Progressive deformation led to the development of serpentinite-matrix ophiolitic mélange within the abyssal realm. This (Kaweah) serpentinite mélange constitutes the majority of the ophiolite belt and encases fragments of both disrupted Early Ordovician oceanic lithosphere and crustal igneousmetamorphic assemblages that were deformed and disrupted as they formed by diffuse spreading along the fracture zone. An ~190 m.y. hiatus in abyssal magmatism cannot be readily accommodated in the current configuration of Earth’s ocean basins, but it was possible during the mid- to late Paleozoic Panthalassa regime, when the proto-Pacific basin occupied over half of the Earth’s surface. The transform history of the ophiolite belt can be directly linked to the late Paleozoic transform truncation of the SW Cordilleran passive margin. Following juxtaposition of the transform ophiolite belt with the truncated margin a change in relative
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[email protected] Saleeby, J., 2011, Geochemical mapping of the Kings-Kaweah ophiolite belt, California—Evidence for progressive mélange formation in a large offset transformsubduction initiation environment, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 31–73, doi:10.1130/2011.2480(02). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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J. Saleeby plate motions led to the inception of east-dipping subduction, and the en masse accretion of the ophiolite belt to the hanging wall of the newly established subduction zone. Structural relations and isotopic data on superimposed igneous suites show that the ophiolite belt was not obducted onto the SW Cordilleran continental margin. The accreted ophiolite belt formed the proto-forearc of the newly established active margin. The ophiolite belt never saw high-pressure/temperature (P/T) metamorphic conditions. Rare small blocks of high-pressure metamorphic rocks were entrained from the young subduction zone by serpentinite diapirs and emplaced upward into the ophiolitic mélange within a proto-forearc environment. An Sm/Nd garnet-matrix age on a high-pressure garnet amphibolite block suggests subduction initiation at ca. 255 Ma. This timing corresponds well with the initiation of arc magmatism along the eastern Sierra Nevada region. In Late Triassic to Early Jurassic time proximal submarine mafic eruptions spread across and mingled with hemipelagic and distal volcaniclastic strata that were accumulating above the accreted ophiolite belt. These lavas carry boninitic to arc tholeiitic and primitive calc-alkaline geochemical signatures. By Middle Jurassic time siliciclastic turbidites derived from early Paleozoic passive margin strata and early Mesozoic arc rocks spread across the primitive forearc. In late Middle to Late Jurassic time tabular plutons and dike swarms of calcalkaline character invaded the ophiolite belt in a transtensional setting. Deformation fabrics that developed in these intrusives, as well as cleavage that developed in the cover strata for the ophiolite belt, imparted components of superimposed finite strain on the ophiolitic mélange structure but did not contribute significantly to mélange mixing. By ca. 125 Ma, copious gabbroic to tonalitic plutonism of the western zone of the Cretaceous Sierra Nevada batholith intruded the ophiolite belt and imparted regional contact metamorphism. Such metamorphism variably disturbed U/Pb systematics in rare felsic intrusives of the ophiolite belt but did not significantly disturb whole rock Sm/Nd systematics. Age constraints gained from the Sm/Nd and U/Pb data in conjunction with Nd, Sr, and Pb isotopic and trace element data clearly define the polygenetic abyssal magmatic history of the ophiolite belt. The variation of Nd and Sr radiogenic isotopes over time from the Paleozoic abyssal assemblages, through early Mesozoic supra-subduction zone volcanism to Early Cretaceous batholithic magmatism, record the geochemical maturation of the underlying mantle wedge without the involvement of SW Cordilleran continental basement.
INTRODUCTION Mélanges are characterized by the encasement of coherent blocks of rock within a pervasively deformed less competent rock matrix (Hsü, 1968). The blocks may be exotic or native to the hosting matrix, although the inclusion of exotic blocks is the most commonly cited feature interpreted to indicate extreme disruption in the formation of mélange. The common occurrence of mélanges as tectonically bounded units within exhumed active margin assemblages has led to the common assumption that tectonic disruption leading to mélange formation is, a priori, the direct result of subduction megathrust movements. In colloquial usages the term subduction mélange is commonly used in place of what should be the descriptive term mélange. Subduction megathrust environments are well suited for mélange formation, but this chapter takes the position that a number of alternative mechanisms for mélange formation exist, and that such mélange units may form and then be subsequently accreted into active margin assemblages en masse.
This chapter integrates structural, geochronological, and geochemical data for the resolution of the temporal relations and environment of formation of the regional mélange structure of the Kings-Kaweah ophiolite belt of the southwestern Sierra Nevada Foothills metamorphic belt. Similar Paleozoic ophiolitic mélange and related tectonite units occur along the entire length of the Sierra Foothills (see Saleeby, 1990, for a review). Mesozoic overprints have obscured many of the critical relations and have been emphasized in the literature, thereby obscuring the importance of Paleozoic ophiolitic tectonics in terms of the first order shaping of the SW Cordilleran margin. Within the Kings-Kaweah ophiolite belt, several different mechanisms for late Paleozoic and possibly Early Triassic mélange formation are recognized that are not directly related to subduction megathrust motions. Moreover, these processes in conjunction with components of superimposed finite strain account for the entire mélange structure, with no compelling evidence that subduction megathrust motions directly contributed to the mélange mixing process. An abundance of literature documents such disruptive processes, yet they are rarely
Geochemical mapping of the Kings-Kaweah ophiolite belt, California taken into account in the literature on mélanges. For example, vast tracks of severely disrupted abyssal crust and upper mantle are known to have developed at spreading center–transform intersections and are thereby packaged into abyssal crust along fracture zones (Aumento et al., 1971; Bonatti et al., 1971, 1973; Van Andel et al., 1971; Melson and Thompson, 1971; Melson et al., 1972; Bonatti and Honnorez, 1976; Fox et al., 1976; DeLong et al., 1977; Schreiber and Fox, 1977; CAYTROUGH, 1978; Karson and Dick, 1983; MacDonald et al., 1986; Pockalny et al., 1988; Johnson and Dick, 1992). Furthermore, thick and expansive accumulations of severely disrupted sediment that formed by massive submarine landslides are known to form in modern transform valley troughs as well as along both active and passive margins (cf. Moore et al., 1970, 1976; Jacobi, 1976; Johnson and Dick, 1992; Deplus et al., 2001) but are only rarely recognized as such after their emplacement and metamorphism within active margin belts (cf. Cox and Pratt, 1973; Schweickert et al., 1977). Thirdly, serpentinite diapirism with the inclusion of exotic metamorphic blocks and related seafloor extrusion of serpentinite debris flows are known to be important processes in oceanic transform and forearc environments (Lockwood, 1971; Bonatti et al., 1973, 1974; Fryer et al., 2000; Johnson and Dick, 1992). A major challenge in the analysis of active margin belts lies in the recognition of these less familiar mélange formation mechanisms, and the resolution of their role within the regional tectonic history. The ophiolitic mélange problem in the Sierra Foothills belt, and arguably at global scale, encompasses transform tectonics of the abyssal realm as well as subduction zone initiation and primitive forearc evolution. Intimately related to this problem is the possibility that true abyssal lithosphere evolves in geodynamic settings that are not likely to render ophiolite emplacement into continental margin orogens (cf. Stern, 2004). The position taken in this chapter is that subduction initiation along a large offset transform juncture is a viable mechanism for such abyssal ophiolite emplacement along continental margin orogens. This mechanism does not necessarily entail ophiolite obduction onto continental crust, in that fragments of abyssal lithosphere may be accreted to the hanging wall of the newly established subduction zone and thereby form parts of the resulting proto-forearc. For example, tectonic relations along the Izu-Bonin-Mariana arc system indicate that the system originated in the Eocene by subduction initiation along a large offset oceanic transform (Hilde et al., 1977). Fragments of the Pacific plate were accreted to the hanging wall of the young subduction zone (De Bari et al., 1999), to be subsequently joined by proto-forearc assemblages formed by in situ igneous rifting with the eruption of boninitic to arc tholeiitic magma series rocks (Stern et al., 1991; Stern and Bloomer, 1992). The Macquarie Ridge–Puysgur trench–Fiordland plate juncture system represents a case of ongoing subduction initiation along a transform system where abyssal lithosphere is being actively accreted to part of the newly established hanging wall (Varne and Rubenach, 1972; Casey and Dewey, 1984). The Kings-Kaweah ophiolite belt records a petrotectonic evolutionary sequence and
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lies in a paleogeographic setting that entails elements of both of these Cenozoic subduction initiation systems. Specific challenges in interpreting the geologic history of the Kings-Kaweah ophiolite belt, beyond its mélange structure, arise from its intruded and contact metamorphosed state. Subsequent to its tectonic disruption and emplacement into the SW Cordilleran margin the ophiolite belt was invaded by copious gabbroic to tonalitic plutons of the Sierra Nevada batholith (Mack et al., 1979; Saleeby and Sharp, 1980; Chen and Moore, 1982; Clemens-Knott and Saleeby, 1999). Contact metamorphism in albite-epidote to hornblende hornfels facies is near pervasive, leaving only minor lower-grade domains as areas to focus on ophiolite protolith features. Regardless, pre-batholithic features are locally preserved well enough to decipher the geochemical heritage and early tectonic development of the belt. Previous detailed structural and petrographic studies characterized the structural and metamorphic state of the ophiolite belt in detail (Saleeby, 1975, 1977, 1978, 1979), and geochronological studies have provided a number of age constraints on the igneous development of the belt and the complexities of polyphase igneous and thermal overprints (Saleeby and Sharp, 1980; Saleeby, 1982; Shaw et al., 1987). This paper presents additional isotopic and geochronological data, and coupled major and trace element data from sample sites representative of the principal units of the ophiolite belt exhibiting the lowest grade metamorphic overprints observed, as well as on sample sites preserving critical structural and stratigraphic relations that unfortunately possess relatively high-grade metamorphic overprints. Special focus is placed on isotopic and trace element data in terms of the discrimination of MORB igneous suites from subsequent suites having formed within a supra-subduction zone environment. Major and trace element data are further used as a means to evaluate the potential severity of disturbances in the radiogenic isotopic systems of Nd, Pb, and Sr, as these systems most clearly record the early MORB history of the belt. An overview of the critical structural relations of the ophiolite belt is presented first, and then the geochemical data are integrated with the structural relations as a mapping tool. The results of this integrated structural-geochemical mapping procedure are then integrated with regional tectonic relations in the formulation of a model for progressive ophiolitic mélange development followed by continental margin emplacement in a large offset transform-subduction initiation environment. GEOLOGIC OVERVIEW Regional Structure The Kings-Kaweah ophiolite belt extends for ~130 km along the western Sierra Nevada Foothills between 35.9° N and 37° N (Fig. 1). It constitutes the southern segment of the Foothills metamorphic belt, where the Foothills belt is more highly intruded by the Sierra Nevada batholith than to the north. Basement core data (May and Hewitt, 1948; Wentworth et al., 1995; Saleeby, 2007, and unpub. data) indicate that rocks typical of
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Index map Foothills metamorphic belt
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Petrotectonic unit map of Kings-Kaweah ophiolite belt Legend
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ea
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Foothills metamorphic belt Late Jurassic (152 -148 Ma) Owens Mountain complex, basalt-diorite-trondjhemite dike swarm with ophiolitic and siliciclastic screens Middle to Late Jurassic (170 -156 Ma) Mill Creek intrusive complex, gabbroic to tonalitic sheets
ah we Ka River
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Upper Triassic to Jurassic western facies Kings sequence: siliciclastic and volcaniclastic schists and slates, local marble Permian to mid-Triassic Calaveras complex: chert-argillite with local marble and siliciclastic schists and slate Permo-Carboniferous Kaweah serpentinite mélange-KSM (296 +15 Ma) Serpentinized peridotite, pillow basalt, chert, and ophicalcite blocks Disrupted gabbroic intrusives and basaltic pillows + chert blocks Serpentinite matrix dominated with subordinate ophiolitic crustal blocks
Fig. 3
Basalt-gabbro blocks derived exclusively from Kings River ophiolite Early Ordovician Kings River ophiolite-KRO (484 +18 Ma) Basaltic pillows and sheeted dikes, and static and cumulate gabbros Peridotite tectonites Subcrop map units
Geochemical samples for Kings River ophiolite not shown in Figures 2 and 3:
Early Cretaceous Sierra Nevada batholith
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Antelope Mountain
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Smith Mountain
190 e Tul ver Ri 36O 00′ N
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Figure 1. Petrotectonic unit map of the Kings-Kaweah ophiolite belt and related metasedimentary rocks. Geology after Saleeby (1977, 1978, and 1979), and subcrop mapping of Great Valley margin from gravity-magnetic modeling (Saleeby, 1975; Oliver and Robbins, 1978) and basement core data (Saleeby, 2007, and unpub. data).
Geochemical mapping of the Kings-Kaweah ophiolite belt, California the Kings-Kaweah ophiolite belt and its metamorphosed lower Mesozoic cover strata, as well as crosscutting Sierra Nevada batholith plutons, occur at least as far west as ~30 km beyond the Foothills basement exposures, and that such rocks continue in the Great Valley subsurface at least as far south as 35.4° N. Figure 1 depicts the Kings-Kaweah ophiolite belt in terms of its constituent petrotectonic units: some units characterized as large polylithologic slabs, and others as mélange units characterized by distinct block assemblages. Pre–Sierra Nevada batholith metamorphic wall rocks that lie directly east of the Kings-Kaweah ophiolite belt consist of chert-argillite and marble of the Permian to Middle Triassic Calaveras complex, which are locally intergradational with siliciclastic schists, mafic to felsic metavolcanic rocks, and marble, which constitute the western facies of the Upper Triassic–Jurassic Kings sequence (Saleeby et al., 1978; Saleeby and Busby, 1993). These metasedimentary rocks are in ductile-brittle fault contact along the east margin of the KingsKaweah ophiolite belt (Figs. 1 and 2), but also, at least locally, sit nonconformably above the belt (Fig. 3). In this structural setting these rocks are referred to as cover strata for the ophiolite belt. Cleavage that is axial planar to meso- and map-scale folds within the cover strata crosses the folded nonconformity with the ophiolite belt, and re-deforms the basement mélange structure by imparting additional components of finite strain. Structural and temporal relations of cleavage development in the cover strata, and ductile deformation fabrics in relatively low-volume early Mesozoic intrusions that crosscut the ophiolite belt, are discussed in Saleeby (1979) and Saleeby and Dunne (2011). The central segment of the Kings-Kaweah ophiolite belt (36.5° N to 36.7° N) is in part cut out by a series of Early Cretaceous gabbroic to tonalitic plutons of the Sierra Nevada batholith (Saleeby and Sharp, 1980; Clemens-Knott and Saleeby, 1999) and is in part buried beneath Quaternary deposits of the southeastern Great Valley (Fig. 1). A buried pendant of the ophiolite belt lies directly west of the edge of the Foothills between 36.45° N and 36.7° N. The pendant and adjacent buried Sierra Nevada batholith plutons are shown in subcrop in Figure 1 based on gravity and magnetic modeling (Saleeby, 1975, 2007; Oliver and Robbins, 1978), geologic observations of small isolated hills that expose basement rocks, and on basement core data (Saleeby, 2007, and unpub. data). The central segment of the ophiolite belt sits between two distinct, yet closely related, segments that are relatively well exposed along steep sided grassy hills. The northern segment consists of the Kings River ophiolite, and the southern segment consists of the main exposures of the Kaweah serpentinite mélange. The Kings River ophiolite and the Kaweah serpentinite mélange are closely related. Serpentinite mélange zones of Kaweah affinity extend into and disrupt the Kings River ophiolite, whereas mélange blocks derived from this ophiolite occur as one of the distinct inclusion assemblages within the Kaweah serpentinite mélange. Serpentinite-matrix mélange zones of the Kings-Kaweah ophiolite belt appear to have formed by at least four distinct mechanisms (Saleeby, 1978, 1979). These consist of the following:
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1. Focused ductile shear strain that developed under highgrade, retrogressing to medium-grade, metamorphic conditions, and that resulted in the progressive disruption and mixing of the original ophiolite stratigraphy. Where such strain penetrated upper mantle levels the resulting peridotite mylonites underwent domainal hydration to form antigorite ± talc ± tremolite schist zones that concentrated subsequent shear strain, resulting in the entrainment of blocks derived from crustal rocks. 2. The progressive disruption and abrasion of completely serpentinized peridotite blocks with the marginal entrainment of exotic blocks, probably during diapiric rise. Vestigial ultramafic blocks in these cases commonly possess a distinct blocky fracturing, which, based on textural relations, developed prior to Sierra Nevada batholith contact metamorphism. Such blocky fracturing is common in the core areas of serpentinite diapirs (Moiseyev, 1970; Lockwood, 1971). 3. The emplacement of foliated serpentinite sheets along extensional and dilational shear fractures in otherwise coherent mafic rocks. 4. Mixing of debris flow and detrital serpentinites with various ophiolitic slide blocks on the seafloor. Such sedimentary mixtures readily acquired a strong shape fabric during the superposing of finite strain, thereby attaining a structure resembling that of tectonic mélange. Direct links between the later three mechanisms are suggested by their intergradational relationships, and by the logic of sourcing detrital and debris flow serpentinites from surfaced diapiric serpentinites. The four serpentinite mélange formation mechanisms described above account for the majority of the mélange structure of the Kings-Kaweah ophiolite belt. An additional fifth mechanism was locally important, consisting of deeply sourced diapiric serpentinites that entrained rare high-pressure metabasite blocks that are exotic relative to the mafic crustal rocks of the Kings-Kaweah ophiolite belt. Structural, metamorphic, and age data presented below show that this subordinate mechanism significantly postdates the principal activity of the four main serpentinite mélange formation mechanisms. Kings River Ophiolite The Kings River ophiolite is exposed in a series of tectonic slabs that are as long as ~20 km and are separated by serpentinite-matrix mélange zones and crosscutting plutons of the Sierra Nevada batholith (Figs. 1 and 2). The principal slabs of the Kings River area consist of the depleted peridotite-cumulategabbro Red Mountain–Tivy Mountain slab, sheeted dike–pillow lava slabs of Hughes Mountain and Dalton–Bald Mountain, and the depleted peridotite-mafic tectonite slabs of Hog Mountain and upper Hughes Creek. Additional slabs of cumulate and static textured gabbro and pillow basalt with local sheeted dike sets occur at Smith and Antelope Mountains, respectively, within the highly intruded central segment of the ophiolite belt (Fig. 1). Within the tectonic slabs, various intervals of the original ophiolite succession are preserved. Homogeneity within the various
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Figure 2 (Continued on facing page). Geologic map of the Kings River ophiolite, showing localities of geochemical samples (geology after Saleeby, 1978, and unpub. data).
Geochemical mapping of the Kings-Kaweah ophiolite belt, California
Middle to Late Jurassic (170-156 Ma) Mill Creek intrusive complex: wehrlite - clinopyroxenite, hornblende pyroxene gabbro, diorite, tonalite, and leucotonalite dikes
Figure 2 (Continued).
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Quaternary non-differentiated Early Cretaceous tonalite and granodiorite of Sierra Nevada batholith (SNB)
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High angle brittle-ductile fault Dominant trend of Mesozoic composite foliation
Cover strata sequence Middle to Upper Jurassic siliciclastic turbidites and black slate with sedimentary serpentinite/ophiolitic debris flows. Western Unconformity facies Upper Triassic to Lower Jurassic tuffaceous and siliceous argillite and Kings sandstone, pebble-cobble conglomerate and mudstone, rare felsic tuff sequence and mafic volcaniclastic and pillowed volcanic lenses Gradational contact - possible unconformity Permian to mid-Triassic (?) Calaveras complex: bedded radiolarian chert, chaotic chert-argillite and diamictite, with shallow water Permian limestone blocks, and local siliciclastic/volcaniclastic turbidite infolds (?) Nonconformity with all above units
Valley
Permo-Carboniferous ductile shear fabrics within gabbro-peridotite slabs
i o
Inner/outer contact metamorphic aureole boundary
Geochemical sample sites:
9
4 11
Cover strata (C suite) Ophiolitic ductile shear zones and mélange (M suite) Kings River ophiolite blocks (O suite)
Kaweah Serpentinite Matrix Mélange (KSM): Permo-Carboniferous MORB magmatism and mixing Sedimentary serpentinite and ophicalcite derived matrix Downplunge viewing direction for cover strata nonconformities
Bedded radiolarian chert (Pennsylvanian - Permian) Basaltic pillows, pillow breccia, and hyloclastite Foliated d serpentinite matrix with chert-ophicalcite blocks (red) and metabasite blocks (black) Hornblende-clinopyroxene gabbroids and rare diorite, plagiogranite, and gabbroic pegmatites Hornble en serpentinized peridotite blocks Blocky fractured fr
E. Jurassic diorite dike
Variably y serpentinized mylonitic peridotite blocks blocks derived from Early Ordovician Kings River ophiolite: Mafic b blo Pillowed Pillow we and massive basalt and mafic dike complex mélange blocks Clinopyroxene gabbro and anorthositic gabbro mélange blocks Clinopy Serpentinite mélange colluvium
* Protholith terms refer to state prior to Sierra Nevada batholith contact metamorphism
Figure 3. Geologic map of the Yokohl Valley area of the Kaweah serpentinite mélange, showing critical relations with cover strata and localities of geochemical samples (geology after Saleeby, 1979, and unpub. data).
Geochemical mapping of the Kings-Kaweah ophiolite belt, California lithologic zones, as well as overlap in the various lithostratigraphic intervals of the original succession that are preserved in adjacent slabs, facilitates the reconstruction of the original lithostratigraphic section (modified after Saleeby, 1978). From the base upward the reconstructed section consists of (1) a variably serpentinized tectonitic harzburgite, clinopyroxene harzburgite, and dunite that are >~6 km thick; (2) cumulate to static textured gabbros (~2 km thick); (3) mafic sheeted dikes (up to ~1 km thick); and (4) pillow basalt (up to ~2 km thick). The pillow basalt locally includes meter-scale lenses of metalliferous radiolarian chert, although a stratigraphic top has not been determined for the pillowed section. All principal exposures of pillow basalt (Hughes, Dalton-Bald, and Antelope Mountain slabs) consist of exceedingly homogeneous pillows, local pillow breccia, and local basaltic feeder dikes. Likewise the sheeted dikes are exceedingly homogeneous basaltic dikes with rare diabase dikes and screens. Basaltic pillows and dikes are either aphyric or modestly plagioclase phyric. The gabbroic rocks are also quite homogeneous with coarse locally layered clinopyroxene gabbro and anorthositic gabbro strongly dominating, and layered troctolite and minor plagioclase ± olivine clinopyroxenite present near the base of the gabbroic section. Wehrlitic and resolvable dunitic cumulates are rare, as are chromite pods concentrated along what is interpreted as the base of the cumulate section, although strong plastic deformation has obscured the primary relations. Much of the Tivy Mountain section is composed of coarse-grained, massive clinopyroxene-plagioclase rock suspected to be adcumulates, although primary textural relations are obscured by contact metamorphism. Igneous textures appear to fine southeastward with transitions into finer static textured and coarse diabasic zones, which, along with the basal ultramafic cumulates to the northwest, suggest a general facing to the cumulate section in the direction of the Dalton–Bald Mountain sheeted dike-pillow basalt sequence (Fig. 2). Smith Mountain (Fig. 1) preserves rocks lithologically identical to the coarse cumulates and finer static-textured gabbros of Tivy Mountain. Gravity-magnetic and basement core data indicate that a large mass of the Smith Mountain gabbro extends in the subsurface northwestward toward Tivy Mountain (Saleeby, 1975, 2007; Oliver and Robbins, 1978). The Kings River ophiolite contrasts from many or perhaps most ophiolites (cf. Coleman, 1977) by an apparent lack of significantly fractionated rocks. The crustal section is composed entirely of basalt and various gabbroids. No intermediate to felsic-composition plutonic, hypabyssal, or volcanic rocks as part of the principal ophiolitic igneous suite have been discovered. Such fractionated rocks are common in the younger igneous suites that have been superposed across the exclusively MORB affinity Kings River ophiolite mafic crustal section. Near pervasive contact metamorphism has inhibited detailed petrogenetic studies of the Kings River ophiolite. All primary minerals of the basaltic pillows and dikes are completely overprinted, and only traces of olivine, pyroxenes, and spinel are preserved in the ultramafic rocks. Remnants of igne-
39
ous plagioclase and clinopyroxene are more widely preserved in gabbros and locally in coarse diabases. Olivine is completely replaced by chlorite, serpentine, and magnetite in troctolites and gabbros but is commonly preserved as grain core remnants in ultramafic rocks. One of the outstanding features of the Kings River ophiolite is the transposition of the base of the gabbroic section and the continuation of such high plastic strain fabrics downward through much of the depleted peridotite section (Saleeby, 1978). Such strain is of much greater magnitude and encompasses a wider range of retrogressing conditions than what is typical of depleted peridotite sections of many well-preserved ophiolites (cf. Coleman, 1977). Structural form lines are shown in Figure 2 for pervasive mylonitic fabrics in the peridotites. Such widespread ductile shear in the mantle section extends into the mafic crustal section along more discretely defined ductile shear zones, shown in generalized form in Figure 2. Small mafic intrusions consisting of transposed dikes, and folded and boudinaged lenses that are petrologically distinct from the Tivy and Smith Mountain type gabbros, were magmatically emplaced into the shear zones during the principal phase of high-temperature plastic deformation. The shear zones coalesce and penetrate through much of the Red Mountain peridotite, leaving only local vestigial domains where higher temperature, upper-mantle-flow fabrics remain (i.e., Carter and Ave Lallemant, 1970; Boudier and Coleman, 1981). In the Hog Mountain peridotite slab the shear zones are more domainal, leaving larger vestigial domains with the remnants of high-temperature, mantle flow fabrics. The upper Hughes Creek slab is pervasively mylonititc along steep foliation surfaces. The shear zones represent high distributed shear strain within the sub-oceanic mantle that penetrated the crust along more concentrated zones, but they are not mélange. There are two types of shear zones on the basis of their orientations relative to the original ophiolite lithostratigraphy: (1) a basal zone that runs along the sub-oceanic Moho level of the ophiolite, as exposed along the lower Kings River Valley (Fig. 2); and (2) steeply dipping, longitudinal shear zones that cut at high angles across the ophiolite section, and which commonly grade into serpentinite-matrix mélange zones. Deformation fabrics of the basal shear zone are both cut by and merge into fabrics of the longitudinal shear zones, indicating that the two types of shear zones are partly coeval. The basal shear zone is characterized by strong constrictional fabrics with stretch factors typically >~10 that render a rodding structure to many of the transposed intrusions within the peridotites as well as corrugated separation planes in the mylonitic fabric of the peridotites. A relatively well-defined maximum in the principal stretch direction along the basal shear zone plunges at intermediate angles southward beneath the Tivy Mountain gabbro section (Fig. 2). Conflicting non-coaxial shear fabrics distributed within this shear zone appear to integrate into a coaxial component of the principal stretch (Saleeby, 1978). The longitudinal shear zones, where expressed in the peridotite sections, grade into serpentinite mélange zones with the inclusion of rock types exotic to
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J. Saleeby
the mantle section, such as pillow basalt, chert, and ophicalcite. Deformation fabrics in the longitudinal shear zones and their transitions into serpentinite mélange zones typically reflect a stronger flattening component than the constrictional fabrics that characterize the basal shear zone. Ductile shear zones of the Kings River ophiolite developed during the intrusion of a series of distinct brown hornblendebearing plagioclase-clinopyroxenites, clinopyroxene gabbros with local dioritic fractionates, and sets of solitary diabase dikes. High-temperature hydrothermal veins characterized by mixtures of antigorite, Cr-chlorite, talc, tremolite, and rare anthophyllite developed in the peridotite host during shearing and intrusive activity. These veins were variably transposed along with the mafic intrusive bodies. The intrusive bodies are pervasively mylonitic to ultramylonitic and blastomylonitic with some grading into coarse-banded clinopyroxene amphibolite. They underwent limited fractionation during deformation and crystallization to form variably transposed veins of mylonitic diorite. Embrittlement of the mafic bodies late in the ductile deformation regime is recorded by hornblende ± Na scapolite veining along extensional fractures that are oriented normal to the principal stretch direction. Locally such veins are also transposed into the principal stretch direction by late-stage ductile shear bands. Veining and hydration of the ultramafic host and mafic intrusive lenses are interpreted to have resulted from the influx of ocean water during ductile shearing. Attempts to confirm this by stabile isotope data are foiled by the effects of near pervasive Sierra Nevada batholith contact metamorphism. Mélange units and ductile shear zones of the Kings River ophiolite are crosscut by a series of Middle to Late Jurassic tabular plutons and dike swarms (Saleeby and Sharp, 1980; Wolf and Saleeby, 1995; Saleeby and Dunne, 2011). These intrusions consist of the Mill Creek and Owens Mountain complexes (Figs. 1 and 2) and numerous small isolated dikes that are most readily observed where they cut ultramafic rocks. The Mill Creek complex consists of NW-trending lenticular gabbroic, dioritic, and tonalitic plutons containing modest to strong ductile deformation fabrics. The Mill Creek complex is structurally coherent. The Owens Mountain complex consists of basaltic, gabbroic, dioritic, and trondhjemitic dikes and small lenticular stocks that intrude and include as screens the western serpentinite mélange zone of the Kings River ophiolite (Fig. 1). The Owens Mountain complex is also deformed but structurally coherent, like the Mill Creek complex. A critical feature of the Mill Creek and Owens Mountain complexes is that they crosscut and thus postdate serpentinite mélange formation, yet they are plastically deformed. Thus they provide age constraints on both mélange development and on finite strain superimposed across the mélange structure. The ductile deformation fabrics of the Mill Creek and Owens Mountain intrusive complexes, as well as those of the Kings-Kaweah ophiolite belt mélange structure, are truncated by Lower Cretaceous plutons of the Sierra Nevada batholith (Figs. 1 and 2), which, for the most part, lack ductile deformation fabrics.
Kaweah Serpentinite Mélange The Kaweah serpentinite mélange consists primarily of a foliated serpentinite matrix in which ophiolitic blocks of a wide range of sizes, shapes, and compositions are dispersed. The serpentinite matrix was derived from tectonitic peridotite, diapiric serpentinite, and serpentinitic debris flows and detrital rocks. The mélange blocks consist mainly of serpentinized peridotite, gabbro, diabase, basalt, chert, and ophicalcite. The blocks in many localities are oriented with long axes parallel to the matrix foliation, and they range in size from centimeters to kilometers. Relict primary features are preserved within the blocks, such as bedding, pillows, dikes, and cumulate layering. Ophicalcites occur as bedded carbonates with dispersed serpentinite ± basalt ± chert clasts, as submarine regoliths developed on serpentinites and serpentinized peridotites, and as vein networks along submarine faults and extensional fractures within ultramafic rocks. Outcrop mapping reveals clustering of mélange blocks into several lithologic associations (Saleeby, 1977, 1979). The associations are defined as mélange units (Fig. 1), some of which appear to represent internally mixed vestiges of once intact ocean-floor sections. Four principal types of mélange units are recognized: (1) serpentinite matrix-rich, with highly dispersed blocks typically of mafic tectonite, chert, and ophicalcite; (2) dense block clusters derived almost solely from the mafic crustal rocks of the Kings River ophiolite; (3) block and blocklike bodies of distinctive gabbroids and diabases that were intruded into variably deformed and serpentinized peridotite, and adjacent blocks consisting of pillow basalt, pillow breccia, and hyloclastite, as well as local chert; and (4) completely serpentinized peridotite, basalt, and chert, commonly with ophicalcite blocks and/or matrix disseminations. The first two units represent variably disrupted remnants of the Kings River ophiolite peridotite section and its crustal section, respectively. The third type of mélange unit represents “abnormal” abyssal crust composed of crustal-level serpentinized peridotite that was intruded by variably fractionated gabbroids and diabases, and which sat beneath pillowed eruptions and pelagic sediments. The fourth type also represents “abnormal” abyssal crust consisting of serpentinite diapirs and debris flows, detrital serpentinite, ophicalcite, chert, and pillowed eruptions. Relationships discussed below indicate that serpentinite diapirism and surface level extrusion were progressive through time, spanning residence in the abyssal realm as well as in the proto-forearc realm following accretion of the Kings-Kaweah ophiolite belt to the SW Cordilleran margin. A severe problem in interpreting these later mélange units is that hypabyssal dikes and pillow lava–breccia mounds having primary contacts with surfaced serpentinites readily broke into mélange block-like inclusions, and hosting detrital serpentinite textures were severely deformed upon superimposed folding and cleavage development. Also, examination of Figure 3 reveals a number of chert “blocks” that are oriented at high angles to the matrix foliation. These bedded bodies are commonly crosscleaved parallel to the matrix foliation, and many could be the
Geochemical mapping of the Kings-Kaweah ophiolite belt, California remnants of chert intervals that were interbedded with detrital serpentinite matrix materials. These “mélange” units may in fact not be mélange but instead be highly strained remnants of bedded and diked primary sequences for which strain localization in hosting serpentinites resulted in a serpentinite matrix, mélangelike assemblage. Age constraints and geochemical data, covered below, as well as stratigraphic relations, reveal two distinct series of MORB pillowed eruptions that constitute the Kings-Kaweah ophiolite belt. The older series formed the upper crustal levels of the Kings River ophiolite abyssal lithosphere section, and sat above a coherent interval of basaltic sheeted dikes. The younger series was erupted during the tectonic disruption of the seafloor and the extrusion of serpentinite. Locally such pillow basalts, pillow breccias, and hyloclastites are interbedded with ophicalcite and sedimentary serpentinite, but in most localities such flows into serpentinitic sediments were broken into blocks or intergradational lenses with matrix material by subsequent deformation. Pillow lavas erupted over the highly disrupted ophiolitic substrate are commonly interbedded with and overlain by radiolarian chert, with the thickest and most continuous chert section increasing upward in siliceous argillite laminae. Such cherts and interbedded pillow flows form diagonal and transverse outcrop belts that cross the structural trend of the mélange. One such outcrop belt that is well preserved runs NE-SW across the central part of the Yokohl Valley map area (Fig. 3). The chert section of this belt is faulted and cross-cleaved parallel to the matrix foliation. The remnants of radiolarian tests are common in thin sections from this belt but are too recrystallized for paleontologic study. Chert and siliceous argillite of the Calaveras complex lies along a complexly deformed nonconformity above serpentinite mélange, where the matrix was derived primarily from serpentinitic debris flows and detrital rocks (Fig. 3). Much of the chertargillite possesses a chaotic fabric, which commonly grades into diamictites in areas of relatively low superimposed strain. Limestone clasts and blocks ranging up to ~500 m in diameter are encased in the chaotic chert-argillite. These rocks are most readily interpreted as olistostromes or submarine landslides (Saleeby 1979). Some of the larger limestone olistoliths contain Permian shallow-water benthic foraminifers (Saleeby et al., 1978). The remnants of radiolarian tests are common in thin sections of the chaotic cherts but are everywhere too recrystallized for paleontologic study. The stratigraphic thickness of the Calaveras complex is poorly constrained, whereas structural thickness ranges up to ~3 km where it sits on the Kaweah serpentinite mélange (Fig. 3), and up to ~5 km in the outcrop belt east of the KingsKaweah ophiolite belt (Fig. 1). It seems likely that the Calaveras complex of the study region in its entirety once sat above the disrupted Kings-Kaweah ophiolite belt, but currently much of it is dislodged from its original basement by Sierra Nevada batholith magmatism, or alternatively, Kings-Kaweah ophiolite belt basement remnants lie concealed beneath the eastern outcrop belt. The upper stratigraphic levels of the Calaveras chertargillite grade into siliciclastic, volcaniclastic, and mafic flows
41
of the Upper Triassic–Jurassic western facies Kings sequence (Saleeby, 1979; Saleeby and Busby, 1993). These strata possess a cleavage but are structurally intact except where they are adjacent to local basement-derived serpentinite dikes (Fig. 3). Except for such dikes and their derivative submarine debris flows, the lack of ophiolitic lithologies mixed into the Calaveras or western Kings sequence places a pre–early Mesozoic relative age constraint for mélange mixing, analogous to the structural chronology posed by Jurassic intrusive complexes that cut the Kings River ophiolite ductile shear and serpentinite mélange zones. Structural complexity and severity of contact metamorphic overprinting require an iterative approach in the structural and petrologic analysis of the Kaweah serpentinite mélange and its relations with the Kings River ophiolite. We now turn to the use of geochemical techniques as a mapping tool in this analysis. GEOCHEMICAL MAPPING AS A STRUCTURAL AND TECTONIC TOOL The Kings-Kaweah ophiolite belt presents a unique problem in its tectonic analysis relative to many ophiolites in that numerous generations of mafic intrusive and submarine volcanic units are associated with the ophiolite belt with ages that span much of Paleozoic and Mesozoic time. As demonstrated below, only the Paleozoic assemblages constituted consanguineous abyssal lithosphere, with a variety of Mesozoic assemblages having formed in a supra-subduction zone environment. The Kings-Kaweah ophiolite belt is also unique relative to numerous ophiolites (cf. Mattinson, 1976; Hopson et al., 1981, 2008; Tilton et al., 1981; Harper et al., 1994) in that cogenetic felsic intrusives are exceedingly rare. The paucity of such intrusives, the severity of structural and metamorphic overprints, and multiple generations of crosscutting dikes and plutons have inhibited high-precision dating of the ophiolite belt’s principal igneous suite. Zircon-bearing felsic rocks that can be shown by structural relations and geochemical data to be consanguineous with the Kings River ophiolite crustal sequence have not been discovered. Meter-scale plagiogranite bodies occur in three gabbroic mélange blocks from the southern segment of the Kings-Kaweah ophiolite belt. Zircon from these, along with a decimeter-scale diorite-mylonite layer from a transposed gabbro rod within the Kings River ophiolite basal shear zone, yielded disturbed U/Pb zircon systematics, indicating Permo-Carboniferous igneous generation (Saleeby and Sharp, 1980). These rare felsic rocks and their hosting mafites were erroneously interpreted as remnants of the Kings River ophiolite crustal section, leading to an initial erroneous Permo-Carboniferous age assignment for this ophiolite section. Likewise, early Mesozoic (ca. 200 Ma) U/Pb zircon ages on additional felsic dikes interpreted to be part of the Kings River ophiolite crustal section are shown to be members of a superimposed supra-subduction zone igneous suite (Saleeby and Dunne, 2011; and below). The polygenetic nature of the ophiolite belt, and the age range over which magmatism has
42
J. Saleeby TABLE 1. INFORMATION ON FIELD SETTINGS AND PROTOLITHS* OF GEOCHEMICAL SAMPLES † Sample Protolith Latitude (°N) Lo n g i tu de ( ° W ) O1 Pl phyric pillow basalt 36.84122 119.32083 O2 Pl phyric pillow basalt 36.75606 119.21613 O3 Aphyric pillow basalt 36.74280 119.16774 O4 Pl phyric pillow basalt 36.43381 119.06870 36.75947 119.23226 O5 a. Pl phyric basaltic dike b. Ophitic diabase screen O6 Aphyric basaltic dike 36.76591 119.25207 O7 Cpx gabbro layered cumulate 36.79389 119.36019 O8 Cpx gabbro 36.78779 119.35648 O9 Cpx gabbro 36.78092 119.39213 O10 Cpx gabbro layered cumulate 36.58867 119.34355 O11 Cpx gabbro 36.31176 119.04746 36.81609 119.38025 O12 Pl-ol clinopyroxenite layered cumulate O13 Troctolite layered cumulate 36.79847 119.36898 M1 Aphyric pillow basalt 36.26122 119.05173 M2 Aphyric pillow basalt 36.26874 119.04552 M3 Hb-cpx diabase dike 36.79389 119.33611 M4 a. Hb-cp x gabbro mylonite 36.91908 119.41852 b. Hb diorite mylonite 36.81626 119.39012 M5 a. Hb-cpx anorthositic gabbro mylonite b. Hb-cp x gabbro mylonite 36.8162 5 119.39011 c. Hb-qtz diorite mylonite M6 a. Hb-cp x gabbro mylonite 36.26424 119.07985 b. Hb-qtz diorite mylonite M7 a. Plagiogranite screen 3 6. 28 16 4 1 19 .0 9 6 80 b. Hb-cpx diabase dike M8 Pegmatitic hb gabbro pod 36 .2 03 2 2 11 9. 04 89 9 M9 Gt amphibolite 36.24600 1 1 9 .0 3 9 7 6 C1 Uralitic diorite dike 36.26199 119.03341 C2 Px phyric basaltic tuff breccia 36.26875 119.02265 a. Px phyric basaltic tuff breccia 36.25436 119.01481 C3 b. Pl phyric dacite block in basaltic breccia 36.25438 119.01483 C4 Px phyric basaltic broken pillow breccia 36.25401 119.01088 C5 Px-pl phyric massive basaltic-andesite 36.22072 119.01531 C6 Px phyric pillowed basaltic-andesite 36.22425 119.01029 C7 Px phyric basaltic pillow breccia 36.22175 119.00516 C8 Aphyric pillow basalt 36.23575 119.03929 C9 Graded siliciclastic turbidite 36.23172 119.03088 *Protolith refers to pre–Sierra Nevada batholith contact metamorphic state. Key mineral symbols: Cpx—clinopyroxene; Gt—garnet; Hb—hornblende; ol—olivine; Pl—plagioclase; Px—pyroxene; qtz—quartz. † O—Kings River ophiolite; M—ophiolitic ductile shear zones and Kaweah serpentinite mélange; C—cover strata.
affected it, require detailed iteration between field, geochronological, and geochemical techniques for its proper age analysis. The strategy of this study is to build on the existing geochronological and radiogenic isotopic data (Saleeby and Sharp, 1980; Shaw et al., 1987), and to use it in conjunction with geological relations to guide additional sampling and analyses to (1) constrain the igneous generation age of the abyssal lithosphere section preserved in the KRO; (2) constrain its age of ductile shearing and its disruption to form ophiolitic mélange; (3) constrain in time the emplacement of the ophiolite belt into the SW Cordilleran continental margin; (4) utilize geochemical discriminate analyses to further test the MORB origin of the Kings-Kaweah ophiolite belt, as suggested by Shaw et al. (1987); (5) constrain in time the transition from abyssal MORB magmatic growth to supra-subduction zone magmatic growth; and (6) use the above in conjunction with structural and stratigraphic relations to better resolve the tectonics of ophiolite generation, disruption, and emplacement into the SW
Cordilleran margin. The principal new data sets that are presented include Nd and supporting Sr and Pb isotopic data, major and trace element data for the same sample suite, and additional U/Pb zircon ages. The field setting and protolith information on the samples studied are presented in Table 1. The samples are separated into three suites that are differentiated by letter modifiers such that O represents those of the Kings River ophiolite crustal section; M represents those associated with serpentinite mélange formation, including igneous emplacement into high-temperature, ophiolitic ductile shear zones; and C represents those associated with the lower Mesozoic cover sequence. Analytical procedures used for the new data sets are presented in the Data Repository1 along with
1
GSA Data Repository Item 2011260, Geochemical data tables and presentation of analytical techniques, is available at www.geosociety.org/pubs/ft2011 .htm, or on request from
[email protected], Documents Secretary, GSA, P.O. Box 9140, Boulder, CO 80301-9140, USA.
Geochemical mapping of the Kings-Kaweah ophiolite belt, California 0.5128
0.5125
A (cpx)
0.5124
B
0.5124
(hb)
0.5123
143
Nd/144Nd
0.5126
43
(pl)
Age = 484 ± 18 Ma εNd (484) = +10.7 ± 0.4
0.5122
2
Age = 299 ± 32 Ma
0.5122
εNd (299) = +10.0 ± 0.7 2
errors, MSWD = 3.2
errors, MSWD = 6.8
(pl) 0.5120 0.08
0.12
0.16
0.20 147
pillow basalts troctolite
0.24
0.28
0.5121 0.32
144
0.12
0.16 147
Sm/ Nd
cpx = clinopyroxene sheeted dikes gabbros pl-ol clinopyroxenite
0.08
0.20
0.24
0.28
144
Sm/ Nd
pl = plagioclase hb= hornblende plagiogranite diorites hornblende pegmatite
gabbros diabase dikes pillow basalts
Figure 4. Isotopic evolution diagrams of 143Nd/144Nd. (A) Kings River ophiolite samples. (B) Ophiolitic ductile shear zone and Kaweah serpentinite mélange samples. Isochron solution after Ludwig (2001). Data in Table DR1 (see footnote 1).
the tabulation of the new and previously published data. The focus is first on the Kings River ophiolite, followed by the Kaweah serpentinite mélange, and then topics regarding ophiolite emplacement and supra-subduction zone residence are pursued. Kings River Ophiolite: Early Ordovician Abyssal Lithosphere Homogeneity of the lithologic zones that make up the Kings River ophiolite facilitates the broad characterization of the ophiolite by geochemical techniques. Based on petrography, field relations, and geochemical data presented below, the principal slabs of the KRO, and the Smith Mountain and Antelope Mountain slabs as well as the cluster of petrographically similar mélange blocks of the northern Yokohl Valley area (Figs. 1 and 3), are all considered to be consanguineous members of the Kings River ophiolite. Fifteen samples were selected from these rocks for Sm-Nd whole rock and mineral isochron techniques, as well as major and trace element abundance studies. The samples include plagioclase-olivine clinopyroxenite, troctolite, anorthositic gabbro, clinopyroxene gabbros, ophitic diabase, sheeted basaltic dikes, and pillow basalts. Except for the clinopyroxenite, sample selection was based on the field sites as being highly representative of the principal crustal rocks of the Kings River ophiolite, and the sites being non-proximal to crosscutting plutons. Clinopyroxenites are somewhat rare in this ophiolite, and one was chosen for analysis in the hope of generating a large Sm-Nd spread to the whole-rock analyses. Coarsegrained gabbroids, including the clinopyroxenite sampled, are the only Kings River ophiolite crustal rocks that have not been
completely recrystallized by Sierra Nevada batholith contact metamorphism. Careful attention was given to gabbros in search of relatively fresh clinopyroxene and plagioclase remnants for Sm-Nd mineral analyses as a means of better constraining isochron age relations. Our focus first is on the isotopic systematics, and then on the elemental abundance relationships. The Kings River ophiolite Sm-Nd data are presented in Table DR1 (see footnote 1) and are plotted on a 143Nd/144Nd isotopic evolution diagram in Figure 4A. For most samples, metamorphic recrystallization prohibited mineral separate analyses, but for samples O8, O9, O11, and O12, igneous plagioclase, and for samples O9, O11, and O12, igneous clinopyroxene, these were sufficiently preserved for separation and analysis. The Sm-Nd array for the Kings River ophiolite mineral–bulk rock data define a fairly well constrained isochron age of 484 ± 18 Ma, with εNd (484) = +10.7 ± 0.4. The inclusion of clinopyroxene and plagioclase mineral data provides a considerable 147Sm/144Nd spread, facilitating age significance to the isochron, and rules out the possibility of such an array representing a mixing line. The MSWD (mean square of weighted deviates) of 3.2 indicates that scatter about the best-fit line is larger than expected on the basis of analytical errors alone, and indicates minor isotopic heterogeneities or minor disturbance during metamorphism. These results confirm the preliminary work of Shaw et al. (1987), suggesting that the Kings River ophiolite crustal section is distinctly older than the Permo-Carboniferous U-Pb zircon ages originally reported for plagiogranites and diorite mylonite of the ophiolite belt (Saleeby and Sharp, 1980). The initial εNd value of +10.7 ± 0.4 indicates derivation of the Kings River ophiolite crustal section from a source that lies near the end member composition of the depleted MORB mantle
44
J. Saleeby (Tables DR3 and DR4; see footnote 1). Major element variations as a function of SiO2 (Fig. 6A) show limited scatter, suggesting only modest alteration by seafloor and/or supra-subduction zone metamorphism. Basaltic pillows and dikes from the Kings River ophiolite plotted as CaO versus MgO indicate minimal effects of seafloor metamorphism (after Humphris and Thompson, 1978; Mottl, 1983). Alterations discussed below for mobile trace elements are thus considered to be dominated by suprasubduction zone fluid fluxing and/or Sierra Nevada batholith contact metamorphism. Focusing first on major element data, the four Kings River ophiolite pillow basalt samples that were taken from widely spaced localities are tightly grouped around SiO2 = 50% and show major element compositions that are typical of N-MORB (normal mid-oceanic-ridge basalt) except for possible K2O enrichment, although K2O contents are within the range of geographic variation patterns of modern MORB (Sun and McDonough, 1989; Hoffman, 2004). Sheeted dike and gabbro samples show more scatter, but without K2O enrichment. The gabbro compositions appear reasonably grouped, considering their widely spaced sample sites and implicit fractionated state, whereas the plagioclase-olivine clinopyroxenite and troctolite cumulates show expected dispersions from the gabbro cluster. Major element variation patterns permit a common liquid line of descent for the basalts and gabbroids (Fig. 6A). Not represented in the data, but of importance, are dunite cumulates. On Al2O3-MgO and MgO-SiO2 variation diagrams, tie lines between Fo90 olivine and the centroid of the basaltic dike data points are shown, and on Al2O3-MgO and CaO-SiO2 diagrams,
(Zindler and Hart, 1986; Hoffman, 2004). The range of initial 87 Sr/86Sr (Sri) and initial Pb isotopes determined for a subset of the Kings River ophiolite Sm-Nd samples further suggest such an end member composition for this ophiolite’s mantle source. Sri ranges over 0.7023–0.7030 (Table DR1), and for initial Pb α = 17.14–17.82, β = 15.38–15.52, and γ = 36.80–37.38 (Table DR2). The dispersion of these values is notably greater than that which is recorded by the coherence of the Sm-Nd systematics. This is evident in initial εNd-Sri, initial 206Pb/204Pb-Sri, initial 207 Pb/204Pb-206Pb/204Pb, and initial 208Pb/204Pb-206Pb/204Pb variation diagrams (Fig. 5). These plots show the Kings River ophiolite data in relation to the field of East Pacific Rise MORB (Hoffman, 2004). As discussed below, trace-element concentration data further suggest that the Sr and Pb systems have been disturbed relative to Nd, and thus the range of initial Sr and Pb values shown in Figure 5 is a range of maximum values, with the minimum values for each system taken as approximations of primary values. The extreme positions of the initial Pb fields are consistent with primitive mantle-normalized La/Sm ratios (data in Table DR3) versus initial εNd for Kings River ophiolite basalts, also plotted relative to modern Pacific MORB in Figure 5E (after Sun and McDonough, 1989; and Hoffman, 2004). The normalized La/ Sm values of <1 at high initial εNd values parallel the initial Pb data, showing the effects of pronounced long-term LILE (large ion lithophile element) depletions in the Kings River ophiolite mantle source regime. Major and trace element abundance data for the Kings River ophiolite sample suite further define a MORB association
Pb/ 204 Pb
ε(Nd)t
206
+5 0 -5
B
20 19 18
207
A +10
15.9
21
17 0.702
0.703
16
0.704
0.702
Sr/ 86 Sr
0.704
Sr/ 86 Sr
E
39 10 38
15.6 15.5
15.3
17
18
19
20
Pb/ 204 Pb
East Pacific Rise MORB (observed) Permo-Carboniferous ductile shear zones and Kaweah mélange (initial ratios @ 295 Ma)
5
37 36
15.7
206
15
ε(Nd)t
Pb/ 204 Pb 208
0.703 87
D
C
15.8
15.4
87
40
Pb/ 204 Pb
+15
Early Ordovician Kings River ophiolite (initial ratios @ 484 Ma)
0 17
18 206
19
Pb/
204
Pb
20
0.1
1
2
3 4
(La/Sm)n
Figure 5. Initial radiogenic isotopic variation diagrams for the Kings-Kaweah ophiolite belt in relation to modern Pacific MORB compositions (after Hoffman, 2004). (A) εNd-87Sr/86Sr. (B) 206Pb/204Pb-87Sr/86Sr. (C) 207Pb/204Pb-206Pb/204Pb. (D) 208Pb/204Pb-206Pb/204Pb. (E) εNd-primitive mantle normalized La/Sm (after Sun and McDonough, 1989). Data in Tables DR1, DR2, and DR3 (see footnote 1).
Geochemical mapping of the Kings-Kaweah ophiolite belt, California 22
pl
1.6
10
1.2
8
Al 2 O 3
TiO 2
FeO
20
0.8
18 16
0.4
46
50 SiO 2
0 42
54
50 SiO 2
0 42
54
20
0.15
16
54
5
20
fresh
12
pl
se al a f te lo ra or tio n
8 6 4
46
50
0 42
54
46
50 SiO 2
4 42
54
46
50
2
54
4
6
8
10 MgO
SiO 2
0.4
12
12
14
Kings River ophiolite
10
A
Pillow basalts
0.3 8
K 2O
Sheeted dikes
6
Gabbros
0.2 0.1
2 46
50
pl-ol clinopyroxenite
0 42
54
Olivine and plagioclase tie lines 46
SiO 2
50
B
Ductile shear zones and serpentinite mélange
Troctolite
4
54
Pillow basalts
Mylonitic gabbros
Diabase dikes
Mylonitic diorites
Garnet amphibolite
Plagiogranite
SiO2 16
22
2.5
20
12
2
TiO 2
Al 2 O 3
16
FeO
1.5
18
1
46
50
54 SiO 2
58
62
0 42
66
8 4
0.5
46
50
54 SiO 2
58
62
0 42
66
0.4
10
46
50
46
50
54 SiO 2
58
62
66
54
58 SiO 2
62
66
16 14
8
0.3
12
4
10
0.2
CaO
MnO
6
MgO
15
MgO
8
SiO 2
8 6
0.1
2
4
0 42
46
50
54 SiO 2
58
62
0 42
66
10
0.4
8
0.3
46
50
54 SiO 2
58
54
58
62
66
2
42
FeO
6
K 2O
Na 2 O
10
14
12
0.1
0.05
4
Na 2 O
50
CaO
MnO
8
14 42
SiO 2
10
12
0 42
46
ol
16
MgO
46
ol
15
0.2
20
0 42
4
CaO
12 42
pl 20
6
2
14
45
Al 2 O 3
24
4
0.1
2 0
0.2
42
46
50
54 SiO 2
58
62
66
0
42
46
50
62
66
K2O + Na2O
MgO
SiO 2
Figure 6. Major element variation diagrams for the Kings-Kaweah ophiolite belt. (A) Kings River ophiolite samples. (B) Ophiolitic ductile shear zones and Kaweah serpentinite mélange samples. Data in Tables DR4 and DR6 (see footnote 1). Fresh MORB and seafloor alteration fields in A, after Humphris and Thompson (1978) and Mottl (1983).
46
J. Saleeby
tie lines between An90 plagioclase and the centroid are shown. Perhaps most definitive are the Al2O3-MgO relations, which suggest a parental liquid composition similar to the dikes, with plagioclase and olivine fractionation rendering the cumulates and pillow basalts. The potential role of clinopyroxene fractionation is unclear, and of possible second order importance, although concentration of clinopyroxene in cumulate gabbros could have slightly depleted the residual melt in SiO2 as observed in the pillow basalts. Enrichments of Mn, Ti, Fe, and possibly K in the pillow basalt data can also be explained by minor fractionation of a sheeted dike-like parental liquid. Trace element data for the pillow basalts and sheeted dikes of the Kings River ophiolite are normalized to N-MORB in Figure 7A (after Sun and McDonough, 1989). The general displacement of the pillow basalt data array from the sheeted dike array parallels the fractionation pattern suggested by the major element data. Also shown in the plot is the compositional trend of E-MORB (enriched mid-ocean-ridge basalt) relative to N-MORB (normal mid-oceanic-ridge basalt) (after Sun and McDonough, 1989). Much of the Kings River ophiolite basalt data fall into the field bounded by N-MORB and E-MORB, although distinct positive spikes occur for Ba, Pb, and Sr, and, for some samples, Cs and K. These trace elements are known to have considerable compositional variation with geographic position in modern MORB (Hoffman, 2004), although the spread in the Figure 7A data is greater than that observed and indicates that these trace elements have undergone at least some enrichments relative to MORB. In that CaO-MgO relations in the basalts show no evidence for significant seafloor metamorphic alteration (Fig. 6A), such enrichment is interpreted to have occurred in the supra-subduction zone environment by a combination of fluid fluxing through the mantle wedge and/or more proximally
A
100
E - MORB 10
1
0.1
B E - MORB
N - MORB normalized
N - MORB normalized
100
driven Sierra Nevada batholith contact metamorphism. The trace elements of Cs, Ba, K, Pb, and Sr are typically enriched in suprasubduction zone fluids (cf. Pearce et al., 1995). In contrast, Nb of the Kings River ophiolite basalts has retained coherency with the MORB pattern, consistent with its negligible mobility in supra-subduction zone fluids (Pearce et al., 1995). The Figure 7A data also show that the rare earth elements (REEs) have not been enriched by such fluid interactions, which is consistent with the minor mobility of Nd in supra-subduction zone fluids, lending further confidence to the Sm-Nd isochron age and initial εNd values of Figure 4A. The range of initial Sr and Pb isotopic ratios determined for the Kings River ophiolite, in light of the coherency of the REE data, are interpreted to reflect differential alteration of the samples by the influx of exotic Sr and Pb by supra-subduction zone fluid fluxing and/or Sierra Nevada batholith contact metamorphism. Figure 8 shows Rb/Sr and 235U/207Pb evolution diagrams for samples for which initial isotopic ratios were determined. The samples consist of bulk rock data for basalts and gabbros, and plagioclase data for two coarse-grained gabbros exhibiting only grain-boundary metamorphic overprints. The Rb/Sr and 235 U/207Pb systems were chosen for this analysis because that during alteration, potential parent nuclide enrichments for these systems are minimal relative to daughter nuclide enrichments, in contrast to the 238U/206Pb and 232Th/208Pb systems. In each diagram the plagioclase data plot closest to the origin and yield the lowest initial ratios, calculated for 484 Ma, based on the Sm-Nd isochron age (Fig. 4A). A 484 Ma reference isochron is passed through the plagioclase data points in Figures 8A and 8B, and on each plot the isochron also passes through the O7 gabbro sample, which shows little or no enrichment of Sr or Pb relative to N-MORB. Comparison of trace element abundances of
10
1
0.1 Cs Rb Ba Th U Nb K La Ce Pb Pr Sr P Nd Zr Sm Eu Ti Dy Y Yb Lu
Cs Rb Ba Th U Nb K La Ce Pb Pr Sr P Nd Zr Sm Eu Ti Dy Y Yb Lu
Figure 7. Trace element normalization plots for basaltic rocks of the Kings-Kaweah ophiolite belt normalized to idealized N-MORB and in comparison with idealized E-MORB (after Sun and McDonough, 1989). (A) Kings River ophiolite basaltic pillow lavas and sheeted dikes. (B) Basaltic pillow lavas and diabase dikes related to the Kaweah serpentinite mélange. Data in Table DR3 (see footnote 1). Symbols are the same as in Figure 6.
Geochemical mapping of the Kings-Kaweah ophiolite belt, California
and Tilton, 1991) show a range of radiogenic compositions for which similar composition fluids could have profound alteration effects superposed over the relatively non-radiogenic ratios of the Kings River ophiolite. The Figure 8 relationships also lead to the interpretation that the plagioclase and/or the most primitive gabbro initial isotopic ratios most accurately reflect those of the Kings River ophiolite mafic crustal section, indicating Sri = 0.70225, and for Pb α = 17.139, β = 15.384, and γ = 37.384. In Figure 5 it is clear that these initial Sr and Pb isotopes lie at, or even beyond, the extreme long-time integrated LILE depletions typical of the Pacific MORB mantle source regime. In Figure 7A it is clear that the removal of the secondary spikes for Cs, Ba, K, Pb, and Sr results in an N-MORB to slightly enriched MORB profile for the Kings River ophiolite basalts. The major and trace element abundance data in conjunction with Nd, Sr, and Pb isotopic data clearly point to an N-MORB, or perhaps a slightly enriched MORB, source for the Kings River ophiolite. Reconstructed stratigraphic thicknesses for the various lithologic zones of the ophiolite agree well with the seismic velocity structure of “normal” abyssal crust. The widespread occurrence of sheeted basaltic dikes, virtually devoid of significantly fractionated members, is also consistent with an abyssal ridge-crest origin for the ophiolite (cf. Robinson et al., 2008). Based on geochemical data, lithostratigraphic relations, and
O7 with those of N-MORB is provisional in that O7 is a cumulate of probable MORB derivation, and depletions of LIL trace elements relative to erupted MORB is expected on the basis of crystal-liquid elemental partitioning and the implicit fractionated state of a cumulate. Nevertheless, in comparison with the O9 cumulate gabbro data point in each diagram, enrichments in Sr and Pb correlate with divergence from the 484 Ma isochron. This is shown to the right of the two isotopic evolution diagrams by plots of δ, defined as the divergence of given 87Sr/86Sr and 207 Pb/204Pb ratios from the respective reference isochron versus (Sr or Pb) ENM, defined as the respective elemental enrichments relative to N-MORB (after Sun and McDonough, 1989). In the δ versus ENM plots the position of the O9 gabbro is shown relative to N-MORB by the light cross, and relative to Sr and Pb concentrations of sample O7 by the darker cross. This is done in order to help see through the effects of O7 and O9 as cumulates. There is clearly a positive correlation between δ and ENM, and such a correlation, albeit noisy, is indicated for the basaltic samples as well. The correlation between the divergence of the initial isotopic ratios and the elemental enrichments of Sr and Pb argue strongly for a secondary origin for the enrichments of Sr and Pb, relative to N-MORB, and by analogy the trace element enrichments of Cs, Ba, and possibly K in Figure 7A. Lead and Sr isotopic data for the Sierra Nevada batholith (cf. Chen
A
0.7032
47
8 7
0.7030
87
Sr / 86 Sr
6 0.7028 5
δ
0.7026
87 86
Figure 8. Isotopic evolution diagrams for 87Sr/86Sr (A) and 207Pb/204Pb (B), each showing 484 Ma reference isochron (after Fig. 4A) forced through samples with most primitive initial ratios at t (time) = 484 Ma. Symbols are same as in Figure 6A, but with an encircled p symbol denoting gabbro plagioclase separates. Shown to the right of each isotopic evolution diagram are plots of the divergences of the respective isotopic ratios off of reference isochrons (δ) versus respective elemental enrichments (ENM) of samples relative to N-MORB (after Sun and McDonough, 1989). These plots show that secondary Sr and Pb enrichments correlate with disruption of primary initial isotopic ratios. Data in Tables DR1, DR2, DR3, and DR4 (see footnote 1).
4 3
0.7024 0.7022
x 10-4
484 Ma 2
P P
0.7020 0
∼2σ: 0.002
0.004
87
0.006
Rb /
86
1
0.008
NM
0
0.010
0
1
2
3
Sr E NM
Sr 1.6
B 15.65
1.4
δ 15.45
207
Pb / 204 Pb
1.2 15.55
484 Ma P
0.6
P
0.4
15.35
15.25
1
x 10-1 207 0.8 204
∼2σ:
0.2
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0 0
0.02
0.04
235
U/
0.06
204
Pb
0.08
0.10
0
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4
6
Pb E NM
8
48
J. Saleeby
petrography, the Kings River ophiolite is the clearest example of a complete abyssal crust and upper mantle lithosphere sequence preserved in the North American Cordillera. We now focus on a distinctly younger abyssal igneous event resolved within the Kings River ophiolite, and its relationships with the development of the Kaweah serpentinite mélange. Permo-Carboniferous Ductile Shear Zones and Kaweah Serpentinite Mélange The Yokohl Valley area (Fig. 3) was chosen for the principal analysis of the Kaweah serpentinite mélange because of its relatively low thermal metamorphic overprint, as compared with all other areas of the Kings-Kaweah ophiolite belt, and for the preservation of a number of important primary relationships. The Yokohl Valley map area is unique to the Kings-Kaweah ophiolite belt in that a textural (±index mineral) “isograd” may be resolved, which corresponds to what is defined here as inner versus outer contact metamorphic aureole positions. The inner aureole position is defined by near-pervasive actinolitic hornblende + sodicplagioclase ± clinozoisite or epidote static overprinting in mafic assemblages, black-wall skarn domains and vein networks, and nematoblastic antigorite overprints in ultramafic assemblages, and the occurrence of andalusite ± sillimanite in cover strata pelites. Outer aureole positions are defined by a lack of these indices and also the widespread preservation of primary igneous phases and/or pre–Sierra Nevada batholith metamorphic assemblages. As displayed in Figure 3, a large central area lies within an outer aureole position, encompassing a number of important primary relationships. Primary relationships that are relatively well preserved in the Yokohl Valley map area include the intrusion and eruption of abyssal igneous rocks through disrupted and actively shearing Kings River ophiolite–related abyssal lithosphere, pelagic sedimentation across the resultant polygenetic basement complex, and nonconformable overlap by hemipelagic strata of the Calaveras complex and in sequence siliciclastic and volcanogenic strata of the Kings sequence. This area also displays the internal complexity of the Kaweah serpentinite mélange, as it covers three well-defined mélange units (Fig. 1) characterized by assemblages derived as tectonic blocks from the Kings River ophiolite, variably deformed intrusive bodies emplaced into serpentinized peridotites that had ascended to crustal levels, and surfaced serpentinites that mingled with pillowed eruptions and pelagic sediments. Pillow basalts sampled from the Kaweah serpentinite mélange (samples M1 and M2) were taken from the NE-SW–trending belt of interbedded basalt and chert that extends across the center of the Yokohl Valley area (Fig. 3). This partially disrupted stratified belt was deposited on a serpentinite substrate. It contrasts in field setting from the thick homogeneous pillow basalt sections of the Hughes, Dalton-Bald, and Antelope Mountain slabs of the Kings River ophiolite by the intercalation of radiolarian chert, hyaloclastite, and ophicalcite, and the absence of an underly-
ing sheeted dike complex. All pillowed flows of the observed Kaweah serpentinite mélange are aphyric. Local, isolated feeder dikes for the pillowed flows are typically highly altered, having interacted with serpentinite host rocks, and most have been broken into relatively non-deformed, angular “mélange blocks.” Some feeder dikes are pepperitic, suggesting that they were emplaced into non-lithified sedimentary serpentinite (cf. Kokelaar, 1982). The Kaweah serpentinite mélange–related mafic dikes grade in texture from aphyric basalts to coarser diabases, microgabbros, and gabbros, with the coarser varieties not having been observed to be directly linked to pillowed basalts. Where not highly recrystallized from contact metamorphism the coarser rocks appear distinct from intrusive rocks of the Kings River ophiolite by the common occurrence of subordinate brown igneous hornblende along with clinopyroxene. Solitary and nonsheeted sets of igneous hornblende–bearing diabases are widespread along the Kings-Kaweah ophiolite belt, cutting mafic and ultramafic levels of the Kings River ophiolite, dispersed as “mélange blocks” and within intrusive complex blocks within the Kaweah serpentinite mélange. One such dike, with the best remnants of an igneous texture observed (sample M3), was taken directly adjacent to the margin of the longitudinal ductile shear zone that cuts the eastern margin of the Tivy Mountain gabbro (Fig. 2). This dike was chosen for its texture and its field setting of lying proximal to a shear zone that variably transposes dikes of similar composition. More extensive sampling was performed on gabbroic mylonites and their dioritic fractionates that occur within the ophiolitic ductile shear zones. Based on contradictory intrusive and mylonitization relationships, igneous emplacement within the shear zones is shown to have been coincident with hightemperature ductile shearing (Saleeby 1977, 1978). Because ductile shearing instigated tectonic disruption of the Kings-Kaweah ophiolite belt, the age of such shearing places important time constraints on the initial phases of serpentinite mélange development. The gabbroic and dioritic intrusives of the shear zones are distinct from the thick cumulate sections of Tivy and Smith Mountains, which are characterized by intermediate to calcic plagioclase and coarse diopsidic augite with local olivine. The mylonitic gabbros of the ductile shear zones are characterized by a similar compositional range in plagioclase and clinopyroxene, but they commonly contain brown hornblende that is igneous in origin and that appeared early in the crystallization sequence. Protolith compositions of the deformed intrusive bodies include hornblende ± plagioclase-clinopyroxenite, hornblende-clinopyroxene gabbro, and anorthositic gabbro, and small volumes of hornblende ± quartz diorite. Large clinopyroxene and hornblende grains commonly form porphyroclasts in a finely laminated blastomylonitic matrix, locally with millimeter to decimeter scale dioritic veins that are variably transposed. Olivine bearing gabbroids have not been recognized as part of this distinctive suite of ophiolitic intrusives. Sampling of the shear zone gabbros and diorites focused on the upper Hughes Creek slab (samples M4a and M4b), and samples M5a, M5b, and M5c from the Kings River ophiolite basal
Geochemical mapping of the Kings-Kaweah ophiolite belt, California
with substantial Cretaceous thermal disturbance (Saleeby and Sharp, 1980). Zircon was extracted from the sample M4b and M6b mylonitic diorites and re-extracted from the M5c mylonitic diorite and M6a plagiogranite, and analyzed in conjunction with this study (Table DR5 [see footnote 1]). These samples were subjected to stepwise chemical leaching procedures similar to those used in Mattinson (1994) in order to remove labile components of U and Pb (GSA Data Repository1). The new data are plotted along with the original data for samples M5c and M6a in the Figure 9 concordia diagram. The new data form a tight array adjacent to the original published upper intercept, reflecting the removal of highly labile discordance bearing U and Pb components. The two original samples included in this analysis yield much lower precisions, reflecting improved analytical techniques, and cascade down concordia considerably reflecting the isotopic signature of their labile components. The upper intercept age of 296 ± 15 Ma for the entire suite corresponds well with the isochron “age” of the Kaweah serpentinite mélange sample suite, which is scaled onto concordia in Figure 9. Also scaled onto concordia are K/Ar hornblende ages from Kings-Kaweah ophiolite belt basites, differentiated as lying in inner versus outer contact-aureole field positions, and concordant U/Pb zircon ages from dike swarms and Sierra Nevada batholith plutons that crosscut and contact metamorphosed the Kings-Kaweah ophiolite belt. The K/Ar ages from inner contact aureoles are completely reset to proximal
0.0525
0.0515
7a 6b 4b 5c
Intercepts at 296 +15 and 117 + 13 Ma MSWD = 15, 2σ errors Sm-Nd isochron age (M suite)
280
Hornblende K/Ar ages outer aureole inner aureole Concordant U/Pb zircon ages SNB plutons Dike swarms
Pb/ 206Pb
7a
207
shear zone along the lower Kings River (Fig. 2). Vestigial lenses and blocks of serpentinized mylonitic peridotite with mylonitic gabbros are dispersed through the Kaweah serpentinite mélange, the largest of which recognized is the Lindsay Ridge slab in Figure 3. Samples M6a and M6b were taken from mylonitic gabbro and diorite of this slab. Serpentinized peridotite also forms the host for shallowlevel mafic intrusive swarms that are not penetratively sheared along with their hosts. These intrusives consist of basaltic dike rock, hornblende-clinopyroxene diabase and static textured gabbro, and rare pegmatitic hornblende gabbro and plagiogranite. Sample M7a is from a plagiogranite body, and sample M7b is from a set of diabase dikes that crosscut the plagiogranite and its gabbroic host. The plagiogranite and diabase dikes are subordinate members of vari-textured gabbroids that form a series of composite intrusive sheets lying within an antigorite ± talc ± tremolite schist matrix within the Yokohl Valley area (Fig. 3). The Elephant Back slab of the Yokohl Valley area (Fig. 3) consists of composite hornblende-clinopyroxene gabbro and microgabbro intrusive bodies that are marginally sheared along with serpentinized peridotite host rocks. Local pegmatitic hornblende gabbro segregations are well preserved within the Elephant Back intrusive bodies, one of which yielded a 250 Ma K/Ar hornblende age (Saleeby and Sharp, 1980) and which is included here as sample M8. Figure 4B is a 143Nd/144Nd isotopic evolution diagram for the entire Kaweah serpentinite mélange suite of gabbros, diorites, basalts, diabases, and plagiogranites, including those that intrude the Kings River ophiolite. All of the data on the neodymium evolution diagram are bulk rock, except for the coarse hornblende from the pegmatite and its partially recrystallized plagioclase matrix. The widespread tectonitic fabrics and near pervasive metamorphic overprints of the Kaweah serpentinite mélange sample suite otherwise prohibit mineral separate analyses. The Figure 4B data define an apparent linear array corresponding to an isochron age of 299 ± 32 Ma with εNd (299) = +10.0 ± 0.7. The isochron age is consistent with U/Pb zircon ages on plagiogranites and the mylonitic diorites discussed further below, although the scatter about the isochron line is considerable (MSWD = 6.8). The scatter could reflect small disturbances related to metamorphism and/or supra-subduction zone fluid fluxing, or to minor variations in the initial isotopic compositions of the various samples. Some component of the later mechanism seems likely, considering the large array of field settings for the samples compounded by the large area over which they were taken. The isochron age is thus considered a rough approximation for the igneous age of the suite. One of the essential features of the Sm-Nd results for the Kaweah serpentinite mélange suite of samples is the correspondence of the isochron “age” with U/Pb zircon ages on plagiogranites and mylonitic diorites. As mentioned above, published zircon ages for the three dispersed plagiogranite-bearing blocks from the Kaweah serpentinite mélange, and the sample M5c mylonitic diorite, suggest a Permo-Carboniferous igneous age
49
240 0.0505
5c
200 0.0495
160 0.0485
120 0.0475 15
25
35
45
55
65
75
238
U/ 206Pb
Figure 9. Concordia diagram based on 207Pb/206Pb versus 238U/206Pb (after Terra and Wasserburg, 1972) for plagiogranite and mylonitic diorites of ophiolitic ductile shear zones and Kaweah serpentinite mélange (data in Table DR5; see footnote 1). Also plotted along concordia are concordant U/Pb zircon ages from dike swarms and batholithic plutons that crosscut and contact metamorphosed the ophiolite belt, K/Ar hornblende ages from ophiolite-belt mafic rocks differentiated as from inner versus outer contact aureole positions (after Saleeby and Sharp, 1980; Clemens-Knott and Saleeby, 1999; Saleeby and Dunne, 2011), and M suite Sm-Nd isochron age with a 2σ error envelope (after Fig. 4B). MSWD—mean square of weighted deviates; SNB—Sierra Nevada batholith.
50
J. Saleeby
Sierra Nevada batholith ages, whereas those from outer aureole positions are partially disturbed. The 117 ± 13 Ma lower intercept for the plagiogranite and dioritic mylonites clearly corresponds with the thermal metamorphic maxima as reflected in the Sierra Nevada batholith U/Pb zircon and reset K/Ar hornblende ages. Scatter in the more discordant ophiolitic zircon data points, as well as in the data for the other plagiogranite zircon populations not presented here (Saleeby and Sharp, 1980), are interpreted to reflect multiple pulses of thermal overprinting, which began in the Jurassic and intensified into the Early Cretaceous. Considering both the Sm/Nd isochron age and the U/Pb zircon ages, and the overlap of their uncertainties, ca. 295 Ma is considered a reasonable approximation for the age of the Kaweah serpentinite mélange suite, and is used for initial Pb and Sr isotopic composition corrections discussed below. The initial εNd value of +10.0 ± 0.7, derived above for the Kaweah serpentinite mélange suite, also indicates derivation from a source that lies near the end member composition of the depleted MORB mantle (Hoffman, 2004). Moreover, Sri from the Kaweah serpentinite mélange suite ranges over 0.7027–0.7033 (Table DR1), with initial εNd versus Sri for the suite lying within the field of modern Pacific basin MORB (Fig. 5A) except for the highest Sri value, which, as discussed below, is considered to have been altered in the supra-subduction zone environment. Initial Pb values for the Kaweah serpentinite mélange suite are notably more radiogenic than for the Kings River ophiolite suite, with α = 17.962–18.493, β = 15.511–15.600, and γ = 37.663– 38.211 (Table DR2). Nevertheless, the Kaweah serpentinite mélange suite Pb values plot within the field of Pacific MORB on initial 206Pb/204Pb-Sri, 207Pb/204Pb-206Pb/204Pb, and 208Pb/204Pb -206Pb/204Pb variation diagrams (Fig. 5). As discussed below, trace element concentration data further suggest that the Sr and Pb systems have been disturbed relative to Nd, and thus the range of Sr and Pb values shown in Figure 5 is a range of maximum values as with the range of Kings River ophiolite values. As with the Kings River ophiolite basaltic samples, the Kaweah serpentinite mélange basalts yield a primitive mantle normalized La/Sm versus initial εNd field that lies toward the pronounced long-term LILE depletion domain of the field for modern Pacific MORB (Fig. 5E). Major and select trace-element abundance data are presented for the Kaweah serpentinite mélange sample suite in Table DR6 (see footnote 1), except for the extremely coarse-grained hornblende pegmatite. Major element variations as a function of SiO2 (Fig. 6B) show a distinctly wider range than for the Kings River ophiolite sample suite. The two pillow basalt samples (M1 and M2) are at SiO2 = 49–50 wt% and show major element compositions that are typical of N-MORB (Sun and McDonough, 1989). The fact that all of the gabbros analyzed are from scattered, relatively small intrusive bodies that were emplaced into serpentinized peridotite leads to the expectation that coherent fractionation trends would not be evident in the overall majorelement variation patterns, and that a common liquid line of descent is highly unlikely. In aggregate, however, the variation
of Na2O, K2O, and CaO with SiO2, and the array on the AFM plot, point to local silica-alkali enrichment. On the AFM plot, tie lines are shown between proximally related mylonitic gabbros and diorites. The plagiogranite and dioritic samples show Na2O enrichment relative to K2O, typical of oceanic felsic associations (Coleman and Peterman, 1975). The low-volume diorites and plagiogranites appear to represent local closed-system fractionation products derived from adjacent gabbroids. Trace element data for the M1 and M2 basaltic pillows and the M3 and M7b diabase dikes (Table DR3) are normalized to N-MORB in Figure 7B (after Sun and McDonough, 1989). Also shown on the plot is the compositional trend of E-MORB (after Sun and McDonough, 1989). The Kaweah serpentinite mélange data are similar to the Kings River ophiolite basalt data, mainly in the field bounded by N-MORB and E-MORB, and some samples show distinct positive spikes for Cs, Ba, Pb, and Sr. The Figure 7B data also show that REEs for the Kaweah serpentinite mélange basalts were not influenced significantly by such mobilization, lending further confidence to the Sm/Nd isochron age and initial εNd values derived in Figure 4B. The similarity of some positive spikes for Cs, Ba, Pb, and Sr to the Kings River ophiolite data suggests that the implicit enrichments in these elements occurred in the common supra-subduction zone fluid fluxing and/ or Sierra Nevada batholith contact metamorphic environment that the Kaweah serpentinite mélange and Kings River ophiolite underwent. As with the Kings River ophiolite, Pb has clearly been mobilized in the bulk rocks during contact metamorphism, considering that radiogenic Pb is partially disturbed in the plagiogranite and mylonitic diorite zircon populations (Fig. 9). As with the Kings River ophiolite suite, the Sri and initial Pb values for the Kaweah serpentinite mélange suite are interpreted as a range of maximum values, with the least radiogenic ratios conceivably approximating true initial values. This leads to the interpretation of the Kaweah serpentinite mélange suite having an approximate primary isotopic composition of Sri = 0.7027, and for initial Pb α = 17.962, β = 15.511, and γ = 37.663. The radiogenic isotopic data, as well as the major and trace element abundance data, indicate that the Kaweah serpentinite mélange suite of intrusions and pillow basalts is also of MORB affinity. Field relations and the age data indicate that these PermoCarboniferous MORB affinity magmas were emplaced and erupted into the Early Ordovician Kings River ophiolite abyssal crust and mantle lithosphere during high-magnitude shear strain of the hosting lithosphere, and its progressive disruption to form serpentinite mélange on the seafloor. Basaltic magmas of the Permo-Carboniferous suite erupted across the serpentinitic substrate as mélange was forming, and such lavas were interbedded with sedimentary serpentinite, ophicalcite, and radiolarian ooze. The widespread stabilization of early igneous hornblende in many of the Kaweah serpentinite mélange intrusives is consistent with their emplacement along a major transform fracture zone (Melson et al., 1972; Engle and Fisher, 1975; Bonatti and Honnorez, 1976; Honnorez et al., 1984; Johnson and Dick, 1992; Schroeder and John, 2004; Cipriani et al., 2009). One possibility
Geochemical mapping of the Kings-Kaweah ophiolite belt, California is that water was sequestered out of the variably serpentinized peridotites that hosted the intrusive masses. Geochemical studies and modeling of partially serpentinized peridotite of the Feather River massif of the northern Sierra Foothills, a likely correlative of the Kings-Kaweah ophiolite belt ultramafics and mafic tectonites as discussed below, suggest that the respective abyssal serpentinization front reached ~40-km depths, possibly along a transform fault (Li and Lee, 2006). Given a MORB magma extraction depth of 30–50 km (Lee et al., 2009), the Kaweah serpentinite mélange suite of intrusives conceivably ascended through ≥30 km of partially hydrated ultramafic wall rocks prior to stalling out within ductile shear zones, tensile fractures, or relatively small plutons within the disrupted Kings River ophiolite abyssal crust and upper mantle section. This would yield ample opportunity for minor sequestering of water and the stabilization of early igneous hornblende. The ~190 m.y. hiatus in abyssal MORB magmatism indicated by the age data presented above is profound relative to abyssal magmatic patterns of the modern ocean basins, and requires attention. Polygenetic Abyssal Lithosphere The Kings River ophiolite consists of a complete abyssal spreading-ridge crust and upper mantle sequence of Early Ordovician age. The homogeneity of this mafic crustal section, extensiveness of sheeted dikes, and lack of differentiated rocks suggests spreading ridge genesis above a steady-state magma chamber, which further suggests genesis along a fast-spreading ridge like the East Pacific Rise (i.e., Babcock et al., 1998). The coherency of the Sm-Nd systematics and uniformity of bulk compositions for the Kings River ophiolite sample suite suggest a common source and liquid line of descent for the principal slabs and mélange blocks. The along-strike extent of voluminous Kings River ophiolite mafic rocks, minus small transverse zones of crosscutting Sierra Nevada batholith intrusions, is ~50 km (Fig. 1). Geochemical studies of the East Pacific Rise show that distinct domains of magma source regime and liquid line of descent can exceed 40 km along strike (Reynolds et al., 1992). Thus the principal mafic slabs of the Kings River ophiolite and their derivative mélange blocks are interpreted as the products of one source, ascent, and crystallization domain along a fastspreading ridge in the early Paleozoic Panthalassa basin. Pelagic sedimentation across the Kings River ophiolite abyssal basement is feebly recorded in small metalliferous radiolarian chert lenses within the pillow basalt section, and as dispersed blocks of similar metalliferous chert within the Kaweah serpentinite mélange. The greatest stratigraphic thicknesses measured for such metalliferous chert blocks are typically ~30 m. Such blocks are devoid of argillaceous, serpentinitic lute or tuffaceous components, which contrasts with local thin laminae present in thicker chert sections that lie above the Permo-Carboniferous pillow basalts of the Kaweah serpentinite mélange. Sedimentation rates for radiolarian oozes in the early Paleozoic are not constrained, and in conjunction with the disruption of the Kings River ophio-
51
lite chert section into dispersed mélange blocks this ophiolite’s pelagic sedimentation history is obscured. This has resulted in an ~190 m.y. hiatus in the geologic history recorded in the KingsKaweah ophiolite belt. The Early Ordovician abyssal lithosphere of the Kings River ophiolite underwent an intense phase of tectonic disruption in the Permo-Carboniferous. The association of high-temperature ductile shearing coincident with MORB magmatism, the surfacing of variably serpentinized depleted peridotites on the seafloor, and the mingling of serpentinitic detritus with pelagic deposits are typical of the intersection zones of slow-spreading ridges and large offset transforms (cf. Bonatti et al., 1973; Karson and Dick, 1983; Johnson and Dick, 1992; Cannat, 1996; Rommevaux-Jestin et al., 1997; Schroeder and John, 2004; Cipriani et al., 2009). The profound relationship recorded for the Kings-Kaweah ophiolite belt, however, is that this distinctive deformational and magmatic regime was superposed across MORB igneous crust and abyssal mantle lithosphere after the ~190 m.y. hiatus of geologic history. This seems profound, given that the oldest known abyssal lithosphere currently residing in an ocean basin is ca. 168 Ma (Ogg and Smith, 2004). The potential notion that the Permo-Carboniferous petrotectonic regime occurred in a supra-subduction zone environment is rejected on the basis of the pelagic and hemipelagic sedimentation history of the Kaweah serpentinite mélange as well as the geochemical data presented above. The thickest radiolarian chert sections that lie depositionally on Permo-Carboniferous pillow basalts are in several areas ~100 m thick (cf. Fig. 3). Pelagic sedimentation rates for radiolarian oozes in the late Paleozoic are somewhat better constrained than for the early Paleozoic. For example, a post compaction rate for radiolarian cherts studied in the Permian Dalong Formation of southwest China are ~0.4 cm/103 yr (Gu et al., 2007). Such a rate, if applicable, would imply an ~30 m.y. accumulation of predominantly clean radiolarian chert commencing in the Permo-Carboniferous, or ca. 295 Ma, prior to the onset of Calaveras hemipelagic sedimentation. Calaveras chert-argillite does not appear to have much, if any, volcanic input, and as discussed below the products of suprasubduction zone volcanism were not expressed in the cover strata until well into Middle Triassic time, at the earliest. CONDITIONS AND SETTINGS OF METAMORPHISM, DEFORMATION, AND MÉLANGE MIXING Petrographic data and map relations for the Kings River ophiolite indicate that geographically focused thermal metamorphism and cleavage development related to Jurassic plutonism and tectonism, and regional thermal metamorphism related to Early Cretaceous Sierra Nevada batholith plutonism, were superimposed on (1) regional low to medial grade static textures within the coherent mafic crustal section of the Kings River ophiolite and its derivative mélange blocks, (2) hightemperature upper-mantle flow fabrics of the peridotite section of this ophiolite and its derivative mélange blocks, (3) mylonitic fabrics of the Permo-Carboniferous ductile shear zones, and
52
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(4) matrix schistosity in serpentinite mélange zones. Similar textural relations are recorded for the Kaweah serpentinite mélange with thermal overprints affecting matrix schistosity as well as regional low to medial grade assemblages or mylonitic fabrics within mélange blocks. Rocks of the Kings-Kaweah ophiolite belt did not undergo high P-T metamorphic conditions, nor was this ophiolite belt subjected to intermediate- to high-pressure conditions during Sierra Nevada batholith metamorphism. Petrologic constraints on depth of emplacement for Sierra Nevada batholith plutons that intrude the Kings-Kaweah ophiolite belt indicate relatively shallow levels (Putman and Alfors, 1969; Ague and Brimhall, 1988; Liggett, 1990; Clemens-Knott et al., 2011), consistent with the predominance of andalusite and only local andalusite + sillimanite in meta-pelites of the cover strata. The regional low-grade conditions are most readily attributed to a combination of ocean floor metamorphism and subsequent burial metamorphism and cleavage development beneath lower Mesozoic cover strata. Widespread mylonitization and tectonic disruption clearly occurred during the invasion of the Early Ordovician abyssal lithosphere by the Permo-Carboniferous abyssal magma series, and are not related to subduction megathrust tectonics. The association of high-strain mylonitic gabbros and peridotites with serpentinite diapirs and extrusions, sedimentary serpentinite, and ophicalcite is typical of the intersections of slowspreading ridge segments with large transform offsets on the ocean floor (Aumento et al., 1971; Melson and Thompson, 1971; Melson et al., 1972; Bonatti et al., 1971, 1973; Bonatti and Honnorez, 1976; Johnson and Dick, 1992). Metamorphic tectonites that developed in such environments typically exhibit deformation proceeding in a retrogressing P-T regime, commencing with basaltic solidus conditions in parallel with such deformational conditions as are recorded in the Kings-Kaweah ophiolite belt. Submarine studies of ridge-transform intersection environments at slow- to intermediate-rate spreading ridges also indicate that deep mafic crustal and upper mantle rocks are exposed as metamorphic core complexes, particularly when magma production rates lagged behind spreading rates (CAYTROUGH, 1978; Cannat, 1996; Baines et al., 2003; Schroeder and John, 2004; Ildefonse et al., 2007; Tucholke et al., 2008). The sub-oceanic Moho transition zone of the Kings River ophiolite possesses intense structurally concordant constrictional deformation fabrics developed during relatively low flux Permo-Carboniferous MORB magma injection that rendered structural and fabric relations like those developed by upper mantle corner flow off of spreading axis shoulders beneath oceanic core complexes (cf. Schroeder et al., 2007). Mylonitic fabrics developed in the Kings River ophiolite in this structural setting merge into and are cut by steep mylonitic fabrics of the longitudinal shear zones, which are interpreted to be transform shears, or ridge crest discontinuity shears, that evolved into major strike-slip fault and serpentinite mélange zones during large-magnitude transform displacement of the Kings-Kaweah ophiolite belt. The upper age boundary for significant mélange mixing is provided by the nonconformable overlap relations of lower Meso-
zoic volcaniclastic and siliciclastic strata, as discussed below. Such nonconformable overlap occurred in a supra-subduction zone environment within the SW Cordillera. The age and structural setting of rare high-pressure mafic metamorphic blocks that occur within the Kaweah serpentinite mélange further constrain the tectonic events that occurred between mélange development in the abyssal realm, and depositional overlap by the active margin strata. Cryptic High-Pressure Metamorphism As discussed above, the Kings River ophiolite, as well as crustal blocks within the Kaweah serpentinite mélange, record metamorphic and deformational processes of the abyssal realm with the superposing of regional medial-grade burial conditions and thermal metamorphic overprints acquired in its subsequent supra-subduction zone setting. Two monolithologic mélange blocks of 5-m scale have been discovered that contrast with this pattern, demonstrating the existence of a once underlying high-P metamorphic environment. These two blocks are in mélange units from the southern segment of the ophiolite belt and occur in matrices for which textures and structures that suggest a diapiric and/or detrital origin are relatively well preserved. These rare high-P metamorphic blocks are clearly exotic to the KingsKaweah ophiolite belt relative to all other metabasites. Sample M9 is from one of the two garnet amphibolite blocks discovered, located in the Yokhol Valley area (Fig. 3). The field setting of the block is such that it lies in an area where serpentinite diapirism is shown to have been polyphase, as evidenced by serpentinite “dikes” cutting the cover strata and feeding sedimentary serpentinites within the strata. These relations can be visualized by orienting Figure 3 in such a way as to place the clear arrow positioned along the left margin of the figure at the bottom facing up. In this orientation the folded nonconformity between the Kaweah serpentinite mélange and its cover strata may be viewed as down-plunge. A vertical dike-like body of remobilized sedimentary serpentinite penetrates the nonconformity and cuts into the cover strata along the Harvey Ruth fault zone. The “dike” connects to multiple sedimentary serpentinite-ophiolitic debris flow beds within the cover strata, indicating that it was a growth structure. To the southwest, another dike-like serpentinite body extends southward into the cover strata, with its southeast terminus covered by Quaternary sediments of Round Valley. The M9 garnet amphibolite block is situated in the apparent core area of this “dike.” The M9 garnet amphibolite block records a deformationmetamorphic regime that is distinct from that which characterizes the Kings-Kaweah ophiolite belt. The sample possesses a simple mineralogy and a well-preserved crystalloblastic texture. Hornblende is highly annealed and is embayed by euhedral garnet porphyroblasts. Zoning is not present in the hornblende or garnet grains. Small rutile, albite, and apatite grains are dispersed throughout the hornblende-rich matrix. Fine inclusions within garnet porphyroblasts consist of quartz, augite, and
Geochemical mapping of the Kings-Kaweah ophiolite belt, California
±1 8M a
high-P metamorphic derivative of an oceanic ferrogabbro (Mottana and Bocchio, 1975). Trace-element variation patterns of Ti, V, Zr, Cr, and Y for sample M9 (Fig. 11; data in Table DR6) are most consistent with a MORB derivation. Based on Sm-Nd systematics, derivation from Kings River ophiolite–like or Kaweah serpentinite mélange–like abyssal crustal rocks appears unlikely on the basis of the topologic relations of the respective isochrons (Fig. 10), unless sample M9 has undergone significant alteration in the metamorphic environment. Neodymium is enriched in the sample, compared with all other Kings-Kaweah ophiolite belt basalts and gabbroids (Tables DR3, DR4, and DR6), and small degrees of Nd mobilization are known to be important in at least some deep supra-subduction zone settings (Pearce et al., 1995). Small components of Nd enrichment from marine sediments would shift εNd to less radiogenic values (Hoffman, 2004), thereby pulling the M9 Sm-Nd systematics down off a potential primary composition like that represented by the Kings River ophiolite or Kaweah serpentinite mélange Sm-Nd arrays. Phosphorus in sample M9 is also enriched relative to other Kings-Kaweah ophiolite belt basalts and gabbroids (Table DR6). Various degrees of chemical mixing are known to occur along with high-temperature tectonic abrasion and mixing in deep subduction zones (Bebout and Barton, 2002), and subducted hemipelagic sediments are a viable source for phosphorus enrichment (Ingall et al., 1990). Such chemical changes could also entail components of silica depletion, particularly for fluid fluxes subparallel to the subduction megathrust interface (Manning, 1997). Thus the extreme silica depletion of the putative ferrogabbro protolith need not be wholly a primary feature.
48 4
0.51255
Nd/144Nd
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su
ite
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s
te ui
29
9
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a
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albite. Retrograding and/or contact overprinting is restricted to chlorite sheaths lying along some grain boundaries, and apatite grains that are partially replaced by calcite. Calcite occurs as sub-millimeter-thick veins within the mélange block, and is disseminated in the proximal serpentinite matrix as grain boundary coatings and ophicalcite vein networks. Calcite veining and retrograding in the block are interpreted to represent fluid exchange with the serpentinite matrix. The equilibrium pressure and temperature conditions of metamorphism for the M9 garnet amphibolite block were calculated using the internally consistent average P-T mode in THERMOCALC, version 3.26 (Powell and Holland, 1994; Holland and Powell, 1998). Critical phase compositions used in the calculation (garnet, hornblende, and plagioclase) are given in Table DR7 (see footnote 1). A value of 802 ± 65 °C and 15.9 ± 2.5 kb is derived. The thermobarometric results place the garnet amphibolite nominally within the high-temperature eclogite facies field (Spear, 1993). Such is not the case, as this sample contains albite and augite. There is no textural evidence of the amphibolite being a retrograded eclogite unless the albite grains are a product of jadeite breakdown in the presence of quartz. However, no jadeite component was detected within the augite inclusions, and the texture and mineralogy of the sample are interpreted to reflect ascent to, and subsequent rapid descent from, the determined P-T conditions without the growth of eclogite facies phases. Sample M9 was chosen for Sm-Nd geochronology because it shows only minimal textural evidence for a thermal metamorphic overprint, in contrast to the other block that was discovered that lies in an inner contact aureole position and shows a greater textural overprint. Figure 10 is a 143Nd/144Nd evolution diagram that shows data for sample M9 garnet and its hornblende-rich matrix, which yield a two point “isochron” corresponding to an age of 255 ± 20 Ma with εNd (255) = +8.6 ± 0.4. This age is taken as the approximate time of peak metamorphic conditions for the amphibolite. A K/Ar age of 195 ± 5 Ma was reported for the M9 hornblende (Saleeby and Sharp, 1980). This age is interpreted as a minimum age, considering that the amphibolite is partly retrograded; it sits in an outer contact aureole position, and unfortunately it also sits adjacent to a large Early Jurassic dioritic dike that cuts the Kaweah serpentinite mélange (Fig. 3). Major element data for the garnet amphibolite suggest that it was derived from a Ti-rich ferrogabbro (Table DR4 [see footnote 1] and Fig. 6B). The high positive initial εNd value suggests that it could be of MORB derivation, but major element compositions are too rich in Fe, Ti, and K, and too poor in Si, for typical MORB derivation. Titanium-rich ferrogabbros are present along transform fracture zones from all major ocean basins (cf. Engle and Fisher, 1975; Bloomer et al., 1989; Constantin et al., 1996). A number of these ferrogabbros also show alkali enrichment relative to MORB, like sample M9. The silica depletion of sample M9 is comparable to the most primitive ferrogabbros dredged from fracture zones. In terms of bulk composition, sample M9 also resembles the most primitive mafic eclogite recovered from the Voltri Group of the northern Apennines, also interpreted as a
53
0.51205 0.51195 0.51185
mtx
M 9 garnet amphibolite Age = 255 ± 20 Ma ε Nd (255) = + 8.6 ± 0.4 2 errors 0 0.2 0.6 0.10 0.14 0.18 0.22 0.26 0.30 0.34 0.38 147 144
Sm/
Nd
Figure 10. Isotopic evolution diagram of 143Nd/144Nd for garnet and hornblende-rich matrix for sample M9 garnet amphibolite showing a 255 ± 20 Ma “isochron” age (isochron solution after Ludwig, 2001). Also shown are isochrons from O and M suite samples from Figure 4. Data in Table DR1 (see footnote 1); gt— garnet; mtx—matrix.
54
J. Saleeby
A
A rc th o leiitic
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Boni ni ti c
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Figure 11. Trace element discriminate diagrams for Kings River ophiolite basalts (circles), Kaweah serpentinite mélange basalts and diabases (squares), and garnet amphibolite (star), and for cover-strata mafic volcanic rocks (triangles). Discriminate fields after Pearce (1982) and Shervais (1982). Data in Tables DR3, DR4, DR6, and DR8 (see footnote 1). MORB—mid-oceanic-ridge basalt; OIB—oceanic-island basalt.
Given the above rationale for metamorphic alterations of sample M9 in the deep subduction zone environment, the derivation of the sample M9 protolith from a Kings-Kaweah ophiolite belt–affinity ferrogabbro that has lost silica and gained a small Nd isotopic contaminant signal from underthrust sediments is favored. The widespread occurrence of relatively low-volume ferrogabbros along fracture zones may be attributed to lateral injection of relatively low-volume fractionated magmas from spreading axis environments into bounding transform walls (Bloomer et al., 1989). The scenario that is tentatively favored is that the ferrogabbro protolith was generated as part of an intrusive complex in conjunction with Middle Permian spreading adjacent to the outer transform wall that bounded the Permo-Carboniferous abyssal crust that was generated within the leaky transform system. The outer transform igneous complex entered a neo-subduction zone that nucleated beneath the Kings-Kaweah ophiolite belt while the intrusive complex was still warm from primary heat. Fragments of the complex were accreted to the hanging wall, where they underwent high P and T metamorphism as well as metasomatic changes that were driven by fluids liberated from underthrust oceanic crust and hemipelagic sediments. Rather than remaining for an extended period of time in the deep hanging-wall environment, and progressing through a complete counterclockwise P-T trajectory (cf. Wakabayashi, 1990), blocks of the complex were abruptly entrained and transported upward in serpentinite diapirs into the overlying Kings-Kaweah ophiolite belt, and possibly were erupted as clasts in serpentinite mud volcanoes. Rapid extraction from the peak metamorphic environment inhibited the development of zoning in the principal metamorphic phases, as contrasted with zoning patterns in major phases of high-P blocks from subduction complexes that have completed a counterclockwise path (cf. Wakabayashi, 1990).
COVER STRATA OF THE KAWEAH SERPENTINITE MÉLANGE Cover strata related to the Kings-Kaweah ophiolite belt consist of the southern exposures of the Permian to Middle Triassic Calaveras complex, and the western facies of the lower Mesozoic Kings sequence. These cover strata occur primarily as pendant rocks structurally against the eastern margin of the Kings-Kaweah ophiolite belt and also as nonconformable infolds above the Kaweah serpentinite mélange (Figs. 1 and 3). We focus on these infolds, and specifically those of the Yokohl Valley area (Fig. 3). Viewing Figure 3 in down-plunge orientation (clear arrow at bottom, pointing up), an ~1.5-km-long segment of the nonconformity at the base of the Calaveras complex is seen folded about NW-striking axial surfaces. These folds and related cleavage deform the transitional contact with overlying siliceous argillite and siliciclastic-volcaniclastic strata of the Kings sequence, which occurs as a doubly plunging synformal keel within the Calaveras. This relationship is disrupted to the west by the Harvey Ruth fault zone, which bounds the western margin of the Calaveras exposures. The thicker, more intact intervals of bedded chert within the Calaveras complex resemble the upper-level cherts of the diagonal belt of pillow basalt and chert that extends NE-SW across the Yokohl Valley area, in terms of containing siliceous argillite laminar interbeds and overall bedding forms. This, in conjunction with Figure 3 map relations, suggests that bedded cherts of the diagonal belt, and others like it in the Kaweah serpentinite mélange, graded upward to chert and chert-argillite of the overlying Calaveras complex with increasing argillaceous components. During this hemipelagic sedimentation history the surfacing of serpentinite diapirs severely disrupted the then accumulating stratigraphic succession. The locus of diapiric emplacement was along a transverse fault,
Geochemical mapping of the Kings-Kaweah ophiolite belt, California which cuts and defines the southeast margin of the diagonal pillow basalt–chert belt, and along the Harvey Ruth fault zone. The transverse fault has been deformed in conjunction with cleavage development in the cover strata. The Harvey Ruth fault zone is suggested to have formed a major submarine scarp in the Kaweah serpentinite mélange, with uplifted ocean floor mélange and serpentinized peridotite massifs bounding a trough to the northeast in which Calaveras hemipelagic sediments ponded. Calaveras complex rocks are not observed southwest of the Harvey Ruth fault zone, including the numerous basement cores of the southeastern Great Valley (May and Hewett, 1948; Saleeby, 2007). All cover strata exposures and basement cores from regions southwest of the fault zone are western Kings-sequence-affinity turbidites and volcaniclastic rocks. Thus the Harvey Ruth fault zone constitutes an important tectonic boundary within the Kaweah serpentinite mélange; it originated during deposition of the Calaveras hemipelagic strata, and it continued its activity during the supra-subduction zone residence of the Kings-Kaweah ophiolite belt as shown by its having cut arc volcanic–bearing units of the cover strata. Western facies Kings sequence strata appear to have also been partly ponded into the down-dropped wall of the Harvey Ruth fault zone. Aside from the more massive mafic volcanic intervals that are discussed below, these strata consist at lower stratigraphic levels of siliceous and tuffaceous argillite and arenaceous and conglomeratic strata, much of which were derived from a chert-rich source like the Calaveras complex, and at upper levels from distally derived siliciclastic turbidites. The Harvey Ruth fault zone was remobilized during deposition of this sequence, at least in part as a conduit for renewed serpentinite diapirism. As evident in the down-plunge view of Figure 3, the serpentinite “dike” was a growth feature feeding multiple sedimentary serpentinite-ophiolitic debris flows within the turbidite section. In this setting the upthrown wall of the fault constituted a forearc ridge, which hosted the serpentinite diapirs that entrained small high-pressure metamorphic blocks that were emplaced into the Kaweah serpentinite mélange. Supra-Subduction Zone Volcanism Mafic- to intermediate-composition pillowed and volcaniclastic flows, and tuffaceous admixtures, occur within hemipelagic and siliciclastic cover strata for the Kings-Kaweah ophiolite belt. The stratigraphic positions of these volcanic rocks appear to be restricted to the interval between the cessation of Calaveras radiolarian chert deposition and the arrival of distal siliciclastic turbidites into the region. Unfortunately a vent complex for the volcanic rocks has not been discovered, although likely feeder dikes are observed cutting Kaweah serpentinite mélange basement rocks (Fig. 3). A suite of seven samples from mafic flows, a solitary possible feeder dike that cuts the Kaweah serpentinite mélange, and a dacite clast from a mafic flow were studied geochemically as a means of testing for supra-subduction zone petrogenesis. The more mafic rocks consist of aphyric pillow basalt
55
(sample C8), and pyroxene ± plagioclase ± olivine phyric basalt and basaltic-andesite pillowed and volcaniclastic breccias (samples C2, C3a, C4, C5, C6, and C7). Pyroxenes are commonly pseudomorphed by green amphibole with remnants of only clinopyroxene observed. Rare olivine phenocrysts are pseudomorphed by chlorite ± serpentine ± magnetite. A possible feeder dike for the volcanic rocks, consisting of uralitic diorite, was also studied (sample C1). Major and select trace-element abundance data are presented for the suite in Table DR7 (see footnote 1) and Figures 11 and 12. Except for the dacite clast (sample C3b), silica ranges >52%–55%, with MgO ranging >4%–13%. The dispersed stratigraphic positions of the mafic units, lying within primarily argillaceous strata, raise the question as to the time interval over which the units were erupted. This raises further the question as to whether coherent differentiation trends should be expected for the suite as a whole. Major element variation (Fig. 12) shows some coherency, however. There is a general trend of Al2O3, CaO, and TiO2 enrichment, and MgO depletion with SiO2. There is no coherency to alkalis, and with the elevated K2O value of 1.32% for sample C6, at least some alkali mobility is suggested. An iron enrichment trend is suggested for the entire suite on the basis of the FeO/MgO versus SiO2 plot. On this plot the data straddle the boundary between island arc tholeiite and calc-alkaline–boninite lava series, as defined by Reagan and Meijer (1984) for the Island of Guam in the Mariana forearc. The three lowest stratigraphic level samples (C2, C3a, and C4) have some major element similarities to boninites, with MgO ranging from 9% to 13%, and with SiO2 at ~52%. These samples also show elevated Ni and Cr contents, which is consistent with a boninitic affinity. In general, however, the suite more closely resembles an arc tholeiite association. Trace element data for the cover-strata volcanic rocks are consistent with an arc tholeiite–transitional boninite lava series. Select trace element variation diagrams for the mafic volcanic members, as well as basalts from the Kings-Kaweah ophiolite belt, are presented in Figure 11 (after Pearce, 1982; Shervais, 1982). On the V versus Ti plot the cover strata suite spreads along the arc tholeiite–calc-alkaline–boninite field. For Cr versus Y the three high-Mg rocks plot in the boninite field, and the rest in the boninite-arc tholeiite transition. For Zr versus Ti the three high-Mg rocks again plot in the boninite field, and the rest in the arc tholeiite field. All Kings-Kaweah ophiolite belt basalts plot within the MORB field. The two cover-strata volcanic rocks with the lowest SiO2 (samples C3a and C8) were analyzed for Nd and Sr isotopes (Table DR1). Initial isotopic composition corrections were made for a nominal age of ca. 215 Ma, which is justified below in the discussion of the age of the cover sequence. Initial εNd for these samples is +8.2 (±0.4) and 7.9 (±0.4), respectively. The Sri values are 0.7036 and 0.7040, respectively. The Nd data point to a strong depleted mantle component, whereas the Sr data suggest a seawater component added to the depleted mantle component. The initial εNd values fall within the upper range of values typically measured for arc tholeiite-boninite series rocks, whereas the Sri values lie at the lower range (Taylor et al., 1994;
56
J. Saleeby 16
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Figure 12. Major element variation diagrams for cover-strata volcanic rocks. Island arc tholeiite–calc-alkaline–boninite boundary on FeO/MgO versus SiO2 diagram from Reagan and Meijer (1984). Data in Table DR8 (see footnote 1).
Stern et al., 1991). Thus the mantle source for these lavas appears to have been at the more highly depleted end of the spectrum for typical sources of such lavas. The association of early-stage boninitic-affinity lavas, followed by arc tholeiites, is typical of the initial stages of suprasubduction zone magmatism following subduction initiation (Stern and Bloomer, 1992; Bloomer et al., 1995; Stern, 2004). Relatively shallow-level partial melting of partially serpentinized harzburgite, like that of the Kings River ophiolite, is a likely source for the earlier boninitic-affinity lavas with deeper melting of a similar depleted source that gave rise to the tholeiitic lavas. The relatively advanced state of time integrated LILE source depletion, as indicated by the Nd and Sr isotopic data, is consistent with abyssal mantle lithosphere akin to that of the Kings River ophiolite, forming the principal source rock in the mantle wedge that rendered the mafic lavas. Stratigraphic Setting and Age Constraints of Proximal Supra-Subduction Zone Volcanic Rocks The eruption of proximal volcanic rocks of the cover strata began at the cessation of radiolarian chert deposition in the Cala-
veras complex, and ceased by the time of siliciclastic turbidite deposition of the western Kings sequence. These relationships are exhibited for the sample C2 basaltic unit that lies along the Calaveras–Kings sequence contact, and the sample C8 basaltic unit that is mainly faulted against, but appears to be in local depositional contact with, the turbidites (Fig. 3). The local depositional contact is likely an angular unconformity, now transposed, representing an unknown hiatus. Age constraints for the cover strata are only quite broad, and thus the time interval over which proximal volcanism occurred is not well constrained. Accumulation of Calaveras chert-argillite began after a putative ~30 m.y. of radiolarian ooze deposition above ca. 295 Ma Kaweah serpentinite mélange pillow basalt. This suggests that Calaveras chert-argillite sedimentation began at ca. 265 Ma, or in the Middle Permian. The inclusion of Permian shallow-water-limestone slide blocks within chaotically deformed Calaveras chert-argillite (Saleeby, 1979) only constrains hosting chert-argillite deposition to the Permian or younger. Ammonoid remains from several localities within Kings sequence argillites indicate an early Mesozoic age (Saleeby et al., 1978). U/Pb zircon ages of 170 ± 1 Ma on a hornblende andesite dike that crosscuts the sample C3 volcanic unit (Saleeby and Sharp, 1980), and of 152 ± 1 Ma on
Geochemical mapping of the Kings-Kaweah ophiolite belt, California a trondhjemitic dike that cuts the siliciclastic turbidites, further indicate an early Mesozoic age for much of the section (Saleeby and Dunne, 2011). The 255 ± 20 Ma Sm-Nd isochron age for the sample M9 garnet amphibolite putatively further constrains proximal supra-subduction zone volcanism to have initiated in early Mesozoic time. Maximum age constraints for the sample C3 volcanic unit are provided by U/Pb zircon age data on sample C3b, an angular dacite clast of ~30 dm diameter that lies in a basaltic tuff breccia (sample C3a). The size and textural immaturity of the clast indicate that it probably underwent only one cycle of transport within the hosting submarine pyroclastic flow. This further suggests that its source was proximal to the vent complex for the mafic flow. In terms of bulk composition (Table DR7 and Fig. 12) and the presence of plagioclase microphenocrysts, completely recrystallized, the dacite clast is similar to silicic members of the boninite-arc tholeiite suites of Chichijima in the Bonin Islands forearc (Taylor et al., 1994). Possible consanguinity between the dacite clast and its mafic host suggests that the age of the clast could approximate the age of the sample C3 volcanic unit and other similar units proximal to the sample C3 site (Fig. 3). Sample 3b yielded a meager population of fine euhedral zircon. Figure 13A is an age frequency plot for 206Pb/238U ages determined by laser ablation ICP-MS (inductively coupled-plasma mass spectroscopy) techniques on 23 fine euhedral grains extracted from the clast. Analytical data are presented in Table DR9a. The data define a mean age of 219.1 ± 2.9 Ma. Ages for a number of grains range to as young as earliest Cretaceous, similar to the highly discordant zircon from the Kaweah serpentinite mélange plagiogranites (Saleeby and Sharp, 1980). The zircon grains of sample C3b, exhibiting the younger ages, are considered disturbed by Sierra Nevada batholith contact metamorphism and have been omitted from the Figure 13A analysis. The distribution of 207Pb/206Pb ages for those grains included in the age analysis (Table DR9a) suggests that most of these ages are concordant, and with the small scatter of the 206Pb/238U ages (MSWD = 1.4), 219 Ma is taken as the eruption age for the dacite clast source. U/Pb zircon data indicating an age of ca. 200 Ma was presented for the sample C1 dioritic “feeder” dike (Saleeby and Sharp, 1980). These age constraints suggest a Late Triassic to earliest Jurassic age for the principal phase of supra-subduction zone volcanism in the cover strata. The onset of supra-subduction zone volcanism in the Late Triassic in conjunction with stratigraphic relations (Fig. 3) places a Middle Triassic bound on the termination of hemipelagic sedimentation for the Calaveras complex. Approximate age and sediment provenance constraints for the siliciclastic turbidite of sample C9 are provided by U/Pb ages of its detrital zircon population (Table DR9b and Figs. 13B and 13C). Sample C9 consists of a 10-cm-thick sandstone bed graded to medium sand size along its base, where it consists of ~60% quartz, ~20% feldspar, plagioclase dominant, and ~20% lithic fragments, both volcanic and metamorphic. Recrystallization inhibits the precision of the determined modes and may have skewed the analysis away from a higher lithic content. Fig-
57
ure 13B shows an age probability plot for back to 500 Ma on the basis of 37 206Pb/238U ages. In terms of age constraints on the turbidite, the strong peak at 175–160 Ma suggests a late Middle Jurassic maximum depositional age. The minimum age of the turbidite section is constrained by it and its first cleavage being cut by a 152 ± 1 Ma trondhjemite dike swarm (Saleeby and Dunne, 2011). The 175–160 Ma peak corresponds to one of two major early Mesozoic magmatic flux events in the SW Cordilleran arc, the other being during the Triassic (Stern et al., 1981; Chen and Moore, 1982; Fiske and Tobisch, 1978; Busby-Spera, 1984; Barth et al., 1997; Ducea, 2001; Saleeby and Dunne, 2011). The sample C9 detrital zircon clearly reflects detrital inputs from both early Mesozoic arc belts. The full spectrum of 206Pb/238U detrital zircon ages for sample C9 is given in Figure 13C, with comparative plots from regional Paleozoic siliciclastic units that constituted proximal parts of the SW Cordilleran passive margin and overlying lower Mesozoic eolianites, which were possible sources for the ancient detrital zircon populations. Of particular interest are the spectra from continental slope-rise strata that occur in southern Sierra Nevada pendants as basement for the eastern facies of the Kings sequence (Saleeby and Busby, 1993; and unpub. data), and similar pendant rocks from the northern Mojave Desert region (A. Chapman and J.B. Saleeby, unpub. data). The spread of ages between ca. 0.95–2.1 Ga and 2.45–2.8 Ga corresponds to a series of peaks in the C9 turbidite. In contrast, lower Paleozoic passive-margin shelf rocks of the Snow Lake pendant of the eastcentral Sierra Nevada, the central Mojave Desert, and Death Valley (Grasse et al., 2001; Barbeau et al., 2005) do not yield spectra similar to the Late Archean-Proterozoic population of sample C9. Sample C9 peaks in the 390–640 Ma range are not readily explained with the southern Sierra–northern Mojave source spectra. Ages in the 390–400 Ma range could reflect fringing arc rocks emplaced into and erupted over lower Paleozoic (Shoo Fly) slope-rise strata of the northern Sierra Nevada (Saleeby et al., 1987), whereas detrital zircon with ages ranging from 390 to 640 Ma are abundant in lower Mesozoic eolianites that spread across the western reaches of the SW Cordilleran passive margin (Dickinson and Gehrels, 2003; Fig. 13C). The distribution of Paleozoic SW Cordilleran passive margin strata along the California region is pursued further below. In summary, proximal supra-subduction zone submarine mafic volcanic rocks of the Kaweah serpentinite mélange cover strata were erupted in Late Triassic to earliest Jurassic time. Upper Middle to lower Upper Jurassic siliciclastic turbidites appear to rest unconformably on the youngest of these mafic flows. A vent complex has not been identified for the Late Triassic–earliest Jurassic flows. Intercalation of the flows with argillite-rich strata suggests episodic volcanism off the flanks of a vent complex within a basinal setting. Such a vent complex has been identified for Late Triassic–earliest Jurassic arc tholeiitic to transitional calc-alkaline submarine mafic volcanics of the Jasper Point–Penon Blanco Formation ~100 km north of the Kings-Kaweah ophiolite belt along the Sierra Foothills belt (Saleeby, 1982; Snow, 2007). This volcano-plutonic complex
58
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Figure 13. U/Pb zircon age data for cover strata samples. (A) Frequency plot of 206Pb/238Pb ages on single igneous zircon grains from sample C3b dacite block. (B) Probability plot of younger spectrum of 206Pb/238Pb ages for detrital zircon from sample C9 siliciclastic turbidite. (C) Probability plots of 206Pb/238Pb ages for detrital zircon of C9 siliciclastic turbidite in comparison to 206 Pb/238Pb ages for detrital zircon populations from Paleozoic siliciclastic units of the southern Sierra Nevada region and for lower Mesozoic eolianites of the SW Cordillera, which were potential source components for the turbidites (after Grasse et al., 2001; Dickinson and Gehrels, 2003; Barbeau et al., 2005; Chapman et al., 2011). Data in Tables DR9a and DR9b (see footnote 1). MSWD—mean square of weighted deviates; Pz—Paleozoic.
Geochemical mapping of the Kings-Kaweah ophiolite belt, California was constructed on the Permo-Carboniferous Tuolumne ophiolitic mélange. The Jasper Point–Penon Blanco volcanic construct appears to have constituted a broad pillowed shield with hemipelagic sediment intervals followed by submarine volcaniclastic cone construction. The upper volcaniclastics interfinger northward along the Foothills belt with argillite-rich strata, producing a section similar to the principal stratigraphic section in which the Kaweah serpentinite mélange cover strata volcanics lie. It seems likely that the mafic flows in this serpentinite mélange cover strata likewise formed off the flanks of a major constructional center, presently not exposed. Thus the pattern of early Mesozoic supra-subduction zone mafic submarine volcanism constructed on a disrupted and accreted late Paleozoic ophiolitic substrate is a regional pattern along the Foothills belt. Furthermore, the Jasper Point–Penon Blanco construct, and part of its intrusive roots, rest unconformably beneath Upper Jurassic siliciclastic turbidites (Saleeby, 1982; Ernst et al., 2009), further attesting to a common early Mesozoic tectonic history along the Foothills belt. TECTONICS OF MÉLANGE FORMATION IN A LARGE OFFSET TRANSFORM–SUBDUCTION INITIATION ENVIRONMENT In this section a tectonic model is developed that accounts for the polygenetic abyssal magmatic history of the KingsKaweah ophiolite belt, its progressive disruption to form ocean floor mélange, and its accretion into the SW Cordilleran active margin and ensuing residence in a supra-subduction zone environment. A number of initial questions must be addressed in order for such a model to be seriously considered: (1) Was the Kings-Kaweah ophiolite belt “obducted” onto the Cordilleran margin? (2) Given up to an ~190 m.y. hiatus in abyssal magmatism, is the implied residence time within the abyssal realm reasonable within the physical constraints of seafloor spreading rates and characteristic sizes of major ocean basins? And (3) Given a major abyssal transform phase for the Kings-Kaweah ophiolite belt, did the plate kinematics of this plate juncture circuit directly into the plate kinematics of the Cordilleran margin? Or alternatively, was the abyssal transform phase decoupled from Cordilleran tectonics, and the transform assemblage merely accreted en masse, independent of its transform history? This question may also be posed as: Is there evidence within the SW Cordillera, independent of the transform history recorded in the Kings-Kaweah ophiolite belt, for late Paleozoic transform tectonics having affected the region? Gross Emplacement Geometry The Kings-Kaweah ophiolite belt was not obducted onto the SW Cordilleran margin. Regionally pervasive deformation structures and fabrics associated with the ophiolite belt are steep to vertical. West-dipping mid-crustal reflectors along the western Sierra Nevada–Great Valley transition directly north of the Kings River region (Miller and Mooney, 1994) are commonly cited as
59
evidence for eastward obduction of the Foothills belt, including its ophiolitic basement, in the Late Jurassic. However, basement core, as well as surface map data, shows that such reflectors underlie voluminous Early Cretaceous Sierra Nevada batholith rocks and subordinate pendants of Foothills belt rocks (May and Hewett, 1948; Saleeby, 2007). The referenced reflectors are continuous and highly coherent, and would not survive in such a state after the intrusion of copious batholithic intrusions. Furthermore, these reflectors coincide with a patch of recent seismicity (Gilbert et al., 2007), and thus it is much more likely that the reflectors are young and potentially active structures related to ongoing mantle-lithosphere removal processes and related lower crustal flow beneath the region (Zandt et al., 2004; Le Pourhiet et al., 2006; Saleeby et al., 2003). The emplacement geometry of the Kings-Kaweah ophiolite belt appears to have constituted a lithosphere-scale wedge that was accreted to the hanging wall of a newly established subduction zone. Serpentinite diapirs were sourced from deep enough below the accreted wedge to entrain high-pressure metamorphic blocks from underplated abyssal lithosphere, and the diapirs penetrated up through the accreted ophiolite belt without entrainment of hypothetical continental structural basement for the belt, assuming an obduction geometry. Early Mesozoic boninitic to arc tholeiitic volcanics that were erupted through the Kings-Kaweah ophiolite belt were likewise sourced within a mantle wedge composed of hydrous depleted peridotites with strong time integrated LILE depletions. Finally, Early Cretaceous members of the Sierra Nevada batholith that were intruded through the Kings-Kaweah ophiolite belt lack any evidence of having interacted with hypothetical continental structural basement for the ophiolite belt. The above is displayed in the Figure 14 plot of initial εNd verses Sri for principal igneous suites of the southern Sierra Nevada (DePaolo, 1981; Pickett and Saleeby, 1994; Clemens-Knott et al., 2011; and data reported here). The plot shows data fields for the KingsKaweah ophiolite belt, cover strata mafic volcanics, shallow-level Early Cretaceous Sierra Nevada batholith rocks emplaced into the Kings-Kaweah ophiolite belt, deep-level Early Cretaceous Sierra Nevada batholith rocks emplaced south of, but along strike of, the Kings-Kaweah ophiolite belt, and Late Cretaceous Sierra Nevada batholith rocks emplaced into North American crustal rocks of the axial to eastern Sierra Nevada batholith. The plot also shows a representative field for Proterozoic sialic basement of the SW Cordillera region, an important component for the axial to eastern Sierra Nevada batholith suite. The western Foothills suites display a geochemical maturation of the underlying mantle wedge with time, starting with its proto-composition as depleted mantle of Kings-Kaweah ophiolite belt affinity. Progressive addition of slab-derived fluids and minor terrigenous sediment components to the mantle wedge with time progressively shifted the εNd and Sri values progressing from the early Mesozoic suprasubduction zone mafic volcanics to the cross-cutting batholithic units (cf. DePaolo, 1981; Lackey et al., 2005; Stern et al., 1991; Clemens-Knott et al., 2011). In contrast, the entire axial to eastern Sierra Nevada batholith suite shows a strong Proterozoic North
60
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ancy would have worked in favor of a subduction initiation event, given the correct tectonic circumstances. It appears that the SW Cordilleran margin and the adjacent Panthalassa basin presented such circumstances at the end of the Paleozoic.
A B C D
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KKOB (Pz) KSM cover strata volcanics (Lt. Triassic - E. Jurassic) Shallow w. SNB (125-115 Ma) Deep w. SNB (115-100 Ma)
E F
e. SNB (100-85 Ma) SW Proterozoic basement
Figure 14. Plot of initial εNd verses Sri for samples from the KingsKaweah ophiolite belt (KKOB), its cover strata mafic volcanics, crosscutting plutons of the Early Cretaceous Sierra Nevada batholith (SNB), axial to eastern batholithic rocks of the southern Sierra Nevada, and SW Cordilleran Proterozoic basement (DePaolo, 1981; Pickett and Saleeby, 1994; Clemens-Knott et al., 2011; data presented in Table DR1 [see footnote 1]). Also delineated is the geochemical expression of the Foothills suture. Pz—Paleozoic.
American lithosphere or crustal component that is consistent with the North American crustal host rocks for the inner Sierra Nevada batholith zones. The boundary between the distinctive eastern continental affinity and western oceanic affinity Sierra Nevada batholith coincides with a profound tectonic break in metamorphic framework rocks wherein Paleozoic ophiolitic rocks lie to the west, and early Paleozoic Cordilleran passive-margin-affinity rocks lie to the east. The pre-batholithic boundary is named the Foothills suture, shown by its expression in the Sierra Nevada batholith εNd versus Sri in Figure 14, and discussed further below. Aging of Oceanic Plates As noted earlier, an ~190 m.y. hiatus in abyssal magmatism may not seem reasonable within the context of modern plate tectonics. The oldest known abyssal crust is ca. 168 Ma in the western Pacific, and it is nearing subduction into the Mariana trench (Ogg and Smith, 2004). However, the geometry of the continental masses and ocean basins was very different throughout the Paleozoic, as compared with the Cenozoic. During the Pangean supercontinent era the Panthalassa ocean basin occupied nearly two-thirds of Earth’s surface (Murphy and Nance, 2008). Thus such an extended abyssal-realm residence for the Kings-Kaweah ophiolite belt does not appear to pose any problems that cannot be accounted for in our current understanding of plate tectonics through geologic time. An interesting consideration regarding the abyssal magmatic hiatus is the implication of the residence of such old oceanic lithosphere in the Panthalassa basin. Based on simple principles of conductive cooling of oceanic lithosphere as it ages off its respective spreading axis, such old oceanic lithosphere should have carried strong negative buoyancy forces. Such strong negative buoy-
Transform Tectonics and Subduction Initiation along the SW Cordilleran Margin One of the definitive features of the SW Cordilleran margin is a regionally extensive zone of Permo-Carboniferous transform truncation that coincides with the pre-batholithic metamorphic framework of the Sierra Nevada (Davis et al., 1978; Walker, 1988; Kistler, 1990). Evidence supporting this event includes the high-angle truncation of NE-SW–trending facies boundaries and fold-thrust structures in the Neoproterozoic-Paleozoic passive margin sequence along the eastern Sierra Nevada region, and the truncation of the regional Paleozoic “McCloud” fringing arcmarginal basin system that ran outboard of the central to northern Cordilleran passive margin (Rubin et al., 1990). In conjunction with this truncation event was the transpressive accretion of NW-trending strike-slip ribbons along the truncation zone, which constitute the principal Paleozoic metamorphic framework units for the Sierra Nevada batholith. Figure 15 shows the distribution of Paleozoic metamorphic framework ribbons in the Sierra Nevada as well as the southern termination of the McCloud arc in the eastern Klamath Mountains. The Sierra Nevada batholith framework ribbons occur along the western Foothills, where they are variably covered by lower Mesozoic active margin strata as poly-metamorphosed pendants within this batholith, and as the eastern wall of the batholith along the Owens Valley region. The general distribution of the displaced ribbons relative to the zone of Permo-Carboniferous continental truncation is shown in the inset map of Figure 15. Essential features of Figure 15 are the delineation of regional tectonic domains based on facies relations and petrotectonic assemblages of Paleozoic rock exposures. The domains include areas once occupied by Paleozoic rocks that have been intruded out by Mesozoic plutons or covered by Mesozoic or Cenozoic strata. The principal domains are the western North America craton margin, the passive margin shelf, its slope and rise, the McCloud fringing arc with remnants of its subduction complex, and Panthalassa lithosphere that was accreted to the SW Cordilleran margin. The principal Paleozoic rock exposures for the Sierra Foothills belt and the southeastern Klamath Mountains are differentiated within the context of the regional domains. Of critical importance in the western Sierra are exposures of Paleozoic ophiolitic mélange and metamorphic tectonites of the Feather River, Bear Mountain, and Tuolumne complexes, all of which have igneous and metamorphic age, and a number of structural relations that are similar to those of the ophiolitic ductile shear zones and serpentinite mélange of the KingsKaweah ophiolite belt (Saleeby, 1990). Co-extensive with these Paleozoic oceanic basement exposures are belts of Calaveras complex chert-argillite, all broadly constrained in age from
Eastern Klamath Mtns.
TR
a uil ah Co
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t
ial Ax
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a
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ea
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e ing h nic ato r C
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ath De
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on
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ma rgi ato n nh in g e
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Geochemical mapping of the Kings-Kaweah ophiolite belt, California
0
100 km
olf W e t hi W
Mojave Desert
Figure 15. Maps showing principal Paleozoic tectonic elements and facies systems of the SW Cordillera in relation to the Permo-Carboniferous continental truncation zone, Sierra Foothills Paleozoic ophiolite and related (Calaveras) chert-argillite belts, and key Paleozoic tectonic elements of the eastern Klamath Mountains. Some of the larger isolated plutons and basal contacts of the lower Mesozoic overlap units of the Sierra Foothills and the Klamath Mountains are delineated for location purposes. No attempts were made to restore Basin and Range extension nor Middle Jurassic dextral drag along the Sierra Nevada and its proximal backarc region. The strikeslip ribbons of the Sierra Nevada are reconstructed from stratigraphic relations of metamorphic pendant belts. (Sources: D’Allura et al., 1977; Irwin, 1977; Schweickert et al., 1977; Saleeby et al., 1987; Kistler, 1990; Saleeby, 1990; Saleeby and Busby, 1993; Metcalf et al., 2000; Dickinson and Lawton, 2001; Glazner et al., 2005; Stevens and Green, 2000; Stevens et al., 2005; Stevens and Stone, 2005; Nadin and Saleeby, 2008; Saleeby and Dunne, 2011.) Pz—Paleozoic; Cz—Cenozoic.
62
J. Saleeby
Permo-Carboniferous to Permo-Triassic (Cox and Pratt, 1973; Schweickert et al., 1977; Behrman and Parkison, 1978; Watkins et al., 1987; Saleeby, 1979). As with the Calaveras complex that is spatially related to the Kings-Kaweah ophiolite belt, these belts are interpreted as mainly “Permian” hemipelagic accumulations that rested above, and were accreted along with, Panthalassa lithosphere fragments. The locus of accretion is the Foothills suture, the join between North American and Panthalassan lithosphere. The suture is hypothesized to have had a polyphase history, first, of Permo-Carboniferous sinistral transform motion, followed by latest Permian sinistral oblique thrust imbrication during subduction initiation, and then re-deformation within the axis of superposed early Mesozoic arc magmatism. The eastern wall of the suture in the central to northern Sierra consists of lower Paleozoic slope-rise strata of the Shoo Fly complex, and plutons and overlap strata of the southernmost McCloud arc (D’Allura et al., 1977; Schweickert et al., 1977; Saleeby et al., 1987; Rubin et al., 1990; Harding et al., 2000). Similar lower Paleozoic sloperise strata lie in southern Sierra pendants east of the intruded-out Foothills suture, and form basement for the eastern facies of the lower Mesozoic Kings sequence (Saleeby and Busby, 1993; and unpub. data). The northward continuation of the Foothills suture is suggested below to have extended into the southern Klamath Mountains, but it has since been severely overprinted by early Mesozoic subduction-related thrusting (Davis et al., 1978). The inner margin of the truncation zone along the Sierra Nevada and northern Mojave Desert has been completely intruded out by the Mesozoic batholithic belt. Nevertheless, the inner truncation boundary can be regionally delineated by stratigraphic and sedimentary facies discontinuities between NW-trending pendant belts and by regional geochemical discontinuities in the batholithic belt (Moore and Foster, 1980; Walker, 1988; Kistler, 1990; Dunne and Suczek, 1991; Saleeby and Busby, 1993; Greene et al., 1997; Stevens et al., 2005; Lackey et al., 2005; Saleeby and Dunne, 2011). The inner truncation structure is shown as the Axial Sierra fault in Figure 15. The trace of the Permo-Carboniferous Axial Sierra fault has been disrupted by Mesozoic dextral faults (Lahren and Schweickert, 1989; Kistler, 1990; Saleeby and Busby, 1993; Nadin and Saleeby, 2008). The distribution of passive-margin shelf facies rocks along and adjacent to the truncation zone has been further shuffled by Permo-Triassic and mid-Cretaceous thrust faults, and conceivably by early Mesozoic extensional faults (Stevens and Greene, 2000; Stevens and Stone, 2005; Nadin and Saleeby, 2008; Saleeby and Dunne, 2011). The principal superposed Mesozoic structures are shown in Figure 15 in green, some preserved in batholithic and pendant rocks, and some largely cut out by batholithic rocks. In the extreme southwestern Sierra Nevada and adjacent Mojave Desert, large-magnitude extensional structures of Late Cretaceous age that exhumed the Sierra Nevada batholith to lower crustal levels have obscured pre-batholithic pendant relationships (Nadin and Saleeby, 2008; Saleeby et al., 2007). Elsewhere in the Sierra Nevada, exhumation of the batholith has been limited to mid- to upper crustal levels, leaving pendant stratigraphy and batholith petrochemical patterns
intact to the extent that the complex structural relations along the truncation zone can be reasonably constrained (Fig. 15). The accreted Panthalassa lithosphere domain extends across the projected trace of the Foothills suture in the eastern Klamath Mountains region as the Trinity peridotite, which forms depositional basement for the southernmost McCloud arc, and the Klamath Central Metamorphic Belt that consists of Paleozoic MORB metabasites that were partly subducted eastward beneath the Trinity peridotite in mid-Paleozoic time (Wallin and Metcalf, 1998; Metcalf et al., 2000; Barrow and Metcalf, 2006). The Trinity peridotite consists of serpentinized harzburgite, dunite, and lherzolite that are similar to the long-term LILEdepleted mantle that rendered the crustal section for the Kings River ophiolite. Plagioclase lherzolite that equilibrated under shallow upper-mantle, diapiric conditions (Quick, 1981) yields an Sm-Nd mineral–bulk rock isochron and an initial εNd that are indistinguishable from those determined for the Kings River ophiolite mafic crustal section (Jacobsen et al., 1984; Shaw et al., 1987). The high εNd value indicates that the Trinity peridotite was derived from the Panthalassa-depleted MORB mantle. The Trinity peridotite was ascending and undergoing partial melting beneath a Panthalassa spreading center at the same time that the Kings River ophiolite crustal section was generated along a Panthalassa spreading center by the partial melting of a Trinity-like peridotite. Unfortunately, vestigial lenses of peridotites retaining the remnants of high-temperature mantle-flow fabrics in the Hog and Red Mountain slabs of the Kings River ophiolite are serpentinized and contact metamorphosed, so they are not well suited for mineral isotopic studies that could more directly test a Trinity peridotite affinity. The Trinity peridotite was incorporated into a supra-subduction zone environment during the Early Silurian initiation of the southern segment of the McCloud arc system (Metcalf et al., 2000), whereas the Kings River ophiolite remained in the Panthalassa abyssal realm until the Late Permian. This is consistent with the Kings River ophiolite being generated at a fast spreading ridge, as concluded above, which potentially generated large expanses of abyssal lithosphere, including that of the Trinity peridotite, capable of rendering highly divergent tectonic evolutionary paths of potentially derivative ophiolitic fragments. The inset map of Figure 15 shows the southwest Cordillera truncation zone in a regional context. The truncation structure has been equated with the Mojave-Sonora megashear (Anderson and Silver, 1979), a zone of Late Jurassic sinistral shear that has produced little translation that can be documented. The term California-Coahuila transform has been adopted from Dickinson and Lawton (2001) as the Permo-Carboniferous truncation and translation zone, so as not to confuse this plate juncture for the superposed intraplate supra-subduction zone strain of the “megashear.” The Caborca block is shown in the northwestern Mexico region astride the transform. This block corresponds to a fragment of the passive margin shelf that was translated 500–800 km along the transform from the truncated shelf at Sierra Nevada latitudes (Stevens et al., 2005). Stratigraphic relations along the truncation locus indicate that the principal phase of sinistral
Geochemical mapping of the Kings-Kaweah ophiolite belt, California translation was during the Pennsylvanian (Stevens et al., 2005). Isolated exposures of slope-rise strata along the northwest margin of the Caborca block could be facies changes across the margin of the block and/or additional strike-slip ribbons displaced within the transform zone. The inset map also shows the locus of Triassic arc magmatism, which runs inboard and along the trace of the transform zone (Stern et al., 1981; Chen and Moore, 1982; Busby-Spera, 1984; Barth et al., 1997; Saleeby and Dunne, 2011). The actual age range of this “Triassic” arc belt is from the latest Permian (ca. 248 Ma) into the Early Jurassic (ca. 200 Ma), with the highest flux of plutons occurring in Early to Middle Triassic time. The principal phase of transform truncation and displacement along the California-Coahuila system, and the onset of “Triassic” arc magmatism, are in accord with the transform phase of development, and the subduction initiation emplacement of the Kings-Kaweah ophiolite belt. Translation and Emplacement of the Kings-Kaweah Ophiolite Belt The favored hypothesis for the generation, displacement, and emplacement of the Kings-Kaweah ophiolite belt is shown in Figure 16, and an overview of the critical phases of geologic history recorded in this ophiolite belt are summarized in Table 2. Lacking definitive constraints, a simple approach is adopted for the initial plate configuration whereby the genesis of the Kings River ophiolite along a fast-spreading ridge is depicted to have been proximal to a large offset transform roughly aligned with the future California-Coahuila transform (Fig. 16A). The fracture zone is shown extending into the Cordilleran passive margin as a marginal offset in the Neoproterozoic rifted margin. The marginal offset is shown as the outer edge of what was to become the Caborca block in the Permo-Carboniferous. The configuration of the passive margin to the south of the Caborca block native site (present geographic coordinates) is poorly constrained owing to subsequent tectonic overprints (Dickinson and Lawton, 2001; Nance and Linnemann, 2008). The possible existence of additional continental ribbons displaced from the outer edge of the Caborca block is also poorly constrained, with one possibility being the older continental basement elements within the Chortis block of the Central America isthmus (Rogers et al., 2007). Vast tracts of new abyssal lithosphere are shown to have been generated in Early Ordovician time (Fig. 16A), capable of rendering highly divergent evolutionary paths of its potential ophiolitic fragments, as indicated by the contrasting histories of the Kings River ophiolite and the Trinity peridotite. Considering the distribution of the continents and major ocean basins in the early Paleozoic (Murphy and Nance, 2008), the seafloor spreading kinematics of Figure 16A could have circuited directly into Iapetus plate motions. A hypothetical site for the Early Ordovician ascent of the Trinity peridotite beneath a Panthalassa ridge segment is shown adjacent to the “Kings River ophiolite” ridge segment, across a major transform fracture zone. The “Trinity” ridge-transform segments are considered likely nucleation sites
63
for the initiation of east-dipping intra-oceanic subduction that rendered the southern segment of the McCloud arc, which was formed by Early Silurian time as recorded in the Klamath Mountains (Wallin and Metcalf, 1998). Once subduction initiated for the McCloud arc, the plate configuration depicted requires that the oceanic transform zone separating the Kings River ophiolite and Trinity ridge segments became active as a boundary transform, possibly analogous to the modern South Sandwich transform that bounds the southern Scotia Arc (British Antarctic Survey, 1985). Devonian–Early Mississippian fold and thrust belt structures developed along the southern McCloud backarc and adjacent Cordilleran shelf region suggest an early phase of McCloud arc impingement along the Cordilleran margin (cf. Smith and Miller, 1990) the details of which are not treated here. Figure 16B shows the plate geometry in the Late Pennsylvanian (295–290 Ma), entailing the initiation of the CaliforniaCoahuila transform, and intra-oceanic rifting along the KingsKaweah ophiolite belt–hosting transform. The initiation of a transform-spreading geometry similar to the Cayman Trough is adopted (CAYTROUGH, 1978; Rosencrantz et al., 1988) whereby ephemeral short spreading-center segments lie nested between transform walls with metamorphic core complex segments. The b–b′ cross section diagrammatically shows Ordovician MORB crust and mantle lithosphere (Kings River ophiolite) along the edge of the transform deforming into a metamorphic core complex and being entrained into the reactivated leaky transform zone. The geometry and kinematics of the putative core complex differ from oceanic core complexes that have been described in the literature (cf. Baines et al., 2003; Schroeder and John, 2004; Tucholke et al., 2008) in that the driving transtensional deformation is affecting the edge of an aged oceanic plate along an intra-oceanic rift, as opposed to transform plate edges developed proximal to more or less steady-state abyssal spreading centers. In this cross section a detachment fault similar to that imaged at the mid-Atlantic Ridge–Atlantis fracture zone intersection (Canales et al., 2004) is shown disrupting the Kings River ophiolite crustal section and rooting into a neovolcanic rift along which new oceanic lithosphere is generated. The Kings River ophiolite Moho ductile shear zone is shown forming along the asthenospheric corner flow zone off the shoulder of the neo-rift zone (after Baines et al., 2008). Entrainment of the Kings River ophiolite into the rift system requires the nucleation of a transform segment between this ophiolite and its native abyssal plate, as shown in the Figure 16B inset. Figure 16C shows the plate geometry in the Early Permian (285–260 Ma), with the growth of an oceanic microplate within the dilational transform system. The microplate is named the Foothills ophiolite belt microplate for its current expression as the principal ophiolitic exposures of the Sierra Foothills, including the Feather River, Bear Mountains, Tuolumne and KingsKaweah sub-belts and the co-extensive Calaveras hemipelagic belts. The Foothills ophiolite belt microplate was composed primarily of “abnormal” oceanic crust with an abundance of surfaced serpentinized upper mantle rocks, mafic metamorphic
J. Saleeby Mc Cl
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Figure 16. Generalized plate tectonic model for the generation of the Kings River ophiolite along a fast-spreading ridge in the Early Ordovician Panthalassa ocean basin, followed by the generation of the Kaweah serpentinite mélange in a Permo-Carboniferous leaky transform–slow-spreading-ridge system like that of the Cayman Trough, and the emplacement of the Kings-Kaweah ophiolite belt along the Foothills suture during subduction initiation. Insets show diagrammatic cross-sectional relationships at key time intervals. The Foothills ophiolite belt (FOB) consists of the Feather River, Bear Mountains, and Tuolumne sub-belts, in addition to the KingsKaweah ophiolite belt and the overlying and co-extensive Calaveras chert-argillite belts. Color-coding for details of the Kings-Kaweah ophiolite belt is the same as in Figures 2 and 3, and for other regional features (Fig. 15). Paleolatitude and orientation of the southwest North American craton are after Cocks and Torsvik (2002) and Nance and Linneman (2008). See text for details. KRO—Kings River ophiolite; Pz—Paleozoic.
Geochemical mapping of the Kings-Kaweah ophiolite belt, California
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TABLE 2. OUTLINE OF PETROGENETIC AND TECTONIC EVENTS RECOGNIZED IN THE KINGS-KAWEAH OPHIOLITE BELT (KKOB) Geologic time* (Ma) 125–100
Tectonic regime Intrusion of gabbroic to tonalitic plutons of western zone Sierra Nevada batholith and resulting contact metamorphism.
152–125
Second slaty cleavage formed in cover strata turbidites.
157–148
Intrusion of tonalitic intrusive sheets and basalt-trondhjemite solitary and sheeted dike sets with sinistral shear, and local dynamothermal contact metamorphism.
160–152
Nonconformable deposition of siliciclastic turbidites, serpentinite diapiric dike emplacement, and formation of first cleavage in turbidites.
170–157
Intrusion of gabbroic to tonalitic sheets with longitudinal dextral shear and local dynamothermal contact metamorphism, and local basalt to andesite solitary dike emplacement.
190–175
Early cover strata erosion, possible surfacing of serpentinite diapirs.
205–195
Intrusion of dioritic and trondhjemitic solitary dikes.
225–190
Submarine eruption of boninitic to arc tholeiitic pillowed and volcaniclastic flows, local submarine dacite dome-tuff eruption, deposition of tuffaceous and siliceous argillite, and sandstones-conglomerates derived from chert-rich source.
265–225
Deposition of Calaveras chert-argillite, near pervasive soft sediment deformation and inclusion of Permian shallow-water-limestone blocks.
ca. 255
Emplacement of KKOB to hanging wall of neo-subduction zone, and underlying high-pressure † metamorphism with subsequent emplacement of derivative metamorphic blocks into KSM by serpentinite diapirism.
295–255
Oceanic transform displacement of KKOB to SW Cordilleran transform margin.
295–265
Pelagic sedimentation of radiolarian oozes.
ca. 295
Diffuse oceanic spreading, transform generation of ophiolitic ductile shear zones, core complex § deformation, and transform capture of KRO , and initiation of submarine serpentinite diapirism and ocean floor mélange formation.
484–295
Residence of KRO in Panthalassa ocean basin with sparse pelagic sedimentation of metalliferous radiolarian ooze.
ca. 484
Seafloor spreading generation of KRO at fast spreading center in Panthalassa ocean basin.
*Time intervals not shown lack resolvable geologic record. Italics denote approximate upper and lower bounds on time constraints. † KSM—Kaweah serpentinite mélange. § KRO—Kings River ophiolite.
tectonites deformed along transform shear zones and core complex segments, dispersed basaltic-hypabyssal carapaces, mafic and ultramafic rubble piles, and variably deformed pelagic oozes. The c–c′ cross section shows a diagrammatic profile across the transform zone with tectonic slabs of the Kings River ophiolite forming a median ridge that faced into an axial deep along which pelagic and hemipelagic oozes of the Calaveras complex accumulated. During this time interval the Kings-Kaweah ophiolite belt was undergoing progressive deformation into ophiolitic mélange by transform shearing and large-magnitude displacements, and progressive serpentinite diapirism. The Calaveras complex is also being progressively deformed from soft sediment to lithification states by transform shearing as well as by slumping driven by vertical tectonism and conceivably transform seismicity. Viewing Figure 16C in a regional context, displacements related to the growth of the Foothills ophiolite belt microplate
are shown circuiting into the California-Coahuila transform and displacing the Caborca block. Such transform motion conceivably continued farther into the Rheic suture zone that transected the interior of Mexico, and which accommodated the impingement of Gondwana with the southern margin of Laurentia (Nance and Linnemann, 2008). Ribbons of slope-rise facies strata of the Cordilleran passive margin and superimposed southernmost McCloud arc rocks were entrained in the transform zone and accreted along the continental truncation zone inboard of the encroaching abyssal lithosphere. The autochthonous sloperise facies system lying inboard of the truncation zone, which also constituted part of a marginal basin behind the McCloud arc, underwent an unknown amount of shortening along the Golconda thrust belt during the later Permian phases of transform activity. This in part may account for the disproportionate along-strike dimensions of the strike-slip ribbons as compared
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with the current across-strike width of the slope-rise facies belt, compounded by additional dispersal of the ribbons along Mesozoic dextral faults (Fig. 15). Golconda thrust belt convergence is interpreted to have been driven by a reversal in the McCloud arc, which, based on the disruption of eastern Klamath arc activity and depositional overlap by the shallow-water reefal McCloud limestone, occurred in Early to Middle Permian time (Irwin, 1977; Stevens, 2009). A possible mechanism shown for driving the reversal was the collision of a major ocean island chain that formed in Panthalassa low latitudes, the remnants of which are dispersed along the northern Cordillera as the Cache Creek terrane, and from which shallow-water limestone blocks and clasts were derived and regionally reworked into early Mesozoic hemipelagic and volcaniclastic deposits (Davis et al., 1978; Ross and Ross, 1983). A number of Permo-Carboniferous events along the inboard wall of the transform system reflect extensional and/or transtensional deformation interpreted to be kinematically linked to the California-Coahuila transform regime. These events include extensional exhumation of the southern end of the McCloud arc subduction complex (Barrow and Metcalf, 2006), the termination of McCloud arc–related volcanism in the northernmost Sierra Nevada with subsidence that promoted overlapping hemipelagic sedimentation (D’Allura et al., 1977), the development of a borderland facies belt oriented along the truncated edge of the passive margin shelf (Stone and Stevens, 1988), and widespread extensional tectonism in the passive margin east of the truncation zone (Smith and Miller, 1990). This extensional-transtensional regime was abruptly overprinted by Golconda thrust belt deformation, and then later by Permo-Triassic sinistral transpression, interpreted below to mark initiation of subduction along the transform zone. The simplistic plate geometry adopted in Figure 16 offers an explanation for an enigma regarding the Calaveras complex belts of the Sierra Foothills. Permo-Carboniferous reefal limestones that formed on the Cache Creek ocean island chain are biogeographically distinct from the Permian McCloud arc capping reefal limestones, and both are distinct from coeval shelf limestones of the Cordilleran passive margin (cf. Ross and Ross, 1983). Blocks and slabs of the Cache Creek–type limestones are common in Permo-Triassic chert-argillite belts accreted to the margin of the McCloud arc, stretching from the Sierra Foothills to the southern Yukon (Davis et al., 1978). However, in the Calaveras complex belts of the Sierra Foothills blocks and clasts of both Cache Creek and McCloud-type limestones are present. This is the only region where fragments of both limestone types are present in the same rock assemblages, which is in line with the Figure 16C reconstruction whereby blocks of each were readily sourced in submarine landslides from the inboard wall of the transform system and delivered into the Calaveras depositional trough(s). The Foothills ophiolite belt microplate appears to have been juxtaposed with the truncated SW Cordilleran margin in the latest Permian (Fig. 16D). As final juxtaposition progressed, Panthalassa
lithosphere began converging with the truncated Cordilleran margin, and a new subduction zone nucleated. Permo-Carboniferous “abnormal” oceanic lithosphere of the Foothills ophiolite belt microplate was susceptible to accretion to the hanging wall of the new subduction zone owing to its buoyancy that resulted from widespread serpentinization. According to the Figure 16 model the majority of the impinging oceanic lithosphere was old (≥200 Ma) and consisted of “normal” ridge crest–generated lithosphere of Kings River ophiolite affinity. Juxtaposition of aged cold oceanic lithosphere with the buoyant “abnormal” oceanic lithosphere ribbon presented ideal circumstances for subduction initiation (cf. Stern, 2004). The ca. 255 Ma Sm/Nd age for the Kaweah serpentinite mélange garnet amphibolite block is taken as the approximate time of subduction initiation. High-pressure metamorphism is envisaged to have occurred along the neosubduction megathrust with the high temperature (~800 °C) of metamorphism recorded in the garnet amphibolite block possibly reflecting the subduction of a warm mid-Permian segment of the Foothills ophiolite belt microplate outer edge. Fragments of the high-P metamorphic selvage that formed along the neo-subduction megathrust were entrained in serpentinite diapirs and transported up to crustal levels in the proto-forearc wedge, possibly erupting in mud volcanoes. The d–d′ cross section diagrammatically shows the en masse accretion of the Kings-Kaweah ophiolite belt to the hanging wall of a neo-subduction zone and its juxtaposition with the passive margin, para-autochthonous strike-slip ribbons along the Foothills suture. The Kings River ophiolite transform median ridge is shown hypothetically persisting as a bathymetric high forming a proto-forearc ridge, facing inward to the Calaveras axial deep sediment wedge that was likewise accreted en masse along the Foothills suture. Serpentinite diapirs are shown, sourced from the subduction megathrust zone where they entrained fragments of the high-P metamorphic selvage that formed along the neo-megathrust. The upper plate of the neo-subduction zone responded to oblique convergence by sinistral transpression as recorded by Permo-Triassic east-directed thrusting of the Sierra Nevada– Death Valley thrust system (Stevens and Stone, 2005) and similarage thrust structures of the Mojave Desert region (Miller and Sutter, 1982; Walker, 1988), now highly dispersed in the northwest Mojave region by large-magnitude Late Cretaceous extension (Chapman et al., 2011). Partly coincident with Permo-Triassic transpression, relatively low-volume arc plutonism initiated along the trace of the truncation zone (Stern et al., 1981; Chen and Moore, 1982; Barth et al., 1997; Saleeby and Dunne, 2011). Such arc plutonism is recorded for end of Permian–Early Triassic time, whereas Late Triassic and earliest Jurassic arc activity is recorded by widespread voluminous silicic ignimbrites, some of which were ponded in large submarine calderas that were nested in a regional arc graben system (Fiske and Tobisch, 1978; BusbySpera, 1984, 1988; Schweickert and Lahren, 1989). The temporal and spatial relations of supra-subduction zone magmatism, as presented in the Figure 16 model, seem at odds with the generalized subduction initiation model of Stern (2004).
Geochemical mapping of the Kings-Kaweah ophiolite belt, California In the Figure 16 model a relatively low-volume calc-alkaline to locally shoshonitic arc was established along the eastern Sierra region before Late Triassic–earliest Jurassic boninitic-arc tholeiitic volcanism occurred along the Foothills belt. In the Stern (2004) model, which is based on the Izu-Bonin–Mariana arc system, such proto-forearc volcanism develops in response to profuse slab rollback as negatively buoyant lithosphere founders during subduction initiation (cf. Hall et al., 2003) and the upper plate of the new subduction zone undergoes regional extension. This form of subduction initiation is termed spontaneous and in theory arises primarily from buoyancy contrasts across transform faults in abyssal lithosphere. Stern (2004) also presents an end member alternative to spontaneous initiation that is termed induced, which results from far-field plate forces rendering the necessary component of convergence for the initiation and sustenance of subduction. The Macquarie Ridge–Puysgur Trench– Fiordland plate juncture system is cited as an example of ongoing induced subduction initiation, and emphasis is placed on the importance of upper plate compressive deformation, opposed to extension, as the sign of induced initiation. The Permo-Triassic subduction initiation event of the SW Cordillera resembles the induced subduction initiation regime that is occurring today along the Macquarie Ridge–Puysgur Trench– Fiordland system. There the transform system that is evolving into a new subduction zone extends from abyssal lithosphere into passive margin–type lithosphere of the Campbell-Challenger plateau, where transpressive deformation is in large part producing the New Zealand landmass. As of yet, however, this system has not produced any known boninitic-arc tholeiitic volcanic associations in an extensional proto-forearc setting. Returning to the Permo-Triassic event of the SW Cordillera, subduction initiation appears to have been in response to far-field oblique compressive forces, as evidenced by regional transpressive deformation of the upper plate. Perhaps as subduction progressed, the convergence trajectory changed to a stronger normal component, whereupon negative buoyancy forces in the aged downgoing slab began to dominate the dynamics of the system, and a phase of profuse slab rollback ensued. The result of this rollback was the inflow of asthenosphere into the forearc mantle wedge, promoting boninitic magma genesis at shallow levels and arc tholeiitic magma genesis at deeper levels. The already heated axial region of the arc responded to regional extension by changing the mode of arc magmatism to a series of dispersed large-volume submarine calderas, as opposed to a chain of calc-alkaline to shoshonitic plutons, which were conceivably overlain by andesitic strata-cones. The common occurrence of silicic ash components within the siliceous argillites that the Kaweah serpentinite mélange coverstrata mafic volcanics were erupted into is consistent with distal large-volume silicic eruptions occurring intermittently during the forearc region mafic volcanism. The relationships outlined above suggest that the Permo-Triassic subduction initiation regime along the SW Cordillera has elements of both induced and spontaneous initiation, as defined by Stern (2004), and thus appears to be a hybridization of the process as defined.
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Initiation of Franciscan Subduction The regional geology of California is distinguished by the late Mesozoic–early Cenozoic Franciscan subduction complex, which produced the California Coast Range basement regime. Regional tectonic relations, as well as the history of the Coast Range ophiolite and the initial high-pressure metamorphism recorded in the Franciscan complex, call for a subduction initiation event in the Middle Jurassic (Stern and Bloomer, 1992; Anczkiewicz et al., 2004; Shervais et al., 2005). The timing of this event corresponds well with the onset of a high–magmatic flux event of late Middle to Late Jurassic age that is recorded within the early-stage Sierra Nevada batholith and its metavolcanic pendants (Stern et al., 1981; Chen and Moore, 1982; Snoke et al., 1982; Ducea, 2001). Such a subduction initiation event calls for a major change in the plate tectonic regime of the SW Cordillera in the time interval between the production of the Triassic arc belt and the late Middle to Late Jurassic belt. This time interval corresponds to a distinct lull in early Mesozoic arc magmatism of the SW Cordillera (Stern et al., 1981; Chen and Moore, 1982; Barth et al., 1997; Saleeby and Dunne, 2011), dextralsense transpression in the region (Saleeby et al., 1992; Saleeby and Dunne, 2011), and to unconformities in the forearc region as recorded in the Kaweah serpentinite mélange cover strata and the Jasper Point–Penon Blanco sequence. In absolute age this lull occurred between 200 and 175 Ma. The explanation favored here for these ill-defined events is that they record the dextral-sense oblique collision of the Insular Superterrane, a major composite arc terrane of peri-Gondwana affinity that impinged northward and accreted into the northern Cordillera region in the Middle to Late Jurassic, and which was progressively compressed and accreted into the northern Cordillera in the Early Cretaceous (Davis et al., 1978; McClelland et al., 1992; Getty et al., 1993). As the Insular Superterrane slid northward along the SW Cordilleran margin toward its northern Cordillera accretion site it left a leaky transform–spreading ridge system in its wake along which the Coast Range and other southern Cordillera Middle Jurassic ophiolites formed (Saleeby et al., 1992). Nucleation of Franciscan subduction followed in the wake of Insular Superterrane northward migration along the leaky transform system. CONCLUSIONS The Kings-Kaweah ophiolite belt underwent two distinct phases of abyssal MORB magma genesis. The first phase was at 484 ± 18 Ma (Early Ordovician), which generated a complete abyssal crust and depleted upper mantle sequence probably along a fast-spreading ridge in the Panthalassa ocean basin. The second phase is constrained by 295 ± 15 U/Pb zircon ages on rare felsic intrusives, and by a 299 ± 32 Ma Sm-Nd isochron age on a wide range of mafic to rare felsic rocks. This Permo-Carboniferous phase of abyssal magmatism occurred along a large offset transform fracture zone. During this phase of magmatism and tectonism a part of the Early Ordovician abyssal lithosphere that
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was proximal to the fracture zone was disrupted into an oceanic metamorphic core complex and was cut into a series of elongate slabs by transform-related ductile shear zones. Deformation progressed within the transform zone to the state of producing mélange. Mélange mixing was facilitated by copious serpentinization and related diapiric mobilization, some of which surfaced onto the seafloor. Pillowed abyssal tholeiite eruptions, ophicalcites, and radiolarian oozes mingled with serpentinite debris along fault scarps and debris aprons. As deformation and pelagic sedimentation progressed, biogenic oozes were mixed and ultimately replaced by siliceous argillite derived from terrigenous material and distally derived volcanic ashes. The hemipelagic deposits were mobilized by submarine mass wasting and accumulated along axial troughs within the transform system. Regional tectonic relations of the SW Cordillera indicating a late Paleozoic transform truncation of the NeoproterozoicPaleozoic passive margin may be directly linked to the transform history of the Kings-Kaweah ophiolite belt. Direct kinematic linkage between the oceanic transform system and the passive margin truncation zone led to the juxtaposition of the ophiolite belt with a veneer of para-autochthonous strike-slip ribbons that together were accreted along the truncation locus. Emplacement of the ophiolite belt occurred by its en masse accretion to the hanging wall of a new subduction zone along the preexisting transform juncture starting at ca. 255 Ma. Related arc magmatism began in the eastern Sierra by ca. 248 Ma. Fragments of high-pressure oceanic metamorphic rocks that formed along the new subduction zone were locally entrained by serpentinite diapirs that intruded up into the overlying ophiolite belt, which introduced the exotic blocks into the mélange assemblage. In Late Triassic to earliest Jurassic time, fluid-assisted melting of depleted peridotites accreted to the hanging wall of the subduction zone rendered boninitic to arc-tholeiitic mafic magmas that were erupted within and above the ophiolite belts’ hemipelagic cover strata. Later in the Jurassic, siliciclastic turbidites spread across the KingsKaweah ophiolite belt as the forearc region matured. Detritus for the turbidites was supplied mainly by exhumed early Paleozoic passive margin–related strata and by early Mesozoic arc rocks of the eastern Sierra Nevada region. By late Middle to Late Jurassic time (170–148 Ma) small calc-alkaline plutons and dike swarms were emplaced into the Kings-Kaweah ophiolite belt within a transtensional deformation field, and by ca. 125 Ma copious gabbroic to tonalitic plutons of the Sierra Nevada batholith intruded and pervasively contact metamorphosed the ophiolite belt. Regional relations of the SW Cordilleran truncation and subduction initiation event, as well as the geologic history of the Kings-Kaweah ophiolite belt, suggest that subduction initiation was driven primarily by far-field plate tectonic forces and involved, first, highly oblique subduction whose tangential components were directly inherited from the prior transform phase of motion. Once subduction had been ongoing, and as the subduction trajectory evolved to a stronger normal component, aged cold, early Paleozoic abyssal lithosphere that was entering the trench began to founder, leading to a distinct slab rollback phase. This
resulted in regional extension across the arc and forearc environment, leading to a reorganization of arc magmatism to scattered large-volume submarine calderas and to dispersed boninitic-arc tholeiitic volcanism in the forearc region. In this context the early Mesozoic SW Cordilleran subduction initiation event resembles a hybridization between far-field-controlled “induced subduction initiation” and a more local buoyancy force–controlled “spontaneous subduction initiation,” as defined by Stern (2004). ACKNOWLEDGMENTS The author is indebted to C.A. Hopson for decades of inspiration in pursuit of addressing the ophiolite problem. Conversations with P.D. Asimow, J.M. Eiler, M.C. Gurnis, J.W. Hawkins, C. Şengor, J.W. Shervais, and R.J. Stern were of great value for the data interpretation presented. Direct assistance, acquisition of spikes and standards, and helpful tips in radiogenic isotopic geochemistry from J.G. Wasserburg, D.A. Papanastassiou, H. Ngo, and J.C. Chen are kindly acknowledged. Additional assistance in various aspects of the geochemical analytical work presented here by T. Bunch, A. Chapman, and M. Ducea are kindly acknowledged. Assistance in database compilation of Great Valley basement cores and technical drafting by Z.A. Saleeby is acknowledged. Assistance in mineral separation procedures by M. Chaudhry and I. Saleeby was essential for this study. The author thanks Harvey and Bobby Ruth for their hospitality while doing fieldwork, and their assistance in fieldwork logistics. Critical reviews by E.A. Miranda, K.D. Putirka, and J. Wakabayashi were a great asset. A grant from the Gordon and Betty Moore Foundation helped bring this study to completion. (Caltech Tectonics Observatory Contribution no. 119.) REFERENCES CITED Ague, J.J., and Brimhall, G.H., 1988, Magmatic arc asymmetry and distribution of anomalous plutonic belts in the batholiths of California: Effects of assimilation, crustal thickness, and depth of crystallization: Geological Society of America Bulletin, v. 100, p. 912–927, doi:10.1130/0016 -7606(1988)100<0912:MAAADO>2.3.CO;2. Anczkiewicz, R., Platt, J.P., Thirlwall, M.F., and Wakabayashi, J., 2004, Franciscan subduction off to a slow start: Evidence from high precision Lu-Hf garnet ages on high grade blocks: Earth and Planetary Science Letters, v. 225, p. 147–161, doi:10.1016/j.epsl.2004.06.003. Anderson, T.H., and Silver, L.T., 1979, The role of the Mojave-Sonora megashear in the tectonic evolution of northern Sonora, in Anderson, T.H., and RoldanQuintana, J., eds., Geology of Northern Sonora: Hermosillo, Sonora, Sonora Instituto de Geologia, Universidad Autonoma de Mexico, p. 59–68. Aumento, F., Loncarevic, B.D., and Ross, D.I., 1971, Hudson geotraverse: Geology of the Mid-Atlantic Ridge at 45°N: Royal Society of London Philosophical Transactions, ser. A, v. 268, p. 623–650. Babcock, J.M., Harding, A.J., Kent, G.M., and Orcutt, J.A., 1998, An examination of along-axis variation of magma chamber width and crustal structure on the East Pacific Rise between 13°30′N and 12°20′N: Journal of Geophysical Research, v. 103, p. 30,451–30,467, doi:10.1029/98JB01979. Baines, A.G., Cheadle, M.J., Dick, H.J.B., Scheirer, A.H., John, B.E., Kusznir, N.J., and Matsumoto, T., 2003, Mechanism for generating the anomalous uplift of oceanic core complexes: Atlantis Bank, Southwest Indian Ridge: Geology, v. 31, p. 1105–1108, doi:10.1130/G19829.1. Baines, A.G., Cheadle, M.J., John, B.E., and Schwartz, J.J., 2008, The rate of oceanic detachment faulting at Atlantis Bank, SW Indian Ridge: Earth and Planetary Science Letters, v. 273, p. 105–114.
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Geochemical mapping of the Kings-Kaweah ophiolite belt, California Stevens, C.H., and Greene, D.C., 2000, Geology of Paleozoic rocks in eastern Sierra Nevada roof pendants, California, in Lageson, D.R., Peters, S.G., and Lahren, M.M., eds., Great Basin and Sierra Nevada: Geological Society of America Field Guide 2, p. 237–254. Stevens, C.H., and Stone, P., 2005, Structure and regional significance of the Late Permian(?) Sierra Nevada–Death Valley thrust system, east-central California: Earth-Science Reviews, v. 73, p. 103–113, doi:10.1016/j .earscirev.2005.04.006. Stevens, C.H., Stone, P., and Miller, J.S., 2005, A new reconstruction of the Paleozoic continental margin of southwestern North America: Implications for the nature and timing of continental truncation and the possible role of the Mojave-Sonora megashear, in Anderson, T.H., Nourse, J.A., McKee, J.W., and Steiner, M.B., eds., The Mojave-Sonora Megashear Hypothesis: Development, Assessment, and Alternatives: Geological Society of America Special Paper 393, p. 597–618. Stone, P., and Stevens, Ch., 1988, Pennsylvanian and Early Permian Paleogeography of east-central California; implications for the shape of the continental margin and the timing of continental truncation: Geology, v. 16, p. 330–333, doi:10.1130/0091-7613(1988)016<0330:PAEPPO >2.3.CO;2. Sun, S., and McDonough, W.F., 1989, Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes, in Saunders, A.D., and Norry, M.J., eds., Magmatism in the Ocean Basins: Geological Society [London] Special Publication 42, p. 313–345. Taylor, R.N., Nesbitt, R.W., Vidal, P., Harmon, R.S., Auvray, B., and Croudace, I.W., 1994, Mineralogy, chemistry, and genesis of boninite series volcanics, Chichijima, Bonin Islands, Japan: Journal of Petrology, v. 35, p. 577–617. Terra, F., and Wasserburg, J.G., 1972, U/Pb systematics in lunar basalts: Earth and Planetary Science Letters, v. 17, p. 65–78. Tilton, G.R., Hopson, C.A., and Wright, J.W., 1981, Uranium-lead isotopic ages of the Samail ophiolite, Oman, with applications to Tethyan ocean ridge tectonics: Journal of Geophysical Research, v. 86, p. 2763–2775, doi:10.1029/JB086iB04p02763. Tucholke, B.E., Behn, M.D., Buck, W.R., and Lin, J., 2008, Role of melt supply in oceanic detachment faulting and formation of megamullions: Geology, v. 36, p. 455–458, doi:10.1130/G24639A.1. Van Andel, T.H., von Herzen, R.P., and Phillips, J.D., 1971, The Vema fracture zone and the tectonics of transverse shear zones in oceanic plates: Marine Geophysical Research, v. 1, p. 261–283.
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MANUSCRIPT ACCEPTED BY THE SOCIETY 21 DECEMBER 2010
Printed in the USA
The Geological Society of America Special Paper 480 2011
Constraints on the evolution of the Mesohellenic Ophiolite from subophiolitic metamorphic rocks R. Myhill* Department of Earth Sciences, University of Cambridge, Downing Street, Cambridge, Cambridgeshire, CB2 3EQ, UK
ABSTRACT Narrow, discontinuous bands of high-grade subophiolitic metamorphic rocks, comprising predominantly amphibolite facies metabasites with rare metasediments, are observed at the contact between the complexes and subjacent mélanges of the Mesohellenic Ophiolite exposed in northwestern Greece. Both conventional and pseudosection thermobarometry have been used to yield estimated peak pressure-temperature (P-T) conditions of these tectonic sheets. Toward the leading edge of the ophiolite, subophiolitic rocks of the Vourinos Complex record peak metamorphic temperatures of 770 ± 100 °C. Pressures of 4 ± 1 kbar beneath the Vourinos are estimated on the basis of hornblende composition and are similar to the expected pressures from ophiolitic overburden. Beneath the exposed Dramala Complex, at the trailing edge of the ophiolitic body southwest of the Vourinos, estimated temperatures reached 800 ± 40 °C and 12.00 ± 1.27 kbar at the top of an apparent inverted metamorphic gradient imposed by discrete phases of accretion. High pressure assemblages beneath ophiolitic bodies imply exhumation relative to the overlying ophiolite. Estimated homologous temperatures in the upper plate are similar to those inferred for channeled exhumation during continental collision. Mineral assemblages lower in the Dramala sole indicate reduced temperatures and peak pressures. Similar pressures obtained within lower temperature sole rocks beneath Vourinos and Pindos suggest that a shallowly dipping thrust may have been responsible for obduction. Peak temperatures and pressures are in agreement with those estimated for secondary thrust propagation beneath a proto-arc after subduction in an intra-oceanic setting.
*
[email protected] Myhill, R., 2011, Constraints on the evolution of the Mesohellenic Ophiolite from subophiolitic metamorphic rocks, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 75–94, doi:10.1130/2011.2480(03). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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INTRODUCTION The emplacement of ophiolites onto relatively buoyant continental margins, and the style of deformation beneath these complexes of dense oceanic crust and upper mantle during obduction, are controversial topics. In addition to work on the ophiolitic bodies themselves, constraints can also be derived from subophiolitic rocks. Autochthonous sequences are commonly observed between rocks deformed little by ophiolitic emplacement and the ophiolitic complex itself. These can take the form of tectonically disrupted mélanges composed of a mixture of lithologies. Although many mélanges are described as having chaotic structure (Şengör, 2003), discernible sequences are commonly present. The origin of lithologies contained within the mélange, and the degree of deformation and lithologic mixing imposed by the emplacement of the ophiolite, are often difficult to resolve. Nevertheless, the presence of a mélange is extremely important, as its low strength provides an ideal substrate for ophiolitic movement, and lithologies within the mélange provide constraints on transport of the overlying thrust sheets. Directly beneath the ophiolite, higher grade metamorphic rocks can be observed, separated from the underlying mélange by an abrupt tectonic and metamorphic contact. These rocks are commonly referred to as metamorphic soles, and constitute highly deformed sheets or sequences of sheets (<500 m thick) found structurally beneath many ophiolites (Jamieson, 1986; Williams and Smyth, 1973). Fragments of these sheets of higher grade material may be incorporated into the mélange, some still attached to ophiolitic rocks, separated from the main ophiolitic contact by out-of-sequence thrusts (see subsection on ophiolitic mélanges of the Mesohellenic Ophiolite). The P-T conditions of metamorphic soles have been of significant interest to the academic community. Temperatures elevated above the mean temperature of ambient mantle and the surface suggest formation by dynamothermal heating beneath hot sub-oceanic mantle, analogous to moving a hot iron over cold sheets (Spray, 1980). The corollary of this heating is rapid cooling of the overlying mantle; the inverted thermal anomaly is removed within two million years (Hacker, 1990; Hacker et al., 1996). Recovered Ar-Ar dates from the soles are thus close to the dates for the inception of subduction (Spray, 1984). Early work on the uppermost metamorphic rocks from beneath many ophiolites estimated low peak metamorphic pressures. These studies were conducted in the absence of barometric calibrations, and pressures were calculated on the assumption of no exhumation relative to the ophiolite. With the advent of quantitative geothermobarometry, an increasing number of studies have presented data from metamorphic soles to infer burial depths greatly exceeding those of reconstructed ophiolitic overburden. This observation led Wakabayashi and Dilek (2000, 2003) to suggest that all ophiolitic “soles” have been exhumed relative to the overlying ophiolite. Inverted temperature gradients are often observed. Only a few studies have also shown an accompanying
inverted pressure gradient (e.g., Semail Ophiolite, Gnos, 1998; Jamieson, 1986), a necessary requirement to demonstrate successive accretion of high-grade metamorphic sheets to the base of a moving ophiolitic body during exhumation. The current study applies recent advances in both conventional and isochemical phase (or “pseudosection”) thermobarometry relevant to the metamorphism of basic rocks (Diener et al., 2007; Green et al., 2007), combined with spatial and structural information to more accurately constrain tectonic processes occurring during early emplacement of the Mesohellenic Ophiolite, Northern Greece. GEOLOGICAL SETTING The ophiolites exposed in Greece trend north-northwest in two primary zones (Fig. 1) known as the Innermost Hellenic Ophiolite Belt to the east (e.g., Smith, 1993), and the Western Hellenic Belt to the west. They are thought to have been obducted during closure of the Neotethys Ocean between the Jurassic and Early Cretaceous (Smith and Rassios, 2003). The eastern and western belts, and their extensions into the Mirdita ophiolites in Albania, Kosovo, and Serbia to the north, and farther into the Dinaric ophiolites in Bosnia and Croatia, are known as the Vardar and Pindos Zones, respectively (Dilek et al., 2007). The paired ophiolitic complexes exposed within the Western Hellenic belt are remnants of the Jurassic Pindos oceanic basin, a semi-independent seaway on one margin of the Neotethys (Rassios and Moores, 2006). The surface outcrops are separated by the Mesohellenic Trough, a north-northwest–trending Eocene– Miocene molassic basin (Vamvaka et al., 2006) that appears to mark the final suture created as a result of closure of this basin. On its western side lay the continent of Apulia, and on the east, the ribbon continent known as Pelagonia. NE-verging kinematic indicators imply a Pindos Basin source for the ophiolites (reviewed in Rassios and Moores, 2006). An alternative Vardar source lacks geological evidence other than the absence of a tomographically visible southwest-dipping subduction zone (van Hinsbergen et al., 2005). This observation is probably a result of the lack of a resolvable velocity contrast between isolated Jurassic slabs and the convecting mantle. The Vourinos Ophiolitic Complex rests tectonically on the Triassic and Jurassic passive margin of Pelagonia, represented by a thick, predominantly carbonate sedimentary sequence overlying various basement formations. Units within the complex have broadly supra-subduction zone chemical signatures, including a relatively thick cumulate sequence with abundant dunites, wehrlites, and pyroxenites with proto-arc characteristics (Dilek et al., 2007). The whole ophiolite is tilted about a subhorizontal axis, and differentially sheared and folded so that although the basal harzburgite in the east is paleo-horizontal, the ophiolite sequence dips west at ~70°, and the upper crustal dikes and lava sheets in the west are overturned (Rassios and Dilek, 2009). The ophiolitic complex is capped by a thin Jurassic pelagic carbonate sequence (Moores, 1969) and is unconformably overlain by Cenomanian
Constraints on the Mesohellenic Ophiolite evolution from subophiolitic metamorphic rocks limestone (Brunn, 1956), dipping west at ~40°. The whole complex is highly deformed but relatively undisrupted, displaying ~3 km of cumulates beneath an estimated 2 km of dikes and subsurface or eroded lavas, pseudo-stratigraphically overlying 13 km of mantle harzburgites and dunites (as estimated from a cross section by Rassios and Moores, 2006, their figure 4). The Pindos Ophiolite rests tectonically on the passive margin of Apulia, and is composed of the Dramala and Aspropotamos Complexes. These are far more disrupted than the Vourinos Complex, and are geochemically mid-ocean ridge (MOR) to supra-subduction-zone transitional in nature. The cumulate and lava sequences of the Aspropotamos Complex are structurally overlain by the Dramala Complex (Jones and Robertson, 1991; Robertson, 2004), an overthrust mantle suite of harzburgite and lherzolite (Rassios, 1991), pseudo-stratigraphically overlain by a 0.5–1.5-km-thick transitional Moho sequence containing
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dunitic and troctolitic cumulates (Rassios and Moores, 2006). Similarities between the Aspropotamos and Vourinos Complexes have led to the suggestion that the Aspropotamos was sourced from a region east of the Vourinos Complex (Jones and Robertson, 1991). A continuation of ophiolitic rocks beneath the Mesohellenic Trough is inferred on the basis of magnetic surveying of the basin (Memou and Skianis, 1993). The modeled body is likely to be a direct continuation of the Vourinos and Dramala Complexes (although deformed and broken by basin-bounding faults), and the ophiolite as a whole was recently referred to as the Mesohellenic Ophiolite (Rassios and Moores, 2006). Spatial variations between the two complexes mirror others within the Western Ophiolitic Belt, and may be used to constrain processes that occurred during supra-subduction zone spreading (Dilek et al., 2007).
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Ultramafic massifs (Jurassic) Ophiolites and sedimentary cover (Jurassic-Cretaceous) Pre-Apulian carbonate platform and flysch (Triassic-Eocene) Pelagonian Platform (Triassic-Cretaceous) Mesozoic continental margin of Eurasia 22°
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Figure 1. Simplified tectonic map of Greece, adapted from Robertson and Shallo (2000) and Dilek et al. (2007). Labels for the Western Hellenic ophiolites and their extensions into Albania (from north to south): M—Mirdita ophiolites; B—Bulqize; Shp—Shpati; She—Shebenik; Kr—Krasta; Vo— Voskopoja; V—Vourinos; P—Pindos; K—Koziakas; O—Othris; E—Euboea.
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Mélanges occur at the base of all the complexes of the Mesohellenic Ophiolite. These mélanges consist of a zeolite to lower greenschist-facies metasedimentary matrix containing blocks of other lithologies, which range in metamorphic grade from zeolite to upper amphibolite facies. Beneath the Pindos Complexes the subophiolitic mélange is the Avdella, found beneath both the Dramala and Aspropotamos Complexes (Fig. 2). Beneath the Vourinos Complex a unit distinct from the Avdella is observed. Whereas Zimmerman (1972) designated that unit a tectonic mélange, Naylor and Harle (1976) suggested that it was essentially a stratigraphic sequence and named it the Agios Nikolaos Complex. Both units have been interpreted as accretionary wedges formed during ophiolitic obduction onto the passive margins of Apulia (Avdella mélange, Jones and Robertson, 1991) and Pelagonia (Ghikas, 2007; Zimmerman, 1972). Metamorphic soles are observed beneath both the Vourinos and the Pindos Complexes, henceforth the Vourinos Sole and Loumnitsa unit (Jones, 1990) respectively. Both exhibit mineralogies indicative of upper greenschist to granulite facies.
THE MESOHELLENIC OPHIOLITIC SOLE AND MÉLANGES Ophiolitic Mélanges of the Mesohellenic Ophiolite The Avdella Mélange is a Jurassic accretionary wedge exposed structurally beneath the Pindos Ophiolitic Complex. In places exceeding 1 km in thickness, it consists of blocks of primarily Late Triassic to Middle Jurassic (Terry and Mercier, 1971) pelagic carbonates and cherts, within-plate basalt (WPB) and mid-oceanic-ridge basalt (MORB) lavas (Kostopoulos, 1989) and ophiolitic material within a matrix of incompetent shale, marl, and sandstone lithologies (Jones and Robertson, 1991). Metamorphism is generally of zeolite facies, with only deepburial diagenesis and minor alteration. Close to the contact with the metamorphic sole and ophiolite, however, lower greenschist facies metamorphism is reached (Jones and Robertson, 1991). Blocks within the matrix are up to 200 m in diameter. In places, more extensive thrust sheets up to 2 km thick are observed,
40°15
Zygosti-Rodiani Complex Skoumtsa
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Dotsikos Strip Complex
Chromio
Vourinos Complex
40°05 Vassilitsa
Frourio Subsurface ophiolitic units inferred from magnetic studies
Liagkouna Avdella Mélange
Zavordhas Mélange
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Agios Nikolaos Grias
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Extrusives Sheeted dikes Dykes Cumulates Mixed zones Zones (restite (Restiteand andcumulates) Cumulates) Mantle Unit units
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Figure 2. Overview map of exposed and modeled subsurface extent of the Mesohellenic Ophiolitic Complexes. Data derived from Rassios and Moores (2006) and from maps courtesy of A. Rassios and the Greek Institute of Geology and Mineral Exploration. A shadow has been applied to the Dramala section to illustrate the overthrust nature of the contact. Locations and units marked are as discussed in the text. The region around the Zavordhas Monastery, shown in Figure 3, is outlined in the southeast corner of the map.
Constraints on the Mesohellenic Ophiolite evolution from subophiolitic metamorphic rocks
Altitude /m a.s.l.
surrounded by little matrix (Jones and Robertson, 1991). Despite this block-in-matrix style, intact sequences can be observed within the thrust sheets, including WPB-affinity lava flows, olistostromes, and turbidites containing clasts of all Avdella lithologies. More details are given in other publications (Jones and Robertson, 1991; Ghikas et al., 2010, and references therein). Beneath the Vourinos Complex, the Zavordhas Mélange varies in thickness from zero to ~1 km (Zimmerman, 1972). It primarily consists of lower greenschist facies metasediments, sandwiched between the ophiolitic complex and underlying carbonates of the Pelagonian ribbon continent. These metasediments have the appearance of a matrix within which blocks of limestones, cherts, WPB-affinity lavas, and metatuffs are observed. In the Zavordhas region (Fig. 3), pervasive
A Emplacement thrust/ metamorphic sole
greenschist facies alteration is observed from actinolite-albitechlorite-epidote meta-tuffs of the upper thrust slices through to sericite-quartz-calcite-chlorite phyllites dominating the matrix, and into hornblende-pyroxene-plagioclase alkaline mafic rocks altering to actinolite, epidote, and albite, observed intruding carbonates and meta-muds of the Pelagonian shelf. The majority of the mélange thus appears approximately isofacial. At the top of the mélange, however, some lensoid blocks of serpentinized harzburgite sourced from the ophiolite are soled by amphibolite facies metamorphic rocks, including garnet-muscovite schists and hornblende-plagioclase-ilmenite amphibolites. Three such blocks are shown on the map in Figure 3 at 39°59′30″N, 21°47′00″ E. The apparent stacking of these blocks may develop either by sedimentary or tectonic processes. If erosion and
21°48
A′′
A′
600
21°47
400
H
200
Emplacement thrust
Basal mélange thrust horizontal scale = vertical scale
N W
79
P
I
40°00
40°00
21°46
E
H
Basal mélange thrust
S
C S
Zavordhas Monastery P
H ?
S
River Aliakmon
?
?
?
39°59
V
A
C
Se
ct
39°59
io
n
lin
e
21°48
21°47 A′′
P
Quaternary units 21°46
Subophiolitic units I Alkaline metabasic intrusion V Interbedded tuffs and sediments P Predominantly phyllitic sediments
Distances in meters
C Carbonates
0
500
1000
Amphibolites observed
Part-lithified carbonate debris Course pebble conglomerate Coarse Aliakmon alluvium
Ophiolitic units S Serpentinite H Harzburgite and dunite Dunite (Serpentinised) (serpentinized)
Figure 3. Simplified map and cross section of the Vourinos subophiolitic mélange in the region of the Zavordhas Monastery. In this area the mélange is bounded to the northwest by the discontinuous metamorphic sole and overlying ophiolitic complex, and to the southeast by the base of the predominantly carbonate rocks of the Pelagonian margin.
80
R. Myhill
sedimentation, followed by shearing, were the only processes resulting in the presence of exotic blocks within the mélange, then the presence of similarly oriented blocks (serpentinite soled by both metabasite and metapelite) in a small area is unexpected, because erosion of mantle and sole material requires removal of a large quantity of overlying ophiolitic material not seen within the mélange. In this particular situation, tectonic disruption at the base of the ophiolite appears more likely. On a broad scale, the Zavordhas Mélange comprises a series of narrow, stacked thrust sheets up to 10 km in length as determined by Naylor and Harle (1976), though the sheets themselves have also undergone a high degree of shearing and folding so that original relationships are somewhat obscured. Intense mylonitization was followed by tight to isoclinal folding from centimeter to hundred meter scales within the Zavordhas region. In the region north of the Skoumtsa Mine isoclinal folding is picked out by millimeter-scale laminations. Based on these observations, lithologic mixing beneath the Vourinos Complex appears to have both a sedimentary and tectonic nature. The broadly stratigraphic sequence of shales, tuffs, cherts, limestones, and olistostromes (siltstone with siltstone and limestone blocks), and an east-west increase in volcanic units of Naylor and Harle (1976), have been affected by intense shearing and out-of-sequence thrusting on 100-m-length scales. The preservation of intact sedimentary sequences, and the presence of thrust sheets and character of blocks within the distinct Avdella and Zavordhas Mélanges, suggest that the same processes controlled their formation. For more details, see Ghikas et al. (2010). Metamorphic Soles High-grade metamorphic rocks ranging from upper greenschist to granulite facies are found adjacent to the contact with the ultramafics of the ophiolite bodies. A thrust contact separating these high-grade rocks from the underlying greenschist facies rocks of the mélange can usually be accurately constrained in the field, and in the case of high-grade rocks found as blocks within the underlying mélange, sharp contacts can be observed on all sides of the block. These high-grade rocks have been previously observed and identified as the metamorphic sole of the overlying ophiolitic complexes (Jones, 1990; Moores, 1969; Spray, 1980). They are predominantly metabasic in composition, although metasediments in the form of garnet-mica schists and quartzites are also found. In the Vourinos, high-grade metamorphic units reach maximum thicknesses of ~20 m and lateral dimensions of ~500 m in the southern Zavordhas area, greater than the 4 m observed by Zimmerman (1972) but significantly less than the ~120 m thickness and ~1 km extent exposed at Liagkouna in the Pindos. There is no consistent stratigraphy preserved within the sole. Faulting is commonly observed as planes subparallel to parallel with both the axial planes of tight to ptygmatic folds and the ophiolitic contact, continuing well into greenschist facies rocks. The sole
can therefore be considered to be the uppermost disrupted thrust slices of the mélange. Thick soles commonly exhibit multiple sheets of metamorphic rocks ranging in thickness from sub-meter to tens of meters. Where observed, these show a general increase in grade toward the base of the ultramafic ophiolitic rocks, displayed in systematic changes in mineral assemblage and compositions, and in mechanisms of deformation (e.g., Gnos, 1998; Jamieson, 1986; Spray and Roddick, 1980). For the rocks of Liagkouna beneath the Dramala Complex of the Pindos Mountains, several internally deformed and folded thrust slices can be observed. Faulting and kink folding is prevalent within metasedimentary phyllites at the base of the section. Moving toward the ophiolitic contact, isoclinal folding and evidence of ductile deformation in boudinage and other shear fabrics have developed within both garnet-mica schist and amphibolite, revealing an apparent inverse metamorphic gradient supported by mineralogical changes. This inverse gradient was previously noted by others (Pichon and Brunn, 1985; Spray, 1980). Similar structures, including complex isoclinal and ptygmatic folding, are observed within metamorphic rocks beneath the Vourinos Complex. 40 Ar-39Ar dating from these amphibolite soles are within error of each other, at 165 ± 3 Ma for the Pindos and 171 ± 4 for the Vourinos (Spray, 1984). K-Ar dating (Thuizat et al., 1981) assigns an age of 176 ± 5 Ma to the Pindos sole. Zircon dating from plagiogranites within the ophiolite suite have yielded Bajocian ages of 171 ± 3 Ma for the Pindos and 168.5 ± 2.4 Ma and 172.9 ± 3.1 Ma for the Vourinos (Liati et al., 2004). Latestage igneous activity was therefore synchronous with the thrusting that formed the ophiolitic soles of the Vourinos and Pindos. Structures clearly indicate early displacement to the northeast, followed by later northerly thrusting, as observed by Jones (1990). These are in good agreement with emplacement directions of both the Vourinos and Pindos as analyzed by Rassios and Moores (2006). ANALYTICAL TECHNIQUES Compositions were determined by wavelength dispersive electron probe spectrometry, using a Cameca SX100 electron probe microanalyzer at the University of Cambridge. Operating conditions were 15 kV accelerating voltage, a 10 nA beam current, and a beam diameter of 10 μm. Representative mineral compositions are given in Tables 1 and 2. Bulk compositions were taken from previously published work (Spray, 1980; Spray and Roddick, 1980) for use in thermobarometric modeling. SAMPLE DESCRIPTIONS General Observations Sampled amphibolites are dominantly medium to coarse grained and schistose in nature (Fig. 4). Imperfect mineralogical banding is common (Figs. 4B and 4E), with separation of
2.45
13.3
0.11
0.92
18.18
0.26
8
10.96
2.22
1.04
97.71
TiO 2
Al2O3
Cr 2O 3
Fe2O3
FeO
MnO
MgO
CaO
Na2O
K2O
Totals
97.9
1
2.26
10.84
8.24
0.28
18.02
0.92
0.08
13.17
2.49
40.59
aw0
98.25
0.92
2.4
11.05
10.54
0.2
14.76
0.92
0.08
13.18
2.62
41.58
6.167
15.72
0.176
0.689
1.765
2.365
0.024
1.791
0.103
0.008
2.268
0.299
6.228
Pargasite
15.74
0.176
0.692
1.762
2.338
0.025
1.838
0.103
0.01
2.312
0.293
6.19
97.84
0.92
2.38
11.04
10.64
0.19
14.36
0.92
0.07
12.9
2.67
41.75
15.878
0.241
0.645
1.861
2.594
0.037
1.869
0.106
0.006
2.015
0.074
6.429
95.54
1.23
2.17
11.32
11.34
0.29
14.56
0.92
0.05
11.14
0.64
41.89
Pargasite
15.88
0.288
0.637
1.821
2.559
0.033
1.845
0.105
0.006
2.121
0.12
6.345
96.16
1.48
2.15
11.14
11.26
0.25
14.46
0.92
0.05
11.8
1.05
41.6
aw4
Pindos–Liagkouna
Notes: Locations are as found in Figure 1. Samples are as described in the text.
Ferropargasite
Nomenclature
15.74
15.75
0.194
0.667
1.765
1.866
0.036
2.289
0.105
0.01
2.36
0.284
Sum
0.202
K
0.033
Mn
0.656
2.318
Fe2
Na
0.105
Fe3
1.817
0.013
Cr
1.79
2.391
Al
Ca
0.281
Ti
Mg
6.143
Si
CATIONS calculated on the basis of 23 oxygens
40.29
SiO2
Sample
Locality
Mineral
15.91
0.29
0.654
1.836
2.553
0.034
1.857
0.106
0.006
2.138
0.101
6.335
95.72
1.48
2.2
11.17
11.17
0.26
14.47
0.92
0.05
11.82
0.88
41.3
Pargasite
15.611
0.091
0.549
1.745
2.796
0.025
1.533
0.103
0.006
2.213
0.069
6.48
95.83
0.48
1.9
10.9
12.56
0.2
12.27
0.92
0.05
12.56
0.61
43.38
15.439
0.051
0.45
1.762
2.936
0.029
1.408
0.102
0
1.881
0.055
6.765
95.93
0.27
1.57
11.13
13.33
0.23
11.39
0.92
0
10.8
0.5
45.79
Edenite
15.553
0.076
0.525
1.777
2.707
0.028
1.546
0.104
0.016
2.17
0.071
6.532
95.28
0.4
1.8
11.04
12.09
0.22
12.31
0.92
0.14
12.25
0.63
43.48
aw7
Amphiboles
15.56
0.269
0.331
1.901
1.98
0.041
2.024
0.106
0.005
2.451
0.105
6.351
95.99
1.38
1.12
11.61
8.69
0.32
15.84
0.92
0.04
13.6
0.91
41.56
Ferropargasite
15.52
0.248
0.311
1.972
1.961
0.035
2.012
0.106
0.007
2.341
0.135
6.395
96.03
1.27
1.05
12.05
8.61
0.27
15.75
0.92
0.06
13
1.18
41.87
vl6
15.73
0.171
0.675
1.78
2.343
0.026
1.849
0.104
0.016
2.276
0.288
6.205
97.52
0.89
2.32
11.07
10.48
0.2
14.73
0.92
0.13
12.87
2.55
41.36
Pargasite
15.71
0.169
0.67
1.762
2.399
0.024
1.783
0.104
0.009
2.258
0.267
6.262
97.2
0.88
2.31
10.97
10.73
0.19
14.22
0.92
0.08
12.78
2.36
41.76
bh0
Vourinos–Zavordhas
Magnesiohornblende
15.48
0.026
0.464
1.881
2.878
0.028
1.274
0.102
0.009
2.218
0.09
6.51
95.71
0.14
1.62
11.87
13.05
0.22
10.29
0.92
0.08
12.72
0.81
44
bh20
Magnesiohornblende
15.438
0.088
0.377
1.81
2.34
0.045
1.89
0.102
0.002
2.08
0.074
6.628
97.58
0.47
1.31
11.39
10.59
0.36
15.25
0.92
0.02
11.9
0.67
44.7
dn1
Chromio
TABLE 1. REPRESENTATIVE COMPOSITIONS AND DESIGNATED NAMES FOR AMPHIBOLES FOUND WITHIN THE SUBOPHIOLITIC METAMORPHIC ROCKS OF THE MESOHELLENIC COMPLEXES
Constraints on the Mesohellenic Ophiolite evolution from subophiolitic metamorphic rocks 81
99.80
100.82
0.00
0.00
0.00
7.73
MnO
MgO
CaO
aw0
99.69
0.12
7.28
7.74
0.00
0.01
0.00
0.14
0.00
25.58
0.00
58.82
98.71
0.10
9.10
4.41
0.00
0.00
0.00
0.05
0.00
22.85
0.01
62.19
aw7
98.30
0.08
9.32
4.27
0.00
0.00
0.00
0.02
0.00
22.65
0.01
61.94
99.56
0.00
1.28
21.59
11.50
0.32
8.61
3.21
0.03
3.13
0.19
49.70
100.26
0.06
11.86
0.24
0.01
0.00
0.00
0.08
0.00
19.20
0.00
68.81
100.05
0.01
0.01
9.57
5.13
1.19
23.21
1.75
0.01
20.94
0.12
38.29
bh0
aw0
100.45
1.17
11.44
0.18
0.14
0.00
0.00
0.42
0.00
19.87
0.01
67.22
vl6
99.50
0.05
11.81
0.03
0.00
0.01
0.00
0.21
0.00
19.23
0.00
68.16
Vourinos–Zavordhas
99.17
0.03
0.00
6.67
3.07
6.60
23.39
1.53
0.03
21.03
0.04
36.93
dn1
99.50
0.04
11.86
0.06
0.00
0.00
0.00
0.30
0.00
19.16
0.00
68.08
99.71
0.01
0.03
5.71
3.34
7.08
23.99
1.48
0.03
20.92
0.05
37.24
95.26
0.00
0.00
23.33
0.09
0.13
9.38
0.00
0.02
24.87
0.20
37.25
aw4
94.01
0.00
0.02
22.69
0.09
0.14
9.78
0.00
0.02
24.11
0.20
36.97
99.30
0.00
0.01
13.19
2.52
1.49
22.09
1.09
0.07
21.28
0.05
37.63
97.45
0.01
0.40
22.02
0.03
0.20
0.12
12.70
0.03
23.15
0.19
38.59
dq0
97.74
0.00
0.00
22.90
0.04
0.13
0.12
13.64
0.00
22.68
0.08
38.13
Vourinos–N of Skoumtsa
99.92
0.00
0.00
13.52
2.65
1.17
22.16
1.08
0.05
21.40
0.09
37.90
vl6
Vourinos–Zavordhas
Epidotes
100.52
0.00
0.01
7.70
1.76
10.29
21.86
0.00
0.02
20.84
0.20
37.84
Pindos– Liagkouna
100.30
0.00
0.02
7.87
1.76
10.35
21.41
0.36
0.02
20.69
0.19
37.66
ci3
Garnets Vourinos–Chromio Vourinos–Frourio
Albites
100.23
0.00
0.00
9.42
4.64
1.25
24.64
0.80
0.02
20.72
0.53
38.29
Pindos–Liagkouna
Notes: Locations are as found in Figure 1. Samples are as described in the text.
99.80
0.00
FeO
Totals
0.19
Fe2O3
7.28
0.00
Cr 2O 3
0.12
25.76
Al2O3
K2O
0.00
TiO 2
Na2O
58.71
SiO2
Sample
Pindos–Liagkouna
99.67
Totals
0.00
1.32
21.75
11.80
0.34
9.16
3.21
0.05
3.03
0.15
50.02
Locality
0.00
K2O
1.13
20.85
12.18
0.28
7.77
1.69
0.00
4.33
0.53
51.20
aw4
Feldspars
1.10
Na2O
aw0
Clinopyroxenes Pindos–Liagkouna
Mineral
21.05
7.34
FeO
CaO
2.06
Fe2O3
0.29
0.00
Cr 2O 3
12.23
4.21
Al2O3
MgO
0.51
TiO 2
MnO
51.06
SiO2
Sample
Mineral Locality
96.47
0.01
0.02
28.09
0.00
0.05
0.74
0.00
0.03
1.43
36.86
29.25
aw4
95.83
0.00
0.02
27.76
0.00
0.05
0.74
0.00
0.03
1.21
36.97
29.04
92.78
8.86
0.60
0.09
2.26
0.03
0.78
2.03
0.07
30.05
0.59
47.39
97.34
0.00
0.01
27.73
0.00
0.05
0.36
0.26
0.09
1.30
37.56
29.99
au1
98.89
0.01
0.49
27.21
0.05
0.03
0.29
0.21
0.03
1.83
37.73
31.02
Pindos–SW of Perivoli
93.90
10.12
0.81
0.08
1.84
0.04
1.44
1.76
0.12
31.91
0.37
45.57
Titanites
93.81
9.98
0.75
0.11
2.10
0.05
2.11
0.00
0.03
30.90
0.33
vl6
Micas Vourinos– Zavordhas
47.42
Pindos–Liagkouna
93.88
9.32
0.45
0.14
3.74
0.11
1.12
2.91
0.06
28.37
0.23
47.41
ci3
Vourinos–Frourio
TABLE 2. REPRESENTATIVE COMPOSITIONS OF MINERALS FOUND WITHIN THE SUBOPHIOLITIC METAMORPHIC ROCKS OF THE MESOHELLENIC COMPLEXES
82 R. Myhill
Constraints on the Mesohellenic Ophiolite evolution from subophiolitic metamorphic rocks
A
B
grt
83
amp cpx
cpx pl pl qtz ilm
0.5 mm
amp
ilm
0.5 mm
C
D
rt
amp
pl
amp cpx
ttn
ab ep
0.5 mm
0.5 mm
E
F
pl amp
ilm amp grt
chl
ab 1.0 mm
1.0 mm
Figure 4. Photomicrographs of assemblages and their structural and textural features within subophiolitic metamorphic rocks of the Mesohellenic sole. A–D: Liagkouna, beneath the Dramala Complex, Pindos: (A) Sample aw0, at the upper contact with mylonitic peridotites. Part of a garnet-pyroxenite boudin with large garnet porphyroblast separated from the amphibole-clinopyroxene-plagioclase-ilmenite assemblage by a shear zone of quartz-plagioclase with fragments of surrounding clasts. (B) Sample aw0: Plagioclase-clinopyroxene and ferropargasite-rich banding. (C) Sample aw4, ~80 m below the contact. Pargasite-plagioclase-epidote-sphene-amphibolite, with a narrow band of plagioclase-pyroxene-epidote. (D) Sample aw7, ~120 m below the contact. Pargasites and edenites are in equilibrium with plagioclase. Rutile and titanite are also present. (E) Sample dn1, Chromio, Vourinos. Intense isoclinal and ptygmatic folding in an amphiboleplagioclase-ilmenite-garnet (not pictured here) assemblage adjacent to the contact. The vergence and refolding of these folds indicates transport to the northeast. (F) Sample vl6, Zavordhas, Vourinos. A large garnet porphyroblast rimmed by ferropargasitic amphiboles. Plagioclase, mica, and ilmenite are also present, with chlorite as late alteration. Mineral abbreviations: grt—garnet; cpx—clinopyroxene; pl—plagioclase; ilm—ilmenite; amp—amphibole; qtz—quartz; ab— albite; ep—epidote; rt—rutile; ttn—titanite; chl—chlorite.
84
R. Myhill
amphibole-rich and plagioclase-rich bands on a millimeter to centimeter scale. The most common metabasic assemblage within the subophiolitic metabasic rocks is hornblende + plagioclase ± epidote ± titanite. Rarely, higher grade assemblages are observed, including the upper amphibolite assemblage clinopyroxene + hornblende + plagioclase + ilmenite (Fig. 4C). Rarely, garnet is observed within the higher facies rocks (Figs. 4A and 4F). Where the rocks have undergone little brittle deformation, garnet is associated with plagioclase-rich bands. Within the metasedimentary samples a common assemblage of garnet, white mica (muscovite and phengite), and quartz is commonly accompanied by chlorite. The samples exhibit varying degrees of deformation and post-deformational annealing, apparent in the partial overprinting of compositional banding and recovery of crystals from strain (especially for amphiboles within the metabasites). Brittle deformation is widespread throughout the studied units, apparent as minor shears to major faults. Alteration within the metamorphic rocks is pervasive. Reaction of plagioclase to albite (Figs. 4C and 4E) has occurred in almost all samples. Prehnite formation has been observed replacing calcic plagioclase. Amphibole alteration is also common, as green edenitic rims around brown hornblendes and pargasites, and chlorite, biotite, and actinolite growth around these crystals. Retrogressive chlorite and mica growth are common along shear planes. Overall, the amphibolite and granulite facies subophiolitic rocks observed exhibit many similarities to those found beneath other Tethyan-zone ophiolites (Dimo-Lahitte et al., 2001; Elitok and Drüppel, 2008; Gartzos et al., 2009). Amphibolite Petrography Although the majority of metabasic rocks from the Vourinos Sole are of lower amphibolite facies as previously described, higher grade rocks are observed. Sample vl6 (Fig. 4F) from the Zavordhas region contains a peak metamorphic assemblage of brown amphibole-plagioclase-garnet-titanite-ilmenite-quartz. Amphibole is observed as euhedral to subhedral crystals growing within bands and in garnet pressure shadows. Chlorite, biotite, and phengite are retrogressive phases, observed rimming amphibole and infilling cracks in garnet. Plagioclase is now present as albite, altering to muscovite. Sample dn1, from above the Agios Nikolaos valley, also contains sub- to euhedral garnets (2%) within a dominantly quartz (50%) and green amphibole (40%) matrix. Equant anhedral ilmenite and sphenoidal titanite strongly associated with amphibole are also observed as sub–100 µm crystals. From the Pindos region, collected samples are broadly similar to those observed from the Vourinos. Amphibole + plagioclase ± epidote ± titanite assemblages are observed at the Vassilitsa, Grias, and Perivoli localities (Fig. 2). Three samples from Liagkouna have been selected for further
analysis, collected 0.5 m, 80 m, and 120 m from the ophiolitic contact. The uppermost sample (aw0; Figs. 4A and 4B) contains a peak metamorphic assemblage of brown amphiboleclinopyroxene-plagioclase-ilmenite. Amphibole composes ~83% of the matrix, forming bands with ilmenite (1%) and clinopyroxene (4%) separated by plagioclase (12%). These feldspathic bands (which in other samples contain microcline) may attest to partial melting of amphibolite prior to or during shearing, similar to metamorphic rocks found beneath other ophiolites (e.g., Hacker, 1990; Searle and Cox, 1999). One plagioclase band formed by simple shear contains a boudin of garnet-pyroxenite. Clinopyroxene rims within this boudin have altered to amphibole similar in appearance to those in the hydrous matrix assemblage. Plagioclase within the boudin is altering to prehnite, but retrogressive zoning is limited. At 80 m below the contact (aw4; Fig. 4C), the peak metamorphic assemblage amphibole-plagioclase-titanite-ilmenite-quartz is observed. Amphibole again dominates the assemblage (58%), with equant plagioclase (37%) completely retrogressed to albite, altering to micas. Titanite is associated with amphibole. Apatite is present as a euhedral accessory mineral. Orthopyroxene and epidote are observed within a single plagioclase-rich band low in amphibole, and are not present within the rest of the matrix. Chlorite is present as a retrogressive phase along fractures within the sample. The third sample (aw7; Fig. 4D), structurally ~120 m beneath the ophiolitic complex, contains a peak assemblage of amphibole (75%)–plagioclase (20%)–rutile (3%)–quartz. Plagioclase is heavily, but not completely, altered to albite and mica. Rutile is rimmed by ilmenite, and included by plagioclase. Titanite, ilmenite, and epidote are also present in trace amounts as grain boundary phases and as inclusions within both amphibole and plagioclase. MINERAL CHEMISTRY Selected mineral compositions are provided in Tables 1 and 2. Further chemical characteristics of selected minerals are given below. Amphibole Amphibole in the metamorphic samples covers a range of compositions (Table 1). Clear trends are observed in compositional plots (Fig. 5), and where thick metamorphic sequences are observed compositional variations are a function of distance from the ophiolitic body. Amphiboles with consecutive sheets toward the base of the ophiolites generally increase in Ti content and decrease in Si. For example, at Liagkouna, amphiboles farthest from the ophiolitic contact are edenitic, becoming pargasitic and then ferropargasitic adjacent to mylonitic peridotites of the Dramala Complex. Some retrograde edenitic rims are observed around the pargasites.
Constraints on the Mesohellenic Ophiolite evolution from subophiolitic metamorphic rocks Clinopyroxene Diopsidic clinopyroxenes are rarely found within the amphibolite sequences. Where present at Liagkouna, they have low jadeite contents (Na = 0.07–0.1 atoms per formula unit [a.p.f.u.]). With increasing distance from the base of the ophiolite, Al and Mg contents decrease, and those of Ca and Fe increase.
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and Ti contents range from 6.08 to 6.76 and 0.00 to 0.11 a.p.f.u., respectively. The higher Si and lower Ti are associated with higher spessartine contents in their respective garnets, consistent with lower peak temperatures. P-T CONDITIONS AND METAMORPHIC HISTORY P-T Analysis
Garnet Garnet is rare within the sampled amphibolites, and absent within low-grade sole rocks. Only samples from Chromio beneath the Vourinos Complex, and Liagkouna beneath the Pindos, contain garnet. Both of these samples are relatively almandine rich (Fe2+ = 1.520–1.610 a.p.f.u. for both samples). Chromio garnets are rounded to angular fragments with moderate spessartine contents (Mn = 0.448–0.478 a.f.p.u.), whereas sub-euhedral Liagkouna garnet is lower in spessartine (Mn = 0.073–0.093 a.f.p.u.) and correspondingly higher in pyrope and grossular content. Sole metasediments also contain garnets. At the Frourio sole, garnets are observed within garnet-mica-chlorite schist. High almandine and moderate spessartine and grossular contents typify the rock, but garnets can have high spessartine contents (Mn to 1.675 a.f.p.u.) with significant reduction in grossular and almandine content. Beneath Zavordhas, high almandine garnets (Fe2+ = 1.918–2.114 a.p.f.u.) with low spessartine and moderate pyrope and grossular content (Ca = 0.214–0.466, Mg = 0.400– 0.527 a.p.f.u.) are present. Plagioclase Plagioclase is predominantly retrograde and highly albitic (XAb = 0.87–1.00). Rarely, more anorthitic plagioclase is observed. Beneath Liagkouna, fresh plagioclase increases in anorthite content toward the ophiolitic contact (XAb = 0.80–0.63). At the contact a boudin exhibits prehnite replacing plagioclase. Phengite Phengitic mica within the metasediments exhibits a wide range of compositions even where included within garnet. Si
Vourinos Samples In the Zavordhas region, both mid-amphibolite facies metabasites and metasediments are observed. The majority of metabasic rocks are deemed inappropriate for detailed thermobarometric study, owing to complete albitization of plagioclase. Equally inappropriate are many of the garnet mica schists because of reequilibration and alteration of micas. The semiquantitative metabasic thermobarometer of Ernst and Liu (1998) has been used to estimate conditions of amphibole formation. These result in pressures between 5 and 21 kbar and temperatures of 550–1000 °C. However, these are to be viewed with suspicion, as increasing pressures are positively correlated with the presence of titanite within hornblende crystals. Modeling
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Pressures and temperatures were estimated using both conventional thermobarometry and isochemical phase diagram modeling. The latter was undertaken with THERMOCALC 3.3 (tc330, Powell et al., 1998, recent upgrade) and an internally consistent thermodynamic data set (Holland and Powell, 1998, data set 5.5, November 2003 upgrade) in the system Na2O-CaO-FeOMgO-Al2O3-SiO2-H2O-TiO2-Fe2O3 (NCFMASHTO). Advances in accurately modeling activities of amphiboles and pyroxenes for this system (Diener et al., 2007; Green et al., 2007) mean that estimates of P-T conditions should be improved over previous studies. It should be noted that melting has not yet been satisfactorily modeled in metabasites (Powell and Holland, 2008), and that in such a complex system mineral stabilities may vary significantly with small variations in equilibration volume composition. P-T estimates are accompanied by ±1σ uncertainties. An excess-fluid pseudosection is provided for the Liagkouna locality (Fig. 6, sample 124659 from the Harker Collection, Spray, 1980).
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Figure 5. Compositional plots of Mg, Ti, and sum(A) against Si content of amphiboles from metabasic rocks exposed beneath the Mesohellenic Ophiolite, on the basis of 23 oxygen atoms per formula unit (a.p.f.u.).
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the compositional effect of 1%–3% retrogressive titanite growth from hornblende reduces pressures by up to 16 kbar. Samples without retrogressive titanite growth yield significantly lower pressures than those in which titanites are present using this thermobarometer (around 6 kbar), but they produce temperatures on the order of 100 °C greater than those estimated by other thermometers.
Low pressures (below 6 kbar) are implied by the semiquantitative hornblende barometer of Brown (1977). For rocks in which the buffering assemblage is present, strong clustering of amphibole compositions is seen ~4 ± 1 kbar, in agreement with overburden estimates for the Vourinos Complex. It is worth noting that the empirical calibration of this barometer may yield
14 hb di g ep ru
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hb di pl sph ilm hb ep pl n retio sph ilm ru aw4 g acc g durin n i l o o C
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T (°C) Figure 6. P-T pseudosection constructed in the system NCFMASHTO + H2O for sample 124659 from the Harker Collection, Cambridge, UK, collected from Liagkouna, beneath the Dramala Complex, Pindos (Spray, 1980). Composition (in mol%) is SiO2 = 48.61, Al2O3 = 8.74, CaO = 11.89, MgO = 11.82, FeO = 12.73, Na2O = 2.60, TiO2 = 2.43, O = 1.74 (Fe3+ directly from analyses). Mineral abbreviations: q—quartz; chl—chlorite; g—garnet; pl—plagioclase; hb—hornblende sensu lato; gl—glaucophane; o—omphacite; act—actinolite; ep—epidote sensu lato; sph—sphene; di—diopside; ab—albite; ilm—ilmenite; ru—rutile. Best fit P-T estimates for Liagkouna samples are shown.
Constraints on the Mesohellenic Ophiolite evolution from subophiolitic metamorphic rocks higher pressures if recalibrated against revised P-T estimates; in several other studies it has been shown to yield pressures lower than those of more modern techniques. Peak temperatures of 580 ± 35 °C are attained for garnethornblende pairs within the Chromion sole (dn1), and 670 °C (±35 °C) for the Zavordhas region (vl6), using the thermometer of Dale et al. (2000). Garnet-phengite temperatures have wide intra-sample spreads within most metapelitic samples, suggesting varying degrees of mica alteration and retrogression. However, mineral pairs within a ptygmatically folded Frourio garnetmica schist yield consistent temperatures of 770 ± 100 °C using the formulation of Hynes and Forest (1988). Slightly higher average temperatures of 820 ± 100 °C are obtained from Zavordhas samples. Observed alteration includes pyroxene hydration to pargasite, actinolite, and chlorite growth around amphiboles, and albitization and mica crystallization in plagioclase suggestive of reactions continuing at temperatures below ~550 °C (see Fig. 6 for albite in reactions). Pindos Samples The Pindos exposures of subophiolitic metamorphic rocks are predominantly of greenschist and low-amphibolite facies (e.g., at the Perivoli and Grias localities, Fig. 1). Beneath crustal rocks of the Aspropotamos (e.g., within the Agios Nikolaos Valley in the Pindos; Fig. 2) observed peak metamorphic conditions do not exceed greenschist facies. Metabasic sole rocks of the Pindos generally have amphiboles of edenitic composition, and plagioclases are completely albitic. Nevertheless, higher grade sole rocks are observed. At Liagkouna beneath the Dramala Complex, detailed analysis of several rocks through the sole has revealed an apparent inverse thermal gradient. This is most clearly displayed in amphibole compositions, which evolve from edenite at the base of the sole to magnesiohornblendes and pargasites, and then to ferropargasites adjacent to the mylonitic peridotites of the mantle sequence (Table 1). These chemistry changes correspond to an increase in grade from upper greenschist to upper amphibolite facies. The uppermost sole (aw0), sampled 0.5 m from mylonitic peridotite at the base of the Dramala Complex, preserves calcic plagioclase and is suitable for both conventional and pseudosection thermobarometry. Hornblende-plagioclase thermometry gives temperatures of 800 ± 40 °C (Holland and Blundy, 1994). Numerous indicators of partial melt extraction, including quartzofeldspathic veins and shear zones within otherwise quartzfree rocks, suggest temperatures >700 °C (Green, 1982) and in some cases a loss of hornblende within these zones (Liagkouna samples aw0 and aw4), suggestive of temperatures approaching 900 °C (Green, 1982; Rushmer, 1991). The high proportion of mafic minerals within the sample suggests that several percent of the rock mass may have been lost as a felsic melt. Using the modeled pseudosection for the locality (Fig. 6), the low plagioclase content of the sample (12%) is suggestive
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of pressures of ~12 kbar, greatly in excess of the ~5–7 kbar estimated by the semiquantitative thermobarometer of Ernst and Liu (1998), despite the assumed increased reliability owing to the lack of titanite. High pressures are also modeled using the average P-T mode of THERMOCALC (not bulk composition dependent) for a garnet-bearing pyroxenite boudin present within a quartzofeldspathic shear zone. This anhydrous assemblage yields an optimum P-T estimate of 650 °C and 9.70 ± 0.78 kbar for a water activity of 0.1. For a more reasonable temperature of 800 °C based on the hornblende-plagioclase thermometry, pressures of 12.00 ± 1.27 kbar are attained. Fits are worsened by including hornblende, as expected from the lack of prograde hornblende within the boudin. Increasing water activity also worsens the fit, while increasing the optimum temperature and pressure. A sample recovered from 80 m below the contact (aw4) is indicative of lower pressures than those undergone by the uppermost sole. The general lack of epidote within the pargasiteplagioclase-titanite-ilmenite peak assemblage, and the relative abundance of plagioclase, indicate that peak pressures did not exceed 8 kbar (Fig. 6). A thin metabasic unit at the base of the section (aw7) unusually contains remnants of unaltered plagioclase, enabling hornblende-plagioclase thermometry based on the equilibrium equation edenite + albite = richterite + anorthite (Holland and Blundy, 1994). For a nominal pressure of 5 kbar, this formulation yields a temperature of 625 ± 40 °C. The presence of rutile within the sample is also extremely uncommon, being commonly viewed as a high-pressure indicator within metabasic rocks. The modeled phase diagram provided for this locality (Fig. 6) shows a joint presence of rutile and titanite with plagioclase between 4 and 10 kbar. If the modeled bulk composition can be viewed as an accurate representation of this sample, the lack of significant epidote and ilmenite in the sample suggests that pressures remained ~5.5 kbar. The lack of clinopyroxene suggests that peak temperatures were below ~600 °C. Combining the temperature estimates for the high- and low-grade Liagkouna soles yields an apparent inverse metamorphic gradient of 175 °C over 120 m. Observations of albite and prehnite formation from initial plagioclase in the low- and hightemperature samples suggest that retrogressive reactions continued to temperatures <550 °C and 350 °C, respectively. Comparison with Previous Data from the Tethyan Ophiolites The first detailed thermobarometric analysis of the Tethyan soles was conducted by Spray (1980) at a time when techniques for metamorphic analysis of basic rocks were limited. Nevertheless, peak conditions from Othris were estimated to be 860 °C and ~7 kbar (Spray and Roddick, 1980). Reanalysis of one of the garnet-clinopyroxene-pargasite-prehnite (once plagioclase) samples studied (sample no. 124751 within the Harker Collection, Cambridge) has yielded peak temperature estimates of 870 ± 35 °C at a nominal pressure of 5 kbar based on garnet-clinopyroxene
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thermometry (Ravna, 2000b). Garnet-hornblende thermometry using the high-temperature calibrated formulation by Ravna (2000a) yields a temperature of 910 ± 50 °C, again within error of the garnet-clinopyroxene temperature estimate, and in agreement with semiquantitative hornblende thermometry. Modern pressure estimates cannot yet improve on the barometric estimate, but isochemical phase diagram analysis suggests that if the 7 kbar estimate is accurate, water activity was lowered during garnet growth, as for the Liagkouna sample. Hornblende remains stable to significantly lower water activities than for the chosen Liagkouna composition. Recent work has also revealed granulite facies peak temperatures up to 860 °C beneath the Albanian ophiolites (DimoLahitte et al., 2001). Peak pressure estimates of ~11 kbar in this work again greatly exceed those likely to be found beneath an ophiolite some 15–20 km thick, although other studies (M. Shallo, 2008, personal commun.) suggest that pressures of ~5 kbar may be more typical of these rocks. These results exhibit a striking similarity with results published here, in that while many of the metamorphic rocks record low pressures, some high-grade samples contain mineralogies indicative of significantly higher pressures. Petrographic and geochemical work has been conducted on the Lesvos and Evia metamorphic rocks (Gartzos et al., 2009). Although metamorphic analysis has not yet been conducted on these rocks, the lack of garnet and rutile within these metabasic rocks is suggestive of low pressures. Summary High temperatures beneath mantle sequences of the Tethyan ophiolites appear common and well constrained. Temperatures are similar for high-grade sub-mantle sequence rocks of the Vourinos and Pindos soles, which in turn are comparable to those beneath other Tethyan ophiolites. Metamorphic rocks at the base of ophiolitic crustal sequences in the Pindos are of greenschist facies, recording significantly lower temperatures than the metamorphic soles beneath mantle sequences. A steep inverse thermal gradient is observed beneath wellpreserved sub-mantle sequences, indicated by both thermometry and by the increase of titanium content within the amphibole cores (Searle and Malpas, 1982). These changes appear to be step-wise—that is, abrupt changes are observed in the composition and texture of the rocks, commonly delineated by faults and shear zones, suggesting that (as for many other soles) the inverse gradient observed is significantly steeper than the true metamorphic gradient once present beneath the ophiolite. These observations suggest that consecutive sheets within the ophiolite were accreted to the base of the ophiolite at different times and under different conditions. Although pressure estimates are significantly more uncertain than temperature estimates, observed assemblages and mineral chemistry suggest that pressures within the lower temperature sole remained low. Vourinos samples are broadly consistent with
4–6 kbar pressures, with garnets mostly confined to metasedimentary samples. Lower temperature samples beneath the Dramala Complex are also indicative of ~5 kbar pressures. However, the presence of the garnet-bearing pyroxenite within the sole is intriguing. High-pressure estimates of >10 kbar from this neargranulite-facies sample require an explanation other than that of “static” overburden of the ophiolitic complex. DISCUSSION Possible Setting of Sole Protolith Formation The dominance of metabasaltic rocks within the high-grade metamorphic units is in marked contrast to the Zavordhas and Avdella mélanges, where sediments are significantly more common, and so determining the setting of these rocks will help constrain large-scale geotectonic reconstructions. Vourinos and Pindos subophiolitic amphibolitic rocks can be divided into two distinct geochemical groups, which can also be distinguished on the basis of metamorphic grade (Jones and Robertson, 1991; Spray, 1980). Rocks of alkalic, within-plate-basalt (WPB) affinity are generally of lower metamorphic facies than those of MORB character. Evidence for a strong supra-subduction zone signature is lacking. Trace element compositions of Vourinos metamorphic sole rocks have WPB signatures, similar to greenschist facies metabasites within the Zavordhas mélange (Engwell, 2008). The implied setting of the high-grade subophiolitic metamorphic rocks is within an oceanic basin, away from significant terrigenous input, with lower grade metabasites derived from closer to the continental margin. The high-grade metasedimentary subophiolitic rocks are commonly quartz rich and may represent metamorphosed semi-pelites and cherts also derived from a sediment-starved basin environment. Similar affinities beneath the Vourinos and Pindos suggest the presence of a single fault above which the two complexes were initially obducted, prior to later disruption. P-T Conditions and Exhumation Temperature estimates from the high-grade metamorphic rocks studied here greatly exceed half the temperature of oceanic mantle above a subduction zone, so overriding of the mantle wedge during stable subduction cannot be the only cause of heating. To create these high temperatures invoking significant shear heating, recent upwelling is required, coupled with dynamothermal metamorphism from the moving of a hot ultramafic body over the tip of a newly initiated thrust (Spray, 1984), in agreement with the short time gap between ophiolite formation and cooling of the sole (Liati et al., 2004; Roddick et al., 1979). The low pressures recorded by most of the metamorphic sole are similar to those estimated by reconstructing Vourinos Ophiolite thickness. A folded and sheared Vourinos section (Rassios and Dilek, 2009) can be reconstructed, and, coupled with reasonable density assumptions can be used to obtain a pressure of
Constraints on the Mesohellenic Ophiolite evolution from subophiolitic metamorphic rocks ~4–5 kbar. The present-day thickness of the Pindos is only ~4 km (Rassios and Dilek, 2009), but tectonic disruption makes reconstruction of original thickness difficult. Nevertheless, it may be that all of the Liagkouna sole has undergone some exhumation to higher levels. The rarely observed higher pressure assemblages could not have been attained by simple overthrusting, and do appear to have required exhumation. High pressures within ophiolitic soles are not uncommon, and indeed it has been suggested that pressures exceeding those of ophiolitic overburden are ubiquitous in the geological record (Wakabayashi and Dilek, 2000, 2003). The mechanism for this exhumation is still poorly understood. Detailed analysis is beyond the scope of this paper, but observations from the Mesohellenic Ophiolite are suggestive of a regime similar to that of Himalayan exhumation between ca. 24 and ca. 18 Ma (White et al., 2002). The intense ductile deformation observed and related to early emplacement within the Vourinos Complex (e.g., Rassios et al., 1994) has been related to processes relevant at ca. 30 km depths (Ross et al., 1980), suggestive of lower mantle section exhumation. High homologous mantle temperatures indicated by peak temperatures within the sole, and the propagation of a thrust juxtaposing cold against overlying hot material, are both reminiscent of rocks exhumed beneath the South Tibetan Detachment System (e.g., Beaumont et al., 2001; Caddick et al., 2007; Jamieson et al., 2002). The tectonic geometry and scale of both systems are also similar. Inverse pressure gradients inferred from Liagkouna samples, and previously from the Semail Ophiolitic Sole (Gnos, 1998), suggest that exhumation was focused in the upper plate as for Himalayan exhumation, and high temperatures would place the uppermost sole in a similar deformational regime to the overlying peridotite. If exhumation at the base of ophiolitic complexes is indeed similar to channel flow, like that proposed for the Himalaya, erosion is necessary. There is ophiolitic material beneath the Cretaceous limestone on the Pelagonian margin, but no detailed studies have so far been conducted to determine its origins. Possible Tectonic Scenarios The presence of hot, shallow mantle above the lower plate of the thrust sampled by the metamorphic rocks is possible if the thrust propagated within the hanging wall of a preexisting subduction zone, where sub-arc upwelling resulted in temperatures sufficient (Kelemen et al., 2003) to induce near-granulite-facies metamorphism. In this case, burial to 20–40 km is not required to obtain amphibolite-facies metamorphism in the mantle wedge, as would be the case directly above an initiating subduction zone (Wakabayashi and Dilek, 2003). The alternative explanation of ridge subduction (Brown, 1998; Sisson and Pavlis, 1993) has also been proposed as a general model of sole formation (Shervais, 2001) and could similarly explain the high temperatures recorded. However, this explanation fails to account for the high pressures observed within many subophiolitic metamorphic rocks worldwide, or the presence of supra-subduction zone signatures within
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some soles (Encarnacion, 2004; Parlak et al., 2006; Pomonis et al., 2002; Saha et al., 2005; Wakabayashi et al., 2010). In the specific case of the Mesohellenic Ophiolite, it also fails to account for a crustal-level ophiolitic complex with supra-subduction zone signatures represented by the Aspropotamos Complex. Furthermore, pressures indicated by much of the Vourinos and Pindos soles suggest that mantle flow within the wedge above the thrust would have been strongly inhibited on time scales relevant to ophiolite obduction. Consequently, the observed volumes of supra-subduction zone-nature crust by 10%–20% melting of peridotite could not have been created. On this basis, a single scenario is favored, where a nascent arc developed above a subduction zone until reaching the buoyant continental margin, after which subduction stalled and a secondary thrust propagated through the overlying crust and upper mantle, facilitating obduction. The temperatures recorded by the sole suggest that this thrust propagated down through the crust and mantle before significant cooling had occurred. The high T, low P metamorphism common beneath the Mesohellenic Ophiolite is a good fit for subduction near a rift, and this is reflected in the favored model, with the emergent thrust located ~40 km from a proto-arc rifting environment. IMPLICATIONS Development of the Mesohellenic Ophiolite Processes occurring during the initiation of intra-oceanic subduction have recently been studied with numerical models. Although these models are rheology dependent, and as such prone to uncertainties, their application is of use in complex multivariable systems. Hall et al. (2003) showed that the initiation of subduction may result in rapid rollback of the lower plate after a few million years. This rollback may be favored further when initiated in warm regions, such as close to mid-oceanic ridges (Patel et al., 2007). The implications of this are that soon after the initiation of subduction, mantle melting can create a proto-arc and a marginal proto-arc above the subducting slab. Further combined metamorphic, thermodynamical, and rheological numerical modeling (Nikolaeva et al., 2008) showed that growth of new crust of a volume and areal extent comparable to that observed within the Mesohellenic Ophiolite could be formed within 5 m.y. of thrust inception and after ~500 km of subduction. Importantly, these models do not suggest significant mantle flow where the thickness of the mantle wedge is <40 km, and temperatures adjacent to the slab remain <500 °C to depths >50 km. In contrast, temperatures within the shallow mantle beneath newly forming crust attain granulite facies conditions within 5 m.y. In fact, temperatures at the depths of secondary thrusting calculated for the Vourinos and Dramala Complexes agree closely with those produced by the modeling by Nikolaeva et al. (2008). The observations are used to constrain a possible model for ophiolite evolution (Fig. 7). Rifting started ca. 240 Ma, and the Pindos Basin opened until it reached a width of perhaps 1000 km.
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Figure 7. Schematic cross sections of the favored mechanism for formation, obduction, early emplacement, and deformation of the Mesohellenic Ophiolite onto the Pelagonian Margin. There is no vertical exaggeration. The effects of erosion have not been shown, as these are currently not well constrained. (1) Passive rifting within the Pindos Basin continues until far-field effects put the entire basin into compression. Subduction then initiates near the cooling ridge. (2) During early subduction, upper plate extension induces mantle upwelling. Mantle melting creates the proto-arc-like crust of the Vourinos and the supra-subduction zone nature of the other complexes. On collision with the passive margin of Pelagonia, subduction stalls, and continued compression initiates a further thrust within the hanging wall close to the axis of the supra-subduction rift. Temperatures developed within the mantle wedge are taken from published numerical models (Nikolaeva et al., 2008). Exhumation above the thrust transports high-pressure rocks from >30 km depth (farther west than shown) to beneath Dramala. MORB— mid-oceanic-ridge basalt; HT—high temperature. (3) The complexes are thrust over the oceanic crust and onto Pelagonia. The Aspropotamos is overridden by a Tertiary thrust after significant cooling of the ophiolite. Further compression leads to polarity reversal of subduction, after which closure of the oceanic basin continues until the trailing end of the Mesohellenic Ophiolite is back-thrust onto the passive margin of Apulia.
Constraints on the Mesohellenic Ophiolite evolution from subophiolitic metamorphic rocks The cessation of sea floor spreading was probably related to the collision of Africa with Europe (Smith, 2006). Convergence resulted in intra-oceanic thrusting and westward-dipping subduction, possibly initiating close to the center of the basin. The usual arguments against subduction initiation by ridge collapse—that the sole is of a different composition from the overlying ophiolite (Searle and Cox, 1999; Shervais, 2001)—cannot be invoked here, as the subduction zone developed is not the source of the metamorphic sole. Extension in the upper plate during subduction resulted in the formation of thick crustal cumulates in a nascent-arc setting, with flanking regions developing MORB-island arc transitional crustal sequences. Vourinos is proposed to be the site of the proto-arc, with Dramala further to the west. Within the lava sequences, supra-subduction zone signatures increase stratigraphically upward and toward the developing arc. This evolution is observed both in the Mesohellenic Ophiolite and farther north into the Albanian Mirdita Complexes (Beccaluva et al., 2005; Dilek et al., 2007). Subduction continued until proto-subduction of Pelagonia transferred convergence to a secondary thrust in the upper plate, largely responsible for ophiolitic emplacement. Underthrust material is a potential source of fluid responsible for late-stage igneous activity. For example, boninitic dikes are thought to be produced by fluid sourced from a high-temperature slab, causing melting of already depleted mantle (Ishikawa et al., 2005). The metamorphic rocks studied exhibit abundant evidence for fluid flow within the uppermost underthrust slab, and this may have imparted a prominent chemical signature on the overlying ophiolitic rocks. A later emergent thrust resulted in underthrusting of the Aspropotamos, as proposed by Jones et al. (1991). The original location is constrained by the presence of only a crustal section. The supra-subduction zone chemistry of the upper lavas and low Ti dikes (Jones et al., 1991; Rassios and Moores, 2006) are more akin to the Vourinos Complex than the Dramala. This later thrust is specific to the Mesohellenic Ophiolite, and is not required to explain features of other Tethyan ophiolites. Formation of the Sole and Mélange The secondary thrusting in this model results in basaltic material being overridden by hot mantle sequences, causing mylonitization in the lowermost upper plate and hightemperature metamorphism in the uppermost lower plate where it is juxtaposed against mantle material. This material becomes the metamorphic sole and represents the earliest obduction of the overlying ophiolitic bodies. The broad division of “sole” compositions into two groups, the higher grade amphibolites of MORB and greenschist facies basic rocks of within-plate basalt (WPB) composition (Jones, 1990) may reflect their environments of formation and times of accretion. While the MORB rocks were accreted before the ophiolite overrode the subduction zone, WPB rocks may have been sourced from rift-related volcanics of
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the Pelagonian continental margin. The underlying greenschist facies mélange contains sandstone pebbles within a matrix absent in sandstones (Ghikas, 2007), implying that even the uppermost mélange may not have been transported from deep within the Pindos Basin. As such, the variations in metamorphic grade with composition record the transport and cooling of the overlying ophiolitic body. Kinematic indicators within the lowermost mantle sequence are in good agreement with those within the sole and the underlying mélange. Beneath the Vourinos ophiolite, ductile structures within the metamorphic “sole” are consistent with emplacement to the northeast. For the Liagkouna outcrop of the Pindos, previous workers undertook detailed structural analysis (Jones, 1990; Jones and Robertson, 1991). Ductile deformation in the form of strong isoclinal folding here indicates early motion of the ophiolite toward the northeast, along a vector of 040°, much the same as the Vourinos and in agreement with ductile-brittle fabrics within the Dramala Complex itself (Rassios, 1991). Thus, assuming that relative rotations about a vertical axis have not resulted in exactly opposite apparent motions, it can be providing the metamorphic rocks beneath both the Dramala and Vourinos Complexes were formed beneath the same fault system. Similar fabrics to those observed in the sole are also present within the underlying mélange. In the subophiolitic exposures beneath Vourinos, and within the ophiolite itself, folds verge away from the ophiolite exposures to the southeast near the Zavordhas Monastery, and northeast to the north of Skoumtsa, apparently related to footwall compression and emplacement of the ophiolitic body onto the margin of Pelagonia. Emplacement of the Ophiolitic Bodies Given that kinematic indicators within the lowermost mantle sequence are in good agreement with those within the sole and mélange, it is likely that the continued motion of the ophiolitic bodies was facilitated by mylonitic zones in peridotite and shearing and faulting within overridden lavas that formed the hightemperature sole. After sole formation, further emplacement onto the Pelagonian Platform was enabled by deformation within deposits on the passive continental margin. These deposits, which would have been increasingly dominated by unconsolidated sediments, became the sub-ophiolitic mélanges. Faulting would have been facilitated within these sediments by dewatering, which would have led to high pore-fluid pressures and encouraged brittle failure. This may explain how such narrow bands of high-grade metabasics could be preserved beneath the ophiolite during emplacement onto the continental margin. CONCLUSIONS Important constraints on the emplacement of ophiolitic complexes can be provided by field work coupled with modern metamorphic studies. The Mesohellenic Ophiolite provides a useful
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starting point for these studies, given its large spatial extent and the breadth of work previously carried out. The supra-subduction zone affinity Mesohellenic Ophiolite requires formation above a westward-dipping slab (Smith and Rassios, 2003) to form the proto-arc Vourinos and transitional Dramala Complex crust and upper mantle sequences. Although the slab subducting beneath these complexes would appear to be an ideal source for metamorphic soles, both the presence of a supra-subduction zone affinity crustal section complex and high peak temperatures recorded within lower pressure metamorphic rocks are strong evidence of sole formation in the uppermost lower plate of a secondary thrust propagating within the subduction mantle wedge. This thrust may have formed as a result of upper plate stresses developed during attempted subduction of the passive margin of Pelagonia. Initiation marginward of the Aspropotamos Complex facilitated much of the initial emplacement of the Mesohellenic Ophiolite from the Pindos Basin northeast onto the Pelagonian Platform. Later thrusting then resulted in the Zygosti-Rodiani, Vourinos, and then the Dramala Complexes overriding the Aspropotamos (Jones, 1990). The previously observed MORB-WPB and amphibolite-greenschist division of rocks beneath the complex (Engwell, 2008) can be explained by continued sole accretion during oceanic obduction and earliest margin emplacement. Some metamorphic samples collected are indicative of high pressures similar to those previously documented from beneath many other ophiolitic complexes (Wakabayashi and Dilek, 2000, 2003). High homologous mantle temperatures indicated by peak temperatures within the sole and the propagation of a thrust juxtaposing cold against overlying hot material are both reminiscent of rocks exhumed beneath the South Tibetan Detachment System (e.g., Beaumont et al., 2001; Caddick et al., 2007; Jamieson et al., 2002). It is therefore possible that the processes of exhumation may be similar in both tectonic environments. It is hoped that future studies will be able to further test these hypotheses beneath ophiolites with better preserved soles. Future advances in metamorphic study, and in particular the incorporation of titanium into hornblende models and development of a metabasic melting model (Powell and Holland, 2008), will greatly enhance our ability to obtain detailed pressure-temperature-time (P-T-t) histories from subophiolitic metamorphic rocks. ACKNOWLEDGMENTS This project would not have been possible without the help and encouragement of T. Holland of the University of Cambridge, who supervised this master’s project. The author is also indebted to A. Rassios for her invaluable assistance and many enjoyable discussions on the evolution of the Hellenic ophiolites. D. Ghikas, R. Sparkes, and the other participants in the Aliakmon River Project provided much appreciated support during fieldwork. K. Gray and C. Haywood gave valuable assistance in sample preparation and obtaining microprobe data. I thank J. Wakabayashi and E. Moores for detailed and thoughtful
reviews, which substantially improved this contribution. Further thanks are due to J. Wakabayashi for follow-up discussions. This research was supported by the Aliakmon River Project under the auspices of the Greek Institute of Geology and Mineral Exploration (IGME). Fieldwork was financially supported by Peterhouse, Cambridge, and the Public Power Corporation of Greece; and electron probe microanalysis was made possible by the Department of Earth Sciences, University of Cambridge. REFERENCES CITED Beaumont, C., Jamieson, R.A., Nguyen, M.H., and Lee, B., 2001, Himalayan tectonics explained by extrusion of a low-viscosity crustal channel coupled to focused surface denudation: Nature, v. 414, p. 738–741, doi:10.1038/414738a. Beccaluva, L., Coltorti, M., Saccani, E., and Siena, F., 2005, Magma generation and crustal accretion as evidenced by supra-subduction ophiolites of the Albanide-Hellenide Subpelagonian zone: The Island Arc, v. 14, p. 551– 563, doi:10.1111/j.1440-1738.2005.00483.x. Brown, E.H., 1977, The crossite content of Ca-amphibole as a guide to pressure of metamorphism: Journal of Petrology, v. 18, p. 53–72. Brown, M., 1998, Ridge-trench interactions and high-T low-P metamorphism, with particular reference to the Cretaceous evolution of the Japanese Islands: Geological Society [London] Special Publication 138, p. 137–170. Brunn, J.H., 1956, Etude géologique du Pinde septentrional et de la Macédoine occidentale: Annales Géologiques des Pays Helléniques, v. 7, 358 p. Caddick, M.J., Bickle, M.J., Harris, N.B., Holland, T.J., Horstwood, M.S., Parrish, R.R., and Ahmad, T., 2007, Burial and exhumation history of a Lesser Himalayan schist: Recording the formation of an inverted metamorphic sequence in NW India: Earth and Planetary Science Letters, v. 264, p. 375–390, doi:10.1016/j.epsl.2007.09.011. Dale, J., Holland, T., and Powell, R., 2000, Hornblende-garnet-plagioclase thermobarometry: A natural assemblage calibration of the thermodynamics of hornblende: Contributions to Mineralogy and Petrology, v. 140, p. 353– 362, doi:10.1007/s004100000187. Diener, J.F., Powell, R., White, R.W., and Holland, T.J., 2007, A new thermodynamic model for clino- and orthoamphiboles in the system Na2O-CaOFeO-MgO-Al2O3-SiO2-H2O-O: Journal of Metamorphic Geology, v. 25, p. 631–656, doi:10.1111/j.1525-1314.2007.00720.x. Dilek, Y., Furnes, H., and Shallo, M., 2007, Suprasubduction zone ophiolite formation along the periphery of Mesozoic Gondwana: Gondwana Research, v. 11, p. 453–475, doi:10.1016/j.gr.2007.01.005. Dimo-Lahitte, A., Monie, P., and Vergely, P., 2001, Metamorphic soles from the Albanian ophiolites: Petrology, 40Ar/39Ar geochronology, and geodynamic evolution: Tectonics, v. 20, p. 78–96, doi:10.1029/2000TC900024. Elitok, M., and Drüppel, K., 2008, Geochemistry and tectonic significance of metamorphic sole rocks beneath the Beyşehir-Hoyran ophiolite (SWTurkey): Lithos, v. 100, p. 322–353, doi:10.1016/j.lithos.2007.06.022. Encarnacion, J., 2004, Multiple ophiolite generation preserved in the northern Philippines and the growth of an island arc complex: Tectonophysics, v. 392, p. 103–130, doi:10.1016/j.tecto.2004.04.010. Engwell, S., 2008, The geology of the Aliakmon Valley, Greece [B.Sc. thesis]: Edinburgh, Edinburgh University, 32 p. Ernst, W.G., and Liu, J., 1998, Experimental phase-equilibrium study of Al- and Ti-contents of calcic amphibole in MORB—A semiquantitative thermobarometer: American Mineralogist, v. 83, p. 952–969. Gartzos, E., Dietrich, V.J., Migiros, G., Serelis, K., and Lymperopoulou, T., 2009, The origin of amphibolites from metamorphic soles beneath the ultramafic ophiolites in Evia and Lesvos (Greece) and their geotectonic implication: Lithos, v. 108, p. 224–242. Ghikas, C., 2007, Structural and tectonics of a subophiolitic mélange (Zavordhas Mélange) of the Vourinos Ophiolite (Greece) and kinematics of ophiolite emplacement [M.S. thesis]: Miami, Ohio, University of Miami, 35 p. Ghikas, C., Dilek, Y., and Rassios, A.E., 2010, Structure and tectonics of subophiolitic melanges in the western Hellenides (Greece): Implications for ophiolite emplacement tectonics: International Geology Review, v. 52, p. 423–453, doi:10.1080/00206810902951106.
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MANUSCRIPT SUBMITTED 22 NOVEMBER 2008 MANUSCRIPT ACCEPTED BY THE SOCIETY 21 DECEMBER 2010
Printed in the USA
The Geological Society of America Special Paper 480 2011
Role of plutonic and metamorphic block exhumation in a forearc ophiolite mélange belt: An example from the Mineoka belt, Japan Ryota Mori* Master’s Program of Science and Technology, University of Tsukuba, Tsukuba 305-8572, Japan Yujiro Ogawa† Doctoral Program in Earth Evolution Sciences, University of Tsukuba, Tsukuba 305-8572, Japan Naoto Hirano§ Laboratory for Earthquake Chemistry, University of Tokyo, 7-3-1 Hongo, Bunkyo, Tokyo 113-0033, Japan Toshiaki Tsunogae Masanori Kurosawa Tae Chiba Doctoral Program in Earth Evolution Sciences, University of Tsukuba, Tsukuba 305-8572, Japan
ABSTRACT We investigated the field relations, metamorphic and deformation conditions, age, and chemistry of basaltic, plutonic, and metamorphic blocks in the Mineoka ophiolite mélange belt, Boso Peninsula, central Japan, to clarify their emplacement mechanisms. We considered internal and external deformation of the blocks in the context of the complicated processes by which the ophiolite mélange belt was formed in a forearc setting. A two-stage history leading to the present-day forearc sliver fault zone was revealed: an early stage of deep ductile deformation followed by an episode of brittle deformation at shallower levels. Both stages were the result of transpressional stress conditions. The first stage produced subduction-related schistosity with microfolding and mylonitization and then brecciation during exhumation in the intraoceanic subduction zone, from a maximum depth of garnet-amphibolite facies or eclogitic facies. The second stage was characterized by strong, brittle shear deformation as the rocks were incorporated into the present-day fault zone. The first incorporation of the oceanic plate to the side of the Honshu arc might have occurred during the Miocene, and was followed by right-lateral oblique subduction that has continued ever since the Boso triple junction arrived at its present-day position, thus forming the paleo-Sagami trough plate boundary.
*Current address: Masuo 3-8-25, Kashiwa 277-0033, Japan;
[email protected]. † Corresponding author, current address: Yokodai 1-127-2 C-740, Tsukubamirai 300-2358, Japan;
[email protected]. § Current address: Center for Northeast Asian Studies, Tohoku University, Sendai 980-8576, Japan. Mori, R., Ogawa, Y., Hirano, N., Tsunogae, T., Kurosawa, M., and Chiba, T., 2011, Role of plutonic and metamorphic block exhumation in a forearc ophiolite mélange belt: An example from the Mineoka belt, Japan, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 95–115, doi:10.1130/2011.2480(04). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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INTRODUCTION Places where ophiolitic mélanges are associated with plutonic and metamorphic blocks provide a key to understanding the emplacement mechanism of such blocks during mélange formation. However, little information on these matters has been published, except for part of the Franciscan (e.g., Cowan and Page, 1975; Cloos, 1984; Wakabayashi, 2004, and others) and Ankara mélanges (Dilek and Thy, 2006). Karig (1980) presented the idea that high-pressure metamorphic blocks (“knockers,” blocks in a mélange belt) are brought to the shallow part of the continental margin along a forearc sliver fault. Mann and Gordon (1990) discussed examples of such strike-slip faults in relation to the emplacement of the Guatemalan, Franciscan, and Sambagawa metamorphic rocks. Although these concepts demonstrate that exhumation of deep metamorphic rocks as a consequence of strike-slip faulting is possible, they do not explain the deformation history of those rocks. Cloos (1984) presented an interesting model of exhumation of high-pressure metamorphic blocks via a subduction channel as a return flow, but not in relation to an ophiolite mélange belt. This return flow model was followed by digital experiments by Gerya et al. (2002), but a detailed connection between such a model and field relationships was not attempted. Cowan and Silling (1978), Pavlis and Bruhn (1983), and Iwamori (2003), on the other hand, used an analogue and digital simulation to propose a corner-flow model for exhumation of subduction-related metamorphic rocks, and noted that a rapid uplift occurs along a backstop in the forearc area, especially in the case of oblique subduction. The models proposed by Cloos (1984), Cowan and Silling (1978), Pavlis and Bruhn (1983), and Iwamori (2003) are particularly plausible for a forearc in an oblique subduction setting in which deep-level rocks are selectively exhumed to shallow levels. The examination of relatively young forearc mélange belts, with associated metamorphic and plutonic rocks, may provide an important test to these models of mélange exhumation. However, the deformation of such units has received little study. Previous studies of the rock associations of the Mineoka belt, which is the chief concern in this paper (Figs. 1 and 2), suggested that the emplacement mechanism of the ophiolite sequence can be elucidated by consideration of the field relations and conditions at the time of emplacement, particularly with reference to the mechanism by which the associated plutonic and metamorphic rocks were exhumed (Ogawa and Taniguchi, 1988; Sato et al., 1999; Sato and Ogawa, 2000; Hirano et al., 2003; Takahashi et al., 2003; Ogawa and Takahashi, 2004; Mori and Ogawa, 2005). The general field relations of such blocks, as well as the basaltic and sedimentary rock blocks, are given in Ogawa et al. (2009) with chemical data of major, trace, and rare earth elements without giving new radiometric age data and any tectonic interpretation. This chapter provides new radiometric ages and chemical data from basaltic and other igneous rocks and evaluates these
data in the context of chemical data and field relations of previously described rocks to propose a tectonic model for exhumation of the plutonic and metamorphic blocks of the Mineoka belt. TECTONIC SETTING OF THE MINEOKA BELT The Mineoka belt lies in the Boso triple junction area. The Boso triple junction is the only known Trench-Trench-Trench– type (TTT-type) triple junction in the NW Pacific (Fig. 1), and formed when the Izu island arc moved to its present position in the Miocene (Seno and Maruyama, 1984; Ogawa et al., 1989; Takahashi and Saito, 1997). The Izu arc (the easternmost boundary of the Philippine Sea plate) then collided with the Honshu arc, and the NE boundary of the Philippine Sea plate has since been subducting obliquely under the Eurasia or Northeast Japan plate, forming successive accretionary prisms in the early Miocene and late Miocene to Pliocene, and it is still doing so today (Ogawa and Taniguchi, 1988; Ogawa et al., 1989; Saito, 1992; Hanamura and Ogawa, 1993; Yamamoto and Kawakami, 2005; Ogawa et al., 2008). During the middle Miocene, clastic rocks were deposited in the area of the present-day Mineoka belt. They are composed of three types of rock fragments: ophiolitic rocks, continental rocks, and volcanic island arc rocks (Ogawa, 1983; Ogawa et al., 2008, 2009), which shows that those three distinctive types of crustal material were locally available at that time. In addition, this area includes part of the Cretaceous to Paleogene Shimanto accretionary complex, which is widely developed in SW Japan as the shallow expression of the Sambagawa metamorphic rocks (Isozaki and Maruyama, 1991). The mechanism of emplacement of the ophiolitic rocks in this area during the middle Miocene is thought to have been a kind of incorporation that is an extrusion of ophiolitic rocks into the forearc area, bringing several different kinds of rocks together (Ogawa and Taniguchi, 1988; Sato et al., 1999). Thus, during the middle Miocene, the plate boundary was an ophiolite-bearing fault zone (the Mineoka ophiolite). Since then, the Mineoka belt has been an active fault system in a right-lateral (dextral) fault belt (Nakajima et al., 1981; Ogawa, 1983; Mori and Ogawa, 2005), as well as providing the backstop between the accretionary prism and the forearc basin (Fig. 2). The total strike-slip displacement along the fault belt may be more than 100 km (Lallemant et al., 1996). SUMMARY OF FIELD RELATIONS OF OPHIOLITIC AND OTHER ROCKS IN THE MINEOKA BELT The lithologies of the ophiolitic, plutonic, and metamorphic blocks and the surrounding rocks of the Mineoka belt are distributed in different ways in the three sub-belts (northern, central, and southern) of the Mineoka belt, which, in total, is 5 km in width (Fig. 2) (Taniguchi, 1991; Chiba, 2008). During the early stages of research, the ophiolitic rocks of the Mineoka belt were formerly thought to form an ophiolite mélange, as blocks of various
Role of exhumation in a forearc ophiolite mélange belt, Mineoka belt, Japan
Mineoka belt
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Figure 1. Index map of the localities in this paper. Shown are the Mineoka ophiolite belt, the Sofugan Tectonic Line (S.T.L., arrow) in the Izu arc (Shichito Ridge) collision to the Honshu arc, and the Ohmachi Seamount and Hahajima Seamount (arrow). Adopted and modified from Yuasa (1985).
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Figure 2. General lithologic map of the middle Boso Peninsula, showing the northern, central, and southern sub-belts of the Mineoka belt. The forearc basin (Miura Group on the north) and accretionary prism (Emi Group on the south and farther southward) of Miocene and Pliocene time (adopted and modified from Takahashi et al., 2003). Cross-section lines in Figures 3 and 4 are shown as A, B, and C. T-T-T—trench-trench-trench.
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Role of exhumation in a forearc ophiolite mélange belt, Mineoka belt, Japan rock types were identified within sheared serpentinite matrix (Fig. 3) (Ogawa, 1983; Ogawa and Taniguchi, 1987, 1988). However, recent studies show that the structural style is not that of block-in-matrix, but rather one of a collage of unsheared blocks of various lithologies bounded by discrete fault zones (Takahashi et al., 2003; Ogawa and Takahashi, 2004) (Fig. 4). The serpentinite, mostly composed of lizardite and chrysotile from harzburgite tectonite, does not form a matrix but is one of the main block lithologies. Most of the lithologies can be found in fault contact with each of the other lithologies present. It is only within rela-
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tively narrow (~1 m wide) fault zones that the serpentinite and clastic rocks are sheared. Some such faults cut plutonic and metamorphic blocks (Takahashi et al., 2003; Chiba, 2008). The field relationships of these blocks and their major, trace, and rare earth element chemistry are described in Ogawa et al. (2009). In the northern sub-belt, alkalic basalt blocks are dominant (Ogawa et al., 2009), and they are in fault contact with clastic rocks of either the Eocene Shimanto Supergroup or the Miocene Hota Group (Kawakami, 2004). Alkali basalt blocks are Miocene in age (Hirano and Okuzawa, 2002; Hirano et al., 2003).
pillow lava & dolerite sheet complex Kojima Formation diorite
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Figure 3. Previous model of profiles of the eastern part of the Mineoka belt. (A) Central and southern sub-belts. (B) Mineoka-Sengen area (central sub-belt). (C) Heguri-Naka area (central sub-belt). Note that these sections are drawn as if all of the ophiolitic blocks are surrounded by sheared serpentinite, but the present model is not the case (see details in text). Adopted and modified from Ogawa (1983).
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Mineoka Hill
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Figure 4. New model of the profile of the central and southern sub-belts of the Mineoka belt (eastern part along line C of Fig. 2). Adopted and modified from Takahashi et al. (2003). Note that most of the ophiolitic blocks are shown as faulted, with massive serpentinite in the north of the profile, but with sandstone-mudstone in the south.
In the central sub-belt of the Mineoka Hills area, both large and small blocks of tholeiitic basalt are in fault contact with massive serpentinite (mostly completely serpentinized harzburgite tectonite) (Takahashi et al., 2003). These blocks are mostly Paleogene in age, from basalt, chert-limestone, and clastic rocks (Mohiuddin and Ogawa, 1998). In the southern sub-belt, the blocks are almost aligned on the southern boundary of the Mineoka belt in fault contact with the Hota Group, and the basaltic rocks are of Cretaceous age as far as measured and as mentioned below. Every block of the Mineoka belt is bounded by a right-lateral (dextral) fault with a thrust component, or by serpentinite and/or sedimentary rocks. The large-scale fault pattern in the belt is that of a typical dextral Riedel shear zone (Ogawa, 1983; Mori and Ogawa, 2005), suggesting that dextral shears prevail within the 5-km-wide Mineoka belt. The size of most blocks is in the order of tens of meters, but blocks or slabs of hundreds of meters in scale are not uncommon (Ogawa et al., 2009). Blocks are generally oblate or prolate in shape, but spherical blocks are also common. If elongated, the long axis of blocks would tend to be parallel to, or slightly oblique to, the trend of the Mineoka belt (WNW). Therefore, the Mineoka belt might represent an intermediate model between the previous mélange models depicted by Figure 3 and the fault belt model by Figure 4. This intermediate model differs from that of Figure 3 in which only some of the serpentinite boundaries are sheared, and from that of Figure 4 in which individual blocks have rounded ends. It is important to note that some plutonic and metamorphic blocks have an oblate or prolate shape and lie within narrow sheared serpentinite fault belts surrounded by massive serpentinite or sedimentary rocks.
AGE OF OPHIOLITIC AND OTHER IGNEOUS ROCKS AND THEIR COVER OF PELAGIC SEDIMENTARY ROCKS Igneous blocks include tholeiite, which is variably altered but retains its igenous textures, relatively fresh alkali basalts, and unaltered high-Mg andesite and diorite (tonalite). Alkali basalt is relatively fresh, and the high-Mg andesite and diorite (tonalite) are fresh. One of the basaltic rocks in the middle sub-belt is of late Eocene age (47 ± 10 Ma, Ar-Ar whole rock isochron; Hirano et al., 2003). Such rocks constitute the most common igneous rock type in the Mineoka belt. They are older than the associated plutonic and metamorphic rocks, which are Oligocene, as shown below (ca. 40 Ma or younger). However, considering that these radiometric ages were obtained previously, they are more or less uncertain, without any good plateau ages except for one alkali basalt of ca. 20 Ma Ar-Ar age (Hirano and Okuzawa, 2002). In contrast to the above-reviewed ages, another study showed that a block of radiolarian chert is of Albian age (ca. 100 Ma) (Ogawa and Sashida, 2005). We conducted 40Ar/39Ar incremental heating analyses of two tholeiitic basalts (NH980303-03 from Hashimoto and NH011117-01 from Shinyashiki), a high-Mg andesite (RM050929-05 from Hinata) and a diorite (BM31YD from Yamada) from the southern sub-belt, and an alkali basalt (RM050425-04 from Isomura) from the central sub-belt (details of these localities are given in Ogawa et al., 2009). The spectra of some samples show young ages on the lower temperature gas fractions (Figs. 5A, 5B, 5C) that may be a consequence of Ar loss from weathering and/or alteration of the
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Figure 5. Ar-Ar data plot based on new data of basaltic and andesitic rocks of Cretaceous age: A for Hashimoto, B for Shinyashiki, C for high-Mg andesite at Hinata, D for alkalic basalt at Isomura, and E for diorite at Yamada (for local names of occurrence, see Ogawa et al., 2009). Plateau ages are shown for A through D, and an isochron age for E (Naoto Hirano, original data). Inverse isochrons and age spectra are in the upper part of each figure. MSWD—mean square of weighted deviates. MSWD = SUMS/(n-2) (York, 1968). All errors are 2σ. Fresh and separated groundmass from basaltic samples were crushed to 100–300 μm grains and wrapped in aluminum foil (70 mm in length, 10 mm in diameter) with flux monitors for biotite (EB-1, Iwata, 1998), K2SO4, and CaF2. The samples were irradiated for 24 h in the Japan Material Testing Reactor (JMTR), Tohoku University. During the irradiation, samples were shielded by Cd foil to reduce thermal neutron-induced 40Ar from 40K (Saito, 1994). The Ar extraction and Ar isotopic analyses were done at the Radioisotope Center, University of Tokyo. During incremental heating, gases were extracted in 10 steps between 600 and 1500 °C. The analytical methods are described by Ebisawa et al. (2004).
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samples. Most samples show a well-developed plateau spanning more than 50% of the gas released and including 3–10 heating steps. The 30.6 ± 1.6 Ma plateau age of RM050425-04 (Fig. 5D), may be affected by excess 40Ar as indicated by the 40Ar/36Ar ratio of 304.5 ± 5.8, which is higher than the atmospheric ratio (295.5) in the inverse isochron (Fig. 5D). The 25.5 ± 2.3 Ma age from the inverse isochron diagram thus appears the best age estimate for sample RM050425-04 (Fig. 5D). On the other hand, the plateau ages (80.6 ± 1.7 Ma for NH980303-03; 85.0 ± 3.9 Ma for NH011117-01; 28.6 ± 5.1 Ma for RM050929-05; 37.7 ± 1.2 Ma for BM31YD) should be accepted as the best age estimates because the initial 40Ar/36Ar ratio of each sample corresponds to the atmospheric ratio in the inverse isochrons, respectively (Figs. 5A, 5B, 5C, and 5E). These findings confirm that although one tholeiitic basalt in the central sub-belt is of Paleogene age, those in the southern sub-belt are Cretaceous, and the alkali basalt, andesitic volcanic rocks, and diorite in any of the sub-belts are either Eocene or Oligocene, younger than 40 Ma. These ages demonstrate that rocks of the ophiolite belt formed long before they arrived at their present positions during the Miocene. Another important component of the pelagic cover on some of the basaltic blocks is the Paleogene to Miocene limestonechert sequence, the Kamogawa Group (Mohiuddin and Ogawa, 1998). This sequence is composed of bedded limestone and chert with sporadic ash intercalations and has been dated from Paleocene (ca. 60 Ma) to early Miocene (ca. 18 Ma) on the basis of foraminifers, which means that part of the oceanic plate had existed in pelagic realms. The alkali basalts are much younger, yielding Ar-Ar ages of ca. 20 Ma (Hirano and Okuzawa, 2002; Hirano et al., 2003). In addition to the above-reviewed ages, we obtained K/Ar wholerock dates on three volcanic samples. These analyses were con-
Sample number Kj060921 Hinata51 BM04HG2
ducted by Geochronology and Isotope Chemistry of Ontario, Canada. The host andesitic tuff breccia for the high-Mg andesite fragment we dated (sample RM050929-05; Ar/Ar age of 28.6 ± 5.1 Ma) yielded a 15.6 ± 0.5 Ma age, and andesitic pumice-fall deposits at Kojima at Kamogawa Harbor (Kj-5) yielded an age of 5.8 ± 0.3 Ma. Previously published K/Ar dates from plutonic blocks are as follows: the Yamada knocker (a dioritic body in sheared serpentinite) yielded 40.9 ± 2.1 Ma (whole rock), and a similar diorite block in the Shingan-ji area yielded 27.9 ± 1.8 Ma (hornblende) and 24.1 ± 1.4 Ma (plagioclase) (Hirano et al., 2003). Metamorphic blocks (hornblende schists mentioned above) have K-Ar ages as follows: a hornblende from Byobu-jima gave 33.1 ± 2.3 Ma (Hiroi, 1995a), and that from a block from Heguri-Naka, 39.5 ± 2.2 Ma (Hiroi, 1995b). Those data are summarized in Table 1. The new and published ages indicate that the tholeiitic volcanic rocks formed first, followed by formation of the plutonic and metamorphic rocks, which were in turn followed by alkalic basalts and andesites. Therefore, as Hirano et al. (2003) noted, within the Tertiary rocks, the tholeiitic basalt-bearing ophiolitic suite (ophiolite proper) is the oldest, followed by younger island-arc plutonic and metamorphic rocks, and then still younger alkalic basalts. The youngest rocks are andesitic tuff breccia derived from the Izu island arc during the Miocene. These rocks were probably extruded in the Izu frontal area, and may overlie the older (28.6 ± 5.1 Ma) high-Mg andesite. These relationships suggest that at ca. 40 Ma, after incorporation of Cretaceous chert and Cretaceous to Eocene tholeiitic basalt, the area became an island arc environment with arc plutonism and formation of crust with high Mg andesite. Alkalic basalt was erupted, with no obvious connection to arc activity, and pelagic limestone and chert deposition continued throughout the various stages of development of the Mineoka belt rocks.
TABLE 1. RADIOMETRIC AGES OF IGNEOUS AND METAMORPHIC ROCKS OF THE MINEOKA BELT Location Rock Method, whole rock or wh in. Age (Ma) Reference ht.* Kojima Andesite tuff K-Ar wh 5.8 ± 0.3 This study Hinata
Andesite tuff
K-Ar wh
15.6 ± 0.5
This study
Toge
Alkali basalt
Ar-Ar wh in. ht.
19.62 ± 0.9
RM050425-04
Isomura
Alkali basalt
Ar-Ar wh isochron
25.5 ± 2.3
Hirano and Okuzawa (2002) This study
RM050929-05
Hinata
High-Mg andesite
Ar-Ar plateau
28.6 ± 5.1
This study
Yamada
Diorite
Ar-Ar plateau
37.7 ± 1.2
This study
Kamogawa Harbor Heguri-Naka
Hornblende schist
K-Ar hornblende
33.1 ± 2.3
Hiroi (1995a)
Garnet amphibolite
K-Ar hornblende
39.5 ± 2.2
Hiroi (1995b)
BM31YD Byobu-jima Heguri-Naka BM591(Bentenjima) NH980303-03
Kamogawa Harbor Hashimoto
Tholeiite
Ar-Ar isochron
47 ± 10
Hirano et al. (2003)
Tholeiite
Ar-Ar plateau
80.6 ± 1.7
This study
NH011117-01
Shinyashiki
Tholeiite
Ar-Ar plateau
85.0 ± 3.9
This study
* wh in. ht.—whole rock incremental heating for plateau.
Role of exhumation in a forearc ophiolite mélange belt, Mineoka belt, Japan PETROCHEMISTRY OF BASALTIC, PLUTONIC, AND METAMORPHIC ROCKS Because most of the basaltic rocks are altered without any fresh phenocrysts, we cannot ascertain their tectonic origin by major element chemistry alone. Abundances and ratios of trace and rare earth elements are more useful for elucidating the tectonic setting of magmatism, because they are comparatively immobile during alteration. Basaltic rocks in the Mineoka belt have been divided into tholeiitic and alkali (Ogawa and Taniguchi, 1989a, 1989b; Taniguchi and Ogawa, 1990; Hirano et al., 2003). The tectonic setting of tholeiitic basalts of the Mineoka belt is disputed, although most of these basalts exhibit typical normal mid-oceanic-ridge basalt (N-MORB), or ocean floor basalt chemistry with minor amounts of island arc tholeiite (IAT) (Hirano et al., 2003). New trace- and rare-earth-element data were obtained by inductively coupled-plasma mass spectrometry (Ogawa et al., 2009). The rocks thought to have IAT chemistry by Hirano et al. (2003) show a negative anomaly of Nb and Ta in spider diagrams with a large ion lithophile element–enrichment pattern (Fig. 6), suggesting a kind of island arc setting (Ogawa and Takahashi, 2004). The now-subducted plate, to which these basalts and associated pelagic sedimentary rocks belong, has been called the Mineoka plate (Ogawa and Taniguchi, 1988; Sato et al., 1999). The younger plutonic rocks show a well-developed island arc signature (Figs. 6 and 7). The alkali basalts are within plate basalts (WPB) (Fig. 6), and their younger ages indicate that they erupted independently of arc activity within the Mineoka plate. Besides these basaltic and plutonic rocks, a new andesitic rock was found at Hinata: an andesitic tuff breccia including a high-magnesian andesite clast, as already described (Fig. 7). Considering that most diorite bodies are dated ca. 40 Ma or younger, the area probably became part of an island arc after ca. 40 Ma, which may overlap the previous MORB-type lithology. Metamorphic rocks show mostly island arc affinity. The tholeiitic basalts of the Mineoka plate may be as old as Cretaceous to Paleocene; after 40 Ma most of the igneous rocks show island arc affinities, with some alkali basalt. The implications will be discussed further in the following chapters. DEFORMATION OF PLUTONIC AND METAMORPHIC ROCKS Plutonic Blocks As noted by Mori and Ogawa (2005), the plutonic and metamorphic blocks have undergone several stages of deformation, initially more ductile and later more brittle. The Yamada knocker (a dioritic body in fault contact with serpentinite) in the southern sub-belt of the Mineoka belt provides a typical example of the two stages of deformation (Fig. 8). The initial, more ductile deformation is characterized by mylonitic formation along centimeter to millimeter–wide ductile zones within faults, which are
103
developed inside the blocks. Quartz grains are deformed with wavy extinction, whereas plagioclase is cataclastically broken. Internal shears have various orientations and are characterized by metamorphic minerals such as prehnite and hornblende. The later, more brittle deformation is characterized by brecciation restricted along faults of Riedel shears, associated with extensive cataclastic deformation only in the surface boundaries of the blocks (Fig. 8). Analcime veins within shear planes are near the boundaries with the serpentinite, providing evidence of uplift to relatively shallow levels. However, the shearing does not penetrate the serpentinite body, which is almost massive and contains no shear zones more than tens of centimeters wide. The late shearing in the dioritic blocks was strong enough to create systematic Riedel shears. The internal shear planes are arranged in a small circle associated with strike-slip faults with a thrust component, most of which are dextral, although some are sinistral. This style of deformation suggests that the plutonic and metamorphic rocks underwent several episodes of shearing (conjugate sets) at depth with conical symmetry induced by a transpressional regime (Mori and Ogawa, 2005). Similar two-stage deformation is observed in basaltic rocks from the eastern end of the Mineoka belt (Takahashi et al., 2003; Ogawa and Takahashi, 2004). The first stage is penetrative deformation by shearing and cataclasis within the blocks, together with zeolite, prehnite, or calcite veins that fill near-vertical conjugate shears. This stage of deformation has been suggested to have occurred near a spreading center, coeval with hydrothermal activity that formed the vein fillings (Ogawa and Takahashi, 2004). In the second stage of deformation the basaltic blocks are strongly deformed mostly along the boundaries of the different kinds of blocks apparently by dextral transpressional shearing with flower structures. These shear zones are characterized by Riedel shear systems, as described in the bodies at Shinyashiki (Takahashi et al., 2003) and at Benten-jima (Ogawa and Takahashi, 2004). Metamorphic Blocks Four metamorphic blocks in the Mineoka belt are composed mostly of schistose metabasite, ranging from meters to tens of meters in size with oblate to prolate shapes (Fig. 9). These blocks have all been metamorphosed to amphibolite facies (Ogo and Hiroi, 1991) but show retrograde facies of epidote-amphibolite or greenschist facies. Their chemistry has a characteristic island arc signature, except for the possible MORB-type Sumoba-ishi Island block. The quartz and psammitic (quartz, feldspar, muscovite, garnet, and monazite are the components) schists indicate deposition of the original sedimentary rocks as chert and terrigenous sandstone (Fig. 10). This combination of pelagic sediments and terrigenous sands indicates sedimentation near a trench, close to a continent (Ogo and Hiroi, 1991). Three examples of metamorphic blocks in Kamogawa Harbor and one on land at Heguri-Naka exhibit similar deformation: the Kana-shima block off Byobu-jima (Mori and Ogawa, 2005), the Sumoba-ishi Island block, and the Heguri-Naka block
104 10
10
Mori et al.
3
2
WPB type MORB type
10
10
1
0
-1
10
IAB type -2
10
Rb Ba 10
Th
U
K
Nb Ta
La
Ce
Sr
Nd P
Hf
Sm Ti
Tb
Y
2
MORB type
WPB type
10
Zr
Figure 6. Representative spider diagrams of basaltic rocks, first described by Hirano et al. (2003), and later, some trace and rare earth element (REE) data were added by Ogawa et al. (2009). Trace elements are normalized by normal mid-oceanic-ridge basalt (N-MORB) after Saunders and Tarney (1984) and REE by Primordial Mantle after Sun and McDonough (1989). Analyzed by N. Hirano and M. Kurosawa. WPB—within plate basalt; IAB—island arc basalt.
1
IAB type 10
0
La
Ce
Pr
Nd
Pm
Sm
Eu
Gd
Tb
Dy
Ho
(Fig. 9). The Kana-shima block contains epidote veins, the Sumoba-ishi block contains chlorite veins, and the Heguri-Naka block contains prehnite veins. All of these are products of retrograde metamorphism. Thus the initial schistosity was followed by the development of epidote or chlorite veins, followed by isoclinal microfolding, and then brecciation that preserved the essential traces of the foliation, and finally the formation of cataclasite along faults (Fig. 10) (Mori and Ogawa, 2005). Penetrative schistosity is observed in both metabasite and quartz and psammitic schist layers, but all the lithological boundaries are faults (Fig. 10). The schistosity of these metamorphic rocks (at least three of the four) appears to be a product of compression during subduction,
Er
Tm
Yb
Lu
which is supported by the IAT chemistry of the metabasite as well as by the occurrence of quartz feldspar–bearing psammitic schist. The metamorphic conditions for these blocks were calculated from the chemistry of some of their minerals in Table 2, and are also summarized in Table 2. The average temperature and pressure of metamorphism were estimated for garnet-bearing psammitic schist at Byobu-jma to be 500–550 °C and 500 MPa, respectively, on the basis of the stability of chalcopyrite and the quantity of jadeite components in aegirineaugite within the schist (Ogo and Hiroi, 1991). The pressure-temperature (P-T) conditions are attributed to the intermediate series of regional metamorphism in subduction zones (Ogo and Hiroi, 1991). Our own temperature estimates based on plagioclase and hornblende compositions (Table 1) and the Holland and
Role of exhumation in a forearc ophiolite mélange belt, Mineoka belt, Japan 10
105
2
Andesite tuff breccia (Hinata) 10
10
1
0
-1
10
Diorite (Hinata) Hinata rocks -2
10
High-Mg andesite clast (Hinata)
Diorite, tonalite Kobata andesite
-3
10
Rb Ba
Th
U
K
Nb Ta
La
Ce
Sr
Nd P
Hf
Zr
Sm Ti
Tb
Y
Figure 7. Representative spider diagrams of andesitic, high-Mg andesite, and plutonic rocks at Hinata. Trace and rare earth elements are normalized by N-MORB after Saunders and Tarney (1984) and Sun and McDonough (1989), respectively. Sample lithologies from Hinata and Kobata are shown. For local names of occurrence, see Ogawa et al. (2009). Analyzed by S. Haraguchi, T. Ishii, and M. Kurosawa.
Late shear R1 R2 R1 (S2)
S1
Figure 8. Polished thin section photo (left), and enlargement of the boundary to serpentinite (upper right) of diorite from Yamada, west of Heguri-Naka. Note in the sketch (lower right) that the block boundary has early foliation (S1) that is dislocated by later shear (S2), interpreted as R1 with additional R2. Note that the early foliation is mylonitic of a more ductile deformation (wavy extinction of quartz but cataclasis of plagioclase), whereas the late foliation is more brittle. Strong late shear is mostly dextral in the horizontal frame.
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Blundy (1994) thermometer, suggest metamorphic temperatures of 560–600 °C (Table 2), which is nearly consistent with the results of Ogo and Hiroi (1991). In addition, we applied a semiquantitative geothermobarometer of Ernst and Liu (1998) for a Byobu-jima sample and obtained P-T conditions of 620 °C and 1.5 GPa (Table 3), which is the highest pressure condition so far reported from metamorphic
A
C
rocks in this area. We confirmed that plagioclase, rutile, and titanite are present in the sample, as this method is applicable only under the saturation of Al2O3 and TiO2. As shown in Figure 11A, no exsolution lamella of the Ti-bearing phase has been identified in hornblende, suggesting that the mineral probably preserved the primary composition. This figure also demonstrates that rutile grains are enclosed within
B
Benten-jima
D
Kana-shima
Figure 9. Outcrop photos of four metamorphic blocks. Note that all are of either oblate or prolate shape. (A) Byobu-jima Island (central bar ~ 3 m long); (B) Sumoba-ishi Island (lower circular island); (C) Kana-shima Island (the island is ~10 m in diameter), all in the Kamogawa Harbor area; (D) Heguri-Naka block entrapped within sheared serpentinite.
Role of exhumation in a forearc ophiolite mélange belt, Mineoka belt, Japan texturally retrograde titanite. Because of the higher stability field of rutile (P >1.3 GPa at 600 °C) than that of titanite (P <1.3 GPa) (Liu et al., 1996), the texture probably implies decompression from a P ~1.5 GPa high-pressure stage. One possibility is this higher-pressure condition may correspond to the peak of the metamorphism in the subduction zone, and then it underwent retrograde metamorphism at 500 °C and 500 MPa during exhumation. The P-T conditions are still in the extension of the intermediate series of regional metamorphism of Ogo and Hiroi (1991). We need further research to confirm the metamorphic condi-
A
tions and their age, and to establish the metamorphic implication for tectonics. Garnet-hornblende thermometry of Graham and Powell (1984), conducted on garnet-amphibolite from the HeguriNaka area, indicates a metamorphic temperature range of 560–600 °C (Table 3). Application of the geothermobarometer of Ernst and Liu (1998) on the hornblende with no exsolution texture (Fig. 11B) yielded slightly higher temperatures of 650 °C and also a very high pressure of 2.0 GPa, owing to high Al2O3 and TiO2 contents in the hornblende (~14 wt% and ~0.7 wt%, respectively) within the eclogite field. We further
B
C
107
D
0.5 mm Figure 10. Polished surface of a hornblende schist and the quartz and psammitic parts of the metamorphic block at Byobu-jima Island, Kamogawa Harbor. (A) Microfolds and later brecciation in two stages. (B) Brecciation of quartzose and psammitic components. The boundaries between the two are faulted. (C) Photomicrograph of hornblende schist, showing brecciation (cataclasis) inside. (D) Photomicrograph of a psammitic schist, showing wavy extinction of quartz grains and brecciation.
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examined thin sections to identify relict high-pressure minerals in garnet and hornblende but could not identify any trace of eclogite-facies metamorphism except for rutile inclusions in hornblende (Fig. 11B), which is consistent with the peak metamorphism at P ~2.0 GPa. This is the first report of highpressure metamorphic conditions from the Mineoka area. The tectonic implications will be discussed later in the context of similarity to rocks found in the Ohmachi Seamount area of the Izu arc. The boundary between the metamorphic amphibolites and weathered serpentinite matrix at Heguri-Naka is a fault rela-
tionship with brecciated rocks in both lithologies, and showing dextral faults that are consistent with present whole strike-slip motion of the Mineoka belt. DISCUSSION: EMPLACEMENT OF THE PLUTONIC AND METAMORPHIC BLOCKS Deformation Process and Implication Most of the plutonic and metamorphic rocks show evidence of two stages of deformation, initially more ductile
TABLE 2. REPRESENTATIVE ELECTRON MICROPROBE ANALYSIS OF MINERALS FOR ESTIMATION OF TEMPERATURE Locality Mineral name
Byobu-jima Hornblende
O*
23
SiO2
45.88
Al2O3
Heguri-Naka
Plagioclase 8
Garnet
Hornblende
12
23
64.54
38.0 7
43.73
11.55
22.51
21.34
14.39
TiO2
0.60
0.0 0
0 .2 5
0.69
Cr2O3
0.15
0 .02
0.00
0.12
FeO**
12.73
0.09
26.53
12.79
MnO
0.21
0.00
3.31
0.14
MgO
12.26
0.00
3.01
11.14
CaO
11.40
3.44
7.88
10.51
Na2O
1.87
9.77
0.05
2.32
K2O
0.44
0.06
0.01
0.21
Total
97.07
100.44
100.44
96.02
Si
6.738
2.834
3.002
6.485
Al
1.998
1.165
1.983
2.514
Ti
0.066
0.000
0 .015
0. 076
Cr
0.017
0.001
0.000
0.013
Fe
1.564
0.003
1.749
1.585
Mn
0.025
0.000
0.221
0. 01 8
Mg
2.681
0.000
0.353
2. 46 0
Ca
1.793
0. 162
0. 66 5
1.6 69
Na
0.532
0. 831
0. 00 7
0.6 66
K Total
0.081
0 .0 0 3
0.001
0 .0 3 9
15.495
5.000
7.996
15.527
O*—number of oxygen in molecule. FeO**—total iron as FeO.
†† †
§
Heguri-Naka Garnet amphibolite 1.5 × 1.5 520–600 °C Weak foliation 39.5±2.2 Ma 650 °C, 2.0 GPa † § Temperature estimation based on Holland and Blundy (1994), Graham and Powell (1984), and temperature and pressure on Ernst and Liu (1998) for *, , and , respectively. †† K-Ar ages are from Hiroi (1995a, 1995b) for ** and , respectively.
Brecciation, gouge
Brecciation 5×5
Kana-shima
10 × 10
Chlorite schist
Hornblende schist
Sumoba-ishi
TABLE 3. ROCK NAMES AND SIZES, AND ESTIMATED PALEOTEMPERATURE, PRESSURE, AND AGES OF FOUR METAMORPHIC ROCK BLOCKS IN THE MINEOKA BELT Site Rock Size (m) Temp (°C) Pressure Deformation K-Ar age § Micro-fold, brecciation, gouge 33.1±2.3 Ma** Byobu-jima Hornblende schist 30 × 15 560–600 °C* 620 °C, 1.5 GPa Silicic schist Psammitic schist
Role of exhumation in a forearc ophiolite mélange belt, Mineoka belt, Japan
109
and later more brittle, but even the former occurred at relatively low temperature because the plagioclase is cataclastically deformed. Both stages show small circle patterns of faulting; therefore we consider that the deformation field for both stages might be dextral transpression. The second stage appears to have been at a shallower depth than the first stage, which suggests that after initial emplacement (by intrusion or metamorphism) the blocks were affected by oblique thrusting, and were finally emplaced at shallow levels in a regime of conical symmetry of fracturing induced by transpressional shearing (Fig. 12). Exhumation Process Ages of ca. 30–40 Ma for the plutonic and metamorphic rocks mean that those rocks were emplaced at shallow levels before the Boso triple junction and the Izu arc reached their present positions. The geochemistry of these rocks, their lithologic association, and the pervasive schistosity suggest that the metamorphism occurred in an island arc–trench subduction setting. Plutonic rocks were intruded in an arc setting. This subduction zone could not be the same one that occupies the present-day Sagami trough, to which the Mineoka is currently adjacent, because the subduction in the modern Sagami trough did not begin until ca. 15–20 Ma (Seno and Maruyama, 1984; Ogawa and Taniguchi, 1988). These observations imply that the basaltic blocks, as well as the plutonic and metamorphic blocks, were exhumed from the depths (where more ductile deformation occurred) and then to shallower levels (where more brittle deformation occurred). The age of deformation is not always known, but these two stages might be through Paleogene to Miocene time, considering the metamorphic and stratigraphic controls. As shown by Hirano et al. (2003) and our present trace and rare earth element data, some of the tholeiitic basaltic rocks express MORB signatures. All the alkalic basaltic rocks show WPB affinities (Fig. 6). However, most of the plutonic rocks and andesitic volcanic rocks show developed island arc affinities with markedly negative anomaly Nb and Ta contents (Fig. 7). This combination of MORB, IAT, and WPB rocks and their radiometric ages (Hirano et al., 2003) suggests that the Mineoka belt includes rocks of both early ocean floor and later developed an island arc origin. K-Ar dating of the subduction-related metamorphic rocks shows ages of ca. 33 or ca. 39 Ma. In addition, the ages of the plutonic rocks of island arc affinity are close to the time at which the direction of movement of the Pacific plate changed (42 Ma) as noted by Hirano et al. (2003). This age likely coincides with the initiation of subduction along the Mineoka plate boundary and agrees with the age of the metamorphic rocks presumably formed during subduction of the Pacific plate under the Mineoka plate. It was probably at this time that the boundary between the Mineoka and Pacific plates became a subduction zone (Hirano et al., 2003) (Figs. 13 and 14).
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A
B
Figure 11. Photomicrographs of two representative metamorphic rocks. (A) Hornblende schist from Byobu-jima (sample 030612b-2). The mineralogy of the rock is hornblende (Hbl), plagioclase (Pl), and epidote (Ep). Titanite (sphene; Ttn) contains rutile (Rt) inclusions, suggesting decompression in the retrograde stage (Ernst and Liu, 1998). (B) Garnet amphibolite from Heguri-Naka (sample 040207-10-1), composed dominantly of hornblende and rare garnet. Rutile occurs as inclusions in or along the grain boundaries of hornblende grains, but not as lamellae within the amphibole. This texture indicates that the low TiO2 content (<0.6 wt%) of the hornblende is probably due to high-pressure (~2.0 GPa) metamorphism as discussed in the text.
Age of Deformation and Lithologic Mixing of the Mineoka Belt
Stage 2
Fo
rea
rc
sliv
er
fau
nic
fron
t
lt
Vo
lca
e
nic
to
ec
T ike
-l
an
g ofu
Lin
el
nn
S
on cti
u bd Su ck ro c i h rp k mo roc a t ic Me ton u Pl
a ch
Stage 1
~100 km
Figure 12. Tectonic synthesis (a 3-D perspective view) of the Mineoka plate, in which plutonic and metamorphic rocks (represented by v and x marks, respectively) were formed and exhumed (represented by black triangle mark) at the first stage along a forearc fault zone similar to the Sofugan Tectonic Line in the volcanic island arc setting. The second stage shows incorporation along the forearc sliver fault zone in the oblique subduction zone to the Mineoka belt (represented by rotating cone). Note that both stages are of strong thrust shears in a conical fashion (see Mori and Ogawa, 2005).
As shown by Ogawa and Taniguchi (1988) and Ogawa and Takahashi (2004), the Mineoka ophiolite is much more strongly deformed and altered than the andesite tuff breccia of ca. 15.0 and 5.8 Ma K-Ar age (middle to late Miocene). The 15 Ma deposit is equivalent to the tuffaceous Yabe Group on the Miura Peninsula, in which there are 12 Ma microfossils (Kanie, 1999). These andesite tuff breccias unconformably overlie the ophiolitic rocks, suggesting that deformation and alteration of the ophiolitic rocks had ceased before deposition of the middle or late Miocene tuffaceous deposits (Ogawa and Takahashi, 2004). Therefore, all the ages determined for the pelagic cover (the youngest of which is 18 Ma) and the andesite tuff breccia (15 Ma and 5.8 Ma) overlying the deformed ophiolite strongly support that emplacement of the ophiolite and related rocks into the mélange belt occurred between 18 and 15 Ma; that is, in the middle Miocene. Spreading of the Shikoku Basin ceased at ca. 15 Ma with only a minor amount of later intraoceanic plate volcanism at 12 Ma, which formed the Kinan Seamount (Okino et al., 1994, 1999; Sato et al., 2002). After spreading of the Shikoku Basin ceased, the Boso triple junction arrived at its presentday position as it jumped to the east (sometime between 18 and 15 Ma; Figs. 13 and 14). Then the Mineoka ophiolitic and related rocks were emplaced by obduction or another kind of incorporation at the NW corner of the Pacific into the present Honshu arc (Fig. 13). The presence of ophiolitic rock fragments in middle Miocene conglomerates and sandstones of the Mineoka belt supports such timing (Ogawa and Taniguchi, 1988).
Role of exhumation in a forearc ophiolite mélange belt, Mineoka belt, Japan
18 Ma
40-20 Ma
Mineoka plate incorporation
Eurasian plate
Oblique fault with metamorphic rocks
sm pri y r a Subduction on eti
cr
Small seamount
Shikoku Basin
Previous triple junction
ez
on
e-
Pacific plate Philippine Sea plate
Fr ac
tur
Paleo-Izu arc
Philippine Sea plate
system
Mineoka plate
Jump of the triple junction
ctio n
i
ac
Mineoka plate
ne ws ub du
Sh
nto ma
111
Pacific plate
15-0 Ma
15 Ma
Mineoka ophiolite Izu arc collision
New triple junction New triple junction volc ani ca rc s yste m
Previous triple junction
Pacific plate
Izu
Pacific plate
New
Izu
volc
anic
arc
sys
Previous triple junction
New
tem
Ophiolite belt
Philippine Sea plate
Philippine Sea plate
Subduction
Figure 13. Schematic figures for the development of the Mineoka belt and later incorporation, and the Izu arc collision, modified from Ogawa and Taniguchi (1990). The first figure corresponds to the Mineoka plate, which is confined to the triple junction corner, obducted to the Honshu continental arc in the second stage to provide the Trinity clastics during the middle Miocene. Then the system is incorporated into the Izu arc collision zone, continuing to the present time (fourth stage). The boundary between the Mineoka plate and the Pacific plate may have been, first, a fracture zone, but later converted to a subduction boundary coincident with the beginning of the island arc activity ca. 40 Ma (Hirano et al., 2003). The subducting upper plate is shown by barbs on the trench (see Figs. 12 and 14).
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20 Ma
40 Ma
100-80-50 Ma
18-15 Ma
EU MI
PA
MI
MI incorporation
MI
Jump of triple junction
B PH NNG
SB
B
eq
B
PH
1 Ma 15 Ma
NA
EU
12-5 Ma
5 Ma
EU
EU
Opening of Sea of Japan
EU Tz PA
Granite rock Iz
Beginning of PH collision of lzu arc
Iz
Iz
Iz
PH PH
PH
Figure 14. Tectonic synthetic maps for Philippine Sea plate (PH) history, modified from Ogawa and Taniguchi (1988). Plutonic and metamorphic rocks are the products in the eastern boundary of the Mineoka plate (MI) to the Pacific plate (PA). EU—Eurasia plate; NNG—North New Guinea plate; NA—North American plate; SB—Shikoku Basin; Iz—Izu Peninsula; B—boninite; Tz—Tanzawa Mountains; eq—equator.
The exact age of the deformation of the plutonic and metamorphic rocks is uncertain. There were at least two stages of deformation, but both were the result of strong shearing in response to dextral transpression. The last stage had the same deformation sense as interpreted from fault geometries delineated on aerial photographs (Mori and Ogawa, 2005). Two stages of deformation were also recognized by Takahashi et al. (2003) and Ogawa and Takahashi (2004) in the basaltic rocks of the Mineoka belt. The two transpressional stages occurred after the subduction-related metamorphism (Fig. 13). The retrograde metamorphism observed in the metamorphic blocks may have occurred during exhumation along the first transpressional fault belt (Fig. 12). Rapid exhumation along an oblique subduction zone (Iwamori, 2003) might be the most plausible setting for this. Present-Day Izu Arc as an Analogue for Development and Emplacement of Mineoka Belt Ophiolitic Rocks In the present-day Izu arc, a large-scale fault of the Sofugan Tectonic Line cuts both the old island arc and the forearc body (Yuasa, 1985) (Fig. 1). Ueda et al. (2004) reported
amphibolite-facies metamorphic rocks with serpentinite near the Ohmachi Seamount, which lies just off the present volcanic front. Inclusions in a garnet grain from the rocks indicate that the rocks underwent eclogite facies metamorphism before retrogression (Ueda et al., 2004). This is evidence for exhumation from depth of rocks along faults (such as those associated with the Sofugan-like Tectonic Line as shown in Fig. 13) into the forearc of the Izu subduction zone in its present-day setting. Taira et al. (1989) proposed that the serpentinite-bearing ophiolitic rocks of the Mineoka belt are equivalent to the northern extension of the serpentinite diapirs at the toe of the Izu-Mariana trench slope. These diapirs include high-pressure metamorphic blocks (Maekawa et al., 1992; Fryer et al., 2000). However, because none of the serpentinites of the Mineoka belt are diapiric or highly sheared (Sato and Ogawa, 2000), and also because the only shearing is restricted to fault boundaries (Ogawa and Takahashi, 2004; Mori and Ogawa, 2005), this proposal may be flawed. In addition, many of the ophiolitic blocks of the Mineoka belt are up to 100 m wide and 100 m long, and they appear to form massive blocks within a fault belt in the Mineoka belt. These are not characteristics of diapirs.
Role of exhumation in a forearc ophiolite mélange belt, Mineoka belt, Japan Another possible analogue for the Mineoka ophiolitic rocks in the Izu forearc is the Hahajima Seamount (Fujioka et al., 2005; Ishiwatari et al., 2006). This seamount is composed of MORB to IAT basaltic rocks, and island-arc andesitic rocks including boninite, and many gabbroic rocks and serpentinite, but not of metamorphic rocks. These rocks are near the marginal serpentinite diapirs on the landward side of the trench slope toe, but the materials differ from those of the diapirs. If the Izu arc analogue is extended to Paleogene and Miocene time, the plutonic and metamorphic rocks dated at ca. 40 Ma might be a product of subduction of the Pacific plate below the Mineoka plate (Fig. 13). After undergoing ductile, mylonitic deformation at depth, these rocks were exhumed to shallower levels, probably along the Sofugan-like Tectonic Line. Then, after the emplacement of the Mineoka ophiolite into the Honshu forearc in front of the Izu arc, either by obduction or another kind of incorporation with the Boso triple junction in its present-day position along the oblique subduction boundary of the Philippine Sea plate, it was entrained in a forearc sliver fault belt and subjected to right-lateral transpressional deformation (Figs. 12, 13, 14). The shapes of the plutonic and metamorphic blocks, and their relation to the surrounding rocks, support such a history. Considering the newly found metamorphic conditions in this chapter as the Heguri-Naka garnet-amphibolite, the maximum temperature and pressure of 560–600 °C and 2.0 GPa preferentially favor the model of the Sofugan-like Tectonic Line (Fig. 12). Summarized History of the Mineoka Belt and Possible Existence of a Mineoka Plate The tectonic setting, age, and deformation characteristics of each of the rock types involved is of prime importance for understanding the implications of the emplacement mechanisms of the plutonic and metamorphic rocks for the development of the present-day Mineoka belt. All of the terrigenous rocks in and around the Mineoka belt are of middle Miocene age except for the Shimanto Supergroup (Eocene) and the thin psammitic layers in the hornblende schist block. In contrast, the pelagic sequences, including blocks of limestone-chert (the Kamogawa Group), are Paleocene (60 Ma) to early middle Miocene (18 Ma), including the Eocene– Oligocene (Mohiuddin and Ogawa, 1998). There are also Early Cretaceous bedded chert (ca. 100 Ma) and MORB-type basaltic rocks (ca. 80 Ma), as shown in this chapter. One of the other ophiolite-suite basaltic rocks is dated as Eocene or at the Eocene– Oligocene boundary (50 Ma). Considering these age and chemical controls, the ophiolite suite rocks might represent the Cretaceous to Eocene (or Oligocene) ocean floor or an island arc setting with Cretaceous to early middle Miocene pelagic cover, although those rocks are highly fragmented at present. The fact that the Paleocene to middle Miocene pelagic sequence in Japan is known only in the Kamogawa Group suggests that it is not an equivalent sequence to those incorporated in
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the Cenozoic accretionary prism. That is, it is not part of the Shimanto Supergroup, where claystone overlies the basaltic cover in the Tertiary section (Taira et al., 1989). Therefore, it is plausible that the oceanic rocks of the Mineoka ophiolite, and its pelagic cover of the Kamogawa Group, do not represent the Pacific plate or the Philippine Sea plate but are from an independent, isolated plate that was not exposed to the strong bottom current that elsewhere eroded the Neogene pelagic sequence. Another possibility is that the Neogene erosional event did not affect the Kamogawa Group because it had been already incorporated into the Honshu forearc before the major stage of bottom current erosion that began in the middle Miocene. Regardless of the origin of this plate, its boundary to the Pacific plate was at an island arc. Such pelagic and bedded limestone-chert lithologies are known only in the Mineoka belt in Japan and the NW Pacific Ocean (Mohiuddin and Ogawa, 1998). Similarly, the overlying glauconitic tuffaceous sediments, the Arakawa Formation, are known only in the Mineoka belt (Mohiuddin and Ogawa, 1996). The present-day shallow sequence of the NW Pacific Ocean includes a large depositional hiatus caused by erosion of most of the Tertiary pelagic sequence by a Miocene bottom current, so that Quaternary clay directly overlies the Cretaceous chert sequence (Ogawa and Kawata, 1998). Thus the Cretaceous bedded chert, the Kamogawa Group, and the overlying Arakawa Formation provide a rare example of the original Cretaceous to middle Miocene stratigraphy of the NW Pacific. We note that a similar pelagic to hemipelagic sequence overlying ophiolites is preserved in North New Guinea (Lus et al., 2004). Therefore, it is reasonable to propose the existence of the Mineoka plate in the region of the Boso triple junction in the early Miocene, as well as the Pacific and Philippine Sea plates (Ogawa and Taniguchi, 1988; Sato et al., 1999) (Figs. 13 and 14). The ophiolite stratigraphy discussed above suggests that such a plate might be a counterpart to the North New Guinea plate, as proposed by Seno (1984). ACKNOWLEDGMENTS We thank John Wakabayashi and Yildirim Dilek for inviting our paper to be a part of this Special Paper. The cooperation of observation, analysis, and discussion with Mia Mohammad Mohiuddin (University of Rajshahi), Hiroshi Sato (Senshu University), Hidetsugu Taniguchi (Josai University), Naoki Takahashi (Chiba Museum and Research Institute), Teruaki Ishii, Li Yin-Bin and Satoru Haraguchi (Ocean Research Institute, University of Tokyo), and Akiko Takahashi (Royal Affairs Agency) are greatly appreciated. Special thanks go to Ichiro Kaneoka, Keisuke Nagao, Y. Takigami, and Hirochika Sumino, and the International Research Center for Nuclear Material Science and the Radioisotope Center, University of Tokyo, for Ar-Ar analyses and preparations. We also thank Hayato Ueda, Hirosaki University, and many other researchers who visited the field with us for discussions in front of the outcrops. Part of the research was supported by a Sabo Society Grant. An early
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draft was reviewed by Tsutomu Ota and Tatsuki Tsujimori, to whom we are grateful. Finally we are thankful for the extensive discussions and suggestions of John Wakabayashi, whose many revisions significantly improved the original manuscript. REFERENCES CITED Chiba, T., 2008, Tectonics of Mineoka belt inferred from geologic structures [B.S. thesis]: Tsukuba, Japan, University of Tsukuba, 20 p. Cloos, M., 1984, Flow mélanges and the structural evolution of accretionary wedges, in Raymond, L.A., ed., Mélanges: Their Nature, Origin, and Significance: Geological Society of America Special Paper 198, p. 71–80. Cowan, D.S., and Page, B.M., 1975, Recycling Franciscan material in Franciscan mélange west of Paso Robles, California: Geological Society of America Bulletin, v. 86, p. 1089–1095, doi:10.1130/0016-7606 (1975)86<1089:RFMIFM>2.0.CO;2. Cowan, D.S., and Silling, R.M., 1978, A dynamic, scaled model of accretion at trenches and its implications for the tectonic evolution of subduction complexes: Journal of Geophysical Research, v. 83, p. 5389–5396, doi:10.1029/JB083iB11p05389. Dilek, Y., and Thy, P., 2006, Age and petrogenesis of plagiogranite intrusions in the Ankara mélange, central Turkey: Island Arc, v. 15, p. 44–57, doi:10.1111/j.1440-1738.2006.00522.x. Ebisawa, N., Sumino, H., Okazaki, R., Takigami, Y., Hirano, N., Nagao, K., and Kaneoka, I., 2004, Construction of I–Xe and 40Ar–39Ar dating system using a modified VG3600 noble gas mass spectrometer and the first I–Xe data obtained in Japan: Journal of Mass Spectrometry Society of Japan, v. 52, p. 219–229. Ernst, W.G., and Liu, J., 1998, Experimental phase-equilibrium study of Al- and Ti-contents of calcic amphibole in MORB—A semiquantitative thermobarometer: American Mineralogist, v. 83, p. 952–969. Fryer, P., Lockwood, J.P., Becker, N., Phipps, S., and Todd, C.S., 2000, Significance of serpentinite mud volcanism in convergent margins, in Dilek, Y., Moores, E.M., Elthon, D., and Nicolas, A., eds., Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program: Geological Society of America Special Paper 349, p. 35–51. Fujioka, K., Tokunaga, W., Yokose, H., Kasahara, J., Sato, T., Miura, R., and Ishii, T., 2005, Hahajima Seamount: An enigmatic tectonic block at the junction between the Izu–Bonin and Mariana Trenches: Island Arc, v. 14, p. 616–622, doi:10.1111/j.1440-1738.2005.00488.x. Gerya, T.V., Stockhert, B., and Perchuk, A.L., 2002, Exhumation of highpressure metamorphic rocks in a subduction channel: A numerical simulation: Tectonics, v. 21, 1056, doi:10.1029/2002TC001406. Graham, C.M., and Powell, R., 1984, A garnet–hornblende geothermometer: Calibration, testing, and application to the Pelone Schist, Southern California: Journal of Metamorphic Geology, v. 2, p. 13–31, doi:10.1111/j.1525-1314.1984.tb00282.x. Hanamura, Y., and Ogawa, Y., 1993, Layer-parallel faults, duplexes, imbricate thrust and vein structures of the Miura Group: Key to understanding the Izu fore-arc sediment accretion to the Honshu forearc: Island Arc, v. 2, p. 126–141, doi:10.1111/j.1440-1738.1993.tb00081.x. Hirano, N., and Okuzawa, K., 2002, Occurrence of the Indian Ocean–type isotopic signature in basalts from Philippine Sea plate spreading centers: An assessment of local versus large-scale processes: Journal of the Geological Society of Japan, v. 108, p. 691–700. Hirano, N., Ogawa, Y., Saito, K., Yoshida, T., Sato, H., and Taniguchi, H., 2003, Multi-stage evolution of the Tertiary Mineoka Ophiolite, Japan: New geochemical and age constraints, in Dilek, Y., and Robinson, P., eds., Ophiolites in Earth History: Geological Society [London] Special Publication 218, p. 279–298. Hiroi, Y., 1995a, Metamorphic block in the Kamogawa Harbor, in Chiba Prefecture, ed., Natural Reserve Report for Geology and Minerals: Report from Chiba Prefecture, Japan, Education Board of Chiba Prefecture, 105 p. Hiroi, Y., 1995b, Metamorphic block at Hegurinaka, in Chiba Prefecture, ed., Natural Reserve Report for Geology and Minerals: Report from Chiba Prefecture, Japan, Education Board of Chiba Prefecture, 120 p. Holland, T., and Blundy, J., 1994, Non-ideal interactions in calcic amphiboles and their bearing on amphibole-plagioclase thermometry: Contributions to Mineralogy and Petrology, v. 116, p. 433–447.
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The Geological Society of America Special Paper 480 2011
Mélanges of the Franciscan Complex, California: Diverse structural settings, evidence for sedimentary mixing, and their connection to subduction processes John Wakabayashi Department of Earth and Environmental Sciences, California State University, Fresno, 2576 E. San Ramon Avenue, Mail Stop ST-24, Fresno, California 93740, USA
ABSTRACT The classic mélanges of the Franciscan Complex of California comprise a variety of structural-tectonic settings and give insight into mélange forming processes and material movement patterns within subducting plate margins. Structural settings of mélanges include (1) shale matrix mélanges that occur within, cutting, and bounding coherent nappes, and (2) serpentinite matrix mélanges that occur at the structurally highest levels and commonly cut blueschist and higher grade coherent nappes. Although nappe-bounding zones may have accommodated tens of kilometers or more of movement and represent paleosubduction megathrust zones, exposures at El Cerrito Quarry in the eastern San Francisco Bay area suggest that most displacement is accommodated in a narrow zone (5 m wide in this example) of brittle faults between the nappe-bounding mélange and the coherent nappes. Intranappe mélanges, or their boundaries, accommodated smaller displacements owing to similar or identical units bounding them. Blueschist-facies sedimentary breccias present in the northwest Diablo Range range from nearly undeformed to strongly foliated rocks that appear to be similar or identical to classic foliated shale mélange matrix. Most of the breccia clasts exhibit blueschist-facies metamorphic mineral growth that predated sedimentation, indicating exhumation of source material from blueschist depths, followed by deposition and resubduction to blueschist depth. Published apatite fission-track data suggest that the source of the clasts may have been completely eroded since exhumation. Blueschist-facies sedimentary breccia on one side of a high-grade blueschist block within serpentinite-matrix mélange in southern Sonoma County suggests that the mélange originated as sedimentary serpentinite, following exhumation of blueschist and higher grade rocks early in Franciscan subduction history, prior to deposition of metaclastic rocks (absent as blocks in the mélange). The strongly foliated serpentinite matrix apparently recrystallized during resubduction to blueschist-facies depth, erasing sedimentary textures, in contrast to the unstrained and unmetamorphosed sedimentary serpentinite in the basal Great Valley Group that may be a temporal equivalent. These relationships suggest an early forearc history in which exhumed
Wakabayashi, J., 2011, Mélanges of the Franciscan Complex, California: Diverse structural settings, evidence for sedimentary mixing, and their connection to subduction processes, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 117–141, doi:10.1130/2011.2480(05). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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J. Wakabayashi serpentinite and associated high-grade blocks were shed into the trench and into what later became the forearc basin, prior to forearc high development and significant clastic sedimentation. The occurrence of high-grade metamorphic blocks in structurally low mélanges, as well as their preferential localization along the boundaries of mélanges, also suggests sedimentary mixing. Thus, many Franciscan mélanges, including the nappe-bounding mélanges that might be expected to have accommodated large displacement, show evidence of early sedimentary mixing. Most of the block-in-matrix fabric and introduction of exotic blocks may have resulted from sedimentary processes rather than tectonic strain. In contrast, some mélanges that cut across the bedding or foliation of coherent units apparently had a diapiric rather than a sedimentary origin.
INTRODUCTION In his classic paper, Hsü (1968) used the Franciscan Complex of coastal California (Fig. 1) as the type example to illustrate his principles of mélanges: rock units with a foliated matrix enclosing various blocks. Further studies of various aspects of mélange evolution (e.g., Aalto, 1981; Aalto and Murphy, 1984; Aalto, 1989; Raymond, 1970, 1974, 1984; Maxwell, 1974; Gucwa, 1975; Cloos, 1982, 1983, 1985; Cloos and Shreve, 1988; Cowan, 1974, 1978, 1985; Cowan and Page, 1975; Page, 1978; Jeanbourquin, 2000) have revisited the Franciscan as a type mélange locality. Within the Franciscan and related units, distinct structural settings can be defined on the basis of general field relationships. After an introduction of Franciscan tectonics, I review these settings, then proceed to examine new evidence for an early sedimentary mixing stage that may be common to many Franciscan mélanges, including those that might be considered most likely to have been formed entirely by tectonic processes. The Franciscan Complex of California formed by offscraping and tectonic underplating of rocks (subduction accretion) from an east-dipping subducting plate from 160 Ma to <20 Ma (Wakabayashi, 1992). The Franciscan is well known for its highpressure–low-temperature (HP-LT) metamorphism (Ernst, 1970, 1971). At least 25% of the exposed rocks have undergone HP-LT blueschist or higher grade metamorphism (Wakabayashi, 1999). The Coast Range Ophiolite, consisting of serpentinized ultramafic rocks, gabbro, basalt, and other plutonic and volcanic rocks, structurally overlies the Franciscan (Hopson et al., 1981, 2008). The Coast Range Ophiolite formed in a supra-subduction zone setting (Shervais and Kimbrough, 1985; Shervais, 2001) and is depositionally overlain by well-bedded sandstones and shales of the Great Valley Group that are coeval with clastic sedimentary rocks of the Franciscan (Dickinson, 1970). The Coast Range Ophiolite and Great Valley Group lack HP-LT metamorphism. The three subparallel geologic provinces of the Franciscan, Great Valley Group, and the Sierra Nevada batholith (see Fig. 1) represent, respectively, the subduction complex, forearc basin deposits, and the magmatic arc of an ancient arc-trench system (Dickinson, 1970).
Franciscan lithologies are mostly detrital sedimentary rocks with subordinate basaltic volcanic rocks and chert, and minor limestone. The detrital rocks are offscraped and underplated trench-fill sediments (Dickinson, 1970). The pelagic and volcanic rocks apparently represent fragments of seamounts, other oceanic rise deposits, and the pelagic cover and upper part of the subducted oceanic crust (Hamilton, 1969; MacPherson, 1983; Shervais, 1990), probably with a component of olistostrome blocks from the upper plate (MacPherson et al., 1990; Erickson et al., 2004; Erickson, this volume). The Franciscan comprises “coherent” mappable bodies (referred to as coherent nappes or terranes) and mélange units that consist of a sheared matrix with included blocks (Maxwell, 1974; Blake et al., 1988; Wakabayashi, 1992, 1999a). The volumetric proportion of mélange versus coherent units varies along and across strike in the Franciscan, and this variation is apparent even when post-subduction strike-slip faulting is restored (Ernst, 1975; Wakabayashi, 1992, 1999a) (Fig. 2). Collectively, mélange and coherent units define a series of thrust nappes, and the age of incorporation into the subduction complex youngs structurally downward (Wakabayashi, 1992, 1999a; Wakabayashi and Dumitru, 2007; Ernst et al., 2009; Snow et al., 2010; Dumitru et al., 2010) consistent with models of progressive offscraping and underplating in a subduction complex (e.g., Karig and Sharman, 1975). Franciscan subduction terminated without collision by conversion of a subducting plate margin to a transform margin as a result of ridge-trench interaction (e.g., Atwater 1970). Thus, no other crustal unit structurally underlies the Franciscan on a regional scale. Within the former trench-forearc system, comprising the Franciscan, Coast Range Ophiolite, and Great Valley Group, most of the mélanges have Franciscan affinity. Extensive serpentinite matrix mélange crops out between Coast Range Ophiolite and Franciscan exposures in the northern Coast Ranges (Hopson and Pessagno, 2005; Shervais et al., this volume). Although these serpentinite matrix mélanges were considered Franciscan by Hopson and Pessagno (2005) because of differences in igneous and pelagic sedimentary history between these rocks (called the Tehama-Colusa serpentinite mélange by these authors) and the Coast Range Ophiolite, I will follow the interpretation of
Mélanges of the Franciscan Complex, California Shervais et al. (this volume), who consider part of the mélange to represent the tectonic contact between the Coast Range Ophiolite and Franciscan, with much of the unit having Coast Range Ophiolite affinity, based on the lack of H-P metamorphism and the supra-subduction zone geochemical character of most of the blocks. A sedimentary serpentinite unit forms a mélange unit in
125° W
124°
the basal Great Valley Group (Phipps, 1984), and spatially limited occurrences of mélanges exist that cannot be conveniently fit into either of the three main units crop out locally (Wakabayashi, 2004). In this paper the term mélange will not have a genetic basis, and I use it to connote rock units with a block-in-matrix fabric that have a tectonic, sedimentary, or dual origin. This
121°
122°
123°
119
LEGEND
MODOC
Franciscan Complex Blueschist grade, jadeite rare or lacking in sandstones, not schist
fsch
PLATEAU Blueschist grade, mostly sandstone and shales, blueschist grade,
KLAMATH MTNS
41°N
jadeite common, not schist
41°
Schist, blueschist or higher grade: csch grain size exceeds 0.5 mm; fsch grain size commonly less than 0.3 mm
Mendocino Triple Jct.
GO Fp Fp
Fz
A
C
40°
OM
Other rock units and faults GO Great Valley Group and/or Coast Range Ophiolite Ksb Salinian block
KJN Nacimiento block
Ca Calaveras fault
G Greenville fault
SIE
GV Green Valley fault
H Hayward fault
M Maacama fault
SA San Andreas fault
39°
SGH San Gregorio-Hosgri fault
A fsch
SA
GO
A Transects for which Franciscan nappe columns are shown in
fi
GO GO
GV
RC
H-
38°
Fp
SA
Fz? csch
GO
L
c
D
Figure 2
RA
ci
csch
Approximate western limit of subduction complex rocks
Figs.10B, 15
Ksb
0
100 km
38°
Fb
SA
R
n 37°
GO
120°
Sunol Regional Wilderness B (Figs. 8, 9, 12) Photo, Fig. 7
119°
DA VA
GO
Ca
ea
F Fig. 4
Fp
NE
San Francisco
G
H
Oc 124°
H-RC Healdsburg-Rodgers Creek fault
A
Fp
NT
Pa
E
RR
Fp
39°
CE
S T
Fz
Zeolite grade
40°
fsch Fp?
Prehnite pumpellyite grade, not schist
Fp/b Prehnite pumpellyite or blueschist grade, not schist
37°
GO
N
SGH
C
csch
Ksb?
G
csch
Fp/b
Y
csch
E LL VA
A
G
Ksb
123°
E S
36° 122°
KJN
121°
GO
N
H Fp/b 120°
36°
119°
Figure 1. Distribution of Franciscan and other basement rocks of central and northern California, showing Franciscan of different metamorphic grades. Modified from Wakabayashi (1999).
120
J. Wakabayashi
descriptive, rather than genetic, definition echoes the definition of Silver and Beutner (1980).
or Coast Range Ophiolite. I shall review these five settings, preceded by a discussion of Franciscan mélange unit nomenclature, because it has caused considerable confusion in the past.
GENERAL MÉLANGE SETTINGS WITHIN THE FRANCISCAN AND RELATED UNITS
Franciscan Mélange Terminology: Clarification
Mélanges appear to occur in the following structural tectonic settings in Franciscan and related units (Figs. 2, 3): (1) between (bounding) coherent nappes (alternatively, these may be called terrane bounding mélanges); (2) within otherwise coherent nappes; (3) at the structural top of the Franciscan between the Franciscan and the overlying mantle or Coast Range Ophiolite– Great Valley Group crustal rocks; (4) as sedimentary deposits in the basal Great Valley Group; and (5) in zones that cannot be readily assigned to either the Franciscan, Great Valley Group,
Franciscan unit nomenclature has confused many readers, and the most bewildering nomenclature applies to mélanges. Much of this confusion has resulted from extension of the historic division of the Franciscan into three west-to-east subparallel belts, the Coastal, Central, and Eastern Belts (Berkland et al., 1972) beyond the intended boundaries of this work. This geographically imposed division applies reasonably well in the Franciscan of the northern Coast Ranges, where the zeolitefacies Coastal Belt consists of mostly intact but folded and
North Diablo Range
North Coast Range
Cazadero
C
B
A
South Diablo Range
fsch
Eastern Belt
Pickett Peak terrane
121A 135dz
VS
85-108dz
145-169AL bsj PPmg EM 136-151f
WS/AC SSS
154A 132A
PPss
bs
b
OP 92u
(70-100)
bs
120dz
mel
p
GM
BH bs 85f
SSS 144dz mv
mel p
90f RN
G
95f
95f HPSZ NR
mel p
DDC
95f
AI bsj Ap 95f mel97dz
p
p
PFR
p
MH CCFZ
mel
p SB
NQ 85f 83dz
mel 65f
G mel
H Ophiolite
mel MH p 95f
mel
mel
MP
b
85f S/SN
mel
z CoB
TM 102dz 100dz
52dz HFZ
119A
HM 95f
San Francisco Bay Area Parkfield
F
mel
111dz
YB
WCcsch
mel
119A
bs
E
Ophiolite
Ophiolite SFMS
Clear Lake
D
b P 85f
S 55f
mel p
PRM ECM SMM 95f TCM
p
ERM LM 85f
Central Belt
p
0
mel
65f CoB
z
Coastal Belt
p
10 km
Abbreviations: Column A: SFMS: South Fork Mountain schist, VS: Valentine Springs Formation, YB: Yolla Bolly terrane, HM: Hull Mtn. metagraywacke, PRM: Poison Rocks mélange, ECM: Elk Creek mélange, SMM: Sanhedrin Mtn. mélange, TCM: Tin Cabin mélange, ERM: Eel River mélange, LM: Laytonville mélange, CoB: Coastal Belt. Column B: EM: Eylar Mountain terrane, OP: Ortigalita Peak metagraywacke, GM: Garzas mélange; BH: Burnt Hills terrane. Column C: WS/AC: Panoche Pass Willow Spring slab and Antelope Creek slab, SSS: Skaggs Springs schist, PPmg: Panoche Pass jadeite-bearing metagraywacke, PPss: Panoche Pass jadeite-free metagraywacke. Column D: WC: coherent schists of Ward Creek, SSS: Skaggs Springs schist, mv: metavolcanics, incipient blueschist, RN: Rio Nido terrane, CoB: Coastal Belt. Column E: G: Geysers terrane, DDC: Devil’s Den Canyon terrane, PFR: Pine Flat Road terrane. Column F: TM: Tiburon mélange, AI: Angel Island nappe, A: Alcatraz terrane, HPSZ: Hunters Point shear zone, NR: Nicasio Reservoir terrane, MH: Marin Headlands terrane, CCFZ: City College fault zone, SB/NQ: San Bruno Mtn./Novato Quarry terranes, HFZ: Hillside fault zone. Column G: MH: Marin Headlands terrane, P: Permanente terrane, S: Salinia. Column H: MP: mélange with blocks of Permanente terrane, S/SN: Salinia/Sierra Nevada. For all columns: b: incipient blueschist metamorphism of mostly volcanic rocks, z: zeolite facies; p: prehnite pumpellyite facies, bs: blueschist metamorphism-jadeite absent or rare, bsj: same as bs except jadeite common, fsch: fine grained schists (completely recrystallized), csch: coarse grained schists (completely recrystallized), mel: major mélange zone. 120: estimated age (Ma) of accretion from: A (Ar/Ar dating of metamorphism), AL (Ar/Ar and Lu-Hf dating of metamorphism), U (U/Pb dating of metamorphism), dz (age of clastic deposition from detrital zircon chronology), f (age of clastic deposition from fossils). Note that 119A for Column A is an igneous age of a pre-metamorphic intrusion into graywacke. (70-100): age (Ma) of exhumation from blueschist depths to approximate depth of Great Valley Group from Tagami and Dumitru (1996).
Figure 2. Correlation of Franciscan units as thrust nappes to relative structural position and estimated accretion ages. Blueschist facies units are shaded. Column locations keyed to Figure 1. All contacts are tectonic; incorporation age (not depositional age) decreases downward. Scale is approximate. Figure revised from Wakabayashi (1999) with additional data from Mertz et al. (2001), Anczkiewicz et al. (2004), Dumitru et al. (2010), Wakabayashi and Dumitru (2007), and Snow et al. (2010).
Mélanges of the Franciscan Complex, California imbricated sandstones and shales with some broken formation; the Central Belt is chiefly mélange with prehnite pumpellyite facies matrix, and the Eastern Belt mainly comprises coherent blueschist-facies thrust sheets. The term Central Belt has become nearly synonymous with Franciscan mélanges. However, the proportion of mélange at any given structural zone of the Franciscan varies along strike. For example, in the San Francisco Bay area, structural zones equivalent to (accreted at approximately the same time as) the Central Belt are primarily prehnite-pumpellyite facies coherent nappes separated by discrete mélange zones, and in the Diablo Range equivalent structural zones are blueschist-facies coherent nappes separated by mélange zones (Wakabayashi, 1992, 1999a) (Fig. 2). Similar problems with along-strike variation in Franciscan units
121
have been noted by Aalto (1989) and Erickson (this volume). Further confusion comes from the use of the term Central terrane (Blake et al., 1982, 1984, 1988). This term applies to all mélange zones in any part of the Franciscan Complex other than the Coastal Belt. The “Central terrane” name has caused confusion for readers and researchers who equate it with the “Central Belt,” whereas it refers to mélanges within the Eastern Belt as well as the Central Belt. Owing to the fact that the belt nomenclature poorly fits the entire Franciscan, and the “Central terrane” name causes confusion owing to its (incorrectly) perceived relationship to the belt nomenclature, I prefer structurally- or locality-specific mélange names (e.g., Hsü and Ohrbom, 1969; Raymond, 1974; Cowan, 1974; Maxwell, 1974; Gucwa, 1975; Aalto, 1989; Wakabayashi, 1992).
Tectonic setting of mélanges: Franciscan and related terranes Mélanges within coherent nappes
Franciscan subduction complex
160-90 Ma
Coast Range Ophiolite Serpentinite (CRO) mud volcano: sedimentary serpentinite Mélange at Basal forearc basin structural top deposits: Great of Franciscan Valley Group (GVG)
Mélanges between coherent nappes
Volcanic arc
50 km Mantle
Mantle
Approximate scale 0 0
50 km
Present Coast Ranges
Central Valley Folded Franciscan GVG sedimentary nappes serpentinite Cenozoic post-GVG GVG CRO outlier CRO sediments
Sierra Nevada
Mantle
Figure 3. Diagrams showing tectonic settings of mélanges within the Franciscan and related rocks. Note that the intention of the diagrams is not to infer that Franciscan mélanges were generated by tectonic processes (shearing) only. The diagrams merely show the structural settings of mélanges. An early sedimentary history of mixing followed by partial subduction is compatible with the structural settings illustrated here.
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J. Wakabayashi
Mélanges between Coherent Nappes (Nappe-Bounding Mélanges)
serpentinite is widespread (Coleman, 2000), and most serpentinites of this type occur within shear zones bounding coherent nappes (Wakabayashi, 1992). However, most intra-Franciscan serpentinites in these settings lack included blocks other than rare gabbro, so the serpentinite bodies themselves are not serpentinite-matrix mélanges (Wakabayashi, 2004). Blocks in predominantly shale-matrix mélanges are mostly sandstone
Mélanges that separate coherent nappes (Figs. 2–4) range from tens of meters to several kilometers in structural thickness (a typical example is shown in Fig. 5). The matrices of these mélanges commonly consist of foliated shale. Intra-Franciscan
EXPLANATION Franciscan units 122°30' KJtm: Tiburon mélange: shale and serpentinite Knq: Novato Quarry terrane: 102dz matrix mélange K-feldspar-bearing arkosic sandstone, Photo, Fig.10A Knq Ka Ka KJmh KJai: Angel Island nappe; blueschist facies shale 100dz KJai metavolcanic, metaclastic rocks Esb: San Bruno Mtn. nappe: Jo?sp 83dz 122°15' Ka: Alcatraz terrane (nappe): arkosic K-feldspar-bearing arkosic sandstone, KJtm Knq 153-156AL KJtm Ka sandstone, shale shale KJai KJpb: Rocks of Point Bonita: basalt KJmh KJp: Permanente terrane: basalt, 37° Jo KJmh: Marin Headlands terrane sandstone, shale, limestone, chert KJai 52'30" 37°52'30" KJai (nappe): basalt, chert, Ka KJu: Undifferentiated Franciscan Jo KJu Tu KJai Photo, Fig. Fig.55 Ka sandstone, shale Angel Island KJmh KJmh Non-Franciscan units e Pt. Richmond
KJtm KJai
D
Jo Josp
F
E
N SA
C
95f
t
n lde
Ga
KJpb
95f
Hu
97dz
Jofv
ts .z.
y Cit
95f
C
ol
KJmh
Hil lsid Sf.z. e f. an z. Br 52dz un Esb oM Sa tn. n Br u n KJu o f.
Josp Josp Jogb
KJmh
FS P
lt
fau
KJp
A
KJgv Jogb Josp
N
KJu
122°15'
122°30'
0
Hu
nte
rs
Po in
10 km
KJtm San KJai C D Francisco B E 102dz F KJtm 97dz ar Red Point Ka zo 153-156AL Ka 52dz Jo Rock Richmond 100dz KJmh ne 95f KJai
ts
ef au
KJmh o lt z lle ge fa one u lt zo ne Esb
Ka KJmh ?
?
Knq ?
KJmh Knq 83dz
Hayward f.
Kg
KJu
95f
C
Kg
KJu
ty Ci
KJp
KJp
lsid Hil
P
Kg
San Andreas f.
KJp
he
Point Richmond f.
A
Kg
Major mélange zones
locality (Fig. 11)
Coyote Point KJmh
P
P
83dz: accretion age estimate as given in Fig. 2
KJgv Jofv Jomv Jofv Jogb Hayward KJgv37°37'30"
Photo, Fig. Fig.66
KJp
Jofv Jogb
d
lt au sf ea dr An
37°37'30"
BAY
n Sa
Pacifica
KJgv: Great Valley Group sandstones, shales Kg: Plutonic rocks of the Salinian block P: Paleocene sandstone, shale of the Salinian block Jofv FS: Pre-Eocene Franciscan-Salinian contact Jogb Tu: Tertiary rocks undifferentiated Jofv
KJu KJgv
ar yw Ha
leg e
Jofv Jofv
Po in
KJmh
Coast Range Ophiolite Jo: Undifferentiated; Josp: serpentinite; Jogb: gabbro; Jomv: Mafic volcanics; Jofv: felsic volcanics
Oakland
CO CIS AN FR
San Francisco
B Ka
Ka
nte
rs
Josp KJu Josp
Knq
Ka
Go
Figure 4. Franciscan Complex geology of the San Francisco Bay area with major mélange zones highlighted. Modified from Wakabayashi (2004).
Mélanges of the Franciscan Complex, California with lesser amounts of basalt, chert, and limestone; the block abundance is approximately representative of relative lithologic abundances in the subduction complex as a whole. Many of the blocks in the mélanges appear to have been derived from coherent nappes bounding the mélanges, but all or nearly all of the nappe-bounding mélanges contain blocks absent from the bounding nappes (“exotic blocks”). Sandstones (graywackes), the most common block type in these mélanges, may differ in framework composition from those of the bounding coherent nappes in some mélanges (Jayko and Blake, 1984), whereas in other mélanges the sandstones appear identical to those in the bounding coherent units (Rubin, 2002). Mélange matrix appears to be of the same metamorphic grade as the bounding coherent nappes (Cloos, 1983; Blake et al., 1988; see also new data presented below in the major section that discusses tectonic and sedimentary origins of mélanges). The majority of the blocks in the mélanges appear to be isofacial with the matrix, but some of notably higher grade are present (Cloos, 1983). The metamorphic grade of the mélange matrix ranges from zeolite facies in the youngest, structurally lowest part of the Franciscan (i.e., Coastal Belt) to prehnite-pumpellyite facies in the structurally intermediate levels, to blueschist facies at the structurally highest levels (Cloos, 1983; Blake et al., 1988). The highest grade metamorphic blocks exceed the metamorphic grade of all mélange matrix and include eclogite, garnet amphibolite, and amphibolite, commonly referred to as high-grade blocks (Coleman and Lanphere, 1971). Most high-grade blocks show variable degrees of blueschist overprinting (Moore, 1984; Moore and Blake, 1989; Wakabayashi, 1990). These volumetrically rare
123
but widespread blocks are entirely metabasites and metachert and commonly have “rinds” rich in actinolite and chlorite, indicating that ultramafic rock (variably serpentinized) surrounded them during part of their history (Coleman and Lanphere, 1971; Moore, 1984). Serpentinite crops out in close proximity to highgrade blocks in terrane-bounding mélanges, but it is rarely found in direct contact with the blocks; the blocks are commonly surrounded by a shale matrix (Cloos, 1986). Where large intra-Franciscan serpentinite bodies occur, high-grade blocks are commonly found along their margins but can seldom be observed embedded in matrix owing to inadequate exposure, so it is not clear whether the blocks are (1) concentrated in part of the shale matrix mélange near the contact with serpentinite, (2) in a selvage of serpentinite matrix mélange along the shale-serpentinite contact, or (3) in serpentinite matrix mélange lenses that are themselves blocks in the shale matrix mélange near the serpentinite contact (see below). The highgrade blocks, and their rare coherent equivalents that crop out at the structurally highest level of the Franciscan, yield the oldest metamorphic ages in the Franciscan (Wakabayashi and Dumitru, 2007). These rocks exhibit supra-subduction zone chemistry and have been suggested to be the remnants of a metamorphic sole that formed during the inception of Franciscan subduction within supra-subduction zone crust near a spreading center (Saha et al., 2005; Wakabayashi et al., 2010). The protolith ages of blocks in mélanges mirror the protolith ages of the coherent nappes that range from Early Cretaceous to mid-Cenozoic (e.g., Blake et al., 1988), but protolith ages for basalt or pelagic rock significantly exceed accretion age
Figure 5. Photo of El Cerrito Quarry (see location in Fig. 4), showing a mélange zone that separates a blueschist facies metagraywacke nappe above, from a prehnite-pumpellyite facies metagraywacke nappe (foreground) below. Most of the blocks (visible as areas with blockier texture in the photo) in the mélange are sandstone with minor basalt and chert. One basalt block is labeled.
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J. Wakabayashi
(Wakabayashi, 1992, 1999; Snow et al., 2010; Wakabayashi et al., 2010). The protolith age range of blocks for any given mélange may include older ages than the bounding coherent nappes but does not include any ages younger than those in the bounding nappes, suggesting that the mélanges formed during rather than after the accretion of the coherent nappes (Wakabayashi, 1992, 1999a; Snow et al., 2010). The youngest protolith ages in the mélanges and their bounding nappes (the sandstones) appear to reflect the timing of subduction-accretion (Wakabayashi, 1992, 1999a; Snow et al., 2010). The metamorphic age range for blocks in mélange is not well established, except for the high-grade blocks, which is significantly greater than the age of accretion of the mélange for all but the structurally highest zone (maximum difference of ~80 m.y.), suggesting sedimentary or tectonic recycling of the blocks (Wakabayashi, 1992; Wakabayashi and Dumitru, 2007), as will be discussed later. Exclusive of the high-grade blocks, most of the basaltic blocks in these mélanges exhibit oceanic-island basalt (OIB) or mid-oceanic-ridge basalt (MORB) chemistry, similar to the lower grade coherent metabasalts of the Franciscan (Shervais, 1990; MacPherson et al., 1990; Wakabayashi et al., 2010). Exceptions are blocks of blueschist facies and a lower grade of island arc chemistry that have no coherent equivalents in the Franciscan and may be olistostrome blocks derived from the structurally overlying supra-subduction zone Coast Range Ophiolite (MacPherson et al., 1990; Erickson et al., 2004; Erickson, this volume). The zones separating coherent nappes, but not necessarily the mélanges themselves (see below), appear to have accommodated large tectonic displacement, based on a contrast in accretionary age between the bounding nappes and, in some cases, a contrast
in metamorphic grade between the bounding nappes (Wakabayashi, 1992; Ernst et al., 2009; Snow et al., 2010). Contrasts in apparent accretion age across these mélange zones suggest that these zones (or their borders) represent the former subduction megathrust zone that accommodated active subduction during the time between the accretion of the bounding nappes. For an average plate convergence rate of 10 cm/yr (100 km/m.y.), this suggests very large displacement for a relatively small accretion age difference, so the strain accommodated by nappe-bounding mélanges or their boundaries may be large. A good example of a nappe-bounding mélange crops out at an abandoned quarry in El Cerrito in the eastern San Francisco Bay area (Fig. 5). This mélange separates coherent metasandstones and shales of the blueschist-facies Angel Island nappe above from coherent prehnite-pumpellyite facies sandstones and shales of the Alcatraz nappe below, indicating a minimum of 10–15 km of tectonic transport in a thrust sense, depending on the original fault dip (Wakabayashi, 1992, 2005; also “Upper El Cerrito” and “Lower El Cerrito,” respectively, of Snow et al., 2010). The shale-matrix mélange has a structural thickness of ~30–40 m. Most of the displacement between the coherent nappes appears to be within the uppermost 5 m of the zone between them, or, depending on how the mélange boundaries are defined, along the upper boundary of the mélange. The blocks and matrix are prehnite-pumpellyite facies structurally below this level, whereas the uppermost 5 m consists of blueschistfacies metashale and sandstone blocks and can be considered a broken formation, if not outright coherent. The blueschist-facies sandstone blocks within this 5-m-wide zone contain neoblastic fine sprays (commonly <0.3 mm) of jadeitic clinopyroxene and
Figure 6. Shale matrix mélange within a coherent nappe, Linda Mar (Pacifica). Blocks include graywacke, metabasalt, and chert. View to the north. Abbreviations: ch—chert; mx—shale matrix; ss—sandstone-graywacke; v— metabasalt. Location in Figure 4.
Mélanges of the Franciscan Complex, California
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lawsonite, whereas the coherent metagraywacke above this zone contains neoblastic coarse (to 0.5 mm) blocky fans of jadeitic clinopyroxene, glaucophane, and lawsonite, qualitatively suggesting some degree of metamorphic gradient between the “broken formation” zone and the metagraywacke above. Given that the strain within these rocks is too small to account for the displacement, the displacement must be accommodated on discrete brittle faults cutting this 5-m-wide zone. The lower grade, structurally lower part of the mélange (or the “actual” mélange) contains abundant chert and metabasalt blocks in addition to the more common sandstone blocks in a foliated shale matrix. Although the shale matrix of the mélange and the 5-m-wide zone on top of it exhibit an asymmetric fabric consistent with thrusting of the upper nappe over the lower one (Wakabayashi, 1992, 2005), most of the displacement between the nappes may have been accommodated on a zone of brittle faults between blueschist-facies rocks above and prehnite-pumpellyite facies mélange below rather than within the mélange. Mélanges within Otherwise Coherent Nappes Mélanges occur within coherent nappes of the Franciscan, and they commonly mark internal thrust contacts within them, although in other localities such mélanges cut across the structural grain, truncating bedding or foliation (Wahrhaftig, 1984; Wakabayashi, 1992) (Figs. 3 and 4). Some of the intra-nappe mélanges may thin along strike from a zone a few hundred meters wide to a discrete fault separating thrust imbricates (Wahrhaftig, 1984). Along strike, some intra-nappe mélanges become wider (thicker) and coalesce with other subparallel mélange zones, so that the so-called coherent nappe grades into a mélange itself (in such cases the intra-nappe mélanges merge into nappe-bounding mélanges and reach thicknesses of several kilometers); this mirrors along-strike variation of mélange proportion in the Franciscan as a whole. The Permanente nappe (terrane) is a good example of this variability. It is largely coherent with discrete mélange zones in the northern San Francisco Peninsula region (Figs. 4 and
Figure 7. Photo of mélange zone within an otherwise coherent nappe that includes a lens of serpentinite matrix mélange within shale matrix mélange. This photo was taken looking NW from the Panoche Pass road near the western boundary of Franciscan exposures. Some of the blocks in the serpentinite lens are circled, but many smaller ones are not shown. Most of the blocks are amphibolite partly overprinted with blueschist assemblages except for the large garnet amphibolite–eclogite block that is shown. The metagraywacke on both sides of the mélange is blueschist facies.
6) (Wakabayashi and Moores, 1988; Wakabayashi, 1999b) but has a much higher proportion of mélange to the south in the San Jose area (south of the area of Fig. 4) (McLaughlin et al., 2002). The matrix, block types, metamorphism, and contrasts between blocks and surrounding coherent nappes are the same as for nappe-bounding mélanges, with one exception. Serpentinite matrix mélange crops out in several localities (see examples in Figs. 7–9) where mélange zones cut coherent blueschist facies units situated at relatively high structural levels within the
Figure 8. One of several mélange zones with high-grade blocks that cut coherent nappes in the Sunol Regional Wilderness area (location in Fig. 9). Viewing direction is north. Note that the field relationships and degree of exposure shown are fairly common for this type of mélange, whereas the exposure afforded by the road repair excavation in Figure 7 is exceedingly rare.
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Figure 9. Geologic map of part of the Sunol Regional Wilderness, eastern San Francisco Bay area, showing the locations of photos in Figures 8 and 12. Location shown in Figure 1.
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Mélanges of the Franciscan Complex, California Franciscan Complex. Near Panoche Pass (Figs. 1 and 7), a serpentinite matrix mélange is itself a lens surrounded by foliated shale. Blocks within the serpentinite matrix mélange include amphibolite, garnet amphibolite–eclogite, and blueschist (Fig. 7). The serpentinite matrix is foliated and friable and is composed of mainly lizardite, but selvages of antigorite-bearing serpentinite are present as well as actinolite-chlorite schist that occurs as rinds on some of the blocks and as separate blocks in both the serpentinite and the shale matrix mélange. It is not clear that this type of field relationship (serpentinite matrix mélange as a block in shale matrix mélange) is a common relationship, owing to lack of similarly high-quality exposures, as will be discussed below. The mélange zone itself ranges from ~2 to 30 m thick and separates slabs of blueschist-facies (jadeite-bearing) metagraywacke. The mélange is exposed a few hundred meters east of a contact, marked by serpentinite, between the Franciscan and Great Valley Group (Ernst, 1965). The serpentinite may be related to the Coast Range Ophiolite, although the Great Valley Group–Franciscan contact is apparently marked by a fault that truncates Franciscan units, so the serpentinite may be an intra-Franciscan body (Ernst, 1965). The mélange appears to strike toward the Franciscan– Great Valley Group contact, but it is difficult to trace through thick brush. Similar mélange zones occur in the northern Diablo Range in the Sunol Regional Wilderness area (Figs. 8 and 9). At one locality (Fig. 8) a serpentinite-bearing zone occurs between slabs of apparently identical blueschist-facies metagraywacke. It is not clear whether this mélange actually truncates bedding in the adjacent metagraywacke, but the general branching pattern of mélange zones in this area suggests that they must cut across bedding and foliation locally (Fig. 9). The serpentinite consists of small (<1 m) antigorite-rich serpentinite blocks or lenses surrounded by foliated lizardite-rich serpentinite. The serpentinite is a lens within a broader zone of shale-matrix mélange, although one side of the serpentinite is in direct contact with bounding coherent graywacke (Fig. 9). Blocks have not been found embedded in the matrix, but several eclogite blocks, as well as abundant blocks of actinolite-chlorite schist, crop out near the border of the serpentinite body and graywacke. Similar field relationships of high-grade blocks are associated with other mélanges in this area; these blocks preferentially lie along the boundaries of mélange or serpentinite (localities marked with stars in Fig. 9). Mélange zones containing high-grade blocks and serpentinite also appear to cut across the foliation of the blueschist-facies Skaggs Springs schist (>50-km-long belt of “csch” crossed by transect D in Fig. 1), but, similar to the localities in the Sunol Regional Wilderness, the actual contacts of blocks and matrix have not been documented (Wakabayashi, 1992). The general field relations appear similar to those exposed at the locality of Figure 7, and it is likely that better exposures would reveal essentially identical relationships; the locality at Figure 7 was exposed as a result of landslide repair and had already degraded significantly within a year as a result of slope wash encrusting the outcrop.
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Mélanges at the Structural Top of the Franciscan The structurally highest zones in the Franciscan are rarely preserved, possibly because later faulting has excised them (Wakabayashi, 1992, 1999a; Wakabayashi and Dumitru, 2007). At some localities a mélange zone occupies the structurally highest Franciscan, whereas at others intact thrust sheets (up to 1.5 km in map dimension) of high-grade metamorphic rocks (Ernst et al., 1970; Wakabayashi and Dumitru, 2007) occupy this zone. An excellent example of a mélange at the structurally highest level of the Franciscan, just beneath the Coast Range Ophiolite, is the serpentinite matrix mélange with high-grade blocks described by Shervais et al. (this volume; see also Hopson and Pessagno, 2005). This extensive (map view outcrop width up to 10 km, extending for >80 km along strike) serpentinite matrix mélange includes blocks of both “upper plate” (Coast Range Ophiolite) and Franciscan rocks. It is difficult to assign a specific structural horizon (if there is one) within this unit as the true base of the Coast Range Ophiolite or the structurally highest zone of the Franciscan. Another example of a mélange at the structurally highest level of the Franciscan crops out on Tiburon Peninsula in the San Francisco Bay area (Figs. 3 and 10A). Although this is locally the structurally highest level of the Franciscan, the massive serpentinite (most of it displaying relict harzburgite texture) directly above the mélange is not associated with unmetamorphosed igneous or sedimentary rocks, so it cannot be directly correlated with the Coast Range Ophiolite. This mélange is ~50–70 m thick, has a serpentinite matrix (lizarditerich with some lenses and blocks of antigorite serpentinite), and incorporates an abundance of high-grade metamorphic blocks (Bero, 2003, 2004). It crops out structurally above blueschist facies metagraywacke and chert. A similar but much thinner (~5 m thick) zone is present in El Cerrito (a few hundred meters from and structurally above the locality shown in Fig. 5), but exposures observed as recently as the mid-1980s have been covered by debris rolling downhill from construction above. The matrix was not exposed at this locality, but the zone had a high concentration of high-grade blocks (these were the only visible outcrops) and lay beneath unmetamorphosed volcanic rocks of the Coast Range Ophiolite and above blueschist-facies metagraywacke. A serpentinite matrix mélange with abundant high-grade blocks and clear block-in-matrix contact relationships is well exposed in southern Sonoma County (Allen, 2003; Holland et al., 2009) (Fig. 10B, location in Fig. 1), but the structural level of this mélange is difficult to determine owing to the relatively limited extent of this erosional window through younger (late Cenozoic, post-Franciscan) rocks. The lack of graywacke blocks in this mélange suggests that it may have accreted very early in Franciscan history and represents a structurally high, if not the highest, zone, because graywackes may not have been accreted during the earliest history of Franciscan subduction (Dumitru et al., 2010; Snow et al., 2010; see later discussion).
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Chaotic Unit in the Lower Great Valley Group Phipps (1984) described a chaotic unit within the basal part of the Great Valley Group. This unit extends for >50 km along strike and reaches a thickness of ~1 km. Most of the matrix of this unit is serpentinite, but zones of shale matrix and mixed shale and serpentinite matrix also occur. The matrix exhibits a weak foliation and no post-depositional strain of blocks, in contrast to Franciscan mélanges that have a much more strongly foliated matrix and exhibit far greater apparent strain of blocks (e.g., Cowan, 1985). Much of this unit depositionally overlies the Coast Range Ophiolite, but the unit locally depositionally overlies shales of the basal Great Valley Group shales (such a locality is described in Moores et al., 2006). At such localities the unit includes a basal lag deposit consisting of boulders of gabbro and diabase in a serpentinite matrix. Blocks within the mélange consist almost entirely of ophiolitic lithologies, including chert, basalt, diabase, gabbro, serpentinite, and amphibolite-grade metabasite. Most of the amphibolite blocks appear to lack the blueschist overprint commonly observed in amphibolite (high-grade) blocks in the Franciscan mélanges (Phipps, 1984) and their rare coherent equivalents, but rare blueschist blocks and blueschist-overprinted amphibolite blocks have been reported (Carlson, 1981). Sandstone blocks, the most common block type in most Franciscan mélanges, are absent from the Great Valley Group chaotic unit (Phipps, 1984). The matrix of the chaotic unit, like the units above and below its bounding contacts, lacks burial metamorphism. The chaotic
unit may have originated as serpentinite mud-volcano deposits on the seafloor, similar to those present in the modern Mariana forearc (Fryer et al., 2000). Mélanges of Uncertain Affinity (Mixed Franciscan, Coast Range Ophiolite, Great Valley Group) Some mélanges along Coast Range Ophiolite–Franciscan contacts defy a simple classification of tectonic affinity. These mélanges include blocks of Franciscan, Coast Range Ophiolite, and Great Valley Group affinity (Wakabayashi, 2004) (local field relations shown in Fig. 11). In the Hayward Hills of the eastern San Francisco Bay area, the mélange or shear zones range in thickness from centimeters to ~50 m, and the thicker units persist along strike for several kilometers. These mélange zones separate slabs of unmetamorphosed Great Valley Group sandstone; the sandstone shows little contrast across these zones, indicating minimal displacement. The matrix appears to vary locally from serpentinite to shale, and includes serpentinite and shale mixed at scales of centimeters. Blocks include serpentinite, which commonly exhibits evidence of moderate to high-temperature metamorphism, with metamorphic minerals such as talc and tremolite, along with gabbro, of apparent Coast Range Ophiolite affinity (lacking in burial metamorphism), fine- and coarse-grained blueschist, amphibolite, and garnet amphibolite–eclogite (Wakabayashi, 2004). The latter high-grade metamorphic rocks have blueschist facies overprints typical of Franciscan high-grade rocks. Sandstone blocks are
Figure 10. (A) Photo of a mélange at Tiburon Peninsula that occupies the structurally highest position in the Franciscan Complex. View to the north. Above the mélange is massive serpentinite, as shown in the photo. The mélange appears to be a few tens of meters thick, with a serpentinite matrix and includes high-grade metamorphic blocks. The largest block in the photo is ~15 m long. (B) Sheared serpentinite matrix mélange with amphibolite block, Sonoma County, east of Petaluma. View to the southeast.
Mélanges of the Franciscan Complex, California common, and most of them are identical to the Great Valley Group wall rocks of the shear zones. Mélanges, composed of Franciscan material, are bounded on both sides by Great Valley Group rocks in the York Mountain area of the southern Coast Ranges (southwesternmost part of Fig. 1 within area marked KJN), but these units appear to be tectonic windows into Franciscan rocks structurally beneath Great Valley Group rocks (Seiders, 1982) rather than repeated units within Great Valley Group rocks, as detailed above.
FRANCISCAN AND RELATED MÉLANGES: TECTONIC OR SEDIMENTARY ORIGINS, OR COMBINATION? Much debate has focused on the comparative importance of sedimentary versus tectonic processes in generating Franciscan mélanges. One end-member view is that mélanges were generated entirely, or almost entirely, by tectonic processes (e.g., Cloos, 1982, 1984, 1986; Jeanbourquin, 2000; see also review
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on mélanges in general in Şengör, 2003), whereas others have presented evidence that suggests that much lithologic mixing occurred in a sedimentary environment as olistostromes, and that this mixing preceded deformation associated with burial (partial subduction) of these units (e.g., Cowan, 1978; Page, 1978; Aalto, 1981, 1989; see also Erickson, this volume). It is not disputed that most mélange zones in the Franciscan appear to occupy apparent fault zones (the example given in Erickson, this volume, appears to be an exception), both as zones between coherent nappes and as shear zones–faults within coherent nappes. The mélanges between coherent nappes or their margins have accommodated very large displacements (many kilometers), as noted previously. Accordingly, the amount of tectonic strain accommodated by many Franciscan mélanges may be expected to be very large. What is more controversial is whether some part of the development of block-in-matrix fabric of mélange preceded tectonic strain. In addition to structural and sedimentary evidence presented for a sedimentary mixing phase of mélange development (Cowan and Page, 1975; Cowan, 1978; Aalto, 1981, 1989; Aalto and Murphy, 1984), other evidence points to the sedimentary reworking of exotic material into mélanges. MacPherson et al. (1990) and Erickson et al. (2004; Erickson, this volume) showed that rare but widespread supra-subduction zone igneous blocks (not high grade) within mélanges were most likely derived from the Coast Range Ophiolite because they have no equivalent in the coherent units of the Franciscan. In the structural position of those mélanges (not at the structural highest zone in the Franciscan), it is difficult to envision Coast Range Ophiolite blocks being incorporated by tectonism alone; initial deposition into the trench by olistostromes seems more likely. Franciscan Sedimentary Breccia The following will describe newly identified Franciscan sedimentary breccia localities, whose characteristics and field relationships appear to support an early sedimentary mixing stage of development for some deformed Franciscan mélanges. Sedimentary breccia has been identified in the Franciscan before, both as breccias within a large coherent sandstone slab (Cowan and Page, 1975) and variably deformed breccias along the northern California Coast (Aalto and Murphy, 1984; Aalto, 1989). Aalto (1989) proposed that such breccias may represent the early stage of development of many Franciscan mélanges that otherwise may be considered tectonic in origin. In contrast to the prehnitepumpellyite facies grade of the units described by Aalto (1989), the localities described below are blueschist facies; this and other features provide additional insight into mélange-forming processes and Franciscan exhumation and sedimentation history. In the Sunol Regional Wilderness area of the northern Diablo Range, mélange zones separate and cut coherent metagraywacke and metavolcanic units of blueschist facies (Figs. 1, 8, 9). Most mélanges shown in Figure 9 have sedimentary breccia associated with them, including particularly extensive expo-
sures along a gully locally called the W-Tree Scramble (SE corner of Fig. 9), where the deposits had been called diamictite in a master’s thesis by Christie (1985). The extensive nature of this mélange and differences between the Franciscan units on either side of it suggest that it may be a nappe-bounding mélange. In detail, field relationships appear to show within-nappe mélanges in the north that merge with more extensive mélanges in the south (southern part of Fig. 9), creating a nappe-bounding mélange. Most of the breccias are shale-matrix supported with angular metagraywacke, metavolcanic, and metachert clasts (Fig. 12). Sandstone-matrix breccias also occur, as do clast-supported breccias, consisting mainly of metavolcanic clasts with subordinate metagraywacke and metashale clasts (Fig. 12E). The breccias show a textural gradation from nearly undeformed breccia to strongly foliated material that takes on the appearance of scaly mélange matrix (Figs. 12A–12D). Clast size ranges from centimeters to an observed size of at least 7 m. Much larger blocks (up to hundred meter scale or larger) crop out in the area, but the block-matrix relationships are not exposed. The proximity of larger blocks to breccia localities, such as the 100-m-scale metavolcanic block shown directly east of Figure 12 breccia localities in Figure 9, suggests that the breccia may include blocks of such size. The presence of aragonite veins that crosscut the breccia and growth of lawsonite in the breccia matrix indicates burial of the breccias to blueschist depth, or conditions likely isofacial with adjacent coherent units. The surrounding coherent units or large blocks (the difficulty in defining coherent slabs versus large blocks can be seen in Fig. 9) all exhibit blueschist-facies assemblages with the most common lithology, graywacke, containing neoblastic lawsonite, glaucophane, and jadeitic clinopyroxene. Jadeitic clinopyroxene and glaucophane are common in clasts of the breccia, and some of this metamorphic mineral growth may also postdate deposition, but I have not found a sample where such minerals grow across breccia clast boundaries, and many samples exhibit mineral growth that appears to be truncated at clast boundaries, as noted below. Many of the breccia clasts exhibit internal penetrative fabrics that predate deposition as indicated by abrupt truncation of internal clast fabrics along clast boundaries (Fig. 12G). The growth of syndeformational blueschist-facies minerals in the clasts, including aligned sodic amphibole and jadeitic pyroxene, indicates that the clasts underwent blueschist-facies metamorphism prior to their deposition (Fig. 13C). Thus the clasts are derived from rocks that were subducted to blueschist-facies depths, exhumed, deposited as breccia clasts, then resubducted to blueschist-facies depths. Higher grade metamorphic material is present in scattered clasts as large as 1 m in diameter. An example was collected from the locality shown in Figure 12B. This 15-cm-long clast is quartz rich with glaucophane, lawsonite, and phengite growth to 2 mm in size (versus metamorphic grain sizes of <0.2 mm for lower grade blueschist facies clasts in the breccia and the surrounding large blocks or slabs). High-grade (higher grade than the rest of the breccia) mineralogy occurs as
Mélanges of the Franciscan Complex, California relics: cores of epidote in lawsonite, hornblende in glaucophane, and rutile in titanite (Fig. 13B). This and similar higher grade clasts can be considered an equivalent of the high-grade blocks that exhibit similar metamorphic grain size and overprinting relationships. Serpentinite clasts are locally common in the breccia as well (Fig. 13A).
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In addition to high-grade material that is identifiable as breccia clasts, high-grade blocks occur in mélanges in the Sunol Regional Wilderness area (Fig. 9), but it is difficult to ascertain whether they are clasts in breccia owing to the lack of complete exposure of the surrounding matrix. The high-grade blocks include amphibolite, garnet-amphibolite, and eclogite up to 10 m
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Figure 12 (Continued on following page). Breccia and mélange matrix from Sunol Regional Wilderness exposures (see Fig. 9 for location). Some examples of chert (c) and metabasalt clasts (v) are labeled. (A) Primarily angular sandstone and shale clasts with some chert and metabasalt. (B, C) These areas are richer in shale and are very much like the shale matrix associated with “tectonic” mélanges, but the primary sedimentary breccia textures are not obscured by the comparatively small amount of strain. C is rather rich in volcanic and chert clasts, only a few of which are labeled. The high-grade clast H (shown in Fig. 13) is the pale clast just above the camera case. (D) This photo is the same scale as C, and shows a welldeveloped shear fabric, but domains of weakly deformed sedimentary breccia can be discerned (perhaps best in the center and lower righthand corner). One volcanic clast is labeled; most of the other clasts are sandstone and shale.
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in exposed long dimension. Most of the blocks found occur on the boundaries of shale matrix mélange or on the boundaries of serpentinite bodies (Fig. 9). In contrast, the smaller high-grade breccia clasts (nearly all <1 m in size) found in the W-Tree Scramble transect do not appear to be localized along the border of the unit. Also notable is the apparent scale dependence of orientation of internal fabric within blocks in the mélanges (breccias) of this area (Fig. 9). For smaller (~1 m and less in long dimensions) internal planar fabric (bedding or foliation) may be at high angles to the matrix foliation (Figs. 12 and 13), but larger blocks tend to have internal planar fabric subparallel to the matrix
foliation. The orientation relationship for larger blocks appears to be the general case for Franciscan mélanges, possibly owing to the fact that block shapes preferentially reflect their internal fabric, and such blocks are likely to have been (1) deposited with the short axis at high angles to the depositional surface and/or (2) the longer dimensions of the block will have been rotated toward parallelism with the flattening plane in the mélange with progressive deformation. In summary, the field and textural relationships of the Sunol Regional Wilderness breccias appear to record (1) earlier burial of most of the source material to blueschist depths with
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Figure 12 (Continued). (E) Shale matrix breccia (left) and mafic breccia (right). Height of view is ~7 m. (F) Shale matrix breccia includes a large pillow basalt block (beneath camera case); the exposed long dimension of this block (righthand boundary off edge of photo) is >2 m. Abbreviations in black text highlight some of the clasts-blocks: v—metabasalt; s—sandstone-graywacke. (G) Breccia showing the metamorphic fabric in sandstone clasts at high angles to clast grain boundaries.
Mélanges of the Franciscan Complex, California
including quartz-rich metacherts and omphacite-epidote-quartz clasts. This breccia can be distinguished from cataclastic zones found in such blocks by the presence of breccia clasts exotic to the block. Petrographic examination shows evidence of significant burial after deposition in the form of pressure solution contacts and locally weak development of foliation. In thin section, some minerals such as glaucophane and stilpnomelane appear to overgrow clast boundaries, but it is difficult to be sure whether these minerals are overgrowing clast boundaries or whether there were some internal lithologic variabilities (such as distribution of quartz) within clasts. A macroscopic view of polished slabs of the breccia, however, shows preferential growth of sodic amphibole between clasts, suggesting burial of the breccia blueschist-facies depth (Fig. 14C). No obvious serpentinite clasts are recognized in thin sections of this breccia. The finer matrix of the breccia is pervasively altered to clays and iron oxides, and it is possible that this material may have once been serpentinite. Locally, small clasts consisting of fine-grained chlorite are present (Fig. 14D); these may have been derived from serpentinite. In summary, the evidence from the Sonoma County mélange and breccia locality record early subduction of clast parent
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associated strain, with rare clasts recording the early Franciscan high-grade metamorphism; (2) exhumation of the source material; (3) deposition of the breccia; (4) burial of the breccia to blueschist (lawsonite, aragonite) depths associated with (5) variable deformation of the breccia, with the most highly strained breccia having developed into what appears to be “typical” strongly foliated shale matrix of mélange. I interpret the blueschist-facies breccias of the Sunol Regional Wilderness to reflect the sedimentary precursor to highly deformed mélanges that may otherwise be interpreted as tectonic mélanges. A different type of sedimentary breccia crops out on the side of a 100-m-scale blueschist block in a serpentinite matrix mélange in southern Sonoma County (Fig. 14) (Holland et al., 2009). This mélange contains numerous blueschist, amphibolite, and eclogite blocks in a pervasively sheared serpentinite matrix. The structural thickness of this mélange is difficult to constrain from the field relationships, but the exposure width (several hundred meters) along with apparently steep foliation suggests a thickness of a few hundred meters. The breccia, up to 5 m thick, consists mainly of angular clasts of material identical to the blueschist block, but also clasts of different rocks,
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Figure 13. Photomicrographs of clasts from Sunol Regional Wilderness breccia. (A) Clast of serpentinite, here composed mainly of lizardite. (B) High-grade blueschist clast that contains relics of higher T and P metamorphism, as typically found in the highgrade blocks of the Franciscan. ep—epidote; gln—glaucophane; law—lawsonite; phen—phengite; ttn, rt—titanite with rutile cores. (C) Various clasts, here mostly metavolcanic, in a breccia. These clasts exhibit syntectonic blueschist facies mineral growth that is abruptly truncated at the clast grain boundary. The best example is the large, dark clast below center to the right that contains abundant syntectonic glaucophane (long axes of amphiboles subparallel to the long axis of the grain).
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materials, exhumation, sedimentation, and resubduction, similar to the Sunol Regional Wilderness breccias. The relationships suggest that the enclosing serpentinite matrix mélange may have been a sedimentary serpentinite that was recrystallized on subduction to blueschist-facies depth. High-Grade Blocks in Different Structural Levels of the Franciscan Complex: Evidence of Sedimentary Recycling The distribution of high-grade tectonic blocks in mélanges at multiple structural levels in the Franciscan suggests a sedimentary input as well, and this is important for assessing the
history of many of the nappe-bounding, high-strain mélanges from which preserved sedimentary breccia has not been found, possibly because of the high degree of strain. The high-grade blocks are the oldest metamorphic rocks in the Franciscan and are equivalent to coherent rocks found at the structurally highest level of the Franciscan, as would be expected by their metamorphic age (Wakabayashi and Dumitru, 2007). For the high-grade material to crop out in structurally lower mélanges, some of which were not accreted until 80 m.y. after initial high-grade metamorphism and accretion, appears to require one of the following alternatives (Wakabayashi, 1992) (Fig. 15): (1) Plucking of blocks from the structurally high zone by mélange return flow
Figure 14. (A) Large blueschist block in serpentinite matrix mélange, Sonoma County. View to the east. This locality is a few hundred meters north of that of Figure 10B. On the left flank of the blueschist block are outcrops of breccia (shown in B). (B) This breccia is composed mostly of clasts of the blueschist block similar to the block with clasts of what appear to be blueschist facies metachert and possible serpentinite. (C) Cut slabs of the breccia. Note the dark areas around the clasts, especially apparent in the righthand slab. These dark areas are rich in sodic amphibole, suggesting sodic amphibole growth after deposition. (D) Photomicrograph of the breccia, showing some of the finer grained parts of the breccia and possible serpentinite clasts (arrows) replaced by chlorite.
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A. Sedimentary-tectonic origins of Franciscan mélanges Submarine sliding Coast to trench, including Range Ophiolite serpentinite, high-grade Serpentinite blocks mud volcanoes: sedimentary serpentinite Mélange at Inter-nappe mélange structural top forms along subduction of Franciscan interface by subduction of Great Valley Group trench-fill olistostrome
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Source of most blocks from exhumed Franciscan material
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Deep region for potential plucking of high-grade blocks from structurally highest (originally accreted) zone
Some blocks from forearc mud volcanoes (early in subduction history)
Most displacement on lower contact after accretion
Subducted/offscraped olistostrome: mélange
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B. A tectonic alternative for origin of exotic blocks, Franciscan mélanges Accreted coherent unit Accreted mélange with blocks offscraped from downgoing plate
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Older accreted coherent unit Accreted mélange with blocks offscraped from downgoing plate
Figure 15. (A) Diagrams illustrating the development of Franciscan inter-nappe mélanges from trench-fill olistostromes. Also illustrated is the depth of the region that later mélange–shear zones could have potentially plucked high-grade blocks from their original structurally high position. Note that at the time of emplacement of many mélanges containing high-grade blocks, abundant Franciscan material lay at blueschist depths, as attested to by present exposures. This diagram combines different time stages in Franciscan mélange development, and the schematic seafloor topography is more representative of the very earliest period in Franciscan subduction history prior to deposition of significant clastic sediment in the forearc at 120 Ma. After significant accretion (represented by the accretionary wedge) a forearc high developed and separated the forearc region into trench and forearc basin depositional domains. (B) Diagrams illustrating how subduction and offscraping of rocks from the downgoing plate may have contributed exotic blocks to a mélange not present in either of the coherent nappes bounding the mélange. If the blocks are of higher grade than the matrix and/or adjacent units, return flow must be invoked if blocks are incorporated by tectonic means alone. As discussed in the text, higher grade blocks appear to have been shed into the trench as olistostromes (as in A) prior to subduction-accretion of the mélange, and even lower grade exotic blocks likely have sedimentary origins.
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(Cloos, 1982, 1984), or (2) exhumation of high-grade blocks and redeposition of the blocks into the trench. Evidence of early exhumation and redeposition of a high-grade block has been found, as noted above. The plucking of blocks by return flow means that the highgrade material resided at depth until the plucking occurred. In the case of mélanges that formed at or around 80–100 Ma, the area where plucking had occurred must have been deep, at blueschist-facies depths, for blueschist-facies metagraywacke (now exhumed) was present at structurally intermediate levels (and at blueschist-facies depth) above the mélange zone and structurally below the “parent” or “source” zone (Wakabayashi and Unruh, 1995; Tagami and Dumitru, 1996) (Fig. 15). Thus blocks destined for structurally lower mélange zones would have resided at blueschist depths for progressively longer times with lower structural levels (the younger and structurally lower the mélange, the longer the blueschist-depth residence time of the block). This would predict progressively greater degrees of blueschist overprinting of high-grade blocks with progressively lower structural levels of mélanges (that accreted at a progressively later time). No systematic variation in the degree of blueschist overprinting exists in Franciscan mélanges. Some of the least overprinted amphibolite blocks (including some that do not have a visible blueschist overprint) are present in some structurally low mélange zones that could not have accreted until after 85 Ma or so (80 m.y. after initial high-grade metamorphism). Accordingly, high-grade blocks in structurally lower mélange zones were more likely introduced by sedimentary means, consistent with evidence for exhumed and sedimented high-grade blocks or clasts from the Sunol Regional Wilderness and Sonoma County localities noted above. Additional evidence for early exhumation and sedimentation of high-grade blocks in the Franciscan comes from the presence of high-grade block clasts in blueschist-facies conglomerates (Moore and Liou, 1980) as well as a high-grade block that may rest in apparent depositional contact on blueschist-facies metagraywacke (Moore, 1984). Early exhumation of high-grade blocks is also supported by their occurrence in the basal olistostromes of the Great Valley Group (Carlson, 1981; Phipps, 1984). The presence of high-grade blocks suggests that many of the nappe-bounding mélanges, which may have accommodated high strain and representing the most “tectonic” of Franciscan mélanges, had an early sedimentary history. Angular clasts are common in these nappe-bounding mélanges, but without the gradational exposures seen at the Sunol Regional Wilderness locality it cannot be determined whether these angular clasts are the result of breakage by strain or if these were originally sedimentary breccia clasts given the pervasive strong foliation of the exposed matrix. In addition, the presence of lower grade exotic blocks that do not match anything in either of the bounding nappes is also consistent with sedimentary mixing. However, because internappe zones may represent the paleomegathrust, it is possible that exotic blocks were scraped off the otherwise completely subducted oceanic plate prior to accretion of the next nappe below (and migration of the megathrust zone to the level
beneath the subjacent coherent nappe) (Fig. 15). Accordingly, the high-grade blocks are the one type of exotic block that strongly indicates a sedimentary introduction to the mélange, whereas lower grade exotic blocks permit either tectonic or sedimentary introduction. If the mélange at El Cerrito Quarry is representative of nappe-bounding mélanges, it is possible that most of the displacement between nappes is accommodated along the borders of mélanges rather than within them (Fig. 15). Accordingly, sedimentary processes may have contributed the exotic blocks and much of the block-in-matrix fabric even to those mélanges where high strain may be most expected. Finally, the localization of many high-grade blocks along the borders of mélange units is difficult to explain by a tectonic mechanism (why should strain move the blocks to the edges of shear zones?) but may suggest a sedimentary origin as well, with the blocks along one contact representing a specific sedimentary zone such as a basal lag deposit. Evidence of Sedimentary Mixing in Franciscan Mélanges and Recycling of High-Pressure Metamorphic Material: Summary Based on the above discussion, many different types of evidence support some degree of sedimentary mixing in the history of many Franciscan mélanges that occurred before these mélanges or their borders became zones of displacement. The generation of mélanges from a combination of sedimentary and tectonic processes has been proposed by a number of researchers conducting research on other paleo-convergent plate margins (e.g., Jacobi, 1984; Pini, 1999; Alonso et al., 2006; Dela Pierre et al., 2007; Festa, this volume; Osozawa et al., 2009; Osozawa et al., this volume), in addition to having been previously proposed by some for Franciscan mélanges (e.g., Cowan and Page, 1975; Cowan, 1978; Page, 1978; Aalto, 1981, 1989). The occurrence of recycled blueschist-facies and higher grade material in sedimentary breccia suggests early exhumation of some high-pressure metamorphic rocks in the Franciscan. In addition to the evidence cited earlier (both for this and previously published studies), recycled blueschist-facies minerals were identified by Brothers and Grapes (1989) in Diablo Range metagraywackes, although the main conclusion of that paper—that all jadeitic clinopyroxene in Franciscan metagraywackes is detrital—has been strongly rebutted by abundant textural evidence of in situ (neoblastic) growth of jadeitic clinopyroxene in these rocks (Raymond, 1991; Ernst, 1993). The conclusion that some blueschist facies minerals in Franciscan metagraywackes are detrital appears to be supported by the occurrence of abundant blueschist facies clasts in the breccias described in this study. Although it is difficult to tightly constrain exhumation rates associated with the recycled metamorphic clasts, existing geologic and geochronologic relationships place some limits on exhumation history and patterns of forearc exposure. The oldest high-volume metaclastic rocks in the Franciscan were deposited at ca. 123 Ma (Dumitru et al., 2010), and the blueschist-facies
Mélanges of the Franciscan Complex, California metagraywackes, broken formation, and mélange of the Diablo Range range from ca. 85 to ca. 110 Ma in depositional age (Unruh et al., 2007; Ernst et al., 2009). The source of the blueschistfacies metagraywacke clasts may no longer be preserved (eroded away), because apatite fission-track data of exposed Franciscan jadeite-bearing graywackes suggest that they resided below the depth associated with apatite fission-track annealing until the latest Cretaceous to early Cenozoic (Dumitru, 1989). The oldest metaclastic rocks in the Franciscan are found in the Skaggs Spring schist, which has a maximum depositional age from U/Pb detrital zircon ages of 144 Ma (Snow et al., 2010) and Ar/Ar white mica metamorphic ages of ca. 132 Ma (Wakabayashi and Dumitru, 2007). The detrital zircon ages from the Skaggs Spring schist are comparable to the fossil and detrital zircon age determinations for the oldest Great Valley Group strata (Surpless et al., 2006; Wright and Wyld, 2007), indicating that clastic deposition did not begin in the forearc-trench system until that time. Thus the source for the blueschist facies graywacke clasts in the breccia was deposited after 144 Ma, and perhaps after 120 Ma, and was exhumed and redeposited by ca. 85–110 Ma. High-grade blocks were subducted and metamorphosed at ca. 158–169 Ma (Shervais et al., this volume; Wakabayashi and Dumitru, 2007; Anczkiewicz et al., 2004), exhumed to the seafloor and redeposited as early as ca. 135–147 Ma on the basis of the apparent Valanginian–Tithonian depositional age of the sedimentary serpentinite in the Great Valley Group that includes high-grade blocks (Phipps, 1984) and the new detrital zircon ages for the basal Great Valley Group strata (Surpless et al., 2006; Wright and Wyld, 2007). The Sonoma County serpentinite matrix mélange of Figures 10B and 15, apparently a sedimentary serpentinite, may have been deposited relatively early in Franciscan accretionary history on the basis of its block population, which consists almost entirely of high-grade blocks with no metaclastic rocks. A block of Skaggs Springs schist is present in shale matrix mélange west of the serpentinite matrix mélange. These field relationships indicate deposition of the serpentinite matrix mélange before 123 Ma and possibly before ca. 140 Ma, but after 158–169 Ma. As such it may be a Franciscan temporal equivalent of the Great Valley Group sedimentary serpentinite, and both units may have been deposited as serpentinite mud volcano deposits (Fryer et al., 2000) before the development of the forearc high that divided the forearc sedimentary domains into trench and forearc basin (Fig. 15). Resedimented high-grade material is present in the Sunol Regional Wilderness (likely depositional age range of 85–110 Ma) and in Franciscan mélange zones that accreted from ca. 100 Ma to ca. 80 Ma or younger; the youngest such units may be those that cut the Permanente terrane (Wakabayashi, 1992; Snow et al., 2010; Wakabayashi et al., 2010). It is possible that exhumed Franciscan high-grade material was a minor trench sediment source from as early as ca. 150 Ma until at least 80 Ma. Some Franciscan mélanges may not have had a sedimentary phase of development. The serpentinite matrix mélange at the structurally highest zone described by Shervais et al. (this vol-
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ume) may owe its block-in-matrix structure mostly to tectonic strain. Alternatively the serpentinite-matrix mélange in this structural position may be a sedimentary serpentinite deposit analogous internappe mélange, or it may be at least partly a product of diapirism. Some intranappe mélanges appear to cut across the bedding or foliation of a coherent unit, suggesting minimal displacement across the zone owing to apparently identical rocks on either side. Because they cut across bedding, these mélanges could not have been emplaced by a sedimentary process. The field relations suggest diapiric emplacement, but confirmation of this, from opposing sense-of-shear on opposite boundaries of the mélange, has not been documented in Franciscan mélanges as it has in other localities in the world (e.g., Orange, 1990; Dela Pierre et al., 2007; Festa, this volume). It is also possible for diapiric mélanges to tap an already disrupted zone that had earlier been mixed by sedimentary processes, but no field evidence has been found to support such a mechanism. STRIKE-SLIP VERSUS DIP-SLIP DEFORMATION The dominant role of subduction in the assembly of the Franciscan accretionary complex is not disputed, but the significance of intra-Franciscan, strike-slip faulting in the generation and deformation of mélanges has generated different views. For example, Blake et al. (1988) proposed that most Franciscan mélanges formed at approximately the same time as a consequence of a Late Cretaceous episode of intra-Franciscan strike-slip faulting. Karig (1980) proposed that entrainment in strike-slip fault zones may explain the presence of high-grade blocks in otherwise lower grade mélanges. In contrast, Maxwell (1974), Cloos (1984, 1985), and Wakabayashi (1992) suggested that mélanges formed as a result of progressive subduction-accretion at various times in Franciscan history, associated with large-scale (many kilometers) thrust displacement for internappe mélanges. The refolded nappe structure of the Franciscan, including the inter-nappe mélanges, appears difficult to reconcile with a strike-slip origin, as does the progressive younging structurally downward of the youngest rocks in any given mélange; both features appear much more consistent with progressive subduction-accretion associated with thrusting (Wakabayashi, 1992). Recent U/Pb dating of detrital zircons has vastly improved the geochronologic framework and confirmed (and refined) the downward-younging accretion age progression within the Franciscan (Ernst et al., 2009; Snow et al., 2010; Dumitru et al., 2010). In addition to syn-subduction strike-slip faulting, the significance of post-subduction strike-slip faulting and deformation requires attention because several hundred kilometers of rightlateral strike-slip displacement has cut the Franciscan since cessation of subduction and inception of a transform plate margin (e.g., Seiders, 1988; Wakabayashi, 1999b). This strike-slip displacement does not appear to be associated with development of mélanges, because Franciscan mélanges are present many kilometers north of the Mendocino triple junction (e.g., Blake
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et al., 1985, 1988; Aalto, 1989) in an area that has not undergone strike-slip faulting associated with the transform regime. In addition, late Cenozoic sedimentary and volcanic rocks dating from the early stages of the transform regime unconformably overlie many mélange exposures and do not exhibit blockin-matrix deformation (a good example of this is the overlap shown in the southwestern part of Fig. 9). In fact, no geologic unit whose formation postdates the development of the transform fault regime exhibits or hosts mélange structures, and no blocks of post-subduction rocks are present in any Franciscan mélange. Accordingly, it appears that neither syn-subduction nor post-subduction strike-slip faulting had a significant role in Franciscan mélange deformation. CONCLUSIONS An inspection of Franciscan mélanges shows that they occur in a variety of structural settings, giving insight into the relationship of these mélanges to subduction zone processes (Figs. 3 and 15). New field evidence from sedimentary breccias suggests that many Franciscan mélanges, even the ones in which high strain may be most likely, had an earlier history of sedimentary mixing that may have contributed the exotic blocks therein. Sedimentary serpentinites may be present in the Franciscan, based on the presence of the sedimentary breccia on the side of a block in serpentinite matrix mélange as well as the presence of serpentinite detritus in sedimentary breccia. Mélanges between coherent nappes appear to occupy the position of the paleo-subduction megathrust zone, and were originally deposited as olistostromes on the trench floor prior to subduction and incorporation into the accretionary wedge (Fig. 15). Serpentinite matrix mélange found locally at the highest structural level in the Franciscan may have an early sedimentary origin similar to the internappe mélanges, or this zone may have formed as a shear zone and/or by diapirism. Mélanges within coherent nappes that strike parallel with the structural-stratigraphic grain in such nappes may represent former chaotic sedimentary deposits within the coherent nappes, whereas those that cut across the structural grain may have a diapiric origin. These blueschist-facies sedimentary breccia deposits show evidence of exhumation of material from blueschist facies depths followed by deposition and resubduction to blueschist depths accompanied by variable degrees of strain. It appears that major fault zones within the accretionary prism—both the subduction megathrust itself, and other faults cutting up through coherent nappes—exploited the breccia deposits or perhaps their borders, and that probably represented the weakest zones (Fig. 15). Deformation that affected the mélanges appears to have occurred as a result of thrusting during subductionaccretion, and possibly diapirism, with the possible exception of the structurally highest mélange, which may have accommodated normal fault movement owing to the juxtaposition of unmetamorphosed Coast Range Ophiolite over high-pressure Franciscan units (e.g., Platt, 1986).
The evidence present in the mélanges themselves does not directly constrain the exhumation of the blocks prior to sedimentation as olistoliths. Some of the high-grade blocks, particularly those associated with serpentinite-matrix mélanges, may have been exhumed in diapirs associated with serpentinite mud volcanoes. Lower grade meta-igneous blocks of Coast Range Ophiolite affinity may also have been derived from such deposits, as well as intact seafloor exposures (Fig. 15). Some high-grade blocks may have been shed into the trench after exhumation of coherent high-grade rocks, and possibly exhumation of older mélanges; this is particularly true of high-grade blocks exhumed and deposited fairly late (say, after 120 Ma) in Franciscan history. For lower grade olistoliths (including the fine-grained, blueschist-facies, metaclastic rocks described in the Sunol locality) the source was likely exhumed Franciscan coherent and mélange units (Fig. 15), possibly with greater contribution from coherent units owing to the greater volume of block lithologies associated with coherent units rather than blocks-in-mélange. ACKNOWLEDGMENTS I thank W.G. Ernst, E.M. Moores, K.R. Aalto, and D.S. Cowan for their reviews of this paper. My studies of Franciscan mélanges have benefited from discussions with many researchers over the years, including the reviewers and many of the authors of the papers in this volume, but I would especially like to acknowledge my debt to the late Clyde Wahrhaftig, who made me and my classmates spend a lot of time mapping mélange during our undergraduate field camp. Without that training I cannot imagine doing fieldwork on these rocks. This research has been supported in part by U.S. National Science Foundation grant EAR-0635767. REFERENCES CITED Aalto, K.R., 1981, Multistage mélange formation in the Franciscan Complex, northernmost California: Geology, v. 9, p. 602–607, doi:10.1130/0091 -7613(1981)9<602:MMFITF>2.0.CO;2. Aalto, K.R., 1989, Franciscan Complex olistostrome at Crescent City, northern California: Sedimentology, v. 36, p. 471–495, doi:10.1111/j .1365-3091.1989.tb00620.x. Aalto, K.R., and Murphy, J.M., 1984, Franciscan Complex geology of the Crescent City area, northern California, in Blake, M.C., Jr., ed., Franciscan Geology of Northern California: Society of Economic Paleontologists and Mineralogists, Pacific Section, v. 43, p. 185–201. Allen, J.R., 2003, Stratigraphy and tectonics of Neogene strata, northern San Francisco Bay area [M.S. thesis]: San José, California, San José State University, 190 p. Alonso, J.L., Marcos, A., and Suarez, A., 2006, Structure and organization of the Porma Mélange: Progressive denudation of a submarine nappe toe by gravitational collapse: American Journal of Science, v. 306, p. 32–65, doi:10.2475/ajs.306.1.32. Anczkiewicz, R., Platt, J.P., Thirlwall, M.F., and Wakabayashi, J., 2004, Franciscan subduction off to slow start: Evidence from high-precision Lu-Hf garnet ages on high-grade blocks: Earth and Planetary Science Letters, v. 225, p. 147–161, doi:10.1016/j.epsl.2004.06.003. Atwater, T., 1970, Implications of plate tectonics for the Cenozoic tectonic evolution of western North America: Geological Society of America Bulletin, v. 81, p. 3513–3535, doi:10.1130/0016-7606(1970)81[3513:IOPTFT ]2.0.CO;2.
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The Geological Society of America Special Paper 480 2011
Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili, central Turkey Anne Dangerfield Ron Harris Department of Geology, Brigham Young University, Provo, Utah 84602, USA Ender Sarıfakıoğlu General Directorate of Mineral Research and Exploration, Department of Geology, TR-06520 Ankara, Turkey Yildirim Dilek Department of Geology, Miami University, Oxford, Ohio 45056, USA
ABSTRACT Structural field studies and geochemical and age analyses of the Eldivan ophiolite, which is dismembered within the Ankara Mélange, indicates that it developed as a supra-subduction zone basin within the İzmir-Ankara-Erzincan Ocean, which later subducted to form the İzmir-Ankara-Erzincan suture zone through continental block collision. Whole-rock and mineral geochemical evidence show a supra-subduction zone tectonomagmatic affinity for the ophiolitic crust and mantle, revealing that this basin formed in the upper plate of an intra-oceanic subduction zone. Structural restoration of the sheeted dike complex reveals that the supra-subduction zone spreading ridge of the Eldivan ophiolite was nearly parallel to the Sakarya-Pontide continental margin. U/Pb age analyses of detrital zircon in sandstone within the mélange and in the unconformably overlying Karadağ Formation indicate maximum depositional ages for the units of 143.2 ±2 Ma, and 105.2 ±5 Ma, respectively. Thus, thrust imbrication of the ophiolite and the development of serpentinite mélange were mostly complete by 105 Ma, as indicated by an angular unconformity between the ophiolitic units and the overlying Karadağ Formation. These results reveal how and when the Eldivan ophiolite was constructed, destructed, and incorporated into the serpentinite Ankara Mélange and İzmirAnkara-Erzincan suture zone. The tectonic evolution of the İzmir-Ankara-Erzincan Ocean is similar to that of the Philippine Sea and Banda Sea ocean basins. INTRODUCTION Suture zones provide some of the most important evidence of how continents form and evolve. Since the discovery of plate tectonics, suture zones have mostly been characterized as remnants of large mid-ocean ridge basalt (MORB)–like ocean basins that were destroyed by subduction and modified by collision
of rafted continental blocks of different affinities. The IndusTsangpo suture zone of the Himalayan collision provides a type example (Gansser, 1964; Dewey and Bird, 1970). Its interpretation traditionally involves subduction of huge tracts of MORBlike oceanic crust beneath a continental arc system mounted on the southern edge of Asia. The subducting plate rafted India from thousands of kilometers to the south, eventually bringing
Dangerfield, A., Harris, R., Sarıfakıoğlu, E., and Dilek, Y., 2011, Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili, central Turkey, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 143–169, doi:10.1130/2011.2480(06). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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Arabian Platform Laurasia- Future Pontide margin Intra-Pontide Ocean
Sakarya Continent
Izmir-Ankara-Erzincan Ocean
Kirsehir block
Izmir-Ankara-Erzincan Ocean
Int
Northern Neo-Tethys
e
r- T
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ne l Zo re u b n tu
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one rya z Saka
The İzmir-Ankara-Erzincan suture zone in northern Turkey is a remnant of the İzmir-Ankara-Erzincan Ocean branch of the Neo-Tethys that formed during collision of Gondwanaderived micro-continents (Kırşehir block and Tauride-Anatolide platform) with the Sakarya-Pontide terrane of northern Turkey (Şengör and Yılmaz, 1981) (Fig. 1). The suture zone is made up of ophiolitic material and forms the Ankara Mélange, in the
-Tau ride Ta uri su de t blo ck
Ista
GEOLOGIC EVOLUTION OF THE İZMIR-ANKARAERZINCAN SUTURE ZONE
35°E
30°E
RhodopeRhodope Strandja Massif
and structural reconstructions from near Hançili in central Turkey to present a view that differs from the classical or Himalayan model of suture zone development. This model not only provides important constraints for the tectonic evolution of the central İzmir-Ankara-Erzincan Ocean, and significance of the Ankara Mélange, but also advances our understanding of suture zone diversity.
tu
it northward into continental collision with the Asian continent, which formed the Himalaya and the Tibetan Plateau. New evidence from Himalayan sutures, including the Indus-Tsangpo suture (Yin and Harrison, 2000), and other modern convergent boundaries (Harris, 2003), reveal that this classic model of suture zone development is one end member of diverse types of sutures that involve a variety of tectonic processes that lead to stitching continents together. This study uses the Ankara Mélange within the İzmirAnkara-Erzincan suture zone of Turkey to reconstruct tectonic processes and events associated with continental accretion of the eastern Mediterranean region and the origin of its narrow bands of serpentinite mélange. Despite previous studies of the Ankara Mélange, its age, origin, and relations to other tectonic events in the Mediterranean region are still poorly understood. This study uses crustal and mantle geochemistry and petrography of ophiolitic blocks in the Ankara Mélange, U/Pb age analyses of detrital zircon in mélange and overlying sandstone,
su
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A Southern Neo-Tethys
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Arabian-African Platform
Figure 1. (A) Map of the major tectonic units discussed in this chapter (after Okay and Tüysüz, 1999). The star shows the location of the Eldivan ophiolite and the study area. (B) Reconstruction of the İzmir-Ankara-Erzincan suture zone (after Şengör and Yılmaz, 1981), showing the relationship between ocean branches and continental blocks discussed in this paper.
Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili center of the suture zone between the Kırşehir block and the Sakarya-Pontide terrane. The Ankara Mélange was first named by Bailey and McCallien (1950) and consists of three major tectonic units. Structurally from top to bottom these units are (1) a metamorphic block mélange, also called the Karakaya Formation (Koçyiğit, 1991) and Karakaya Group (Floyd, 1993); (2) a limestone block mélange; and (3) an ophiolitic mélange (Norman, 1984; Koçyiğit, 1991; Tüysüz et al., 1995; Dilek and Thy, 2006). The tectonically lowest ophiolitic mélange unit consists of blocks of basaltic and rhyolitic volcanic rocks, pillow basalt, serpentinized peridotite and ultramafic rocks, and radiolarian-bearing limestone and chert, with minor shale and sandstone in a serpentinite or tuffaceous matrix (Norman, 1984; Tankut et al., 1998). Dike complexes, where present, are commonly doleritic, cutting sequences of serpentinized peridotite and isotropic gabbro, and including plagiogranite (Dilek and Thy, 2006). During convergence, the İzmir-Ankara-Erzincan Ocean crust was imbricated along northward-dipping thrust faults (Şengör and Yılmaz, 1981) and overlain by flyschoidal (Norman, 1984) and forearc basin (Koçyiğit, 1991) deposits that were subsequently imbricated with the İzmir-Ankara-Erzincan Ocean crust during the collision of the Kırşehir block and the Sakarya-Pontide continent. The geodynamic history of the İzmir-Ankara-Erzincan Ocean is threefold, beginning with rifting, transitioning to subduction, and finally ending with continental collision. It opened in the Triassic as a MORB-type ocean basin, as determined by geochemical results from basalt and associated Triassic radiolarian limestone found along the İzmir-Ankara-Erzincan suture zone (Tankut, 1984; Tankut et al., 1998; Floyd et al., 2000; Göncüoğlu et al., 2006). Jurassic and Cretaceous alkali basalt is also found in this suture zone and is interpreted as seamount fragments accreted to an accretionary wedge (Tüysüz et al., 1995; Floyd et al., 2000; Rojay et al., 2001; Göncüoğlu et al., 2006), possibly in the form of hot-spot–generated volcanic ridge systems (Floyd, 1993; Tankut et al., 1998). Also documented along the suture are island arc tholeiites (Göncüoğlu et al., 2006; Sarıfakıoğlu, 2006) and calc-alkaline volcanics (Tankut, 1984; Tüysüz et al., 1995; Göncüoğlu et al., 2006), suggesting subduction-related magmatism. The subduction influence is seen in the western and central parts of the suture zone where supra-subduction zone basalt has been documented, indicating intra-oceanic subduction with upper plate extension (Yalınız et al., 1996; Floyd et al., 1998, 2000; Yalınız et al., 2000b; Dilek and Thy, 2006; Sarıfakıoğlu, 2006; Sarıfakıoğlu et al., 2009). This supra-subduction zone signature is similar to other Cretaceous eastern Mediterranean ophiolites (i.e., Pindos, Troodos, Antalya, Hatay (Kızıldağ), Baer-Bassit, and Semail) and has led to correlations of the İzmir-Ankara-Erzincan suture zone with the Vardar suture in Greece (Yalınız et al., 2000b; Göncüoğlu et al., 2006; Sarıfakıoğlu et al., 2009). Granitoids and exhumation in the Kırşehir block and indentation of the Sakarya-Pontide terrane document collisional closure of the basin (Okay et al., 2006; Önen, 2003, Kaymakci et al., 2003).
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Although evolutionary phases of the ocean basin are known, their ages, particularly the transition between rifting and subduction leading to closure, are poorly understood. The oldest ages from the basin come from Carnian (Tekin et al., 2002) and Norian (Bragin and Tekin, 1996) radiolarians in limestone and chert blocks that indicate the ocean had rifted open by the Late Triassic Period. Post-collisional granitoid ages in the Kırşehir block suggest that continental collision began in the Late Cretaceous Period (Boztuğ et al., 2007), finally closing the ocean basin. Between these events the transition from rift construction to subduction closure is interpreted to have begun as early as 179 Ma, based on ages of plagiogranite with supra-subduction zone characteristics in the ophiolitic mélange by Dilek and Thy (2006), and as late as ca. 85 Ma by Göncüoğlu et al. (2006), using a compilation of metamorphic sole and radiolarian ages. Thus, although the general range of construction of the İzmirAnkara-Erzincan Ocean is established as Late Triassic to Late Cretaceous Periods, individual events of its geodynamic history are weakly constrained. COMPOSITION AND GEOCHEMISTRY OF THE ELDIVAN OPHIOLITE The Eldivan ophiolite is part of the ophiolitic mélange of the Ankara Mélange and lies on the mélange’s western side, south of the city of Eldivan (Fig. 1). It is mostly dismembered and exists as imbricated fragments of ophiolitic components within a serpentinite mélange matrix (Fig. 2). There is little continental clastic material within the mélange. Ophiolitic units present are serpentinized mantle peridotite, massive gabbro, sheeted dikes, pillow basalt, and sheet flows, and epi-ophiolitic limestone and chert. The units are hydrothermally altered but are otherwise unmetamorphosed. Unconformably overlying this imbricated ophiolitic mélange is the younger Karadağ Formation, composed of radiolarian-bearing limestone and chert. In the mapped area the Karadağ Formation tectonically overlies the ophiolitic mélange, but outside the mapped area an angular unconformity is present between the Karadağ Formation and the underlying mélange, suggesting a multiphase structural history. Serpentinized peridotite, volcanic rocks, diabase dikes, and gabbro were analyzed for whole rock major, trace, and rareearth elements (REE) and mineral chemistry. Major and selected trace element X-ray fluorescence (XRF) analysis was conducted at Brigham Young University. Trace and rare earth elements were analyzed by inductively coupled-plasma mass spectrometer (ICP-MS) at ALS Chemex Laboratories, Vancouver, British Columbia (method ME-MS81). Mineral chemistry analysis was conducted using a Cameca SX-50 Electron Microprobe at Brigham Young University. Results of analyses are given in Tables A1 and A2 in the Appendix. Owing to the high degree of alteration indicated by widespread secondary mineralization, most samples are hydrous, resulting in elevated LOI (loss on ignition) values (up to 5% for mafic rocks and 12% for ultramafic rocks). Calcined samples were thus used for analyses,
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and only relatively immobile elements (Ti, Zr, Y, Hf, Th, Ta, Cr, Ni, and REE) were used for discrimination of tectonomagmatic setting (Pearce and Cann, 1971, 1973; Pearce and Norry, 1979; Wood, 1980; Pearce et al., 1981 1984a; Pearce, 1982; Shervais, 1982; Mullen, 1983; Meschede and Casey, 1986; Cabanis and Lecolle, 1989; Floyd et al., 1991).
indistinct serpentine groundmass that occupies fault zones and separates various crustal fragments of the Eldivan ophiolite. These include blocks of basalt, silicic volcanic rocks, limestone, chert, and gabbro, which range in size from a few meters to hundreds of meters, basalt blocks being the largest found. Much of the serpentinite matrix exhibits a scaly shear fabric with anastomosing slickensides. Less altered parts of the peridotite consist of (1) 80%–90% serpentinized olivine, (2) up to 5% spinel in the form of primary chromite and secondary magnetite, and (3) 5%–10% pyroxene that is altered to serpentine but still displays low birefringence and mostly parallel extinction.
Mantle Sequence Serpentinized peridotite forms the matrix of the Eldivan ophiolite. There are small coherent blocks of less altered ultramafic rock, but they are found within a much more massive and 40.44°N
To Eldivan (~12 km)
50
48
53
16 22
75
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64 64 60
Karadag Formation, pelagic limestone with minor chert and sandstone
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Bedded marine chert Serpentinized peridotite Basaltic agglomerate 25 41
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Sheeted dikes
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Road Thrust fault Bedding and dike attitudes
Hançili
Meters
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1000
2000
70 40.37°N 33.40°E
35
40
Figure 2. Geologic map of the Çankırı H30-b2 quadrangle (1:25,000 scale), universal transverse Mercator (UTM) zone 36. See star in Figure 1 for location. Latitude, longitude, and UTM coordinates are included. Dashed lines are inferred contacts.
Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili
Crustal Sequence Massive Gabbro and Plagiogranite Massive gabbro occurs in screens within sheeted dikes and is intruded by plagiogranite. Some bodies of plagiogranite are up to several decimeters in diameter. Plagiogranite dikes mostly lack chilled margins, tentatively suggesting they could be an immiscible liquid phase within a gabbroic magma chamber or intrusions into a non-solidified gabbro. Gabbro contains 50% mostly albitized plagioclase and 30%–50% clinopyroxene. Rocks show typical hydrothermal metamorphism with actinolite, chlorite, and epidote as common replacement phases of clinopyroxene and plagioclase. Fe-Ti oxides, mostly composed of secondary magnetite, make up ~5% of the rock. Most gabbro samples are 45–54 wt% SiO2 and are chemically classified as the plutonic equivalents of basalt and basaltic andesite (Fig. 5). A few samples have higher concentrations of alkalis, which could be a reflection of secondary albitization
Brooks Range
Oman
Eldivan
Subduction influenced
6
Orhaneli
7
Ocean basin
8
Continental
of plagioclase. Gabbro has the lowest abundances of rare earth and trace elements of the crustal sequence, with elemental concentrations 1.5–5 times below normal mid-ocean ridge basalt (N-MORB). REE patterns (Fig. 6A1) show lower LREE relative to heavy rare-earth elements (HREE), similar to N-MORB. Additionally, Th/Ta = 1.25–4.80, La/Nb = 0.54–1.91, and La/Yb = 0.52–1.18 element ratios are generally low, similar to N-MORB. Sample 126a (Fig. 6A1, bold line) is the exception, with higher LREE abundances relative to HREE, perhaps from a subduction source, and high Th/Ta = 9.10, La/Nb = 3.00, and La/Yb = 2.38 ratios. A few samples show flat REE patterns that could be a result of differentiation within the magma chamber. Trace elements (Fig. 6A2) show scatter in mobile elements, particularly Rb, Ba, K, and Sr, mostly likely due to secondary hydrothermal alteration. Most samples show the similar trace element pattern of large ion lithophile elements (LILE) depleted
Whole-rock Al2O3 (wt%)
Serpentinized peridotite is characterized by low abundances of Si, Al, Ca, Na, K, and Ti and high abundances of Mg, Cr, and Ni. Light rare-earth element (LREE) concentrations are low, indicating that serpentinization has not affected the original peridotite geochemistry. High degrees of serpentinization, especially where the fluid/rock ratio is large, mobilize LREEs in the serpentinizing fluid, resulting in U-shaped REE patterns that obscure the original igneous chemistry (Paulick et al., 2006). Li and Lee (2006) show that primary Al2O3 wt% is still preserved in peridotite, with >90% serpentinization and no enrichment in LREE. Peridotites from the Eldivan ophiolite are extensively serpentinized but still show low LREE abundances and are therefore interpreted to have primary Al2O3 weight percentages. This is important, as these percentages are a proxy for degree of partial melting, as Al2O3 wt% decreases with high degrees of melt extraction, such as above a subduction zone (Bonatti and Michael, 1989). Eldivan ophiolitic mantle has Al2O3 weight percentages that range from 0.78% to 2.45%, consistent with values from modern ocean floor and subduction-related mantle (Fisher and Engel, 1969; Ishii, 1985; Shibata and Thompson, 1986; Ishii et al., 1992; Seifert and Brunotte, 1996; Paulick et al., 2006) (Fig. 3). Chromian spinel accessory minerals are also excellent indicators for degree of mantle extraction (Dick and Bullen, 1984). In Eldivan serpentinized peridotite, Cr numbers (#) (Cr/Cr+Al) of Cr-spinel range from 0.47 to 0.70, which are in the range of type 2 and type 3 peridotites (Dick and Bullen, 1984). Cr#s >0.6 (type 3) are categorized as arc peridotites and show the most depletion, whereas type 2 are transitional between arc and MORB mantle. Eldivan mantle Cr-spinels overlap arc and MORB fields in Dick and Bullen’s diagram for Cr-spinel Cr# versus Mg# (Fig. 4A). They also plot in the supra-subduction zone mantle field defined by Kamenetsky et al. (2001) in a TiO2 wt% versus Al2O3 diagram, with some points in the overlap between supra-subduction zone and MORB mantle fields (Fig. 4B).
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5
4
3
2
1
0 Figure 3. Whole rock Al2O3 weight percentages for mantle rocks from continental, ocean-floor, and subduction-trench settings compared with the Eldivan ophiolite mantle. The range in Al2O3 weight percentages in the Eldivan ophiolite spans the range of ocean-floor and subduction-trench mantle. It is also similar to that seen in the Oman ophiolite and somewhat higher than in the Orhaneli ophiolite in the western İzmir-Ankara-Erzincan suture zone and the Brooks Range ophiolite in Alaska, interpreted to have formed in supra-subduction zone settings. Data sources for other mantle compositions are as follows: continental (Carter, 1970; Frey and Prinz, 1978); ocean floor (Shibata and Thompson, 1986; Paulick et al., 2006; Seifert and Brunotte, 1996); subduction trench (Fisher and Engel, 1969; Ishii, 1985; Ishii et al., 1992); Oman (Takazawa et al., 2003); Brooks Range (Harris, 1995); Orhaneli ophiolite (Sarıfakıoğlu et al., 2009).
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N-MORB, except for sample 126a (Fig. 6A2, bold line), which has a negative Nb anomaly, again suggesting a subduction influence. Diabase Dikes Diabase dikes are found mostly in parallel or sheeted arrays that cut screens of massive gabbro. Many dikes show chilled margins, and some preserve flow fabrics. Shear fractures within the dikes have gouge zones filled with epidote and chlorite, which are common seafloor hydrothermal alteration minerals, indicating that the fractures probably formed during seafloor metamorphism. Dikes show two primary orientations, NNW-SSE and E-W, with dips from 40° to 90°. In one locality, horizontal sheeted dikes feed vertical pillow basalts and sheet flows, indicating a 90° rotation of the units about a horizontal axis. Dikes are primarily basaltic andesite to andesite with an SiO2 range of ~56–58 wt% (Fig. 5). Element concentrations are most similar to N-MORB (2 times above and below N-MORB concentrations) compared with other units in the Eldivan ophiolite, with
REE patterns also similar to N-MORB (Fig. 6B1). Low ratios of Th/Ta = 0.82–4.40, La/Nb = 1.22–1.93, and La/Yb are similar to massive gabbro values and again near N-MORB. One exception is sample 203 (Fig. 6B1, bold line), which shows a slight elevation in LREE with no depletion in the high field strength elements (HFSE), similar to enriched MORB (E-MORB) (Sun and McDonough, 1989). Th/Ta and La/Nb ratios in this sample are similar to those in the other dikes, although La/Yb = 2.42 is higher, reflecting the higher LREE concentrations. Trace element patterns (Fig. 6B2) show characteristic N-MORB and low LILE abundances compared with the HREE, with hydrothermal alteration reflected in varying concentrations of mobile Rb, Ba, Th, U, K, and Sr. As seen in the REE, sample 203 (Fig. 6B2, bold line), shows higher concentrations of LILE, beginning with Nb and sloping downward to HREE, with a slight increase in Hf and Zr. The absence of a negative Nb-Ta anomaly suggests that higher values of LILE are not due to a subduction component but perhaps to a less depleted mantle source.
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Figure 4. (A) Cr# (Cr/Cr+Al)/Mg# (Mg/Mg+Fe2+) of Cr-spinel, the Eldivan serpentinized peridotite. Fields of abyssal (dashed line) and arc peridotites (solid line) are taken from Dick and Bullen (1984). (B) TiO2 wt% vs. Al2O3 wt% in Cr-spinel of the Eldivan ophiolite. Fields of supra-subduction zone (SSZ; solid line) and MORB (dashed line) are from Kamenetsky et al. (2001). Data are plotted with Cr-spinel data from the Orhaneli ophiolite (Sarıfakıoğlu, 2009), Brooks Range ophiolite (Harris, 1995), ocean basin peridotites (Shibata and Thompson, 1986; Morishita et al., 2007), Troodos and Oman ophiolites (Augé and Johan, 1988; Takazawa et al., 2003; Tamura and Arai, 2006), and Mariana peridotites (Ishii et al., 1992) for comparison. The Eldivan ophiolite plots mostly within the field of arc peridotites (A) and supra-subduction zone peridotites (B), indicating its subduction influenced character.
Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili In tectonic discriminant diagrams, dike rocks plot in fields of supra-subduction zones (Figs. 7A, 7F), island-arc tholeiites (IAT) (Figs. 7B, 7C, 7G) and backarc basin basalts (BABB) (Figs. 7D, 7E), and with some values plotting in the overlap of subduction zone influenced and MORB fields (Figs. 7B, 7D, 7G). Dikes plot consistently in the supra-subduction zone and backarc basin field in diagrams E and F (Fig. 7), which use the most immobile trace elements of La, Nb, and Yb to infer a tectono-magmatic setting. Additional ternary discriminant diagrams (Fig. 8) also plot dikes in subduction related fields. In diagrams B and C, dikes plot almost entirely as island arc basalt and island arc tholeiites, respectively. Diagram D has scatter between the BABB and MORB fields, whereas A and E do not discriminate between subduction and MORB basaltic rocks. Volcanic Rocks Volcanic rocks include both basaltic and rhyolitic units as blocks and broken thrust sheets within a serpentinized matrix. Basalt occurs as pillows, sheet flows, and brecciated units up to tens of square meters in area. Large blocks of basalt hundreds of meters in diameter protrude up through serpentinite to form a hummocky landscape typical of eroded mélange units. Silicic volcanic units are found only as small blocks on the meter to decimeter scale.
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Major element chemistry reveals a compositional range of volcanic rocks from basalt and andesite to dacite and rhyolite. Basaltic rocks contain 45–53 wt% SiO2 and plot mainly in the basalt field in the total-alkali silica diagram with some overlap into the basaltic andesite field (Fig. 5). Rhyolitic rocks contain 63–74 wt% SiO2 and fall into rhyolite and dacite fields (Fig. 5). Basalts show three distinct geochemical signatures: 1. The most dominant pattern shows LREE depletion characteristic of N-MORB (Fig. 6C1), with low ratios of Th/La = 0.33–2.05, La/Nb = 0.56–2.10, and La/Yb = 0.49–1.29. Element concentrations are equal to and up to 3.5 times more than N-MORB concentrations, giving these basaltic rocks the highest elemental concentrations when compared to the massive gabbros and sheeted dikes. The trace element diagrams for these basaltic volcanics (Fig. 6C2) show a large amount of variation, particularly with mobile elements of Rb, Ba, K, and Sr, which reflect their alteration. However, patterns still show the low LILE abundance (compared with HREE) that is typical of N-MORB. Discriminant diagrams for N-MORB-like basaltic rocks show more scatter than dike rocks but plot in fields of MORB more often than subduction-influenced fields (Figs. 7A–7G). In diagrams B, C, and E, basaltic rocks plot almost entirely within the MORB field, whereas some samples plot close but somewhat outside MORB fields in B, D, and G. In diagrams E and F, which use the
Rhyolite
72 Rhyodacite Dacite
68
Basaltic volcanics
Comendite Pantellerite
Silicic volcanics Gabbro
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Phonolite
Alkaline basalt
Basanite Nephelite
52 Sub-alkaline basalt 48
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Figure 5. International Union of Geological Sciences (IUGS) total alkali silica classification diagram, illustrating the distribution of rock types in the Eldivan ophiolite. Volcanic rocks are basaltic and rhyolitic, whereas dike rocks plot between the two in the basaltic andesite and andesite fields.
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Sample 126a
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Gd
Tb
Dy
Ho
Er
Yb Lu
0.1 Rb Ba Th U Nb Ta K La Ce Sr Nd P Sm Zr Hf Eu Ti Gd Tb Dy Ho Y Er Yb Lu
Figure 6. REE (1) and incompatible element (2) diagrams for (A) gabbro, (B) diabase dikes, and (C) volcanic rocks. REE elements were normalized to chondritic meteorite compositions (McDonough and Sun, 1995). Normal mid-oceanic-ridge basalt (N-MORB) reference line (Sun and McDonough, 1989) is marked by a thin dotted line on all diagrams. For comparison of the different rock types within the Eldivan ophiolite, white and gray fields are plotted in each trace element diagram: volcanic rocks in white (A2, B2), gabbro and diabase dikes in gray (A2, B2), and alkaline rocks in gray (C2). Special samples mentioned in the text are distinguished by bold solid or dashed lines.
Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili most immobile elements, most basalt samples plot in the MORB field. Ternary discriminant diagrams have similar results (Fig. 8), with basalts falling clearly into the MORB fields in diagrams B, C, and D. Diagrams A and E do not distinguish between MORB and subduction-related basaltic rocks. 2. Contrastingly, the geochemistry of samples 16, 19, 20, 21, and 142 (Fig. 6C1, light gray lines) shows highly elevated immobile LREE up to 17 times that of N-MORB and HREE below N-MORB concentrations, giving steeply sloping patterns typical of alkaline ocean island basalts (OIB) (Fig. 6C), but reflect alteration in high K abundance. Element ratios of La/Yb are correspondingly high (La/Yb = 16.21–22.47). Sample 272 (Fig. 6C1, bold line) is subparallel to these samples, showing low abundances in HREE but only a moderate LREE elevation. Trace element patterns for samples 16, 19, 20, 21, and 142 (Fig. 6C2, gray field) show the same elevation in the LILE as seen in the LREE, with some variation in mobile elements, most notably Ba and K. Sample 272 (Fig. 6C2, bold line) has LILE abundances more elevated than the N-MORBtype basalts but less than the alkaline samples. The absence of a Nb anomaly, combined with the low HREE concentrations, suggests that this sample source was not modified by subduction but perhaps was derived from a more heterogeneous source transitional between those that produce N-MORB and OIB basalts. These rocks consistently plot in fields for OIB, WPB, and alkaline basalt on discriminant diagrams (Figs. 7A, 7B, 7D, 7F). In diagrams where alkaline fields are not present, these rocks plot outside all fields (G) or overlap both the IAT and MORB fields (C), and cannot be discriminated. In diagram E, these samples plot in the E-MORB field. Similar results are seen with additional ternary discriminant diagrams (Fig. 8). In all diagrams, these rocks plot as within plate alkali (A, B, D), ocean island tholeiite (C), and WPB (E). 3. The third geochemical signature is seen in basaltic sample 274 (Fig. 6C1 and 2, bold dashed line), which shows LILE enrichment and HFSE (Nb, Ta) depletion relative to N-MORB, suggesting a subduction-influenced source (Fig. 6C2). A subduction-influenced source is also reflected in a slight LREE enrichment and high ratios of Th/La (6.48), La/Nb (2.85), and La/Yb (3.66) (Fig. 6C1). Rhyolitic samples have two separate trace and REE signatures. Most samples are similar to N-MORB in their lower LREE concentrations compared with HREE (Fig. 6C1) but have overall flatter patterns that could reflect higher degrees of differentiation in the magma chamber. This same pattern is also seen in the trace elements (Fig. 6C2), with LILE slightly lower in concentration than HREE, with the exception of Zr and Hf, which show higher abundances. The effects of secondary alteration are seen in the scatter of mobile elements, especially Ba, U, Th, K, and Sr. Alternatively, rhyolitic samples 280 and 282 (Fig. 6C2, bold dashed line) closely match basaltic sample 274. LREE abundances are elevated, reflected in Th/Ta, La/Nb, and La/Yb ratios similar to sample 274 (Fig. 6C1). LILE (Fig. 6C2) are also more abundant except for negative concentrations of Nb, Ta, and Ti,
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characteristic of subduction zones. Some scatter is still seen in the mobile elements, particularly Rb, K, and Sr. Despite the two signatures seen in REE and trace elements, all of the rhyolitic samples plot in fields for volcanic arc granite (Fig. 7, H, I) in discriminant diagrams of Pearce et al. (1984a) that are based on immobile elements of Nb, Ta, and Yb. Interpretation of Whole-Rock and Mineral Chemistry Three different magma affinities are present in the Eldivan ophiolite: N-MORB, alkaline (OIB), and supra-subduction zone. The occurrence of three distinct geochemical signatures in this small area of exposure (~20 km2) implies a high degree of mixing of either (1) upper and lower plate blocks during tectonic emplacement or (2) magma sources during seafloor formation. Mixing of upper and lower plate units is plausible, considering the current imbricated structure of the ophiolite in the mélange. This has been suggested to account for alkaline rocks within the mélange that are interpreted as seamounts accreted into the serpentine mélange from the downgoing plate (Floyd, 1993; Tüysüz et al., 1995; Tankut et al., 1998). Accretionary mixing of an N-MORB downgoing plate, which included seamounts, with a supra-subduction zone upper plate could explain the geochemical variation in the Eldivan ophiolite, although few modern analogues of this process exist. Similar chemical variations to those seen in the Eldivan ophiolite are found in modern backarc supra-subduction zone basins due to mixing different magma sources rather than upper and lower plate components. Such supra-subduction zones or backarc ocean basins are extensional upper plate basins that form above subduction zones owing to lower plate movement away from the upper plate through slab rollback. The combination of extension and subduction in supra-subduction zone backarc settings creates conditions of both mantle depletion and enrichment, which result in basalts of different compositions (Sinton and Fryer, 1987; Price et al., 1990; Stern et al., 1990; Eissen et al., 1994; Hawkins and Melchior, 1985; Dril et al., 1997; Fretzdorff et al., 2002; Sinton et al., 2003). Basalts in the North Fiji, Lau, Mariana, Manus, and East Scotia backarc basins show an overprint of LILE enrichment on N-MORB geochemical patterns, which increase with proximity to the subducting slab. Compositional zoning in the Lau basin, with LILE enriched basalt on the west edge near the arc and N-MORB types in the young central spreading center, show that LILE enrichment decreases as rifting continues, owing to decreased subduction influence from slab rollback (Hawkins and Melchior, 1985; Pearce et al., 1984b). Likewise, initial rifts in the Mariana trough erupt basalts similar to those of the Mariana arc, where older rift zones erupt N-MORB (Stern et al., 1990). North Fiji and East Scotia spreading ridges erupt basalt transitional between N-MORB and alkaline basalt owing to influence from hotspot volcanism (Price et al., 1990; Eissen et al., 1994; Fretzdorff et al., 2002). This chemical array is similar to that seen in the Eldivan ophiolite and occurs entirely in the upper plate, caused by mixing of variably depleted
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Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili and enriched mantle sources or melts (Sinton and Fryer, 1987; Price et al., 1990; Stern et al., 1990; Dril et al., 1997). In this study the backarc basin or supra-subduction zone mixing is favored for the Eldivan ophiolite. Although incompatible and REE diagrams are dominated by N-MORB patterns, there are noticeable subduction and alkaline influences (samples 126a, 203, 280, and 272), as seen in modern backarc and intraarc settings. Additionally, basalt mostly plots as N-MORB, with some scatter into other fields (Figs. 7 and 8), but dike rocks plot consistently within subduction-influenced fields, including IAT, backarc basin, and supra-subduction zone, with minor overlap into N-MORB fields (Figs. 7 and 8). These dike compositional patterns provide direct evidence that the Eldivan ophiolite was at one time in a supra-subduction setting, as the sheeted dike complex represents ocean floor construction. Additionally, rhyolitic volcanics also plot in volcanic arc fields (Fig. 7), giving more evidence for a significant subduction influence. Evidence for a supra-subduction zone setting is also found in the mantle sequence of the Eldivan ophiolite. Cr-spinel (Cr#s 0.47–0.70) plots within fields for mantle more depleted than ocean rift mantle and closer to transitional supra-subduction settings (Fig. 4) similar to Oman-type ophiolites as defined by Harris (1992). This is also supported in whole-rock Al2O3 wt% of the Eldivan ophiolite, which indicates the degree of partial melt extraction, and could be expected for ocean crust in the complex melting regime of a backarc basin. It also closely matches Al2O3 wt% concentrations from the Oman ophiolite but is slightly higher than the Brooks Range ophiolite and Orhaneli ophiolite (western İzmir-Ankara-Erzincan suture zone), all thought to have formed in supra-subduction zone settings (Fig. 3). Finally, a supra-subduction zone interpretation is consistent with other studies along the İzmir-Ankara-Erzincan suture zone in the Ankara Mélange, Dağkuplu Mélange, and Kırşehir block ophiolitic massifs. Other areas of the Ankara Mélange contain alkaline basalts (Çapan and Floyd, 1985, 1993; Tankut et al., 1998), N-MORBs (Tankut, 1984; Tankut et al., 1998), and IAT basalts (Tankut, 1984; Tankut et al., 1998; Tüysüz et al., 1995) similar to the Eldivan ophiolite. Our discovery of suprasubduction zone dikes is consistent with the discovery of supra-
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subduction zone plagiogranite in the Ankara Mélange (Dilek and Thy, 2006). Similarly, the Dağkuplu Mélange in the western İzmirAnkara-Erzincan suture zone shows the same variety of alkaline, N-MORB, and supra-subduction zone geochemistry (Göncüoğlu et al., 2006; Sarıfakıoğlu, 2006; Sarıfakıoğlu et al., 2009). Suprasubduction geochemistry also characterizes Cretaceous ophiolites from the Kırşehir block (Çiçekdağ and Sarıkarıman massifs) (Yalınız et al., 1996, 2000a; Floyd et al., 2000). Most of these previous studies use geochemistry of crustal volcanic rocks, cumulate sequences, dike complexes, and massive gabbros for their interpretation without data from associated mantle peridotite. Including analyses of the peridotite provides additional evidence for high degrees of melt extraction inconsistent with a MORB tectonic model for the Eldivan ophiolite. Sarıfakıoğlu et al. (2009) used both mineral and whole-rock geochemistry of crust and mantle rocks to interpret the Orhaneli ophiolite in the western İzmir-Ankara-Erzincan suture zone as a supra-subduction zone ophiolite. Cr-spinel data from lherzolite and harzburgite of the Orhaneli mantle sequence closely match the Eldivan ophiolite (Fig. 4). Whole-rock Al2O3 from the Orhaneli ophiolite is generally lower but is still within the range of the Eldivan ophiolite (Fig. 3). These results show the same continuity in mantle composition as seen in the crustal-sequence geochemical data that argue for some supra-subduction influence from the western to central İzmir-Ankara-Erzincan Ocean. EPI-OPHIOLITIC SEDIMENTARY COVER UNITS Epi-ophiolitic sediment occurs as blocks, intercalated sediment in pillow lobes, and layered sediments depositionally overlying pillow basalts. Blocks are generally meter to decimeter sized blocks of pelagic, radiolarian-bearing limestone and minor chert within the serpentinized matrix of the mélange but not in direct contact with ophiolitic units. Some radiolarian-bearing red chert is intercalated within pillow basalt lobes. Layered sediments overlying the ophiolite consist of interbedded chert and limestone, and chert interbedded with pillow basalt and volcanic breccia. Individual beds are centimeters thick, but sediment sequences can reach 10 m. In one locality, shale and minor sandstone turbidites were found within the epi-ophiolitic cover. Karadağ Formation
Figure 7. Nine discriminant diagrams for basaltic rocks of the Eldivan ophiolite. In diagrams A–G, basaltic rocks generally plot in MORB fields, although some show too much scatter to be conclusive. Dikes mostly fall into island arc tholeiite (IAT) fields with some overlap in MORB fields. Alkaline basalts plot consistently in ocean island basalt (OIB) or within plate basalt (WPB) fields. Diagrams H and I are for granitic rocks. Samples from the Eldivan ophiolite all plot in volcanic-arc granite fields. These diagrams are from (A) Shervais (1982); (B) Pearce and Norry (1979); (C) Pearce (1982); (D) Woodhead et al. (1993) and Floyd et al. (2000); (E) Floyd et al. (1991); (F) Pearce et al. (1981); (G) Pearce and Cann (1973); (H) Pearce et al. (1984a); (I) Pearce et al. (1984a). BABB—backarc basin basalt; E-MORB—enriched MORB; N-MORB—normal MORB; SSZ— supra-subduction zone; CAB—continental arc basalt.
The Karadağ Formation overlies the Ankara Mélange along an angular unconformity and, because of multiple deformation phases, also overlies the mélange tectonically as a result of later thrust faulting. It is made up of intercalated volcanic and coarse siliciclastics at its base that grade upward into finer sandstones and mudstones and finally clay-rich limestone (Akyürek et al., 1980; Hakyemez et al., 1986). It is interpreted as flysch deposited in a foredeep setting near a continental margin during the onset of collision. The Kursunluduz Member of the Karadağ Formation contains chert bands alternating with red pelagic limestone (Akyürek et al., 1980; Hakyemez et al., 1986). The presence of
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Figure 8. Discriminant ternary diagrams for basaltic rocks of the Eldivan ophiolite. Basaltic rocks are black diamonds, dikes are gray squares, and alkaline basalts are light gray diamonds. In diagrams A–E, basaltic rocks fall mostly into the field of MORB with little overlap into subduction-influenced fields. In diagrams that do not distinguish between MORB and subduction fields, basalts plot in both fields. Dikes from the Eldivan ophiolite fall into subduction-related fields in diagrams that distinguish between MORB and arc-related rocks. In those that do not, dikes plot in the MORB-arc fields. Alkaline basalts consistently plot in enriched ocean island fields or within plate basalt (WPB) fields. Fields are from (A) Meschede and Casey (1986); (B) Wood (1980); (C) Mullen (1983); (D) Cabanis and Lecolle (1989); (E) Pearce and Cann (1971). WPA—within plate alkaline basalt; WPT—within plate tholeiitic basalt; OIT—ocean island tholeiite; OIA—ocean island alkalic basalt; CAB—continental arc basalt; PMORB—plume MORB; VAB—volcanic arc basalt; IAB—island arc basalt; VAT—volcanic arc tholeiite; OFB—ocean floor basalt; LKT— low-K tholeiite. See Figure 7 caption for additional abbreviations.
Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili
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Praegglobotruncana stephani, Rotaliapora apenninica, Hedbergella sp., Ticinella sp., Globigerina sp., Textulariella sp., Cuneolina sp., and Valvulammina sp. radiolarians suggests the age of the Karadağ is Cenomanian to Campanian (Akyürek et al., 1980). Near the Eldivan ophiolite, the Karadağ Formation consists mostly of pelagic limestone and chert, and some clastic material, including sandstone lenses with graded and cross-beds. The angular unconformity between the Karadağ Formation and the underlying imbricated ophiolitic material of the Ankara Mélange suggests that the Karadağ Formation was deposited after or during imbrication of the ophiolite, and is in part correlative with the overlying Maastrichtian flysch of Norman (1984).
al. (2006). Single point analyses were taken with a 35 µm and a 25 µm diameter beam according to grain size. Common Pb corrections are for 204Pb, using an initial Pb composition from Stacey and Kramers (1975). Uncertainties are 1.0 for 206Pb/204Pb, 0.3 for 207 Pb/204Pb, and 2.0 for 208Pb/204Pb. Detrital zircon age extractor and ISOPLOT 3.00 (Ludwig, 2003) were used to determine and sort reliable age data. This detrital zircon age extractor extracts significant peak ages based on at least three grain analyses and the number of grains constituting each peak age. Results are listed in Table A3.
U/Pb Age Analysis of Detrital Zircon in Sandstone Units
Detrital zircon age populations from sandstone in the mélange and the Karadağ Formation have different minimum, maximum, and peak ages, suggesting that they were sourced from different terranes (Fig. 9). Detailed analysis of the mélange sandstone shows an age distribution from 143.2 ±2 Ma to 164.1 ±1 Ma with a peak age of 153 Ma. The Karadağ sandstone shows an age distribution from 105.2 ±5 Ma to 166 ±3 Ma, with a peak age of 130 Ma. The youngest peak age is used here as a proxy for the maximum age of deposition, which is consistent with the stratigraphic positions of the sandstones. The maximum age of the mélange sandstone, and Eldivan ophiolite, is 143.2 ±2 Ma, whereas the maximum age of the Karadağ sandstone is 105.2 ±5 Ma. Peak ages here are interpreted to represent the average age of the terrane dominantly being eroded at the time of deposition. An inherited fraction of zircon from the Neoproterozoic to Paleoproterozoic is present in both the mélange and the Karadağ sandstones (Fig. 9). Detrital zircons of similar age from the Tauride block in southwestern Turkey were documented by Kröner and Şengör (1990), who attributed them to the southern Angara
Sandstone samples were collected from a block within the Ankara Mélange directly adjacent to basalt and from the Karadağ Formation, which unconformably overlies the mélange. The age and tectonic source region for sandstone samples from the mélange and overlying Karadağ Formation were investigated through detrital zircon and sandstone petrography. Siliciclastic material was scarce, and these two samples represent the only sandstones found within the study area. The entire sample collected was processed for detrital zircons. The error introduced by the limited sample size and small number of zircons found within each sample is recognized, however the results are consistent with age data obtained by other methods throughout the suture zone. Zircon U-Pb age analyses were conducted by laser-ablation multicollector inductively coupled-plasma mass spectrometry (LA-MC-ICPMS) at the Arizona LaserChron Center. Analytical methods follow those described in Gehrels (2000) and Gehrels et
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craton of Siberia. Dilek and Thy (2006) also found Proterozoic zircon in plagiogranite from the ophiolite near Ankara and interpreted them as a subduction recycled component from the Rhodope-Strandja Massif in northwestern Turkey and southeastern Bulgaria. These terranes may also have supplied the Paleoproterozoic to Neoproterozoic zircon grains in the mélange and Karadağ sandstones.
No deep water fauna are recorded. These data suggest that the Karadağ sandstone was derived from the carbonate system of a continental margin and was deposited in a nearby marginal basin. This is consistent with the interpretation of the Karadağ Formation as flysch deposited on the imbricated ophiolite, most likely in a foredeep setting near the continental margin, created as continuing subduction brought the Kırşehir and SakaryaPontide terranes together.
Sandstone Petrography STRUCTURE OF THE ELDIVAN OPHIOLITE Sandstone from both formations is compositionally and texturally immature, with low percentages of quartz and angular to subangular clasts. Despite alteration and secondary authigenic growth, the sandstone samples yielded two very different petrographic provenance results. The mélange sandstone is dominated by volcanic lithic fragments (52.33%) and plagioclase (25.33%), with minor quartz (8.66%), K-feldspar (2.00%), and clay minerals (11.66%). In contrast, the overlying Karadağ sandstone is made up of carbonate mud clasts (with some authigenic clay) (45.33%), plagioclase (28.00%), bioclastic grains (15.00%), and quartz (11.33%), with minor volcanic lithic material (0.33%). The composition of the mélange sandstone with its high percentage of volcanic lithic fragments implies that it was sourced from a nearby volcanic terrane. This idea agrees with the tectonic discriminant diagrams of Dickinson et al. (1983) (not shown). There are a number of sources for lithic fragments in the İzmir-Ankara-Erzincan Ocean, including seamounts, island arcs (Tankut, 1984; Tankut et al., 1998; Tüysüz et al., 1995), and the Pontide continental arc to the north. However, the Pontide arc is younger (Turonian) than detrital zircon grains found in the mélange sandstone, suggesting that it is not the source for volcanic lithics. The other plausible volcanic sources are oceanic, giving more evidence for intra-oceanic subduction away from the continental margin. Sandstone from the Karadağ Formation contains virtually no volcanic lithics, bioclastic grains, carbonate mud, or plagioclase grains. The bioclastic material in this sample is a mix of echinoderm, bryozoan, brachiopod, bivalve, and foraminiferan grains, a compositional variation that suggests a well-developed but relatively shallow carbonate system.
The angular unconformity between the Karadağ Formation and the underlying imbricated mélange, and the subsequent shortening of both, imply multiple phases of deformation (Fig. 10). The first phase involved dismemberment of the Eldivan ophiolite and serpentinite mélange development (Fig. 10A). Where paleohorizontal indicators exist, they show predominantly steep dips. Elongated blocks are also commonly vertical, with gaps between them filled with serpentine. If the mélange is associated with an accretionary wedge, it is likely the sections investigated formed near the backstop region, the area where accreted thrust sheets are progressively rotated to steeper dips by accretion of new material beneath them. Accretionary wedge development commonly produces isoclinally folded units with mostly sub-horizontal fold hinge lines. Although horizontal fold hinge lines are found within the serpentinite matrix and in overlying Karadağ Formation units, no hinge lines were found in the several blocks we investigated. Another way to explain the mostly steep dips in mélange blocks is by strike-slip deformation, which would produce steeply plunging hinge lines. There is no evidence of these in any part of the field area. The second phase of deformation occurred after the Karadağ Formation was deposited above the Ankara Mélange. During this second phase, the Karadağ Formation and underlying ophiolite and serpentine were thrust along southwardverging thrust faults (Fig. 10B). Some age constraints for these events are provided by detrital zircon populations from the mélange and Karadağ sandstones.
Epi-ophiolitic cover Karadag Formation
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Gabbro Serpentinized peridotite Figure 10. Multiple deformational phases are found in the Ankara Mélange. The first phase imbricated and rotated the oceanic units into mostly vertical units, encased in serpentine mélange (A). Subsequent phases involved shortening of the Ankara Mélange and the unconformably overlying Karadağ Formation (B). Subsequent events included serpentine diapirism and thrusting of mélange over late Miocene deposits of the Hançili Formation.
Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili Timing of Ophiolite Imbrication Detrital zircon ages from sandstone within and above the Ankara Mélange provide limits on the timing of mélange formation and dismemberment of the Eldivan ophiolite. Zircon grains as young as 143 ±2 Ma within a sandstone of the mélange indicate that it was incorporated into the mélange after this time and before the youngest age of zircon grains within the unconformably overlying Karadağ sandstone, which yields ages of 105 ±5 Ma. Further imbrication of the ophiolite and the Karadağ Formation occurred after 105 ±5 Ma. Imbrication of the Eldivan ophiolite between 143 ±2 and 105 ±5 Ma is consistent with data from other parts of the suture zone that suggest collapse of the ocean basin had begun about this time. For example, radiolarians in limestone deposits from the Kirazbaşı foredeep complex are as old as ca. 135 Ma (late Valanginian) (Tüysüz and Tekin, 2007). Intra-oceanic thrusting began prior to 90 Ma near the Kırşehir block (Yalınız et al., 2000b) and 93 ±2 Ma in the western İzmir-Ankara-Erzincan Ocean (Önen, 2003). Granitoids of 94.9 ±3.4 Ma, in the Kırşehir block, interpreted as the result of supra-subduction zone ophiolite obduction, also suggest that subduction must have been active before 95 Ma (Boztuğ et al., 2007). Restoration of the Eldivan Ophiolite The orientation of sheeted dikes in ocean crust is commonly used as a proxy for spreading ridge orientation. For the Eldivan ophiolite, it would represent the orientation of a supra-subduction zone spreading ridge. Commonly, sheeted dikes are perpendicular to overlying basaltic flows that they feed. Most dikes in the Eldivan Ophiolite strike NNW-SSE and are steeply dipping (Fig. 11A). A minor component of E-W dikes is also found, but most are horizontal. In one locality, which represents the largest single mélange block of basaltic rock in the Hançili region, a series of horizontal sheeted dikes is in contact with vertical pil-
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low basalts and sheet flows that strike E-W, indicating that the entire igneous section has rotated 90° about a horizontal axis. Sedimentary blocks are also steeply dipping, indicating a similar amount of rotation about a horizontal E-W axis. Assuming that other sheeted dikes in the region underwent a similar horizontal axis rotation, applying this rotation to the NNE-SSW dikes provides the original orientation of the suprasubduction zone spreading ridge that produced the Eldivan ophiolite. The rotation maintains the steep dip of the NNW-SSE dikes but changes the average strike to near N-S (Fig. 11B), which is subparallel to the İzmir-Ankara-Erzincan suture zone. However, oroclinal bending of this suture zone from indentation of the Kırşehir block indicates an additional 90° of counterclockwise vertical axis rotation. Paleomagnetic studies used to test the oroclinal bend hypothesis found that the Ankara Mélange near the Eldivan ophiolite has undergone at least 30° of counterclockwise vertical axis rotation since the Eocene, and that it may have already been rotated counterclockwise by even more before this time (Kaymakci et al., 2003). According to Kaymakci et al. (2003), the Çankırı basin underwent rotation during the Eocene Epoch through the midMiocene, and perhaps as early as the Paleocene Epoch. Near the Eldivan ophiolite the Çankırı basin margin rotated 33° counterclockwise during Oligocene time. A clockwise vertical axis rotation of 33° was applied to correct for these rotations, which moves the average strike direction of most sheeted dikes to 041, which was most likely the orientation in the Eocene (Fig. 11C). To completely restore the İzmir-Ankara-Erzincan suture zone back to its pre-collisional indentation trend, an additional 52° of clockwise rotation is needed. Correcting for the horizontal and vertical axis rotations of the sheeted dikes shows that they are subparallel to the İzmirAnkara-Erzincan suture zone, which implies mostly orthogonal motion of the spreading ocean with respect to the subduction boundary represented by the suture zone, indicating little strike slip motion in the creation of the suture zone. These results are
Figure 11. Stereographs of poles to dike attitudes in sheeted dike units of the Eldivan ophiolite. Solid black lines are average strike of sheeted dikes and most likely the spreading ridge: (A) Prior to any restoration. (B) Restored about a horizontal axis according to paleo-horizontal controls. (C) Partially restored to original orientation by 30° of post-Eocene clockwise vertical axis rotation documented from paleomagnetic data. Oroclinal bending of the İzmir-Ankara-Erzincan suture zone indicates that at least another 60° of clockwise rotation is needed to restore the dikes back to their original orientation, which would be near E-W.
158
Dangerfield et al.
consistent with those found by Fayon et al. (2001) and Whitney et al. (2001), who concluded that the northern part of the Kırşehir block was deformed and exhumed by orthogonal collision. However, whereas the İzmir-Ankara-Erzincan suture zone may have formed through orthogonal motion, Fayon et al. (2001) and Whitney et al. (2001) give evidence for a later oblique collision of the Tauride platforms with the southern Kırşehir block, exhuming the southern Kırşehir block through left-lateral wrench faulting. TECTONIC EVOLUTION Review of the İzmir-Ankara-Erzincan Ocean Understanding the overall tectonic evolution of this ocean is crucial to reconstructing the role played by the Eldivan ophiolite. Age constraints of various events throughout the İzmir-AnkaraErzincan suture zone indicate three main phases of İzmir-AnkaraErzincan Ocean evolution: a constructional phase, a destructional phase, and a suturing phase (Fig. 12A). The constructional phase began with rifting at least as old as the late Carnian–early Norian Stages (ca. 215 Ma), based on radiolarians associated with MORB in the central and western parts of the suture zone (Bragin and Tekin, 1996; Tekin et al., 2002). Other radiolarians suggest that it developed into an ocean basin by late Bajocian time (Tüysüz and Tekin, 2007), and seamounts formed on the ocean floor during the Jurassic and Cretaceous Periods (Rojay et al., 2001, 2004; Tankut et al., 1998). Destruction of the ocean basin by intra-oceanic subduction is documented by supra-subduction zone ophiolites, the oldest in the Ankara Mélange, yielding a U/Pb zircon age of 179 ±15 Ma (Dilek and Thy, 2006). Intra-oceanic subduction continued through the Late Cretaceous Period, creating supra-subduction zone ocean crust in the central İzmir-Ankara-Erzincan Ocean, now part of the Kırşehir block (Yalınız et al., 1996; Yalınız et al., 2000b), and the Dağkuplu mélange (Göncüoğlu et al., 2006; Sarıfakıoğlu, 2006; Sarıfakıoğlu et al., 2009). Late Valanginian (ca. 135 Ma) (Tüysüz and Tekin, 2007) to Paleocene (Koyçiğit, 1991) radiolarians are found in foredeep deposits along the Sakarya-Pontide margin, suggesting that active subduction against the continent began in the Early Cretaceous Period. Other events documenting subduction at this time are the occurrence of accretion complexes along the Pontide margin, which were metamorphosed at ca. 100 Ma (Okay et al., 2006), and Turonian Epoch (ca. 90 Ma) through Paleocene Epoch magmatism in the Ponticles (Yılmaz et al., 1997). It is important to note that the oldest age of supra-subduction zone ophiolites, 179 ± 15 Ma, predate the oldest foredeep deposits (ca. 135 Ma) against the continent, suggesting that intra-oceanic extension occurred prior to subduction against the continental margin. Thrusting and imbrication of the Eldivan supra-subduction zone basin in the central İzmir-Ankara-Erzincan suture zone occurred between 105 and 143 Ma, as shown by detrital zircon ages from this study. In the western İzmir-Ankara-Erzincan Ocean, thrusting began at least by 94 Ma, as recorded by the age of a metamorphic sole
(Önen, 2003). This constrains a destructive phase of subduction that began ca. 179 ±15 Ma and ended as early as ca. 60 Ma. Final closure of the İzmir-Ankara-Erzincan Ocean occurred through continental block collision, collisional indentation, and suturing of the Kırşehir block and the larger Anatolide-Tauride platform with the Sakarya-Pontide terranes during the Late Cretaceous Period to the Miocene Epoch. The first evidence of continental collision comes from post-collisional granitoids of the Kırşehir block, which yielded Rb-Sr whole-rock and 207Pb-206Pb zircon ages from 110 ±14 Ma (Güleç, 1994) to 74.9 ±3.8 Ma (Boztuğ et al., 2007). Exhumation of the collision zone in the Central Pontides, based on stratigraphic constraints, and granitoids of the Kırşehir block, based on apatite fission-track ages, documents collision between 86 and 93 Ma and 57 and 62 Ma, respectively (Okay et al., 2006; Boztuğ and Jonckheere, 2007). Collisional indentation of the Kırşehir block caused at least 90° of counterclockwise rotation, 33° of which is well constrained since the Eocene. Ages of continental block collision young away from the central part of the suture zone, where the Kırşehir block is present. In the western part of the suture, where the Kırşehir block is absent, and collision of the Sakayra-Pontide terrane occurred only within the Anatolide-Tauride block,
Figure 12. (A) Age constraints for the evolution of the İzmir-AnkaraErzincan suture zone through time. Three main phases are identified: (1) an initial construction phase in which the ocean basin was forming through ridge spreading with hotspot volcanism creating seamounts on the ocean floor, (2) destruction of the ocean basin through intraoceanic subduction that resulted in intra-oceanic seafloor spreading above a subduction zone and arc magmatism, and (3) collision and suturing of the Kırşehir and Anatolide-Tauride continental blocks with the Sakarya-Pontide terranes. Numbers in the time line and map correspond with the source of age data (below) and sample locations, respectively. The sample location for detrital zircons of this study is represented with a black star. Sources for data are as follows: (1) Tekin et al. (2002); (2) Bragin and Tekin (1996); (3) Dilek and Thy (2006); (4) Rojay et al. (2004); (5) Rojay et al. (2001); (6) Göncüoğlu et al. (2006); (7) Tüysüz and Tekin (2007); (8) Önen (2003); (9) Yalınız et al. (2000); (10) Koçyiğit (1991); (11) Yalınız et al. (1999); (12) Boztuğ and Jonckheere (2007); (13) Yılmaz et al. (1997); (14) Okay et al. (2006); (15) Kaymakci et al. (2003); (16) Boztuğ et al. (2007) and references therein; (17) Fayon et al. (2001). (B) Schematic cartoon model for the evolution of the Eldivan ophiolite during the Cretaceous, using the Philippine Sea plate and Mariana trough as an analogue. Early Cretaceous time documents the beginning of subduction and upper plate extension, as evidenced by supra-subduction zone (SSZ) basalt, foredeep complexes along the continental margin, and ophiolitic metamorphic soles. In the Late Cretaceous Period supra-subduction zone upper plate subduction began along the Sakarya-Pontide margin, causing active volcanism in the Pontide continental arc. The latest Cretaceous Period through the Oligo-Miocene Epochs was characterized by collision and suturing of the Kırşehir block (KB) and AnatolideTauride platform with the Sakarya-Pontide terrane, as evidenced by post-collisional granitoids and fission-track (FT) exhumation ages of the Kırşehir block. Fission-track ages also indicate that wrench faulting exhumed the southern Kırşehir block through left-lateral, strikeslip motion owing to later oblique collision of the Tauride platform with the Kırşehir block. OIB—ocean island basalt.
Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili
13
6 1
e ya Zo8 n S akar
7 4
4
2 4 10 4 4 12 5
12
Kirsehir 9 block 11
Izm i
Erzincan suture arank A r-
12
Anatolide-Tauride block
40Ar/39Ar regional 40Ar/39Ar age of metamorphic sole (8) metamorphic age (8) Active Pontide magmatic arc (13)
Zircon U-Pb SSZ ocean crust crystallization age (3)
Radiolaria from blocks of pelagic sediment assoc. w/ OIB and SSZ basalt (2,6,9)
Oldest radiolaria from limestone and chert blocks (1,2)
Radiolaria from intercalated sediment on OIB pillow basalts (4,5,6) Crystallization ages (various methods) of post-collisional Central Anatolian granitoids (11, 16)
Rotation in the Cankiri basin (15)
Radiolaria from forearc basin deposits (10)
200
Radiolaria from blocks of pelagic sediment (2)
Radiolaria from the Kirazbasi foredeep complex (7,14) FT 1st phase exhumation age of post-collisional granitoids (12) Metamorphism of accretionary complex (14) FT 2nd phase exhumation age of post-collisional granitoids (12,17)
180
160
Exhumation of accretionary complex (14)
140
Ma
120
80
100
Constructional phase Construction of the ocean basin through ocean ridge spreading and seamount volcanism
60
Destructional phase Destruction of the ocean basin through intra-oceanic subduction, creating SSZ ocean crust, forearc basins, and intra-oceanic thrusting
40
Suturing phase Collision of the Kirsehir and Anatolide-Tauride blocks with the Sakarya-Pontide terrane
20
0
Iz
uture can s -Erzin a r a k n mir-A
220
de onti a-P Intr
Central Pontides
14
sut ure
Ist
on e ul z b ure n a S ut
Inter-Tauride
RhodopeRhodope Strandja Massif
159
Detrital zircon U-Pb age, Karadag Fm, this study
220
200
180
160
140
Ma
120
100
80
60
40
20
0
Detrital zircon U-Pb age, epi-ophiolitic sandstone, Eldivan ophiolite, this study
A N
S
25-100 Ma
65 to 135 Ma
Ankara Mélange
65?-179 Ma
Subduction of SSZ upper plate
S-P
B
KB
S-P
Upper plate SSZ extension
KB
S-P
KB
160
Dangerfield et al.
exhumation is documented by 40Ar/39Ar metamorphic cooling ages of 48 ±12 Ma (Önen, 2003). These age constraints suggest that the collision of the Kırşehir block with the Central Pontides may have occurred significantly earlier than collision between the Anatolide-Tauride block with the rest of the continental margin. The age of final suturing (no more deformation) between the Kırşehir block and Anatolide-Tauride platform with the Sakarya-Pontide terrane is not constrained. Boztuğ and Jonckheere (2007) attribute a second phase of granitoid exhumation in the Kırşehir block at 28–30 Ma to collision of the Arabian-African platform in the east, where Fayon et al. (2001) interpret exhumation of granitoids at 35 Ma to be from collision of the Anatolide-Tauride platform. Shortening continued into the late Miocene and Pliocene, as indicated by thrusting of the Ankara Mélange over the edge of the Hançili basin. Plate Tectonic Setting The Eldivan ophiolite was created in the upper plate of the İzmir-Ankara-Erzincan Ocean during oblique intra-oceanic subduction as part of a backarc basin. This created a suite of geochemical signatures, as supra-subduction zone melting modified an N-MORB mantle, which was mixed with an enriched OIB mantle that had previously created seamounts on the ocean floor (Fig. 12B). The current Philippine Sea plate and Mariana trough supra-subduction zone basins are suggested as modern analogues for the tectonic setting of the Eldivan ophiolite, İzmir-AnkaraErzincan Ocean, and Ankara Mélange. The Philippine Sea plate formed as an upper-plate suprasubduction zone basin caused by intra-oceanic subduction (Harris, 2003). Later, the Philippine Sea supra-subduction zone ocean basin began to subduct to the west, creating the Japan, Ryukyu, and Luzon arcs. In a similar way, formation of the Eldivan intra-oceanic basin began subduction beneath the Pontides, creating the Pontide magmatic arc. Subduction of the Eldivan oceanic basin beneath the Pontides allowed several large fragments of mostly supra-subduction zone oceanic crust and some mantle to accrete to the margin, serpentinize, and produce mélange in the forearc between 105 and 143 Ma (Fig. 12B). Eventually the subduction zone was choked by collision of the Kırşehir block with the Sakarya-Pontide terrane, which further imbricated the ophiolite with the overlying Karadağ Formation. Collision continued to indent the continental margin and rotate the Eldivan ophiolite from its original E-W orientation to its current position on the western edge of the large omega-shaped İzmir-Ankara-Erzincan suture zone. How far the Kırşehir block has traveled is not yet constrained. However, according to the Philippine Sea plate model, as supra-subduction zone basins open along a continental margin, fragments of the margin are rifted off and travel away as the basin opens (Harris, 2003). Closure of the supra-subduction zone basin eventually brings many of these fragments back into collision with parts of the original continental margin from which
they were rifted. These processes are illustrated in many parts of the equatorial Pacific and Indonesian regions (Harris, 2003). CONCLUSIONS 1. The dismembered Eldivan ophiolite is a remnant of the İzmir-Ankara-Erzincan Ocean branch of the northern NeoTethys that evolved as a supra-subduction zone basin between the Gondwana-derived Kırşehir and Anatolide-Tauride blocks and the Sakarya-Pontide margin. 2. Parts of the İzmir-Ankara-Erzincan Ocean were accreted to the southern Asian margin as it subducted beneath it. These fragments were incorporated into the serpentine-rich Ankara Mélange. 3. During accretion, most of the units scraped from the İzmir-Ankara-Erzincan Ocean were imbricated, steeply inclined, and later broken into blocks surrounded by serpentinite. These include fragments of mostly oceanic crustal material, limestone, chert, and rare sandstone. 4. The ages of some blocks in the Ankara Mélange are younger than 143 ±2 Ma, with imbrication and initial destruction of the ocean basin having occurred between 143 ±2 Ma and 105 ±5 Ma. These ages are older than those of imbrication of the İzmir-Ankara-Erzincan Ocean in the west, which is documented at ca. 94 Ma. 5. Intra-oceanic volcanic arcs or seamounts are likely source terranes for sandstone units associated with the Eldivan ophiolite, suggesting that the ophiolite formed in an intra-oceanic subduction zone away from significant continental influence. 6. Studies of sheeted dike orientations indicate that the spreading ridge of intra-oceanic supra-subduction zone basins was most likely subparallel to the southern margin of Asia before indentation of the Kırşehir block, and has since been rotated nearly 90° counterclockwise. 7. The tectonic setting and evolutionary history of the Eldivan ophiolite can be characterized as a Western Pacific–type suture system in contrast to the more classic Himalayan-type suture that involves subduction of large tracts of MORB-like oceanic lithosphere and juxtaposition far-traveled of continental blocks of different affinities. This interpretation may also apply to many other Cretaceous ophiolite-bearing suture zones of the Eastern Mediterranean. ACKNOWLEDGMENTS Funding for this project was provided by the American Association of Petroleum Geologists and U.S. National Science Foundation grant EAR 0337221. We would especially like to acknowledge the Maden Tetkik ve Arama Genel Müdürlüğü (MTA) in Turkey for supporting the fieldwork for this study, and we thank Mustafa Sevin, Esra Esirtgen, and Serdal Alemdar for their tremendous help in the field. Also, we thank Victor Valencia at the University of Arizona, who dated our detrital zircon samples, and Steve Nelson and Mike Dorais for critical reviews of our manuscript.
0.06
0.02
1.79
8.46
0.12
42.20
0.18
0.00
0.04
0.01
97.82
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
Total
0.11
8.75
0.78
0.01
0.00
0.00
0.00
0.10
110
126a
0.75
0.25
0.02
1.22
2.11
9.06
9.24
0.18
0.04
1.03
4.61
5.41
5.73
0.13
11.19 10.32 0.14 5.95
8.48
<0.01 0.1 0.0
0.1 0.1 0.0
0.2
<0.1
Er Yb Lu
Hf
Ta
<0.1
0.2
0.2 0.2 0.0
0.0 0.3 0.1
5.80
0.17
8.78
1.95 0.13 0.03
4.99 0.11 0.03 0.03
0.13
1.96
10.62 10.65
8.96
0.22 6.89
0.33
17.47 17.63
0.33
267
0.02
1.45
3.22
8.72
7.18
0.17
9.01
16.62
0.46
53.59
Massive gabbro
257
296a
297b
0.97
0.71
0.91
0.05
0.12
5.41
5.76
6.42
0.24 6.17
0.17
0.04
0.03
5.32 0.08
0.02
0.57
5.08 19.88
6.05
0.26
0.06
0.45
5.31
5.05
6.20
0.11
11.16 11.06 10.88 11.08
15.48 16.08 16.49 15.84
0.93
54.45 53.26 46.22 54.67
292c 295
3
4
5a
5b
10 Basalt
11
88
220
1.49
2.22
5.05
4.38
5.05
6.42
1.21
1.20
0.15
0.16
5.31
7.23
4.95
0.27
0.18
0.56
5.94
7.23
4.41
0.26
0.26
0.27
4.13
8.47
5.85
0.20
0.35
0.68
3.18
8.19
4.62
0.39
0.40
0.83
3.40
7.97
4.24
0.26
0.08
0.58
3.33
7.78
5.41
0.24
0.19
0.71
2.80
7.50
4.86
0.23
12.11 12.11 12.68 19.57 18.43 16.95 20.69
0.22
0.17
0.19
5.44
2.71
0.09
0.11
2.98
6.52
2.43 14.15
0.15
7.14 10.25
12.23 15.87 13.99 11.50 12.66 11.22 10.84 13.41 15.60
3.01
54.62 52.91 52.17 45.45 46.64 47.72 44.59 66.10 47.53
2
TABLE A1A. WHOLE-ROCK CHEMICAL ANALYSES
54.57 54.24
250
6.35
17.92 18.00 11.67
0.50
48.97 52.65 59.70
103
274
2.02
5.28
0.14
0.16
0.17
3.57
0.41
1.03
4.87
10.23 11.42
7.75
0.20
13.06 11.16
14.02 14.82
2.01
49.51 48.21
272
<0.1
<0.2
0.1 0.1 0.0 <0.1
<0.2
0.0 0.1 0.0
<0.01 <0.01 0.1 0.1 0.0 0.0
0.1
0.9
0.9 0.9 0.1
0.2 1.2 0.3
0.1
1.1
1.1 1.2 0.2
0.2 1.7 0.4
0.1
1.5
1.0 1.0 0.2
0.2 1.5 0.3
0.1
0.9
0.9 1.0 0.2
0.2 1.3 0.3
33 224 120 20 42 2.4 134.5 7.8 27 1.0 0.1 121.5 0.5 2.4 0.3 1.6 0.6 0.3 0.9
0.2
0.8
0.8 0.9 0.2
0.2 1.2 0.3
35 226 120 21 45 2.6 140 7.9 21 1.1 0.2 121.5 0.6 2.3 0.3 1.6 0.6 0.3 0.9
0.1
0.5
1.3 1.4 0.2
0.3 1.9 0.5
0.2
1.5
2.3 2.3 0.4
0.5 3.6 0.8
0.2
1.6
2.5 2.4 0.4
0.5 3.7 0.8
0.1
1.4
1.9 1.8 0.3
0.4 2.8 0.6
0.2
1.6
2.3 2.4 0.4
0.5 3.5 0.8
127 24 31 24 23 277 449 419 383 386 120 20 80 30 20 12 84 <5 132 65 62 119 80 77 107 18.2 1.2 0.5 0.4 2.4 501 67.4 52.8 46.7 71.1 10.8 19.5 20.9 15.9 19.8 14 44 47 44 46 0.5 1.5 1.8 1.1 1.5 0.3 0.1 0.0 0.2 0.1 150 28.3 6.0 7.5 66.8 0.5 2.5 2.8 2.1 2.4 2.3 6.5 7.0 5.6 6.5 0.4 1.0 1.1 0.9 1.1 2.3 5.3 5.6 4.5 5.4 0.9 1.9 2.0 1.5 1.9 0.6 0.8 0.8 0.6 0.8 1.4 2.5 2.7 2.0 2.5
0.5
6.4
8.7 8.3 1.3
2.2 14.3 3.0
85 339 <10 18 71 2.0 110 83.4 219 5.7 0.1 21.7 7.7 26.9 4.4 23.7 8.7 2.2 10.5
0.3
2.8
4.4 3.9 0.6
1.1 7.2 1.6
63 255 250 80 125 15.8 553 51.0 91 3.9 0.4 68.3 5.0 13.2 2.1 11.3 3.8 1.5 5.4
0.3
10.7
7.1 6.7 1.1
1.8 11.5 2.5
58 266 10 35 122 4.8 143 69.0 418 3.9 0.3 28.4 8.2 27.2 4.3 22.4 7.7 2.9 9.3
0.4
2.2
4.6 4.2 0.7
1.3 7.8 1.7
23 510 <10 31 216 14.8 230 43.8 64 4.4 0.1 39.7 2.5 9.9 1.9 11.6 4.7 2.2 6.2
0.3
1.9
4.8 4.2 0.6
1.2 8.2 1.7
18 348 <10 42 193 18.9 279 46.2 55 4.0 0.1 46.7 2.9 11.2 2.1 12.9 5.1 2.3 6.7
0.3
3.0
4.5 4.5 0.6
1.0 7.0 1.5
0.3
3.3
4.8 4.5 0.7
1.2 8.2 1.8
0.2
2.4
3.4 3.2 0.5
0.8 5.3 1.2
0.3
2.4
2.4 2.4 0.3
0.6 4.0 0.9
19 20 <5 85 619 1160 75 214 <10 <10 <10 140 64 63 46 <5 176 182 100 74 22.5 31.1 2.3 1.6 129.5 144 210 161.5 41.7 46.6 33.5 24.1 99 112 78 80 3.9 3.9 2.3 3.7 0.5 0.7 0.4 0.4 20.7 30.6 72.3 58.6 2.2 3.8 3.1 3.7 9.3 13.1 9.9 10.4 1.6 2.3 1.7 1.6 9.7 13.3 8.9 8.1 3.9 4.7 3.1 2.8 1.5 1.8 1.2 1.0 4.9 6.3 4.0 3.3
1.1
2.9
2.0 1.7 0.3
0.7 4.0 0.8
0.5
3.8
3.8 3.7 0.6
1.0 6.1 1.3
115 34 370 264 360 50 84 64 124 128 1.6 5.2 341 73.6 19.2 33.4 97 123 13.1 4.8 0.2 0.0 52.4 183.5 12.0 13.7 25.5 29.3 3.4 4.0 15.0 17.7 3.8 4.6 1.5 1.5 4.1 5.2
98.58 97.89 100.41 98.66 98.70 100.14 99.74 100.43 100.03 98.14 101.21 99.67 100.02 100.96 100.25 98.97 99.21 98.35 98.82 98.95 98.65 100.68 99.35
0.00
0.00
0.00
0.14
42.55 43.35
0.12
8.70
1.53
0.02
45.52 44.78
270
1960 2170 2160 2210 33 26 41 9.0 <5 <5 <5 445 659 133 2450 3230 2820 3740 20 <10 180 21 8 15 5 43 <5 <5 47 65 49 45 65 49 44 1.1 0.2 0.2 0.4 16.0 10.2 0.7 0.7 1.4 4.2 13.1 141.5 139.5 39.8 0.5 1.6 0.6 <0.5 7.8 10.3 9.4 5 <2 <2 <2 26 29 42 <0.2 <0.2 <0.2 <0.2 1.0 0.6 0.8 0.4 0.1 0.1 0.0 0.3 0.3 0.2 4.4 3.7 20.4 4.9 111.5 80.8 9.2 <0.5 <0.5 <0.5 <0.5 0.9 1.0 2.4 <0.5 <0.5 <0.5 <0.5 3.0 3.1 7.0 <0.03 <0.03 <0.03 <0.03 0.4 0.4 0.9 <0.1 0.1 0.1 0.1 2.0 2.0 3.9 <0.03 0.1 0.0 0.0 0.7 0.7 1.1 <0.03 0.1 <0.03 <0.03 0.3 0.3 0.5 <0.05 0.1 0.1 <0.05 0.9 1.0 1.3
Tb Dy Ho
Ni V Cr Cu Zn Rb Sr Y Zr Nb Cs Ba La Ce Pr Nd Sm Eu Gd
97.40
0.01
0.00
0.00
0.06
42.23
0.15
9.03
2.45
43.42
45.01
TiO2
(ppm)
266
Peridotite
265
SiO2
Sample 131
Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili 161
162
Dangerfield et al.
TABLE A1B. WHOLE ROCK CHEMICAL ANALYSES Sample
16
19
20
21
142
114
120
126b
127
128
Alkaline Basalt 54.75
46.65
2.06 0.40
1.56
0.71
Al2O3
12.65 13.67 13.57 13.38 14.69 13.27
14.58
16.80
Fe2O3
11.60 11.57 10.40 11.40 10.95 9.05
SiO2 TiO2
47.77 49.47 50.32 50.05 49.54 57.37 2.31
2.51
2.44
2.35
129
203
210
212
215
248
293
293c
Sheeted dykes 54.54 52.99 53.21 57.00 53.77 1.69
1.01
0.76
0.69
293d
280
282
Rhyolitic rocks 57.97
59.02 74.90 73.57 72.68 71.79 70.55 64.40
1.06
0.86
14.72 16.14 15.94 14.85 15.87
15.48
15.59 11.97 11.28 11.07 11.55 11.91 13.96
1.47
0.29
0.80
0.85
0.83
0.65
0.84
14.02 10.50
13.35
8.66
9.84
8.18 11.54
10.88
11.05
4.45
5.41
5.03
6.32
5.62
8.25
MnO
0.15
0.17
0.22
0.12
0.18 0.20
0.23
0.08
0.29
0.16
0.14
0.18
0.16
0.21
0.16
0.00
0.08
0.08
0.09
0.12
0.27
MgO
11.06
7.58
6.85
6.70
5.46 7.35
4.50
5.39
4.55
6.08
7.78
5.62
7.76
5.43
2.65
0.45
1.22
1.33
1.29
1.96
3.00
CaO
9.11
8.16
7.99
8.24 11.35 5.76
6.20
22.27
5.00
8.91
6.78
6.91
4.15
3.79
5.72
0.87
1.65
1.31
1.89
4.04
4.78
Na2O
2.09
3.50
3.52
3.13
4.80 5.75
4.79
0.46
4.76
3.98
4.33
6.13
5.16
5.86
3.77
5.64
4.94
4.82
4.64
3.40
4.94
K2O
2.37
2.12
2.51
3.02
1.08 0.12
0.21
0.12
0.31
0.05
0.59
0.20
0.07
0.08
0.44
0.24
0.38
0.47
0.63
1.59
0.17
P2O5
0.31
0.31
0.30
0.27
0.41 0.05
0.12
0.08
0.14
0.08
0.05
0.07
0.05
0.04
0.20
0.04
0.15
0.18
0.17
0.10
0.12
Total
99.41 99.06 98.12 98.65 100.52 99.30 100.28 101.21 100.01 99.89 99.41 99.83 99.59 100.59 100.05 98.85 99.48 97.82 99.20 99.95 100.73
(ppm) Ni
164
64
57
68
100
88
<5
38
<5
34
47
29
23
15
<5
<5
<5
<5
<5
18
18
V
242
271
263
268
221
247
402
311
417
231
269
245
287
325
147
14
20
48
40
183
240
Cr
330
40
40
70
210
270
<10
40
<10
100
60
20
20
10
<10
<10
<10
<10
<10
30
20
Cu
86
107
98
93
66
<5
8
343
5
40
51
12
27
10
27
8
5
<5
<5
48
44
Zn
114
130
124
133
134
50
92
42
100
75
85
62
73
36
115
30
101
98
97
84
118
Rb Sr Y Zr
38.4 466 23.0 196
32.3 715 24.0 182
40.0 575 23.9 187
45.9 512 22.3 163
14.3
0.5
2.8
1.1
25.9
45.3
43.7
58.1 165.5 193
24.6 31.5
23.2
17.7
23.6
24.3
60
48
61
91
502 214
29
3.9
0.2
4.8
1.5
0.6
0.4
116.0
90.2
52.9
16.9
19.1
22.7
19.4
37
60
48
39
4.0 193 40.9 135
1.6 77 34.4 191
8.0
9.9
87.3 128.5 35.7 158
31.9 164
14.9
45.9
3.8
121
71.4
76.1
36.5
26.3
26.0
96
84
159
Nb
48.1
41.6
41.4
36.4
55.2
1.4
1.6
1.8
1.8
1.4
1.1
3.7
1.6
0.9
4.3
4.9
3.9
3.9
3.8
4.6
Cs
0.1
0.8
0.1
0.1
0.5
0.0
0.1
0.2
0.1
0.0
0.2
0.1
0.1
0.0
0.2
0.1
0.2
0.5
0.7
1.9
8.4
18.7
5.1
23.9
7.7
65.7
17.5
10.8
16.8
75.6
17.7
26.8
37.9
Ba
610 1105
952
1130
339
39.9 123
2.5 0.3 37.6
La
37.7
31.5
32.0
28.7
42.7
2.1
2.7
2.3
2.6
2.7
1.7
4.8
2.1
1.1
6.8
6.5
7.3
4.7
5.7
12.9
6.7
Ce
74.2
63.2
64.9
57.4
81.4
6.8
7.9
6.7
8.0
8.9
5.4
10.7
6.9
3.9
18.7
15.9
20.7
14.1
16.9
27.5
16.3
Pr
8.2
7.3
7.2
6.6
8.5
1.1
1.2
1.0
1.2
1.4
0.8
1.3
1.0
0.6
2.7
2.3
3.3
2.4
2.8
3.6
2.3
Nd
32.3
29.0
28.9
26.5
32.6
5.9
6.4
4.6
6.5
7.3
4.2
6.1
5.6
3.3
13.8
11.5
16.3
12.8
14.4
15.4
11.0
Sm
6.3
6.2
6.1
5.7
6.3
2.6
2.3
1.5
2.1
2.6
1.5
1.8
2.1
1.5
4.6
4.0
5.0
4.3
4.8
3.9
3.3
Eu
1.9
1.9
1.9
1.8
2.0
0.4
1.0
0.6
0.9
1.1
0.6
0.7
0.7
0.5
1.6
1.5
1.9
1.5
1.7
1.2
1.1
Gd
6.3
6.1
6.0
5.6
6.3
3.4
2.9
2.2
2.9
3.1
2.0
2.4
2.5
2.2
5.6
4.6
5.9
5.2
5.7
4.3
3.6
Tb
0.9
0.9
0.9
0.8
0.9
0.7
0.6
0.4
0.6
0.6
0.4
0.4
0.5
0.5
1.1
0.9
1.2
1.0
1.1
0.8
0.7
Dy
4.4
4.7
4.7
4.3
4.6
5.2
3.8
2.9
4.0
4.1
2.8
3.0
3.5
3.1
6.9
5.8
7.1
6.3
6.9
4.8
4.7
Ho
0.9
0.9
0.9
0.8
0.9
1.1
0.8
0.6
0.9
0.9
0.6
0.7
0.8
0.7
1.5
1.2
1.5
1.3
1.5
1.1
1.1
Er
2.2
2.4
2.4
2.2
2.2
3.5
2.4
1.8
2.6
2.5
1.8
1.9
2.5
2.0
4.4
3.7
4.1
3.7
4.3
3.0
3.2
Yb
1.8
1.9
1.7
1.8
1.9
3.9
2.3
1.8
2.4
2.6
1.8
2.0
2.6
2.2
4.2
3.9
3.9
3.1
4.3
3.0
3.3
Lu
0.3
0.3
0.3
0.3
0.3
0.6
0.3
0.3
0.4
0.4
0.3
0.3
0.4
0.3
0.7
0.6
0.6
0.4
0.7
0.5
0.5
Hf
5.1
4.7
4.8
4.4
5.0
1.2
1.9
1.6
2.1
2.6
1.3
1.7
1.7
1.5
4.0
5.8
4.7
5.0
4.9
3.0
2.8
Ta
3.0
2.5
2.6
2.3
3.5
0.1
0.1
0.1
0.2
0.1
0.1
0.3
0.1
0.1
0.4
0.5
0.5
0.5
0.5
0.5
0.2
Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili
163
TABLE A2A. Cr-SPINEL MINERAL ANALYSES Sample
270
270
270
271
271
271
271
271
271
271
271
271
271
271
271
271
271
271
271
271
SiO2
0.07
0.03
0.13
0.07
0.11
0.06
0.20
0.07
0.08
0.05
0.06
0.04
0.11
0.06
0.01
0.02
0.06
0.05
0.04
0.01
0.01
0.01
0.39
0.32
0.23
0.15
0.00
0.03
0.00
0.00
0.02
0.01
0.03
TiO2
0.02
0.31
0.48
0.19
0.22
0.14
0.13
Al2O3
25.02 28.34 27.23
18.47
21.13
21.29 20.17
18.13
24.74
24.04
17.70 22.63 17.64 22.41 23.92 19.45 20.58 21.69 21.97 22.83
Cr2O3
45.64 43.88 43.49
51.73
49.77
49.86 50.97
52.63
47.41
47.86
52.49 49.28 51.27 47.93 46.06 50.34 49.47 48.46 47.20 47.62
FeO
13.96 14.60 15.28
16.64
14.77
14.99 17.02
17.14
13.15
13.97
15.86 13.90 12.65 16.24 14.79 15.91 15.96 14.79 15.29 13.22
Fe2O3
1.72
0.00
0.00
0.51
0.00
0.73
0.00
0.18
0.34
0.00
1.70
0.67
4.18
0.00
1.38
1.20
0.20
0.97
1.59
0.99
NiO
0.10
0.13
0.00
0.00
0.16
0.15
0.10
0.16
0.08
0.17
0.04
0.00
0.09
0.14
0.07
0.02
0.02
0.07
0.06
0.08
MnO
0.49
0.47
0.47
0.64
0.52
0.50
0.57
0.58
0.52
0.57
0.64
0.54
0.63
0.59
0.56
0.60
0.63
0.61
0.53
0.55
14.33 13.75 13.19
11.77
13.42
13.52 10.94
11.45
14.91
13.46
MgO Total
12.41 14.31 14.32 12.11 13.43 12.07 12.00 13.05 12.73 14.21
101.32 101.21 99.81 100.13 100.35 101.48 100.17 100.56 101.36 100.23 101.22 101.58 101.04 99.48 100.25 99.60 98.91 99.70 99.41 99.54
FeO*
13.37 12.84 12.33
11.22
12.59
12.67 10.41
10.89
13.93
12.68
11.80 13.41 13.51 11.48 12.64 11.46 11.42 12.35 11.98 13.33
Calculated on 32 oxygens Si
0.02
0.01
0.03
0.02
0.03
0.01
0.05
0.02
0.02
0.01
0.02
0.01
0.03
0.01
0.00
0.01
0.02
0.01
0.01
0.00
Ti
0.00
0.00
0.00
0.06
0.09
0.07
0.04
0.04
0.03
0.02
0.06
0.04
0.03
0.00
0.01
0.00
0.00
0.00
0.00
0.01
Al
7.04
7.89
7.73
5.48
6.12
6.10
5.95
5.38
6.94
6.87
5.20
6.42
5.13
6.56
6.86
5.76
6.10
6.32
6.43
6.58
Cr
8.62
8.20
8.28
10.29
9.67
9.59 10.08
10.47
8.92
9.18
10.34
9.37 10.00
9.41
8.86 10.01
9.84
9.47
9.26
9.21
Fe2+
2.79
2.89
3.08
3.50
3.03
3.05
3.56
3.61
2.62
2.84
3.31
2.80
2.61
3.37
3.01
3.35
3.36
3.06
3.17
2.71
Fe3+
0.31
0.00
0.00
0.10
0.00
0.13
0.00
0.03
0.06
0.00
0.32
0.12
0.78
0.00
0.25
0.23
0.04
0.18
0.30
0.18
Ni
0.02
0.03
0.00
0.00
0.03
0.03
0.02
0.03
0.02
0.03
0.01
0.00
0.02
0.03
0.01
0.00
0.00
0.01
0.01
0.02
Mn
0.10
0.09
0.10
0.14
0.11
0.10
0.12
0.12
0.10
0.12
0.13
0.11
0.13
0.12
0.12
0.13
0.13
0.13
0.11
0.11
Mg
5.10
4.85
4.74
4.41
4.92
4.90
4.08
4.30
5.29
4.87
4.61
5.13
5.27
4.49
4.87
4.52
4.50
4.81
4.71
5.18
Cr# [Cr/(Cr+Al)]
0.55
0.51
0.52
0.65
0.61
0.61
0.63
0.66
0.56
0.57
0.67
0.59
0.66
0.59
0.56
0.63
0.62
0.60
0.59
0.58
Mg# 0.65 [Mg/(Mg/Fe2+)]
0.63
0.61
0.56
0.62
0.62
0.53
0.54
0.67
0.63
0.58
0.65
0.67
0.57
0.62
0.57
0.57
0.61
0.60
0.66
Sample
271
271
271
271
271
271
271
271
271
271
271
850
850
850
850
850
850
850
850
850
SiO2
1.24
0.06
0.01
0.05
0.03
0.02
0.02
0.00
0.06
0.04
0.06
0.01
0.03
0.00
0.02
0.00
0.02
0.03
0.04
0.02
0.02
0.03
0.04
0.03
0.02
0.03
0.04
0.02
0.04
0.03
0.04
0.02
0.05
0.06
0.02
0.05
TiO2
0.01
0.01
0.04
0.04
Al2O3
22.82 20.00
24.34
24.04 23.97 24.16 24.47 24.78 23.35 23.89 24.65
17.37
17.25 17.55 16.29 15.73 15.47 16.13 16.81 16.19
Cr2O3
45.18 50.11
46.74
47.06 46.18 45.97 46.07 45.55 47.51 46.47 46.11
49.55
49.91 49.33 51.34 52.22 52.48 52.06 51.30 51.73
FeO
14.51 15.83
13.30
13.79 13.98 13.95 13.73 13.75 13.34 13.19 13.24
19.55
18.91 18.71 18.13 18.17 18.18 18.42 19.15 18.76
Fe2O3
2.30
1.08
1.14
0.55
0.49
0.17
0.70
0.25
1.06
1.17
1.08
3.64
3.45
3.35
2.86
2.80
2.90
2.40
2.14
2.77
NiO
0.05
0.08
0.06
0.15
0.12
0.10
0.09
0.14
0.11
0.05
0.12
0.10
0.06
0.02
0.10
0.03
0.08
0.03
0.02
0.02
MnO
0.55
0.57
0.52
0.55
0.60
0.57
0.52
0.00
0.53
0.51
0.52
0.63
0.64
0.61
0.61
0.65
0.68
0.66
0.68
0.68
MgO
13.14 12.27
14.55
14.10 13.70 13.67 14.09 14.26 14.35 14.40 14.57
9.65
10.00 10.05 10.31 10.23 10.23 10.14
9.69
9.96
Total
99.81 100.03 100.67 100.31 99.11 98.64 99.70 98.77 100.34 99.74 100.39 100.52 100.29 99.65 99.70 99.85 100.08 99.94 99.85 100.18
FeO*
12.38 11.62
13.62
13.24 12.93 12.87 13.19 12.83 13.44 13.46 13.63
9.31
9.64
9.65
9.88
9.86
9.88
9.78
9.39
9.64
Calculated on 32 oxygens Si
0.30
0.02
0.00
0.01
0.01
0.01
0.01
0.00
0.01
0.01
0.02
0.00
0.01
0.00
0.01
0.00
0.01
0.01
0.01
0.01
Ti
0.00
0.00
0.00
0.01
0.01
0.01
0.00
0.01
0.01
0.00
0.01
0.01
0.01
0.01
0.01
0.00
0.01
0.01
0.01
0.01
Al
6.56
5.88
6.90
6.86
6.93
7.00
7.00
7.13
6.67
6.84
6.99
5.24
5.20
5.32
4.95
4.79
4.70
4.90
5.11
4.91
Cr
8.71
9.89
8.89
9.01
8.95
8.94
8.85
8.80
9.10
8.92
8.77
10.03
Fe2+
2.96
3.30
2.68
2.79
2.87
2.87
2.79
2.81
2.70
2.68
2.66
4.19
4.05
4.02
3.91
3.92
3.92
3.97
4.13
4.04
Fe3+
0.42
0.20
0.21
0.10
0.09
0.03
0.13
0.05
0.19
0.21
0.20
0.70
0.66
0.65
0.56
0.54
0.56
0.47
0.42
0.54
Ni
0.01
0.02
0.01
0.03
0.02
0.02
0.02
0.03
0.02
0.01
0.02
0.02
0.01
0.00
0.02
0.01
0.02
0.01
0.00
0.00
Mn
0.11
0.12
0.11
0.11
0.13
0.12
0.11
0.00
0.11
0.10
0.11
0.14
0.14
0.13
0.13
0.14
0.15
0.14
0.15
0.15
Mg
4.78
4.57
5.22
5.09
5.01
5.01
5.10
5.19
5.18
5.21
5.23
3.68
3.82
3.85
3.96
3.94
3.93
3.90
3.72
3.82
Cr# [Cr/(Cr+Al)]
0.57
0.63
0.56
0.57
0.56
0.56
0.56
0.55
0.58
0.57
0.56
0.66
0.66
0.65
0.68
0.69
0.69
0.68
0.67
0.68
Mg# [Mg/(Mg/Fe2+)]
0.62
0.58
0.66
0.65
0.64
0.64
0.65
0.65
0.66
0.66
0.66
0.47
0.49
0.49
0.50
0.50
0.50
0.50
0.47
0.49
10.10 10.02 10.46 10.66 10.70 10.60 10.46 10.53
164
Dangerfield et al. TABLE A2B. Cr-SPINEL ANALYSES
Sample
270
270
270
271
271
271
271
271
271
271
271
271
271
271
271
271
271
271
271
271
SiO2
0.07
0.03
0.13
0.07
0.11
0.06
0.20
0.07
0.08
0.05
0.06
0.04
0.11
0.06
0.01
0.02
0.06
0.05
0.04
0.01
0.01
0.01
0.39
0.32
0.23
0.15
0.00
0.03
0.00
0.00
0.02
0.01
0.03
TiO2
0.02
0.31
0.48
0.19
0.22
0.14
0.13
Al2O3
25.02 28.34 27.23
18.47
21.13
21.29 20.17
18.13
24.74
24.04
17.70 22.63 17.64 22.41 23.92 19.45 20.58 21.69 21.97 22.83
Cr2O3
45.64 43.88 43.49
51.73
49.77
49.86 50.97
52.63
47.41
47.86
52.49 49.28 51.27 47.93 46.06 50.34 49.47 48.46 47.20 47.62
FeO
13.96 14.60 15.28
16.64
14.77
14.99 17.02
17.14
13.15
13.97
15.86 13.90 12.65 16.24 14.79 15.91 15.96 14.79 15.29 13.22
Fe2O3
1.72
0.00
0.00
0.51
0.00
0.73
0.00
0.18
0.34
0.00
1.70
0.67
4.18
0.00
1.38
1.20
0.20
0.97
1.59
0.99
NiO
0.10
0.13
0.00
0.00
0.16
0.15
0.10
0.16
0.08
0.17
0.04
0.00
0.09
0.14
0.07
0.02
0.02
0.07
0.06
0.08
MnO
0.49
0.47
0.47
0.64
0.52
0.50
0.57
0.58
0.52
0.57
0.64
0.54
0.63
0.59
0.56
0.60
0.63
0.61
0.53
0.55
14.33 13.75 13.19
11.77
13.42
13.52 10.94
11.45
14.91
13.46
MgO Total
12.41 14.31 14.32 12.11 13.43 12.07 12.00 13.05 12.73 14.21
101.32 101.21 99.81 100.13 100.35 101.48 100.17 100.56 101.36 100.23 101.22 101.58 101.04 99.48 100.25 99.60 98.91 99.70 99.41 99.54
FeO*
13.37 12.84 12.33
11.22
12.59
12.67 10.41
10.89
13.93
12.68
11.80 13.41 13.51 11.48 12.64 11.46 11.42 12.35 11.98 13.33
Calculated on 32 oxygens Si
0.02
0.01
0.03
0.02
0.03
0.01
0.05
0.02
0.02
0.01
0.02
0.01
0.03
0.01
0.00
0.01
0.02
0.01
0.01
0.00
Ti
0.00
0.00
0.00
0.06
0.09
0.07
0.04
0.04
0.03
0.02
0.06
0.04
0.03
0.00
0.01
0.00
0.00
0.00
0.00
0.01
Al
7.04
7.89
7.73
5.48
6.12
6.10
5.95
5.38
6.94
6.87
5.20
6.42
5.13
6.56
6.86
5.76
6.10
6.32
6.43
6.58
Cr
8.62
8.20
8.28
10.29
9.67
9.59 10.08
10.47
8.92
9.18
10.34
9.37 10.00
9.41
8.86 10.01
9.84
9.47
9.26
9.21
Fe2+
2.79
2.89
3.08
3.50
3.03
3.05
3.56
3.61
2.62
2.84
3.31
2.80
2.61
3.37
3.01
3.35
3.36
3.06
3.17
2.71
Fe3+
0.31
0.00
0.00
0.10
0.00
0.13
0.00
0.03
0.06
0.00
0.32
0.12
0.78
0.00
0.25
0.23
0.04
0.18
0.30
0.18
Ni
0.02
0.03
0.00
0.00
0.03
0.03
0.02
0.03
0.02
0.03
0.01
0.00
0.02
0.03
0.01
0.00
0.00
0.01
0.01
0.02
Mn
0.10
0.09
0.10
0.14
0.11
0.10
0.12
0.12
0.10
0.12
0.13
0.11
0.13
0.12
0.12
0.13
0.13
0.13
0.11
0.11
Mg
5.10
4.85
4.74
4.41
4.92
4.90
4.08
4.30
5.29
4.87
4.61
5.13
5.27
4.49
4.87
4.52
4.50
4.81
4.71
5.18
Cr# [Cr/(Cr+Al)]
0.55
0.51
0.52
0.65
0.61
0.61
0.63
0.66
0.56
0.57
0.67
0.59
0.66
0.59
0.56
0.63
0.62
0.60
0.59
0.58
Mg# 0.65 [Mg/(Mg/Fe2+)]
0.63
0.61
0.56
0.62
0.62
0.53
0.54
0.67
0.63
0.58
0.65
0.67
0.57
0.62
0.57
0.57
0.61
0.60
0.66
Sample
271
271
271
271
271
271
271
271
271
271
271
850
850
850
850
850
850
850
850
850
SiO2
1.24
0.06
0.01
0.05
0.03
0.02
0.02
0.00
0.06
0.04
0.06
0.01
0.03
0.00
0.02
0.00
0.02
0.03
0.04
0.02
0.02
0.03
0.04
0.03
0.02
0.03
0.04
0.02
0.04
0.03
0.04
0.02
0.05
0.06
0.02
0.05
0.01
0.01
0.04
0.04
Al2O3
22.82 20.00
24.34
24.04 23.97 24.16 24.47 24.78 23.35 23.89 24.65
17.37
17.25 17.55 16.29 15.73 15.47 16.13 16.81 16.19
Cr2O3
45.18 50.11
46.74
47.06 46.18 45.97 46.07 45.55 47.51 46.47 46.11
49.55
49.91 49.33 51.34 52.22 52.48 52.06 51.30 51.73
FeO
14.51 15.83
13.30
13.79 13.98 13.95 13.73 13.75 13.34 13.19 13.24
19.55
18.91 18.71 18.13 18.17 18.18 18.42 19.15 18.76
TiO2
Fe2O3
2.30
1.08
1.14
0.55
0.49
0.17
0.70
0.25
1.06
1.17
1.08
3.64
3.45
3.35
2.86
2.80
2.90
2.40
2.14
2.77
NiO
0.05
0.08
0.06
0.15
0.12
0.10
0.09
0.14
0.11
0.05
0.12
0.10
0.06
0.02
0.10
0.03
0.08
0.03
0.02
0.02
MnO
0.55
0.57
0.52
0.55
0.60
0.57
0.52
0.00
0.53
0.51
0.52
0.63
0.64
0.61
0.61
0.65
0.68
0.66
0.68
0.68
MgO
13.14 12.27
14.55
14.10 13.70 13.67 14.09 14.26 14.35 14.40 14.57
9.65
10.00 10.05 10.31 10.23 10.23 10.14
9.69
9.96
Total
99.81 100.03 100.67 100.31 99.11 98.64 99.70 98.77 100.34 99.74 100.39 100.52 100.29 99.65 99.70 99.85 100.08 99.94 99.85 100.18
FeO*
12.38 11.62
13.62
13.24 12.93 12.87 13.19 12.83 13.44 13.46 13.63
9.31
9.64
9.65
9.88
9.86
9.88
9.78
9.39
9.64
Calculated on 32 oxygens Si
0.30
0.02
0.00
0.01
0.01
0.01
0.01
0.00
0.01
0.01
0.02
0.00
0.01
0.00
0.01
0.00
0.01
0.01
0.01
0.01
Ti
0.00
0.00
0.00
0.01
0.01
0.01
0.00
0.01
0.01
0.00
0.01
0.01
0.01
0.01
0.01
0.00
0.01
0.01
0.01
0.01
Al
6.56
5.88
6.90
6.86
6.93
7.00
7.00
7.13
6.67
6.84
6.99
5.24
5.20
5.32
4.95
4.79
4.70
4.90
5.11
4.91
Cr
8.71
9.89
8.89
9.01
8.95
8.94
8.85
8.80
9.10
8.92
8.77
10.03
Fe2+
2.96
3.30
2.68
2.79
2.87
2.87
2.79
2.81
2.70
2.68
2.66
4.19
4.05
4.02
3.91
3.92
3.92
3.97
4.13
4.04
Fe3+
0.42
0.20
0.21
0.10
0.09
0.03
0.13
0.05
0.19
0.21
0.20
0.70
0.66
0.65
0.56
0.54
0.56
0.47
0.42
0.54
Ni
0.01
0.02
0.01
0.03
0.02
0.02
0.02
0.03
0.02
0.01
0.02
0.02
0.01
0.00
0.02
0.01
0.02
0.01
0.00
0.00
Mn
0.11
0.12
0.11
0.11
0.13
0.12
0.11
0.00
0.11
0.10
0.11
0.14
0.14
0.13
0.13
0.14
0.15
0.14
0.15
0.15
Mg
4.78
4.57
5.22
5.09
5.01
5.01
5.10
5.19
5.18
5.21
5.23
3.68
3.82
3.85
3.96
3.94
3.93
3.90
3.72
3.82
Cr# [Cr/(Cr+Al)]
0.57
0.63
0.56
0.57
0.56
0.56
0.56
0.55
0.58
0.57
0.56
0.66
0.66
0.65
0.68
0.69
0.69
0.68
0.67
0.68
Mg# [Mg/(Mg/Fe2+)]
0.62
0.58
0.66
0.65
0.64
0.64
0.65
0.65
0.66
0.66
0.66
0.47
0.49
0.49
0.50
0.50
0.50
0.50
0.47
0.49
10.10 10.02 10.46 10.66 10.70 10.60 10.46 10.53
Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili
165
TABLE A3A. GEOCHRONOLOGICAL ZIRCON ANALYSES Isotope ratios #
U (ppm)
206
Pb 204 Pb
U/Th
Sample AD012 Karadag Fm 2 66 1136 1.0 3 52 828 1.0 4 51 804 1.0 5 95 1228 0.7 6 79 980 1.0 7 472 10,136 1.5 8 552 10,496 1.2 9 128 7416 1.8 10 83 2284 2.1 11 67 1092 1.0 12 55 876 1.2 13 97 1596 0.8 14 1159 15,476 1.7 15 67 872 1.1 17 42 1064 1.5 18 221 2936 0.5 19 56 876 1.0 20 51 660 1.2 21 124 1524 0.4 22 120 1456 0.6 23 70 1404 1.1 24 63 408 0.9 25 57 760 0.8 27 237 2708 0.6 28 162 2108 0.7 29 89 712 0.7 31 64 772 1.0 32 141 2804 0.9 33 323 2560 0.4 35 98 680 0.7 36 155 1856 0.8 37 246 2292 0.6 38 174 3116 1.2 39 65 996 0.9 40 57 932 0.9 41 82 1256 0.9 42 52 756 0.8 43 70 1028 0.8 44 67 1048 0.8 45 128 1864 0.8 46 186 3592 1.1 47 43 484 1.4 48 499 4048 0.4 49 93 1448 0.6 50 79 784 1.0 51 49 712 1.0 52 268 3716 0.5
206
Pb* 207 Pb*
± (%)
17.2828 16.1761 24.4570 25.1848 22.6647 17.6135 20.5296 18.3564 20.4875 22.7613 22.6311 21.3839 20.1553 17.1830 20.7232 21.0379 25.6751 20.4148 22.4847 20.0308 18.7970 18.9393 19.0061 18.5119 23.0425 15.2699 30.8811 21.8494 17.1379 14.9228 21.1470 20.9586 21.3721 20.3217 21.4935 20.8703 23.2365 27.4376 19.8264 22.9935 20.2601 20.3063 20.9745 26.8236 17.9530 23.9027 20.8364
30.1 36.6 22.2 21.4 17.4 6.7 1.8 3.8 10.5 10.4 28.1 12.3 1.7 45.8 40.4 4.7 23.3 27.5 18.8 13.7 17.6 11.9 11.8 10.9 12.1 23.2 36.0 8.2 18.9 17.9 7.9 5.7 8.7 15.7 31.1 16.7 27.7 27.9 32.6 12.2 8.8 35.6 3.5 29.7 21.1 16.6 4.4
207
Pb* 235 U*
± (%)
0.1573 0.1598 0.1155 0.1122 0.1246 0.5104 0.1557 0.5382 0.1664 0.1223 0.1225 0.1303 0.1603 0.1496 0.1425 0.1327 0.1106 0.1441 0.1197 0.1416 0.1400 0.1494 0.1500 0.1533 0.1309 0.1991 0.0913 0.1559 0.1636 0.2048 0.1265 0.1339 0.1361 0.1338 0.1322 0.1351 0.1196 0.0992 0.1408 0.1206 0.1743 0.1121 0.1383 0.1141 0.1537 0.1237 0.1540
30.2 36.8 22.2 21.5 17.5 7.1 2.4 4.7 10.7 11.4 28.3 12.4 3.0 45.9 40.4 4.9 23.4 27.6 19.0 13.8 18.1 12.3 11.9 11.1 12.2 23.2 36.2 8.6 18.9 18.0 8.3 6.2 8.8 16.7 31.2 16.8 28.2 28.2 32.7 12.5 9.0 35.6 3.8 29.7 21.2 16.7 4.7
206
Pb* 238 U
0.0197 0.0188 0.0205 0.0205 0.0205 0.0652 0.0232 0.0717 0.0247 0.0202 0.0201 0.0202 0.0234 0.0186 0.0214 0.0203 0.0206 0.0213 0.0195 0.0206 0.0191 0.0205 0.0207 0.0206 0.0219 0.0220 0.0205 0.0247 0.0203 0.0222 0.0194 0.0204 0.0211 0.0197 0.0206 0.0204 0.0201 0.0197 0.0202 0.0201 0.0256 0.0165 0.0210 0.0222 0.0200 0.0214 0.0233
Apparent ages (Ma) ± error (%) corr. 2.5 3.5 1.0 1.7 1.9 2.4 1.7 2.8 1.7 4.5 2.9 1.8 2.5 3.6 2.3 1.4 1.2 2.2 3.0 1.4 4.1 3.2 1.8 1.6 1.3 1.3 3.3 2.6 1.0 1.5 2.5 2.5 1.4 5.9 1.4 1.4 4.8 4.1 2.4 2.8 1.8 1.2 1.6 1.1 2.5 1.6 1.7
0.08 0.10 0.05 0.08 0.11 0.33 0.69 0.60 0.16 0.40 0.10 0.14 0.82 0.08 0.06 0.29 0.05 0.08 0.16 0.10 0.23 0.26 0.15 0.14 0.11 0.06 0.09 0.30 0.05 0.08 0.30 0.40 0.16 0.35 0.04 0.08 0.17 0.15 0.07 0.23 0.19 0.03 0.41 0.04 0.12 0.10 0.36
206
Pb* 238 U*
± (Ma)
125.8 119.8 130.7 130.8 130.7 407.2 147.8 446.1 157.4 128.8 128.3 129.0 149.3 119.1 136.6 129.3 131.4 136.1 124.6 131.3 121.9 130.9 131.9 131.3 139.5 140.6 130.5 157.3 129.8 141.3 123.9 129.9 134.5 125.9 131.5 130.5 128.6 126.0 129.2 128.3 163.0 105.6 134.2 141.5 127.7 136.7 148.3
3.1 4.2 1.3 2.2 2.5 9.3 2.4 12.2 2.6 5.7 3.7 2.3 3.6 4.2 3.1 1.8 1.6 2.9 3.7 1.8 4.9 4.1 2.4 2.0 1.8 1.8 4.2 4.0 1.3 2.1 3.1 3.2 1.9 7.4 1.8 1.8 6.1 5.1 3.0 3.6 2.8 1.3 2.1 1.6 3.2 2.2 2.4
207
Pb* U
± (Ma)
206
Pb* Pb*
± (Ma)
148.3 150.6 111.0 108.0 119.3 418.7 147.0 437.2 156.3 117.1 117.3 124.4 151.0 141.6 135.3 126.5 106.5 136.7 114.8 134.5 133.0 141.3 141.9 144.8 124.9 184.3 88.7 147.1 153.9 189.2 121.0 127.6 129.5 127.5 126.0 128.7 114.7 96.0 133.7 115.6 163.2 107.9 131.5 109.7 145.2 118.4 145.4
41.7 51.5 23.3 22.0 19.7 24.5 3.3 16.8 15.5 12.6 31.3 14.5 4.2 60.8 51.3 5.8 23.6 35.3 20.7 17.4 22.6 16.2 15.8 14.9 14.3 39.1 30.7 11.8 27.0 31.0 9.4 7.4 10.7 20.1 36.9 20.3 30.5 25.8 41.0 13.6 13.6 36.5 4.7 30.9 28.7 18.6 6.3
524.5 667.8 −294.6 −370.0 −103.9 482.8 133.9 390.8 138.7 −114.4 −100.3 37.2 177.0 537.2 111.8 76.1 −420.2 147.0 −84.3 191.4 337.3 320.2 312.2 371.8 −144.8 790.0 −930.1 −14.6 542.9 838.1 63.8 85.0 38.5 157.7 24.9 95.1 −165.5 −597.3 215.2 −139.4 164.8 159.5 83.2 −536.2 440.4 −236.4 98.9
673.3 809.1 571.8 560.3 430.8 148.9 41.5 84.9 248.2 257.4 702.3 294.7 40.1 1055.6 989.3 110.7 617.7 655.9 463.7 321.1 402.1 270.2 269.3 247.1 300.4 492.1 1078.9 199.6 416.0 375.3 188.1 134.5 208.5 368.2 762.7 397.9 701.8 769.6 774.2 301.9 206.7 857.1 83.0 809.9 474.0 420.9 103.1
235
207
166
Dangerfield et al.
TABLE A3B. GEOCHRONOLOGICAL ZIRCON ANALYSES 206 Pb U/Th 206Pb* U 204 207 (ppm) Pb Pb* Sample AD012 Karadag Fm 53 588 4528 0.3 19.2872 54 76 508 1.2 19.6078 55 200 4080 1.1 20.6207 57 377 2436 1.2 21.1660 59 113 984 1.0 21.0459 60 75 1040 0.9 25.9392 61 66 1084 1.0 25.4259 62 77 1060 0.9 22.5195 63 1509 12,208 1.4 20.2377 64 1789 3204 1.5 16.8013 65 68 472 0.8 22.6108 25-1 75 330 0.7 22.9893 25-2 58 549 1.0 18.0383 25-3 59 735 0.7 15.7366 25-4 131 960 1.1 17.4696 25-5 1731 4161 0.5 20.3773 25-6 188 2292 1.1 22.6023 25-7 279 2790 0.9 20.1748 25-8 699 6546 2.0 20.7225 25-9 929 8340 1.4 20.7391 25-10 731 6039 1.6 18.2844 25-11 485 4314 2.2 19.7491 25-12 132 963 0.5 18.4186 25-13 133 6114 1.5 12.1793 25-14 1066 10,971 1.6 14.0484 25-15 105 912 0.7 15.4967 25-15a 63 747 1.1 21.8308 25-16 202 2004 0.8 13.8938 25-17 111 486 0.5 21.0250 25-18 148 1488 1.0 18.5027 25-20 68 1149 0.9 16.7655 Sample AD009 Ophiolitic Mélange 1 112 42,955 4.0 12.7606 2 505 20,945 1.2 20.0645 3 233 24,905 2.7 19.6550 4 97 5005 1.3 20.8711 5 344 23,360 2.1 19.6607 6 201 9185 1.5 19.4428 7 272 9275 0.9 21.0865 8 711 28,700 1.0 20.5243 9 113 5540 2.2 21.5410 10 175 12,395 1.3 21.3972 11 110 5920 1.3 21.0325 12 525 10,380 1.1 21.0241 13 109 5445 1.5 19.7789 15 177 72,730 3.0 9.9249 16 481 8750 1.6 18.0945
#
± (%)
207
Pb* 235 U*
Isotope ratios 206 ± Pb* 238 (%) U
± error (%) corr.
206
Pb* 238 U*
Apparent ages (Ma) 207 ± Pb* ± (Ma) 235U (Ma)
12.3 63.0 7.3 5.0 13.6 22.6 21.7 11.8 5.4 16.3 28.7 35.5 16.8 37.7 24.2 3.4 10.0 8.4 8.1 1.8 24.8 4.3 11.5 3.2 12.5 15.3 19.7 8.2 25.6 7.9 24.4
0.1343 0.1493 0.1746 0.1760 0.1414 0.1039 0.1038 0.1284 0.1627 0.1951 0.1226 0.1131 0.1617 0.1880 0.1472 0.1685 0.1276 0.1428 0.1063 0.1115 0.1207 0.1910 0.1344 1.9907 0.2483 0.1683 0.1467 0.2047 0.1306 0.1503 0.1545
13.1 63.4 7.3 5.4 13.7 23.2 21.9 12.3 5.5 16.4 28.8 35.7 16.9 37.8 24.5 3.8 10.2 8.8 8.8 2.9 24.9 4.5 12.0 4.0 12.5 16.5 19.8 8.3 25.8 8.0 24.8
0.0188 0.0212 0.0261 0.0270 0.0216 0.0195 0.0191 0.0210 0.0239 0.0238 0.0201 0.0189 0.0212 0.0215 0.0186 0.0249 0.0209 0.0209 0.0160 0.0168 0.0160 0.0274 0.0180 0.1758 0.0253 0.0189 0.0232 0.0206 0.0199 0.0202 0.0188
4.6 7.6 1.0 2.1 1.2 5.2 2.8 3.4 1.0 1.4 2.3 3.4 2.3 2.2 3.6 1.7 1.6 2.6 3.5 2.3 2.5 1.5 3.3 2.5 1.0 6.3 1.3 1.0 3.3 1.6 4.3
0.35 119.9 0.12 135.5 0.14 166.1 0.38 171.9 0.09 137.6 0.22 124.8 0.13 122.3 0.28 133.8 0.19 152.1 0.08 151.5 0.08 128.3 0.10 120.5 0.13 135.0 0.06 136.9 0.15 119.1 0.44 158.5 0.16 133.5 0.30 133.3 0.39 102.2 0.78 107.2 0.10 102.3 0.33 174.0 0.27 114.7 0.62 1044.2 0.08 161.0 0.38 120.8 0.06 148.0 0.12 131.6 0.13 127.1 0.20 128.7 0.17 120.0
5.5 10.2 1.6 3.5 1.6 6.4 3.4 4.5 1.5 2.0 2.9 4.1 3.0 3.0 4.2 2.7 2.1 3.5 3.5 2.4 2.5 2.6 3.7 23.8 1.6 7.6 1.9 1.3 4.1 2.1 5.1
127.9 141.3 163.4 164.6 134.3 100.4 100.3 122.6 153.1 181.0 117.4 108.8 152.2 174.9 139.4 158.1 121.9 135.5 102.6 107.4 115.7 177.5 128.1 1112.4 225.2 157.9 139.0 189.1 124.7 142.2 145.9
1.0 1.0 2.9 8.3 4.7 6.2 3.9 3.2 9.4 6.8 11.5 4.1 7.8 1.0 6.3
2.0815 0.1680 0.3295 0.1529 0.2734 0.1716 0.1508 0.1704 0.1518 0.1548 0.1570 0.1533 0.1802 3.9799 0.1744
1.7 2.0 3.6 8.4 5.1 6.3 5.0 3.4 9.5 6.9 11.9 4.4 7.8 1.4 6.5
0.1926 0.0244 0.0470 0.0231 0.0390 0.0242 0.0231 0.0254 0.0237 0.0240 0.0239 0.0234 0.0258 0.2865 0.0229
1.4 1.8 2.1 1.7 2.0 1.0 3.0 1.0 1.3 1.2 2.9 1.7 1.0 1.0 1.6
0.81 1135.7 0.87 155.7 0.59 295.9 0.20 147.5 0.39 246.5 0.16 154.2 0.61 147.0 0.30 161.4 0.13 151.1 0.17 153.0 0.24 152.6 0.38 149.0 0.13 164.5 0.71 1623.9 0.24 145.8
14.5 1142.7 2.7 157.7 6.1 289.2 2.5 144.5 4.7 245.4 1.5 160.8 4.4 142.6 1.6 159.7 1.9 143.5 1.8 146.2 4.4 148.1 2.5 144.8 1.6 168.2 14.4 1630.1 2.2 163.2
206 207
Pb* Pb*
15.8 278.7 83.9 240.8 11.1 123.5 8.2 61.6 17.2 75.2 22.2 −447.1 20.9 −394.8 14.2 −88.1 7.9 167.4 27.1 586.1 32.0 −98.1 36.8 −139.0 24.0 429.9 60.7 726.5 31.9 500.9 5.6 151.4 11.7 −97.2 11.1 174.7 8.6 111.8 2.9 110.0 27.3 399.6 7.3 224.2 14.4 383.2 27.1 1248.0 25.3 962.6 24.2 759.0 25.7 −12.5 14.2 985.2 30.3 77.5 10.6 373.0 33.7 590.7 11.7 3.0 9.1 11.4 11.1 9.4 6.6 5.0 12.7 9.4 16.4 6.0 12.2 11.5 9.8
1156.2 187.5 235.3 95.0 234.6 260.3 70.6 134.5 19.6 35.7 76.7 77.6 220.7 1638.1 423.0
± (Ma) 282.4 1615.1 171.6 119.3 324.9 602.5 571.6 290.1 127.0 355.9 718.3 903.7 376.6 827.5 540.6 80.8 246.8 195.2 192.3 42.9 563.6 98.4 260.1 61.9 256.0 324.1 480.2 167.0 617.9 177.2 536.8 19.8 23.3 67.2 196.2 108.3 143.6 93.1 76.0 226.5 162.3 274.7 97.5 180.2 18.6 140.5
Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili
167
TABLE A3C. GEOCHRONOLOGICAL ZIRCON ANALYSES 206 U Pb U/Th 206Pb* 204 207 Pb Pb* (ppm) Sample AD009 Ophiolitic Mélange 17 432 13,730 1.5 20.6133 18 124 47,620 1.5 7.9033 13 212 6110 0.9 21.2261 20 544 8540 1.4 18.5369 22 379 13,920 1.5 20.3626 23 116 2320 1.3 14.3073 24 222 10,605 2.1 18.0167 25 127 4445 1.3 19.0159 25-1 256 636 2.6 18.8858 25-2 636 17,583 2.2 16.0144 25-3 179 4185 1.4 17.0760 25-4 222 6018 1.9 20.6592 25-5 540 10,473 1.1 20.4560 25-6 268 8841 1.3 21.2937 25-7 898 50,112 1.7 18.2908 25-8 987 6642 1.0 18.5354 25-9 1067 1734 1.6 15.4770 25-10 168 630 0.8 15.4446 25-11 254 5991 1.1 20.6852 25-11a 239 8184 1.3 19.2827
#
± (%) 4.6 1.0 5.8 13.0 4.5 10.1 2.4 10.3 15.8 1.9 14.9 11.3 3.2 7.4 1.0 8.5 31.7 37.0 10.9 6.8
207
Pb* 235 U*
Isotope ratios 206 ± Pb* 238 (%) U
0.1558 6.2503 0.1461 0.1632 0.1583 0.2300 0.3652 0.1742 0.1766 0.8031 0.2052 0.1639 0.1686 0.1603 0.5066 0.1745 0.1999 0.1821 0.1689 0.1720
4.9 1.5 5.9 13.1 4.8 10.2 3.1 10.5 16.1 2.3 14.9 11.4 3.4 7.9 1.4 8.7 31.7 37.2 11.0 6.9
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0.0233 0.3583 0.0225 0.0219 0.0234 0.0239 0.0477 0.0240 0.0242 0.0933 0.0254 0.0246 0.0250 0.0248 0.0672 0.0235 0.0224 0.0204 0.0253 0.0241
± error (%) corr. 1.7 1.1 1.0 1.5 1.8 1.6 2.0 2.0 3.2 1.2 1.0 1.6 1.3 2.8 1.0 1.8 1.0 3.3 1.1 1.5
206
Pb* 238 U*
0.34 148.4 0.75 1974.0 0.17 143.4 0.12 139.9 0.37 149.0 0.15 152.1 0.64 300.5 0.19 153.0 0.20 154.1 0.54 574.9 0.07 161.8 0.14 156.4 0.38 159.3 0.35 157.7 0.69 419.3 0.20 149.5 0.03 143.1 0.09 130.2 0.10 161.3 0.22 153.2
Apparent ages (Ma) 207 ± Pb* ± (Ma) 235U (Ma)
206
Pb* Pb*
± (Ma)
2.4 19.4 1.4 2.1 2.6 2.4 5.8 3.0 4.8 6.7 1.6 2.5 2.0 4.3 4.1 2.6 1.4 4.2 1.7 2.3
124.3 2050.3 54.9 368.8 153.0 925.2 432.6 311.0 326.6 689.3 550.8 119.1 142.3 47.3 398.8 369.0 761.7 766.1 116.1 279.2
107.6 17.7 139.6 293.6 105.0 207.9 52.9 234.5 361.4 40.5 326.6 267.5 74.2 176.2 23.3 192.9 684.9 806.8 258.7 155.2
147.0 2011.5 138.5 153.5 149.2 210.2 316.1 163.1 165.1 598.6 189.5 154.1 158.2 151.0 416.1 163.3 185.0 169.9 158.5 161.2
6.6 13.3 7.7 18.6 6.7 19.4 8.4 15.8 24.6 10.2 25.8 16.3 5.0 11.0 4.9 13.2 53.7 58.2 16.1 10.3
207
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Tectonic evolution of the Ankara Mélange and associated Eldivan ophiolite near Hançili Price, R.C., Johnson, L.E., and Crawford, A.J., 1990, Basalts of the North Fiji basin: The generation of back-arc basin magmas by mixing of depleted and enriched mantle sources: Contributions to Mineralogy and Petrology, v. 105, p. 106–121, doi:10.1007/BF00320970. Rojay, B., Yalınız, K.M., and Altıner, D., 2001, Tectonic implications of some Cretaceous pillow basalts from the North Anatolian ophiolitic mélange (Central Anatolia–Turkey) to the evolution of Neotethys: Turkish Journal of Earth Sciences, v. 10, p. 93–102. Rojay, B., Altıner, D., Altıner, S.O., Önen, A.P., James, S., and Thirlwall, M.F., 2004, Geodynamic significance of the Cretaceous pillow basalts from the North Anatolian Ophiolitic Mélange Belt (Central Anatolia, Turkey): Geochemical and paleontological constraints: Geodinamica Acta, v. 17, p. 349–361, doi:10.3166/ga.17.349-361. Sarıfakıoğlu, E., 2006, Petrology and origin of plagiogranites from the Daşkuplu (Eskişehir) ophiolite along the Izmir-Ankara-Erzincan suture zone, Turkey: Ofioliti, v. 32, p. 39–51. Sarıfakıoğlu, E., Özen, H., and Winchester, J.A., 2009, Whole rock and mineral chemistry of ultramafic-mafic cumulates from the Orhaneli (Bursa) ophiolite: NW Anatolia: Turkish Journal of Earth Sciences, v. 18, p. 55–83, doi:10.3906/yer-0806-8. Seifert, K., and Brunotte, D., 1996, Geochemistry of serpentinized mantle peridotite from Site 897 in the Iberia abyssal plain: College Station, Texas, Proceedings of the Ocean Drilling Program, Scientific Results, v. 149, p. 413–424. Şengör, A., and Yılmaz, Y., 1981, Tethyan evolution of Turkey: A plate tectonic approach: Tectonophysics, v. 75, p. 181–241. Shervais, J.W., 1982, Ti-V plots and the petrogenesis of modern and ophiolitic lavas: Earth and Planetary Science Letters, v. 59, p. 101–118, doi:10.1016/0012-821X(82)90120-0. Shibata, T., and Thompson, G., 1986, Peridotites from the Mid-Atlantic Ridge at 43°N and their petrogenetic relation to abyssal tholeiites: Contributions to Mineralogy and Petrology, v. 93, p. 144–159, doi:10.1007/BF00371316. Sinton, J.M., and Fryer, P., 1987, Mariana lavas from 18°N: Implications for the origin of back-arc basin basalts: Journal of Geophysical Research, v. 92, p. 12,782–12,802, doi:10.1029/JB092iB12p12782. Sinton, J.M., Ford, L.L., Chappell, B., and McCulloch, M.T., 2003, Magma genesis and mantle heterogeneity in the Manus back-arc basin, Papua New Guinea: Journal of Petrology, v. 44, p. 159–195, doi:10.1093/ petrology/44.1.159. Stacey, J.S., and Kramers, J.D., 1975, Approximation of terrestrial lead isotope evolution by a two-stage model: Earth and Planetary Science Letters, v. 26, p. 207–221, doi:10.1016/0012-821X(75)90088-6. Stern, R.J., Lin, P.-N., Morris, J.D., Jackson, M.C., Fryer, P., Bloomer, S.H., and Ito, E., 1990, Enriched back-arc basin basalts from the northern Mariana Trough: Implications for the evolution of back-arc basins: Earth and Planetary Science Letters, v. 100, p. 210–225, doi:10.1016/0012 -821X(90)90186-2. Sun, S.-s., and McDonough, W.F., 1989, Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes, in Saunders, A.D., and Norry, M.J., eds., Magmatism in the Ocean Basins: Geological Society [London] Special Publication 42, p. 313–345. Takazawa, E., Okayasu, T., and Satoh, K., 2003, Geochemistry and origin of the basal lherzolites from the northern Oman ophiolite (northern Fizh block): Geochemistry Geophysics Geosystems, v. 4, 8605, 31 p., doi:10.1029/2001GC000232. Tamura, A., and Arai, S., 2006, Harzburgite-dunite-orthopyroxenite suite as a record of supra-subduction zone setting for the Oman ophiolite mantle: Lithos, v. 90, p. 43–56, doi:10.1016/j.lithos.2005.12.012. Tankut, A., 1984, Basic and ultrabasic rocks from the Ankara mélange, Turkey, in Dixon, J.E., and Robertson, A.H.F., eds., The Geological Evolution of the Eastern Mediterranean: Geological Society [London] Special Publication 17, p. 449–454. Tankut, A., Dilek, Y., and Önen, P., 1998, Petrology and geochemistry of the Neo-Tethyan volcanism as revealed in the Ankara mélange, Turkey:
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Printed in the USA
The Geological Society of America Special Paper 480 2011
Petrology of a Franciscan olistostrome with a massive sandstone matrix: The King Ridge Road mélange at Cazadero, California Rolfe Erickson* Emeritus, Geology Department, Sonoma State University, Rohnert Park, California 94928, USA
ABSTRACT The King Ridge Road mélange is a unit of the Franciscan Complex, cropping out in an area of at least 50 km2 around the town of Cazadero, coastal California. This unit is an olistostrome with a massive, unfoliated sandstone matrix, containing >232 large meta-igneous and chert blocks of greatly varying size, lithology, and metamorphic history within the study area. This sandstone matrix is litharenite or arkosic arenite and exhibits prograde prehnite-pumpellyite facies and retrograde zeolite facies metamorphism. It is devoid of megascopic textures except for rare simple bedding. No fossils have been found, and no Bouma units or other graded beds are present. Detrital zircon geochronology has established the maximum age of deposition of the sandstone matrix at 83 Ma, whereas apatite fission-track data indicate cooling of the olistostrome below 100 °C at ca. 35–38 Ma. The 232 exotic blocks sampled in the study area are dominantly low- to medium-grade greenstones and cherts, together with fewer high-grade blocks partly composed of blue amphibole and/or omphacitic pyroxene, and some amphibolites. Thus, many of the blocks have higher grade metamorphic assemblages than the matrix. All block types are well mixed together, so none greatly predominate anywhere. Blocks of oceanic-island-arc plutonic rocks, including granitoids and recemented breccias, are particularly distinctive for this mélange. One granitoid block has a zircon U-Pb age of 165 ± 1 Ma. The massive sandy matrix of the olistostrome formed by accumulation of hyperconcentrated sedimentary density flows (grain flows) sourced primarily from the Klamath-Sierra continental magmatic arc. Many of the blocks record a pre-mélange history of metamorphism and exhumation, followed by partial subduction and reburial with the matrix after 83 Ma. Cooling below 100 °C took place at 35–38 Ma, probably associated with partial exhumation of the unit, with subsequent removal of ~10 km of cover.
*
[email protected] Erickson, R., 2011, Petrology of a Franciscan olistostrome with a massive sandstone matrix: The King Ridge Road mélange at Cazadero, California, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 171–188, doi:10.1130/2011.2480(07). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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INTRODUCTION A geological mélange (Fr., mixture) may be defined most generally as blocks in a matrix (i.e., Silver and Beutner, 1980). Two end-member models have been proposed for mélange development: one tectonic and one sedimentary (see review in Raymond, 1984). In some cases an assemblage of rocks, such as alternating sandstone and shale beds, may become deformed by progressive strain, commonly extension (e.g., Cowan, 1985). The more competent rocks are elongated and boudinaged into separate blocks or phacoids, which lie in a more incompetent matrix that has flowed around the more competent material during deformation. Foliated shale commonly constitutes that matrix, and this has been referred to as argille scagliose (It., scaly shale) (e.g., Hsü, 1968). Such units may be termed tectonic mélanges. Other mélanges of sedimentary origin may form a block-inmatrix texture in several ways. In the case to be discussed here, preexistent blocks were deposited on the sea bottom and buried by younger sediments that formed their matrix. Such a mélange is an olistostrome, and the blocks are called olistoliths. The sedimentary matrix is unsheared. If the olistoliths are of a type foreign to the environment of the matrix, such as eclogite blocks in quartz sandstone, they are often called exotic blocks. In a tectonic mélange, blocks, and especially matrix, are sheared and deformed, and both have evolved from material within the original boundaries or along the boundaries of the mélange zone. In an olistostrome, on the other hand, the blocks and matrix have separate origins from separate materials and may not be exactly the same age. Olistoliths may have any pre-mélange character, and the matrix in which they lie is undeformed. It is often the case that olistostromes are tectonically reworked so that they have many of the features of a tectonic mélange (Berkland et al., 1972). Such units have been termed tectonized mélanges (Schuster, 1980), and the original sedimentary features may be largely obscured and difficult to identify. Although the Franciscan Complex of coastal California ranks among the world’s best known mélange localities, controversy persists over the extent to which olistostromal mechanisms contributed to mélange development there. Some advocate an early stage of olistostromal development for many deformed mélanges (e.g., Cowan and Page, 1975; Aalto, 1989; Wakabayashi, this volume), whereas others believe that the mélanges developed entirely as a product of tectonic strain (Cloos, 1985; Şengor, 2003). Some of this controversy has arisen because of apparently severe deformation of most Franciscan mélanges that may obscure original sedimentary features (see Wakabayashi, this volume). This study describes a nearly undeformed sandstone matrix olistostrome of large lateral extent (studied in a 50 km2 area and extending much farther) containing a large number and variety of exotic blocks. Below, I present field, petrographic, and geochronologic evidence that constrains the depositional, metamorphic, and exhumation history of this mélange, followed by a discussion of the mélange in the context of regional and general tectonic and sedimentary processes.
KING RIDGE ROAD MÉLANGE FIELD RELATIONSHIPS Study Area Mélange Name and Location The mélange unit to be discussed, the Franciscan King Ridge Road mélange (Fig. 1), is an untectonized olistostrome named after the road crossing the main outcrop area north of the village of Cazadero, California (Fig. 1). The study area lies ~100 km NNW from San Francisco, California, and 6 km north of the Russian River, on both sides of Cazadero Road, in western Sonoma County, California. The village of Cazadero at its center lies at lat 38°32′30″, long 123°07′30″. In this article the author presents the results of a small-scale study of ca. 50 km2 of the mélange, mapped at a scale of 1:12,000 in and around the village of Cazadero and the nearby classic blueschist locale of Ward Creek (Coleman and Lee, 1963; Erickson, 1992) (Figs. 1,
King Ridge Road
Fg
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WC Fg
n
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Cm KRRm Fg
Fort Ross Road
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Fg KRRm Cm
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ea
s
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Figure 2
38°30′00′N
Fa Fss ult
Figure 3 boundary Cazadero Highway 123°07′3′′W
Russian River
River Road N
0 Fss Fg GVcg Cm KRRm WC
1 2
3 4 km
Franciscan sandstone and mélange Franciscan greenstone Great Valley conglomerate Cazadero Cazadero mélange San Francisco King Ridge Road mélange Ward Creek location
Figure 1. Study area location map. Study area inside the heavy boundary, also the boundary of Figure 3. Outer part of map simplified from Blake et al. (2002). Inside the boundary the King Ridge Road mélange (KRRm) outcrop area is labeled; outside the boundary KRRm is merged with U.S. Geological Survey units KJfs, TKfs, and TKfss as generalized Franciscan sandstone: Fss (author’s term). All metamorphosed mafic rocks in either map are labeled Franciscan greenstone (Fg).
Petrology of a Franciscan olistostrome with a massive sandstone matrix 2, 3). Most mapping was done in 1975. Reconnaissance suggests that most of the mélange lies beyond the study area and remains unmapped in detail. An unpublished 1:12,000 geologic map of the study area covering a larger area than Figure 3, together with an accompanying catalogue of hundreds of sampled exotic blocks, are available as pdf files from the author’s file in the permanent Sonoma State University Library digital archive, which can be accessed freely at the author’s archive Web site: http://hdl.handle .net/10211.1/169. The regional geology was most recently summarized by Blake et al. (2002). For a broad historical background of Franciscan geologic studies the reader is referred to Bailey et al. (1964) and Wakabayashi (1992a, this volume).
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An Exposed Block-Matrix Contact Defines the Mélange as an Olistostrome A ~100 × 500 m block of deformed red radiolarian chert, in contact with massive sandstone, is exposed in a cut on the Cazadero Road (Figs. 1 and 2). Both units are components of the King Ridge Road mélange. Theirs is the only block-matrix contact exposed in the study area. Figure 2 shows several features of this contact zone, which together define these exposures as components of an olistostrome. The mélange was constructed solely by sedimentary depositional processes and not by tectonic ones, although it has been tectonically disturbed somewhat during its later history. So far as
Figure 2. An ~100 × 500 m red chert block in contact with massive sandstone of the mélange matrix. Outcrop location UTM (universal transverse Mercator) 10 S 0492727E 4263330N 1927 datum in a road cut on the Cazadero Road. (A) Near-isoclinal folds in chert block 5 m N of contact; hammer for scale. (B) Open fissure along contact from outcrop slumping; exposed chert surface shows irregular stepped surface, without polish or slickensides, thus not a fault; 10-cm white pipe for scale. (C) Field photograph shows contact location by white posts and overdrawn white line; white clipboard for scale; sandstone in lower right, chert in upper left across contact. (D) Field photo of massive sandstone without bedding; hammer for scale; sandstone sample no. 1 collected here.
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Ward Creek locality
38° 38° 32′ 38 32′ 30′′ 32 30′′
Fort Ross Road
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Figure 3 (Continued on facing page). Geologic map (in two parts) of the King Ridge Road olistostrome mélange (KRR) around the village of Cazadero and the Ward Creek locality. A total of 232 exotic blocks of greatly varied lithologies are located and assigned to one of four lithologic classes. Blocks are contained in a massive sandstone matrix, which has locally sparse bedded intervals. For more detail on blocks, log on to the author’s digital archive of Cazadero documents at http://hdl.handle.net/10211.1/169 at Sonoma State University.
Petrology of a Franciscan olistostrome with a massive sandstone matrix
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MAIN UNITS
8 KRR
King Ridge Road olistostrome mélange. Sandstone matrix with petrographically highly diverse assemblage of exotic blocks; 232 recorded blocks.
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Other Franciscan units. These are a second mélange, the Cazadero mélange, several faultbounded units of various metabasalts, and the Little Black Mountain rhyolite stock.
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EXOTIC BLOCK TYPES IN THE KING RIDGE ROAD MÉLANGE Chert (type I)
Cazadero village
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Figure 3 (Continued).
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the author knows, it is the least tectonically disturbed Franciscan mélange described to date. Perhaps half of all the exotic blocks in the mélange have some type of pre-incorporation internal structure such as folds or metamorphic foliation, whereas the matrix never does. The folding and shearing of the blocks must have preceded their incorporation into the younger, undeformed mélange matrix. A second, strongly metamorphosed mélange unit with a blueschist matrix found in the study area, the Cazadero mélange of Erickson (1992), will not be discussed in this chapter. Other names have been used for this unit by other authors; it was initially correlated with the Pickett Peak terrane by Blake et al. (1984), and later renamed the Cazadero terrane (Blake et al., 2002). It has been called the Skaggs Spring Schist by Wakabayashi (1992a, 1992b), Wakabayashi and Dumitru (2007), and Snow et al. (2010). Sandstone Matrix In the study area (Fig. 3) the mélange matrix is dominantly (>95%) massive, generally medium-grained sandstone that has little bedding or stratification. No Bouma sequences, graded beds, rippled zones, or related turbidite structures such as scour or drag marks are present, and no cross-bedding, so there are no directional or top indicators present. Joints are randomly oriented. There are no water escape structures such as dish structures, nor are there any sandstone dikes. There are no burrows, and no fossils. The sandstone is unsheared, lacks foliation or lineation, and in hand specimen is not obviously metamorphosed (but see below). This sandstone unit is defined here as a massive sandstone facies of the Franciscan Complex. About 50 km2 of this matrix has been studied in detail (Fig. 3). In addition, reconnaissance farther out along local roads in all directions from Cazadero shows that this massive sandstone facies underlies a much larger area than that of Figure 3 alone. Its outer boundary has not been defined, except for a short segment along the Pacific coast (see below). This large body of unfoliated metasandstone, and the block population within it, are unique among all units described to date in the Franciscan Complex. Rare, sharply bounded shale and siltstone interbeds in single beds or sets a few centimeters to a few meters thick are found locally interbedded with the sandstone. None are graded. No other non-siliciclastic sediments are present in the matrix; specifically, no radiolarian chert strata are interstratified with it. A few small lenses of conglomerate are also present. The absence of stratigraphic markers of original horizontality or bed sequence in the mélange matrix prevents assessment of the present 3-D form of the overall sandstone body, or its thickness. Neither the top nor the base of the mélange is exposed. Exotic Block Population and Distribution The King Ridge Road mélange contains a very large number of exotic blocks of all sizes up to ~1 km. In mapping, an arbitrary
lower size cutoff of ~2 m was used to limit the number examined. In all, 232 blocks equal to or larger than this limit were mapped in the mélange map area of ~50 km2 (Fig. 3). The distribution patterns of blocks in U.S. Geological Survey Map MF-2402 (Blake et al., 2002) and the author’s reconnaissance suggest that there are thousands more beyond the area of Figure 3. Those mapped and located in Figure 3 are cherts and metamorphosed mafic rocks, together with several meta-granitoid and meta-dacite blocks. The number of these large blocks diminishes from east to west across the area; at the same time they appear to be randomly distributed in any ~1 km2 area (Fig. 3), and the petrographic types are well mixed. There is no indication of “ghosted stratigraphy” in the blocks as observed in the classical mélanges at Gwna in Wales (Schuster, 1980, p. 68). It became clear as mapping progressed that there never were significant numbers of primary small exotic blocks, <~1m in size, in the initial block population of the mélange in the area of Figure 3. Such blocks never show up in matrix outcrops. Southwestern Contact of the King Ridge Road Olistostrome Mélange with a Younger Tectonic Mélange Detailed mapping in the area of Figure 3 and reconnaissance for 10–20 km along county roads in all directions beyond that area only shows massive sandstone of the King Ridge Road mélange in outcrop. In general, the external contacts of the mélange have not been determined or described, nor can the author provide an estimate of the size of the total olistostrome outcrop. A small part of the olistostrome’s contact with the unit to its immediate west can, however, be seen along the Pacific coast, southwest of the mouth of the Russian River; here its somewhat gradational contact with a younger tectonic mélange is clearly revealed in a segment of continuous exposures in the sea cliffs. Reconnaissance to the southwest of the area of Figure 3 shows typical King Ridge Road mélange cropping out to within a kilometer or two of the ocean beach from Jenner, California, to Wright’s Beach State Park ~6 km to the southeast, along a boundary striking ~N20°W. To the west of that boundary an unnamed tectonic mélange crops out. At the campground at Wright’s Beach State Park, massive medium-grained sandstone with random jointing crops out. At the park it contains exotic blocks of chert and varied metamorphites, and the whole assemblage is typical of the King Ridge Road olistostrome mélange. As an observer moves ~NNW along the beach cliffs from the park, a series of changes occur in the olistostrome matrix exposed in these cliffs that define the local southwestern contact of the King Ridge Road mélange. Over a distance of ~0.6 km, NNW from the entrance to Wright’s Beach State Park, the character of outcropping Franciscan rocks totally changes, and defines the contact. Key exposures along this gradational contact are described and precisely located by UTM coordinates in an unpublished Northern California Geological Society field-trip guide by the author, available in the Sonoma State University Digital Archive at http://hdl.handle.net/10211.1/498.
Petrology of a Franciscan olistostrome with a massive sandstone matrix At the entrance to the beach proper from the campground the massive character of the sandstone changes, and simple bedding striking northwest and dipping northeast appears in it. Then, as the observer continues north, interstratified dark-brown shale beds a few centimeters to 1–2 m thick, and with the same general attitude, begin to appear in the bedded sandstone. Continuing north, as soon as these shale beds become a few percentages of the exposure, they begin to exhibit shear and internal folds, and the sandstone beds between them start to show boudinage. The boudins tend to form shear-bordered diamond shapes with their long axes parallel to each other. Progressing still farther northward, area sandstone-shale stratigraphy is disrupted, and further extension has transformed the boudins into 1–2-m-long phacoids that are now surrounded by a foliated shale matrix: The outcropping assemblage has become a tectonic mélange. This southwestern boundary of the King Ridge Road olistostrome mélange with the adjoining tectonic mélange was influenced by the stratigraphy. When sufficient shale interbeds were present and some external force was applied, the protolith quickly broke up into a tectonic mélange. The tectonic mélange is composed of the olistostrome mélange and must always be somewhat younger than the latter. The original olistostrome contains a number of 2–3-m exotic blocks composed of chert and various metamorphites. These are hard and tough, and went through the transformation just described without being affected. The younger tectonic mélange contains two kinds of blocks—original exotic blocks, and more recent phacoid blocks. All the new phacoid blocks are of the same lithology, sandstone, whereas the original exotic blocks are varied in lithology.
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Each sample was point counted, using a 20 mm × 20 mm 400-point grid with 1 mm spacing, centered on the stainedsample thin section. From these stained thin sections it was easy to determine the mineralogical identity of quartz grains of all types (Qm, single grains; Qp, polycrystalline grains; and Qc, chert) and the identity of feldspar grains of both types (P, plagioclase, and K, potassium feldspar [Kspar]). Polycrystalline siliciclastic lithic grains (L) are also easy to identify as a general class but are often difficult to label uniquely as to their volcanic, sedimentary, or metamorphic origin. They include a large proportion of aphanitic grains that lack distinctive textures that would identify them. In addition, microprobe analysis of many lithic grains shows that they have undergone thorough allochemical metamorphism (see below), so their present composition is no help in determining their original identity. Consequently only total lithic-clast numbers were determined for each mode, with no attempt to break the category down into subtypes. No carbonate grains were observed, and no fossils. The total modal data are given in Table 1. The sandstone classification scheme of Pettijohn et al. (1987) was used to name the sandstone samples. Seven of these, with dominant lithic clasts, are litharenites, and two, with dominant feldspar clasts, are arkosic arenites. The two groups are widely separated in the QFL diagram (Fig. 5). The data in Table 1 show that the smaller group defined by two samples, nos. 4 and 5, has the highest feldspar proportion, the only significant amount of Kspar, and by far the highest amount of plagioclase as well. The other seven samples are dominated by lithic grains. The two groups define separate high feldspar and high lithic petrofacies in the sandstone. The nature of the boundary between them is not known.
PETROGRAPHY OF BLOCKS AND MATRIX Sandstone Matrix Metamorphism Sandstone Matrix Ten samples of the sandstone matrix of the mélange were taken, spaced across its outcrop area (Fig. 3; locations in the GSA Data Repository1). From each sample, paired billets were taken; one was stained for feldspar type and used in modal analysis, and one was polished for potential microprobe analysis (only two were so used). Sample 9 turned out to be extensively microscopically brecciated and veined and was not included in the study. In thin section view, grains are angular but well sorted. Essentially no matrix is present, and none of the samples is a graywacke (Pettijohn et al., 1987). Pressure solution may have affected the grains so that they fit together tightly (Fig. 4A). The grains show no foliation development, and the sandstone belongs to textural zone 1 (TZ-1) of Blake et al. (1988).
1
GSA Data Repository Item 2011263, Table DR1: Zircon data and Table DR2: Sample locations, is available at www.geosociety.org/pubs/ft2011.htm, or on request from
[email protected], Documents Secretary, GSA, P.O. Box 9140, Boulder, CO 80301-9140, USA.
Extensive veins and patches of white laumontite of metamorphic origin cut the brown matrix sandstone of the King Ridge Road mélange all through the study area. A petrographic and microprobe study was made of selected sandstone specimens to better understand this metamorphic event. The dominant matrix texture visible in hand specimens and petrographic and microprobe slides is still that of the sandstone protolith, so the unit is best described as a metasandstone (Fettes and Desmons, 2007). A petrographic and microprobe study was made of representative litharenite (no. 1) and felsarenite (no. 4) sandstone samples (Fig. 3; Table 1). Figures 4A, 4B, 4C, and 4D show typical textures and mineralogies observed. Plagioclase is generally altered to pure albite, associated with clinozoisite and sericite (Fig. 4B); it commonly retains original twin and zone patterns as relicts. Prehnite and pumpelleyite are found in partly altered detrital chlorite grains (Fig. 4C). Former Kspar is replaced by sericite + clinozoisite + albite (Fig. 4D). Detrital Kspar in petrographic thin sections turns bright yellow if exposed to cobaltinitrate stain solution, and under the
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500 µm
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Figure 4. Microprobe backscatter images of sandstone matrix textures and mineralogies. (A) Tightly packed grains with very little matrix; note absence of deformation and/or foliation. (B) Plagioclase grain replaced by albite + clinozoisite during metamorphism of the mélange. (C) Detrital chlorite grain partly replaced by coexisting prehnite and pumpelleyite during metamorphism of the mélange and defining the prehnite-pumpelleyite metamorphic facies. (D) K-feldspar replaced by albite + clinozoisite + sericite during metamorphism of the mélange.
TABLE 1. MODAL POINT COUNTS FROM SANDSTONE MATRIX SAMPLES Sample Qm Qp Qc P K Lith 1 91 0 11 91 0 2 28 2 103 2 3 72 1 2 39 3 74 24 12 66 1 1 88 4 114 9 1 153 55 74 5 100 16 1 140 44 97 6 108 11 8 56 0 2 14 7 68 2 6 81 0 2 44 8 82 9 7 70 0 2 16 10 88 7 5 122 5 1 96 Notes: Sample 9 not used; highly brecciated. See Figure 5 caption for abbreviations in column heads.
Bt 9 12 8 8 6 12 14 4 13
Total 430 432 373 414 404 409 415 388 436
Downloaded from specialpapers.gsapubs.org on January 1, 2012
Petrology of a Franciscan olistostrome with a massive sandstone matrix Q
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rograde zeolite facies and may have developed as a retrograde phase during exhumation of the mélange.
Craton interior
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Figure 5. QFL diagram showing fields for different provenances of noncarbonate sand-grain populations, based on modal point counts. Q = Qm + Qp + Qc, where Qm = monocrystalline quartz, Qp = polycrystalline quartz, and Qc = chert; F = K + P, where K = total Kspar and P = total plagioclase; L = lithics of all types save chert. Redrawn after Dickinson et al. (1982, their figure 1). Nine sandstone samples from the mélange matrix are plotted; modal data from these are in Table 1, and their locations are given in Table DR2 and shown in the map of Figure 3. The light-gray boundaries define three Franciscan sandstone terranes in Jayko and Blake (1984, in Blake et al., 1984). In the text these three terranes are compared with the sandstone of the King Ridge Road mélange.
microscope it looks unaffected by metamorphism. Microprobe analysis shows, however, that what appear to be Kspar crystals are actually complex intergrowths of albite (An0) + clinozoisite + sericite (Fig. 4D) after Kspar. The response to the K stain is actually from the metamorphic sericite. The Kspar was replaced by the same minerals as plagioclase but with different modal proportions. Felsic to mafic aphanitic grains of original sedimentary or igneous rocks have been partially or wholly transformed to prehnite ± pumpelleyite ± sericite ± quartz ± chlorite ± albite. Every reactive grain in these matrix sandstones had been mineralogically and chemically transformed. At the same time the replacement was remarkably isotextural; small details of phenocryst form are preserved, and grain boundaries are sharp and clear. Organic matter in the sediments was converted to graphite. No relicts of original detrital grains of felsic minerals were observed. The coexistence of prehnite and pumpelleyite in the metasandstone (even in the same original parent crystal; Fig. 4C) indicates that it reached maximum prograde metamorphic conditions of the prehnite-pumpelleyite facies. Laumontite, so abundant in fracture veins in the mélange outcrop, is stable under ret-
A hand sample of each of the 232 mapped exotic blocks was collected. Blocks were characterized by field and hand specimen examination, and in 71 cases by thin section analysis. To organize descriptions, the block population has been subdivided into four groups (A, B, C, and D) that are each in turn subdivided into several subgroups (1, 2, 3, etc.). Group A is cherts of all types; group B is wholly recrystallized highpressure metamorphites whose blocks have blue amphibole, lawsonite, or omphacitic pyroxene; group C is wholly recrystallized, medium- to low-pressure metamorphites; and group D is little-recrystallized greenstones and other units with common relict textures (Table 2). Exotic blocks from these four major groups are identified by unique symbols and are individually plotted in Figure 3. Greater detail for many blocks is given in the unpublished map and block catalogue in the Sonoma State University digital archive described earlier. Individual metabasalt blocks at Cazadero come from all the major facies of metamorphism save the zeolite facies (Table 2). The highest grade blocks have eclogite facies mineralogy. A range of metamorphic assemblages from eclogite to zeolite facies may be displayed by any given exotic block. Mineralogies from several facies will typically be represented in any large group of blocks, and the overall assemblage is well mixed. The textural classification of Coleman and Lee (1963), developed in their initial study of the metamorphites at Cazadero, is used in this discussion. They defined four textural grades, types I, II, III, and IV. Type I rocks lack glaucophane; type II have glaucophane but no metamorphic textures; type III have glaucophane and are foliated but uniform, hence are schists; type IV have glaucophane and are foliated and mineralogically banded, hence are gneisses. Both type I and type II blocks commonly display relict features such as pillow margins and/or vesicular zones. Group A, the chert blocks, comprises 67 unfoliated chert blocks, which are all type I. Within this group, subgroup A1 contains 29 red chert blocks of the type so typically found in Franciscan sequences; subgroup A2 contains 22 blocks that are white-gray but otherwise unaffected, and subgroup A3 contains 16 blocks that have more complex color patterns and multiple generations of quartz veins. Red and gray chert blocks commonly show relict bedding, which is often deformed. Group B contains 42 high-pressure metamorphites, of which 38 were examined in thin section, which contain a blue amphibole that is not always glaucophane (see below), and commonly other high-pressure phases as well, and are texturally type II, III, or IV. This first subgroup, B1, contains 10 foliated blocks that are schists or phyllites, with blue amphibole plus several minerals of the group plagioclase ± omphacitic pyroxene ± clinozoisite ± quartz ± muscovite ± sphene ± lawsonite ± chlorite ±
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ALL BLUE BLOCKS
C+D ALL GREEN BLOCKS
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LOCATION ONLY
29
TOTAL 232
TABLE 2. KING RIDGE ROAD MELANGE EXOTIC BLOCK POPULATION Thin section based Number of Texture + mineralogy subtype subtype None n/a Unfoliated; cherts; quartz; red, white, varicolored B1 10 Gneiss/schist/phyllite; plag + cpx + czt + qz + musc + sph + law + chl + blue amph + green amph + op; no garnet B2 7 Isotropic; plag + cpx + czt + qz + musc + sph + law + chl + green amph + op; no garnet B3 5 Retro. eclogite + eclogite; garnet + cpx + musc + law + blue amph B4 10 Isotropic; meta-felsic plutonic; trace blue amphibole + relict assemblage B5 6 Varied breccias C1 C2
9 6
Greenschist; green amph + gar + czt + plag + chl Isotropic; green amph + gar + czt + pl + chl
D1 D2 D3
8 4 6
Greenstone; chl + relict assemblage; 2 brecciated gabbros also Microaugen gneiss foliation; chl + alb + relict biotite dacite mineralogy Unfoliated; six unique blocks
None
n/a
TOTAL 71
Notes: alb—albite; amph—amphibolite; chl—chlorite; czt—clinozoisite; gar—garnet; law—lawsonite; musc—muscovite; op—orthopyroxene; pl—plagioclase; sph—sphene; q—quartz.
green amphibole ± opaque. None of these rocks has garnet. The rocks show a great variety of textures and mineralogies, which suggests a great variety of protoliths and metamorphicdeformational histories. Subgroup B2 contains seven blue amphibole-bearing granofelsic (unfoliated but phaneritic) rocks, with parageneses drawn from the same group of minerals as the foliated rocks of B1, and also with a great variety of textures and mineralogies. One granofels contains ~50% quartz and is clearly metasedimentary, the only one of all the high-pressure rocks. Two granofelses are recemented breccias. Subgroup B3 contains five granofels metamorphites that contain garnet in addition to other high-pressure phases; all have omphacitic pyroxene partly replaced by blue amphibole, and are probably partially retrograded eclogites. Garnet is invariably deeply corroded and surrounded by shells of retrograde chlorite, muscovite, or lawsonite. In addition, one fresh eclogite was included in B3. Subgroup B4 contains 10 blocks composed of weakly metamorphosed felsic igneous plutons, 2 of which are granitoids and 8 of which are metadiorites. One granitoid is discussed in some detail below; the other is wholly brecciated but not otherwise studied. All these metaplutonic blocks contain very small amounts of blue amphibole. In the one case where it was analyzed (block 1503, see below) the amphibole was ferrorichterite, not glaucophane. Subgroup B5 is composed of six blocks of nongranitoid breccias. One of these is discussed in some detail below. Breccia
veins and zones are also commonly found in dominantly solid blocks, as in the case of the granitoid described below. Groups C and D contain between them 94 greenstone or amphibolite blocks that lack glaucophane or other highpressure minerals. C and D are often hard to tell apart in hand specimen, but they are so different in section that the hand specimen group was split into two; 33 were sectioned. Both groups are type I units in the textural classification of Coleman and Lee (1963). Study of the sample population of the 33 sectioned rocks broke out the two petrographic block groups C and D, each with subgroups. Group C rocks have been wholly reconstituted by metamorphism and lack relict textures of any sort. In subgroup C1, nine blocks are amphibolites composed of amphibole ± garnet ± clinozoisite ± plagioclase ± chlorite. C1 blocks lack quartz or muscovite and are probably meta-igneous. In subgroup C2, six blocks are amphibolites composed of amphibole ± clinozoisite ± muscovite ± quartz ± plagioclase. C2 blocks lack garnet or chlorite and are probably metasedimentary. Approximately half of C1 + C2 are foliated and half granofelses. All may be characterized as moderate to low pressure and temperature metamorphites. Two have some lawsonite in plagioclase but did not complete reequilibration at high pressure. Group D is composed of very weakly metamorphosed metavolcanics, dominantly metabasalts, which have not lost their primary microscopic textures, and are here called greenstones. All except D2 lack foliation or lineation. Large-scale relict vesicular zones and pillows are common.
Petrology of a Franciscan olistostrome with a massive sandstone matrix Subgroup D1 is composed of eight blocks of similar, very weakly metamorphosed unfoliated metabasalts with relict randomly oriented plagioclase laths (jackstraw texture) and relict clinopyroxene; only chlorite could be identified as a new metamorphic phase. In the field, two of them were seen to have relict vesicles with amygdule fillings, and four with relict pillows. One is brecciated. Two blocks of strongly brecciated and recrystallized meta-hornblende gabbros or meta-diorites were also placed in this subgroup. Subgroup D2 is composed of four blocks of distinctive, weakly metamorphosed biotite dacite porphyry with biotite, plagioclase, Kspar, and quartz phenocrysts in a leucocratic albite + quartz groundmass, now strongly modified texturally by a microaugen texture found only in these blocks. A microprobe study of one meta-dacite block shows that biotite crystals still have significant but variable K2O in them. The Kspar has a complex reaction to cobaltinitrate stain in microscopic analysis, as noted above. Chlorite is also a common metamorphic phase. Subgroup D3 contains six unrelated unique blocks. One is a strongly brecciated quartz-plagioclase granofels; one is a garnet + amphibole–bearing quartzite (quartz ~80%); one is a brecciated monomineralic plagioclase granofels; one is a granofels of massive radiating fibrous clinozoisite; and one is a pyroxeneplagioclase granofels with curved boundaries between its crystals, suggesting a metamorphic origin. One is a serpentinite. Some block lithologies found in other mélanges are absent in the King Ridge Road mélange. There are no blocks of shale, sandstone, or conglomerate, or their metamorphic equivalents, no blocks of carbonates, and no blocks composed of the lower parts of ophiolites below the volcanic level, save the one serpentinite block. There are no meta-tuffs or metamorphosed felsic flows. All of the exotic blocks, save those in groups B4 and D2, are metamorphosed mafic rocks or marine cherts, and probably have protoliths of oceanic crust, arcs, or seamounts. Very limited geochemical and isotopic data (see below) support this interpretation.
9 8
A
GEOCHRONOLOGY AND GEOCHEMISTRY Depositional Age of the King Ridge Road Mélange by Detrital Zircon Analysis A detrital zircon age study was made of the King Ridge Road mélange, using a sandstone matrix sample from its highfeldspar petrofacies (Figs. 6A and 6B; Fig. 3). The analyses were performed at the University of Arizona LaserChron Center in the spring of 2008. Analytical results are provided in Table DR1 (see footnote 1). Analytical methods are described in Gehrels et al. (2008). A total of 57 zircons were analyzed, and the results are shown in histograms and probability density graphs in Figures 6A and 6B. The four youngest zircons are 83 Ma (Campanian) (Table DR1) and define the maximum depositional age of the exposed part of the sandstone matrix of the mélange. Other 33 zircons have older Mesozoic ages; none are Paleozoic, and 20 are Precambrian. The Mesozoic zircon results, shown in detail in Figure 6A, are concentrated in five age peaks. Four peaks are sharply delineated at 82, 95, 106, and 119 Ma, and the fifth is broader, centered on 160 Ma. These peaks in detrital zircon supply should correspond with peaks of plutonic activity in the zircon source, which the author speculates is most likely the nearby California arc (Ducea, 2001). The peaks at 82 and 95 Ma do correspond with the large Late Cretaceous magmatic pulse in the California arc, and the 160 Ma peak corresponds with the 160–150 Ma Late Jurassic pulse there as well. The peaks at 106 and 119 Ma, however, cannot yet be assigned to any published magmatic pulse in the California arc. The absence of Paleozoic zircons from the Cazadero sample is striking, given the presence of large Paleozoic units in the Sierran arc nearby (i.e., Hanson et al., 1988). The age gap between dated zircons in the present study is 192–581 Ma (Table DR1; Fig. 6B). The 35% of Precambrian zircons in the sample range
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0 70
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Figure 6. (A) U-Pb age versus number histogram and superimposed relative probability curve for detrital zircon (DZ) ages from the King Ridge Road mélange matrix between 0 and 200 Ma. Plot provided by University of Arizona geochronology laboratory. See Table DR1 (see footnote 1) for data. (B) U-Pb age versus number histogram and superimposed relative probability curve for detrital zircon ages from the King Ridge Road mélange matrix between 200 and 2900 Ma. Plot provided by University of Arizona geochronology laboratory. See Table DR1 for data.
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across the whole of Proterozoic and latest Archean time, from 581 to 2900 Ma, with a moderate concentration in the 1400– 1800 Ma interval (Table DR1; Fig. 6B). Precambrian terranes in the continental interior may have provided clastic grains to the Franciscan trench. U/Th ratios in zircons indicate their igneous or metamorphic origin (Rubatto, 2002). A total of 55 of the 57 zircons in this sample have U/Th ratios characteristic of an igneous origin (Table DR1). Several recently published Franciscan studies provide other detrital zircon ages for the complex, i.e., Snow et al. (2010), Ernst et al. (2009), and references therein. A total of 22 detrital zircon ages from various Franciscan localities have been published to date; the maximum depositional ages range from 52 to 144 Ma. The 83 Ma age of the King Ridge Road mélange is the second youngest of these ages. A distinctive feature of all these Franciscan units is that Paleozoic zircons are absent from all but the oldest sample in this group, the 144 Ma Skaggs
Springs Schist (Snow et al., 2010), just as they are absent from the Cazadero sample. History and Origin of Two Blocks I: Meta-Granitoid Block 1503 The protolith character and metamorphic history of one ~50 m granitoid block (block catalog sample 1503) (Fig. 3; Fig. 7) have been studied in detail (Erickson et al., 2004; see the corresponding, more detailed, poster in the Sonoma State University digital archive at http://hdl.handle.net/10211.1/169). In outcrop the block is a tan-red phaneritic rock. No purple amphibole is seen, and the rock is type I of Coleman and Lee (1963). Thin section and microprobe study shows that the dominant fabric of the rock is composed of quartz, plagioclase, magnesiohornblende, chlorite books, and an Fe-Ti oxide with sphene overgrowths, all intergrown in an igneous-appearing
21 21µm µ Figure 7. Photomicrographs and backscatter probe images from meta-granitoid block 1503. (A) 40× crossed polarized image of a relict igneous hypidiomorphic-granular plutonic texture with quartz, magnesiohornblende, plagioclase, and chlorite. (B) 100× plane polarized image of dark-blue ferrorichterite formed in M1 from magnesiohornblende. (C) 40× plane polarized image of ~1-cm-wide breccia vein cutting metagranitoid. (D) Electron backscatter image of pumpelleyite sheets replacing chlorite in an M2 event.
Petrology of a Franciscan olistostrome with a massive sandstone matrix hypidiomorphic-granular texture (Fig. 7A). The magnesiohornblende contains abundant patches and veins of blue ferrorichterite amphibole (Fig. 7B), and the plagioclase is pure albite. This mineral assemblage is cut by a set of abundant breccia veins up to several centimeters wide (Fig. 7C) composed of fragments of all preexistent minerals in a profoundly aphanitic dark matrix. These breccia veins are in turn cut by veins, mats, and sheets of monomineralic pumpelleyite, in the last case partially replacing chlorite (Fig. 7D). The pumpelleyite veins follow different paths than the earlier breccia veins did. No aragonite or other high-pressure minerals were found in the block, and it displays no foliation or lineation. The block has yielded a U-Pb zircon age of 165 ± 1 Ma (Erickson et al., 2004). The block lacks an actinolite or serpentinite rind. Figure 8 is a trace-element spidergram (Pearce, 1983) that compares this granitoid pluton to M-type and I-type plutonic standards. Rb/Sr and Nd/Sm isotopic data were also collected for the pluton (Erickson et al., 2004) but are not repeated here. Rare earth analysis shows (La)n/(Lu)n = 9.3 (Erickson et al., 2004). The quartz, magnesiohornblende, chlorite, plagioclase, and sphene-mantled iron oxide, all with a hypidiomorphic-granular texture (Fig. 7A), are interpreted as a relict plutonic igneous assemblage (biotite-hornblende quartz diorite) that partly survived one or more metamorphic events (Erickson et al., 2004). The rare-earth pattern suggests strong fractionation in its crystallization history. The chlorite is interpreted as recrystallized former biotite, and the sphene-mantled iron oxide is interpreted as once homogeneous igneous Fe-Ti oxide that expelled much of its titanium to form sphene mantles on residual iron oxide during the first metamorphism of the rock, M1. The trace-element (Fig. 8) and isotopic data for this block (Erickson et al., 2004) show that this metaplutonic body was originally an M-type granitoid (Pearce, 1983) from an oceanic island arc with great geochemical similarity to the modern Mariana Islands. 100.0
Sample element/ORG
Continental arc, I-type Mt. Barcroft pluton #WM-791 60/% SiO2 (Ernst et al. 2003)
10.0
Oceanic arc, M-type New Britain granitoid #40920, 60% SiO2 (Whalen, 1985) Cazadero meta-granitoid block 60% SiO2 Cazadero meta-basaltic trachy andesite block 53% SiO2
1.0
0.1 K20 Rb Ba Th Ta Nb Ce Hf Zr
Sm Y
Yb
Figure 8. Pearce diagram (Pearce, 1983) comparing I-type and M-type arc chemistry with that of meta-granitoid block 1503 and meta-breccia block 1500. ORG is a model ocean ridge granite; its composition is defined in the 1983 paper by Pearce.
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The author speculates that the parent arc of this block was subducted <165 Ma, causing the first metamorphism, M1, of the block. Incomplete replacement of magnesiohornblende by ferrorichterite and the absence of any high-pressure phases suggest that it was not subducted deeply or for a long time. The date of this M1 event can only be bracketed between 165 and 83 Ma. The formation of the probably once-fluidized breccia veins with fragments transported from their point of origin suggests hydrofracturing of the rock; its resolidification suggests recrystallization at low pressure (P) and temperature (T), as no new metamorphic minerals visible even to microprobe observation grew in the aphanitic breccia matrix; probably the block had been moved to near the surface at this time. The block was then probably moved to a higher P and T environment and refractured along new paths that did not follow the breccia veins, and monomineralic pumpelleyite veins, patches, and sheets formed in an M2 event. These pumpelleyite veins formed before the quartz diorite was broken into blocks; they do not extend into the mélange matrix and are not related to the much younger metamorphism of the sandstone (see below), which also produced pumpelleyite. The date of M2, like M1, can only be bracketed between 165 and 83 Ma. The meta-granitoid block was exhumed by an undefined mechanism and incorporated in the developing King Ridge Road mélange sand matrix at 83 Ma along with thousands of other blocks. After olistostromal deposition the block was buried and metamorphosed to prehnite-pumpellyite conditions along with the rest of the unit, as noted above. History and Origin of Two Blocks II: Meta-Andesite Breccia Block 1500 Another example of an exotic block with a complex history comes from a 3 m block of meta-andesite breccia (block sample 1500), which was previously reported on in moderate detail (Pearce and Erickson, 2001); a pdf file of this poster is available in the Sonoma State University digital archive at http://hdl .handle.net/10211.1/169. This block is a complexly recemented breccia. About half of it is composed of ~1 cm fragments of a coarse-grained unfoliated omphacite-glaucophane-lawsonite-albite granofels (Fig. 9A). This assemblage has wholly replaced protolith minerals and structures; no textural or mineralogical relicts remain. No foliation or lineation is present. This assemblage is a blueschist facies (high-pressure) assemblage. The remainder of the breccia is a mass of fine-grained (~0.1 mm) omphacite-glaucophanelawsonite-albite granofels essentially identical mineralogically and chemically to the components of the coarse-grained phase (Fig. 9B). There are local irregular patches of purple glaucophane in all fragment sizes, and the rock is a type II specimen of Coleman and Lee (1963). The rock plots on the TAS diagram (Le Maitre, 2002) as a basaltic trachyandesite (Pearce and Erickson, 2001). A Pearce spidergram (1983) (Fig. 8) shows that the rock has a strong arc signature, similar to that of the granite block
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just described (but see the discussion below). Rare earth analysis gives (La)n/(Lu)n = 0.60. The initial basaltic trachyandesite magma formed at an unknown date and location. The rare-earth profile (Pearce and Erickson, 2001)—(La)n/(Lu)n = 0.60—matches that of a typical normal mid-oceanic-ridge basalt (N-MORB) (Wilson, 1989, p. 17–19) The arc signature in the spider diagram (Pearce, 1983) (Fig. 8), on the other hand, is strong. On careful inspection, however, it is noted that the Ce peak usually found in Pearce arc profiles is missing, and only the first four element tracers (K2O, Rb, Ba, and Th), all of which are water soluble, form the large peak in the left part of the diagram; their concentrations are comparatively high. The author speculates that the N-MORB magma chamber was invaded by water that contained high concentrations of these four soluble elements, producing this unusual composition. The trachyandesite magma just described intruded an active arc. Here a portion solidified, was subducted, and then completely mineralogically and texturally reconstituted in an M1 event to an isotropic omphacite-glaucophane-lawsonite-albite granofels (Fig. 9A), probably at >5 kbar pressure (e.g., Evans, 1990; Maruyama and Liou, 1988).
This meta-andesite metamorphite was then brecciated, which the author speculates was by hydrofracturing by a waterCO2 mixture escaping the subducting plate. The finer fragments of the breccia were then recrystallized to a glaucophane-lawsonite granofels much finer grained than the original one, but at the same conditions, forming a granofels that was mineralogically and texturally identical to the initial rock fabric (Fig. 9B). Microprobe analysis shows that identical phases, such as glaucophane in both coarse and fine breccias, have the same composition, suggesting that they formed under the same conditions. The initial granofels formation, the brecciation, and the recrystallization of the fragmented mass took place as part of the M1 event. This breccia was produced hot and at high pressure, unlike the breccia veins in the granitoid discussed earlier. Veins of aragonite cutting both textural phases of the breccia confirm that high-pressure metamorphic conditions existed throughout this interval (Johannes and Puhan, 1971). At some point the healed, brecciated granofels was exhumed to a lower pressure and temperature environment. Unlike the history of block 1503 there was no low-temperature brecciation event in this block. Instead, the next recorded event was formation of monomineralic pumpelleyite veins at lower grade M2
50 µm
Figure 9. Backscattered electron and petrographic images of basaltic trachyandesite breccia block 1500. (A) Backscattered electron image of high-pressure metamorphic phases with granofels texture. (B) 40× plane-polarized (PP) microscopic image of coarse clast and finer grained clasts in breccia; finer grained clasts are recrystallized to the same glaucophane-lawsonite-jadeite mineralogy as that composing the prebreccia granofels. All were produced in the M1 event for this rock.
Petrology of a Franciscan olistostrome with a massive sandstone matrix cutting the breccia. This M2 event may be the same as that which affected granitoid block 1503 just discussed. Following the above events, a ~3 m block of this brecciated meta-andesite was exhumed to the surface, and at 83 Ma buried by an accumulation of sand forming the matrix of the developing Kings Ridge Road mélange. These two detailed studies of arbitrarily selected exotic blocks in the mélange show that they have very different early histories of formation, subduction, metamorphism, and brecciation-rehealing. Their complexities suggest that many other histories would be found in other exotic blocks in the mélange assemblage; the author speculates that perhaps every block has some unique features. Exhumation Ages from Apatite Fission-Track Analyses An apatite fission-track study (Erickson et al., 2004) of the granite block described above and the surrounding sandstone matrix give dates of 35–38 Ma for exhumation cooling of the presently exposed mélange to ~100 °C. DISCUSSION Correlation of the King Ridge Road Mélange with Other Franciscan Units Matching to Subdivisions of the Franciscan Complex from Wakabayashi (1992a) The Franciscan Complex that underlies most of the Coast Range of California can be subdivided into four subregions on the basis of lithologic character and structural development (Wakabayashi, 1992a). His northernmost subregion, the northern Coast Range, extends from the Oregon border south to approximately the region of the study area, which is transitional into the structural style of what he defined as the San Francisco subregion. The northern Coast Range subregion is traditionally split into three parallel belts: the Eastern, Central, and Coastal Belts (Berkland et al., 1972). In the San Francisco subregion the Franciscan geology is best described as a series of underthrust nappes younging to the west (Wakabayashi, 1992a). Wakabayashi (1992a, Fig. 4) assigned the King Ridge Road mélange to an unnamed (“undifferentiated”) mélange nappe zone that structurally overlies the prehnite-pumpellyite sandstone-shale nappe of what was formerly named the Rio Nido terrane (Blake et al., 1984) and which structurally underlies the blueschist facies metavolcanics of the Ward Creek area and the Skaggs Springs Schist–Cazadero mélange. This implicitly correlates the mélange unit with a structural subzone within the Central Belt. The Central Belt of the Northern Coast Range is characterized by a generally sheared sedimentary matrix metamorphosed to prehnite-pumpelleyite facies, containing a great variety of exotic blocks up to many kilometers in size (i.e., Gucwa, 1975). Parts of it contain aragonite veins, indicating blueschist facies
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conditions (Terabayashi and Maruyama, 1998). The King Ridge Road mélange matches some of these characteristics but not the sheared character of the Central Belt matrix and the aragonite veins; it seems that the two units are a poor match. The author speculates that the Central Belt mélange was originally a sedimentary unit like the King Ridge Road mélange and was later tectonized, perhaps in a manner similar to the Las Tablas unit discussed below (Cowan and Page, 1975). Matching to Terranes of the Northern Franciscan Complex Established by Blake et al. (1984, 2002) Blake et al. (1984) used the tectono-stratigraphic terrane model (Jones, 1977) to subdivide a North Bay area, roughly corresponding to the northern part of the San Francisco Block of Wakabayashi (1992a) and the southern part of the northern Coast Ranges region of the three belts of Berkland et al. (1972), into several terrane subunits; each terrane is characterized by a unique stratigraphy, age, and sandstone composition (Jayko and Blake, 1984). The 1984 set of terranes was recently revised by Blake et al. (2002). This revision redefines about half the area covered by Blake et al. (2002) as a single expanded terrane, the Central terrane, described as a tectonic mélange containing Tithonian fossils (ca. 160 Ma) and highly lithic-rich sandstones. This new Central terrane is different than the Central Belt earlier discussed. The King Ridge Road mélange lithic petrofacies has identical sandstone composition to the old Central terrane, as defined by these authors in 1984 (Fig. 5), but it is much younger, at 83 Ma. On the basis of depositional age, the closest match to the Kings Ridge Road mélange is the Novato Quarry terrane, which has yielded Campanian fossils (Blake et al., 1984) and a maximum depositional age from detrital zircon chronology of 83 Ma (Snow et al., 2010). The sandstone compositions are reasonably comparable to the feldspathic petrofacies (Jayko and Blake, 1984). The Novato Quarry terrane, however, does not contain exotic blocks. The Las Tablas Body and Its Significance The exposures north of Wright’s Beach State Park show the incipient development of tectonic mélange from an older olistostrome mélange. There is one other published description of such a system in a more advanced stage of development that provides useful information on the connection between olistostromes and tectonic mélanges. Cowan and Page (1975) describe a Franciscan massive sandstone unit that they called the Las Tablas sandstone, which crops out ~20 km west of Paso Robles, California, and ~250 km SSE from Cazadero in the Franciscan Nacimiento block. The sandstone body is small, 2.0 × 0.75 km, and it is surrounded by a typical Franciscan tectonic mélange of phacoidal blocks in a sheared shale matrix. The sandstone is not referred to by the authors as a mélange, perhaps because it lacks included exotic blocks. Blocks up to 2 m are associated with the sandstone, but they are contained in conglomerate interbeds. In many ways, however, this sandstone body is very similar to the matrix of the King Ridge Road mélange; both are massive sandstone with rare interbeds of
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other clastic lithologies. Both are the same age (Campanian) and are metamorphosed to prehnite-pumpellyite facies, and both lack any foliation or lineation. Both lack megafossils. The sandstone matrix of the King Ridge Road mélange at Cazadero has these same characteristics. The Las Tablas sandstone body was deformed along its borders with an outward-intensifying strain gradient. The boundary zone grades radially into a tectonic mélange of sandstone phacoids in a foliated shale matrix. Other phacoids of greenstone, chert, conglomerate, and blueschist are in nearby exposures. It is clear that the bordering tectonic mélange is but incipiently developed in sandstone on the border of the Kings Ridge Road mélange, but it has nearly completely disrupted the similar sandstone of the Las Tablas assemblage. The geology at Cazadero and at the distant Las Tablas area suggests that these massive sandstones were formerly much more widespread in the Franciscan Complex. Modeling the Depositional System of the Sandstone Matrix of the King Ridge Road Mélange In an earlier part of this chapter the field and petrographic characteristics of the sandstone matrix of the mélange were described. On the QFL diagram of Figure 5 it was shown how all nine sandstone samples of the mélange plotted in areas of the figure labeled for a magmatic arc; seven samples were litharenites, and two were felsarenites. The QFL diagram of Figure 5 suggests that a magmatic arc that was going through at least two levels of erosional destruction was the source of the sands that make up the mélange matrix. The cluster of seven litharenites that plot in the transitional arc area of the figure probably came from one or more parts of the arc where unroofing of the plutonic core was just getting under way, whereas the pair of high feldspar samples probably comes from at least one other area where erosion had cut down to the plutonic core. The simplest hypothesis is to assume that this arc was the nearby KlamathSierran arc established along the western edge of the North American continent in Jurassic–Cretaceous time (Dickinson et al., 1982; Ducea, 2001). These sands were deposited to build the King Ridge Road mélange. The massive sandstone facies being described is modeled here as having been built up from a large number of sandy sedimentary density flows that flowed onto the deep ocean floor from one or more sources on the Cretaceous shelf and built up one or more massive deposits. Mulder and Alexander (2001) provide a four-type subdivision for all such sedimentary density flows; their classification will be used here. Arranged in order of decreasing density, these flow types are: debris flows, hyperconcentrated density flows, concentrated density flows, and turbidites, in that order. Debris flows are fundamentally different from the other three flow types; they contain significant clay and are cohesive, and their deposits (debrites) are characterized by a lack of sorting of clasts. The sands of the King Ridge Road mélange bear no resemblance to debrites.
The other three types in this classification are not internally cohesive and are classified together by the above authors as frictional flows, whose clasts are supported in the moving flow by grain interaction or turbulent fluid flow. Of these, hyperconcentrated density flows are the densest and have ~25% sand grains; concentrated density flows have variable amounts but much less, and turbidites have <9%. These varied grain proportions strongly affect the deposits they form. Concentrated density flows and turbidites have low enough particle concentrations so that grains can settle independently, and simple normal graded beds and Bouma sequences can form from them upon deposition. Hyperconcentrated density flows, on the other hand, have such a high particle concentration that grain to grain interaction and dispersive pressure prevent settling of single grains from the moving particulate bed (often called a grain flow), and during deposition, normal grading and/or bedding does not develop. Rather, as the moving bed slows, deposition of a massive unbedded unit takes place as the hyperconcentrated flow slows and dewaters and becomes ever more concentrated, and the lowest layer of the particulate bed deposits itself on grains settled earlier (Kneller and Branney, 1995, their figure 1). The basal layer of grains in the flow will merge with the deposit below, and simultaneously the flow will be replenished by grains sinking from above. There will not be a sharp interface at the depositional boundary (their figure 4) and no traction deposits (i.e., bedding). As a sediment flow moves downslope it will slow and finally stop; during this process a sequence of successive deposits of hyperconcentrated flows, concentrated flows, and finally turbidites will form, so that graded beds and even Bouma sequences may form toward the end of runout of the flow. In addition, other units, such as pelagic shale, may become interbedded with the sediment flow toward the end of its run. Because the boundaries of this massive sandstone deposit of the King Ridge Road mélange matrix have not even been defined, and only a small part has been studied in detail (Fig. 3), it would be premature to attempt to model the entire depositional system. The author will only speculate briefly on a model for the local part of it. The well-sorted and clean character of the sand indicates that most probably its ultimate source was sand of continental origin from the magmatic arc, drifting onto the Cretaceous shelf, intercepted by a submarine canyon and directed down into the Franciscan trench. Here, successive hyperconcentrated sediment flows, evolving as described above, buried exotic blocks coming from some other source as they accumulated. The whole assemblage gradually built up into a deposit several kilometers thick. Rarely sand supply was cut off, and only hemipelagic clay and silt were deposited for a short time, forming local thin shale and siltstone beds. Single beds or thin layers of strata are locally present in the generally massive sandstone and record highly variable bedding orientation (Fig. 3). This overall structure may be the result of folding or faulting, or, as an alternate hypothesis, these bed-displaying
Petrology of a Franciscan olistostrome with a massive sandstone matrix exposures may be within isolated blocks of older sandstone transported by the hyperconcentrated flows building the mélange (see Mulder and Alexander, 2001). CONCLUSIONS In the study area the King Ridge Road mélange is a sandstone-matrix olistostrome mélange constructed at least in part 83 m.y. ago. It contains many exotic blocks >2 m in size, at present roughly equal proportions of chert, high-pressure mafic metamorphites, and medium- to low-pressure, dominantly mafic metamorphites. Individual blocks have unique multistage metamorphic and tectonic histories. The blocks were exhumed, then transported into the trench. As they accumulated they were buried by long-continued hyperconcentrated sedimentary density flows of well-sorted sand, forming the present olistostrome mélange. Bedding did not generally develop, nor did grading. Toward the end of the runout of the sediment flows and olistostrome formation these deposits became interbedded with pelagic shale, and bedded sequences were formed. When tectonic stress was applied to this zone of the Franciscan deposits, the sandstone beds extended and broke up into phacoids, while the shale beds became foliated and flowed around the blocks; these new bodies formed an overall tectonic mélange at the point of transition from massive sandstone to interbedded sandstone and shale along the southwestern border of the olistostrome. As the mélange was partly subducted, its lower parts were metamorphosed into the prehnite-pumpelleyite facies. At some time after its peak metamorphism, but before 35–38 Ma, the mélange mass began to exhume and cool. Lower grade zeolite facies conditions developed and allowed younger laumontite veins to form. The mélange cooled to ~100 °C at 35–38 Ma and was fully exhumed after that time. ACKNOWLEDGMENTS Special thanks go to George Gehrels and the staff at the University of Arizona LaserChron center for the detrital zircon study. They provided the plots of Figures 6A and 6B. Funding for the U-Pb geochronologic analyses was provided by U.S. National Science Foundation EAR 0443387 and 0732436 grants. Sarah Roeske did the microprobe analyses at the University of California, Davis. Mattie Mookerjee provided valuable instruction on Adobe Illustrator. Gina Voight, Elyse Lord, Cathi Cari-Shudde, Jen Aaseth, and Mike Smith provided valuable advice and technical support; Marc Drucman and Ron Leu made a great many excellent thin sections. Thanks go especially to J. Wakabayashi for his extensive comments on several drafts of the text, and to J. Shervais and an anonymous reviewer, whose comments on an early draft greatly improved the final product. Last, thanks to the many residents of the Cazadero area who gave the author permission to enter their property. As always, all ideas and data presented herein are the sole responsibility of the author.
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Society of America Bulletin, v. 100, p. 446–457, doi:10.1130/0016 -7606(1988)100<0446:CDIABI>2.3.CO;2. Hsu, K.J., 1968, The principles of mélanges and their bearing on the Franciscan-Knoxville paradox: Geological Society of America Bulletin, v. 79, p. 1063–1074. Jayko, A.S., and Blake, M.C., Jr., 1984, Sedimentary petrology of graywacke of the Franciscan Complex in the Northern San Francisco Bay Area, California, in Blake, M.C., Jr., ed., Franciscan Geology of Northern California: Society of Economic Paleontologists and Mineralogists, Pacific Section, v. 43, p. 121–134. Johannes, W., and Puhan, D., 1971, The calcite-aragonite transition, reinvestigated: Contributions to Mineralogy and Petrology, v. 31, p. 28–38, doi:10.1007/BF00373389. Jones, D., 1977, Wrangellia: A displaced terrane in northwestern North America: Canadian Journal of Earth Sciences, v. 14, p. 2565–2577, doi:10.1139/e77-222. Kneller, B.C., and Branney, M.J., 1995, Sustained high-density turbidity currents and the deposition of thick massive sands: Sedimentology, v. 42, p. 607–616, doi:10.1111/j.1365-3091.1995.tb00395.x. Le Maitre, R.W., ed., 2002, Igneous Rocks: A Classification and Glossary of Terms: Cambridge, UK, Cambridge University Press, 232 p. Maruyama, S., and Liou, J.G., 1988, Petrology of Franciscan metabasites along the jadeite-glaucophane type facies series, Cazadero, California: Journal of Petrology, v. 29, p. 1–37. Mulder, T., and Alexander, J., 2001, The physical character of subaqueous sedimentary density flows and their deposits: Sedimentology, v. 48, p. 269– 299, doi:10.1046/j.1365-3091.2001.00360.x. Pearce, J.A., 1983, Role of the sub-continental lithosphere in magma genesis at active continental margins, in Hawkesworth, C.J., and Norry, M.J., eds., Continental Basalts and Mantle Xenoliths: Natwick, UK, Shiva, p. 230–249. Pearce, L.S., and Erickson, R.C., 2001, Petrology of a brecciated glaucophanelawsonite meta-arc basalt block, Franciscan Complex, Sonoma County, CA: Geological Society of America Abstracts with Programs, v. 33, no. 3, p. 47. Pettijohn, F.J., Potter, P.E., and Siever, R., 1987, Sand and Sandstone (2nd edition): New York, Springer-Verlag, 553 p. Raymond, L.A., 1984, Classification of mélanges, in Raymond, L.A., ed., Melanges: Their Nature, Origin, and Significance: Geological Society of America Special Paper 198, p. 7–20. Rubatto, D., 2002, Zircon trace element geochemistry: Partitioning with garnet and the link between U-Pb ages and metamorphism: Chemical Geology, v. 184, p. 123–138, doi:10.1016/S0009-2541(01)00355-2.
Schuster, D.C., 1980, The nature and origin of the late Precambrian Gwna melange, North Wales, United Kingdom [Ph.D. thesis]: Champaign, University of Illinois at Urbana-Champaign, 383 p. Şengor, A.M.C., 2003, Ophiolite concept and the evolution of geologic thought, in Dilek, Y., and Newcomb, S., eds., Ophiolite Concept and the Evolution of Geological Thought: Geological Society of America Special Paper 373, p. 385–445. Silver, E.A., and Beutner, E.C., 1980, Mélanges: Geology, v. 8, p. 32–34, doi:10.1130/0091-7613(1980)8<32:M>2.0.CO;2. Snow, C.A., Wakabayashi, J., Ernst, W.G., and Wooden, J.L., 2010, Detrital zircon evidence for progressive underthrusting in Franciscan metagraywackes, west-central California: Geological Society of America Bulletin, v. 122, p. 282–291, doi:10.1130/B26399.1. Terabayashi, M., and Maruyama, S., 1998, Large pressure gap between the Coastal and Central Franciscan belts, northern and central California: Tectonophysics, v. 285, p. 87–101, doi:10.1016/S0040-1951(97)00194-7. Wakabayashi, J., 1992a, Nappes, tectonics of oblique plate convergence, and metamorphic evolution related to 140 million years of continuous subduction, Franciscan Complex, California: Journal of Geology, v. 100, p. 19–40, doi:10.1086/629569. Wakabayashi, J., 1992b, Metamorphism and tectonic origin of Franciscan metabasites and a field trip to three localities in the San Francisco Bay area, in Anonymous, Special Publication—California Department of Conservation, Division of Mines and Geology, p. 1–11. Wakabayashi, J., 2011, this volume, Mélanges of the Franciscan Complex, California: Diverse structural settings, evidence for sedimentary mixing, and their connection to subduction processes, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, doi:10.1130/2011.2480(05). Wakabayashi, J., and Dumitru, T., 2007, 40Ar/39Ar ages from coherent, highpressure metamorphic rocks of the Franciscan Complex, California: Revisiting the timing of metamorphism of the world’s type subduction complex: International Geology Review, v. 49, p. 873–906, doi:10.2747/0020-6814.49.10.873. Wilson, M., 1989, Igneous Petrogenesis: London, Unwin Hyman, 466 p.
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The Geological Society of America Special Paper 480 2011
Sedimentary block-in-matrix fabric affected by tectonic shear, Miocene Nabae complex, Japan Soichi Osozawa Department of Earth Sciences, Graduate School of Science, Tohoku University, Sendai 980-8578, Japan Terry Pavlis Department of Geological Sciences, University of Texas at El Paso, El Paso, Texas 79968, USA Martin F.J. Flower Emeritus, Department of Earth and Environmental Sciences, University of Illinois at Chicago, Chicago, Illinois 60607-7059, USA
ABSTRACT Mélanges represent a significant part of the Miocene Nabae accretionary complex. Such mélanges show sheath folds with D1 axial plane pressure-solution cleavage, whereas the coherent unit shows asymmetric folding with D1 slaty cleavage. In addition, the mélanges are characterized by D1 asymmetric shearing, which includes both thrust and right-lateral-sense components, in contrast to D1 pure shear that characterizes the coherent unit. Thus, this tectonic style acted on the climax of prism development can be referred to as a tectonic mélange. However, because the D1 shear displacement is almost negligible, D0 normal faults and basaltic dikes, operated when matrix sediments were not consolidated, are not disrupted. Oceanic materials such as basalt and chert cannot be incorporated into terrigenous matrix, given the small displacement associated with the D1 shearing. Exotic blocks of chert and sandstone show D minus 1 (D–1) cleavage, which is apparent in the older, probably Eocene, accretionary prism. When this prism was exhumed, it supplied debris to the Miocene trench, and then underwent additional D1 deformation, which included the above asymmetric shearing. This sedimentary and two-way-street tectonic process was recycled within the prism as the latter developed. Thus, as the block-in-matrix fabric was originally sedimentary and labeled D0, the tectonic mélange process that forms block-in-matrix fabric is only conjectural for the Nabae complex. Also it is suggested that these deformations are not progressive nor distinctive for each other.
Osozawa, S., Pavlis, T., and Flowers, M.F.J., 2011, Sedimentary block-in-matrix fabric affected by tectonic shear, Miocene Nabae complex, Japan, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 189–206, doi:10.1130/2011.2480(08). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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INTRODUCTION Mélanges are large, mainly sedimentary, rock bodies characterized by stratal disruption and the development of blocks of various sizes and/or shapes, typically dispersed in a finer grained sedimentary matrix (Cowan, 1985; Brandon, 1989). When this block-matrix relationship is recognized, the term block-in-matrix fabric is typically applied (Cowan, 1985). Mélanges commonly contain blocks of oceanic materials such as chert, limestone, and mid-oceanic-ridge basalt (MORB) and oceanic island basalt (OIB), and this type of mélange is typically interpreted as a structural feature, formed along a subduction zone during the accretion process. The term mélange was applied for the Franciscan complex in the California Coast Range by Hsü (1968), and the development of the mélange concept was recently reviewed by Şengor (2003). In this review, Şengor (2003) advocated that the mélanges are virtually entirely of tectonic origin, mainly because of their along-strike continuity for thousands of kilometers, coincident with known highly stretched zones. This view of tectonic mélanges was strengthened by modeled studies of Cloos (1982) and his subsequent further development of the subduction channel model (Shreve and Cloos, 1986). Studies in the AlaskanAleutian forearc and Japan further showed that most mélanges have been affected by tectonic shearing, represented by asymmetric shear structures recognized at mesoscopic to microscopic scales (Fisher and Byrne, 1987; Needham, 1987; Needham and Mackenzie, 1988; DiTullio and Byrne, 1990; Kimura and Mukai, 1991; Onishi and Kimura, 1995; Hashimoto and Kimura, 1999). Accordingly, tectonic mélanges have been genetically connected to underplating and related duplexing processes at convergent margins as well as correlated with tectonic processes associated with the seismogenic zone of subduction megathrusts (Ikesawa et al., 2005). Large tracts of the Japanese forearc are also composed of mélange, and like other areas have seen historical interpretations that ranged from olistostromal origins (Taira et al., 1982) to entirely tectonic (Onishi and Kimura, 1995). Taira et al. (1982), for example, recognized a block-in-matrix fabric and distinct differences of radiorarian age between oceanic blocks and a matrix derived from terrigeneous sediment, to conclude a predominantly olistostrome origin for the Shimanto accretionary prism mélange. The lead author studied several field examples and reached a similar conclusion of the predominance of the olistostromal process in these examples (Osozawa, 1984, 1992). Later studies recognized, however, that many minor structures showed clear evidence for asymmetric shear, and as a result Taira et al. (1992) reinterpreted the Shimanto mélange as dominantly a tectonic mélange. There is now widespread agreement that the mélangeforming process is polygenetic, with at least three major processes contributing to the ubiquitous stratal disruption and lithologic mixing (DelaPierre et al., 2007): (1) surficial processes of submarine mass wasting (olistostrome, debris flow,
slumping, etc.); (2) excess fluid-pressure–driven injections, including features ranging from small-scale clastic diking to large-scale diapirism, which may or may not reach the seafloor; and (3) tectonic disruption, including fault-related processes (fracturing, shear-surface development, catastrophic becciation during seismic slip, etc.) and larger-scale structural shuffling through duplexing and out-of-sequence thrusting. In general, it is difficult to distinguish among these processes in field studies, particularly in distinguishing non-tectonic processes in mélanges that contain a strong tectonic overprint. Thus it is not surprising that many studies have emphasized the tectonic component of mélange formation (e.g., Şengor, 2003). Nonetheless, because other processes are clearly involved in mélange formation, it is important to document cases where these sedimentary processes are recognizable. In this chapter we consider such a case in analysis of the mélange formation in the Nabae complex of Japan. We show evidence supporting the concept that the mélange zone was strongly affected by development of a shear zone, but we show evidence that block-in-matrix fabrics originated as sedimentary structures, probably submarine debris flows. This evidence arises from a series of observations from a microscopic to macroscopic level of the structural evolution of the fabric, using techniques similar to those described in recent works on similar mélanges (Osozawa et al., 2008, 2009). GEOLOGIC SETTING OF THE NABAE COMPLEX The Nabae complex constitutes part of the youngest portion of the Shimanto accretionary zone (Fig. 1), located at Cape Muroto, at the end of the Muroto Peninsula, Shikoku, Japan (Taira et al., 1980). The accretionary prism is being drilled just offshore by the Integrated Ocean Drilling Program (IODP) for the purpose of investigating the seismogenic zone of the Nankai Trough. Thus, studies of onshore equivalents to this drill site are important for assessment of the drilling results in terms of seismogenic processes. The Nabae complex contains mélange as well as block-inmatrix fabrics (Hibbard et al., 1992). However, gabbroic and basaltic dikes and normal faults are also encountered in these mélanges (Osozawa, 1993); a layer of probable debris flow is reported as an intercalation (Taira et al., 1980), and a folded chert block with a muddy injection has been observed (Hibbard et al., 1992). These observations are unusual in accretionary complexes and imply that sedimentary and other processes have been important in mélange formation. Thus, this area was chosen for focused studies in an attempt to determine how sedimentary processes might have contributed to formation of these block-in-matrix fabric mélanges. The Nabae complex consists of two mélange units and two coherent units (Hibbard et al., 1992). From northwest to southeast, these include the Hioki mélange, Tsuro assemblage, Sakamoto mélange, and Misaki assemblage (Hibbard et al., 1992; Fig. 1). Gabbroic and basaltic dikes of the Maruyama suite are
Sedimentary block-in-matrix fabric affected by tectonic shear observed to intrude the Nabae complex, and have been interpreted as post-cleavage by Hibbard et al. (1992) and pre-cleavage by Osozawa (1992). They are discussed further in this chapter. The Nabae complex is overlain by the Shijujiyama Formation, which was deposited in a forearc basin (Hibbard et al., 1992). Both the Hioki and Sakamoto mélanges contain small amounts of basalt and chert as exotic blocks <10 cm in diameter (Fig. 2). The age of the chert may be Paleocene and Eocene
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(occurrence of Eocene radiolarians was reported by Taira et al., 1980), although the radiolarians are poorly preserved, challenging clear identification. The chert lithologies resemble those in a Paleocene (Osozawa, 1992, 2006) exotic block in the Sakihama mélange (DiTullio and Byrne, 1990) in the Eocene Shimanto zone. These pale-green cherts appear to have been deformed as soft sediments are soft and not silicified, which contrasts with the hard and siliceous cherts in Cretaceous and older accretionary
Figure 1. Index and geological maps of the Nabae complex. After Osozawa (1993, 2006).
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zones in Shikoku. The mélanges in the Nabae complex characteristically contain many large fragments of sandstone debris. Sandstone debris is relatively rare in the older accretionary complex. The mélange, however, intercalates turbidite, one of which, in the Hioki mélange, is represented by a fragmented barnacle bed (Sakai, 1987). The structural base of these mélanges consists of hemipelagic black mudstone and silicic tuff (Fig. 1), recognized as coherent layers rather than blocks. The age of the muddy matrix is early Miocene, as indicated by the presence of radiolarian species such as Stycocolys cf. wolffii, Calocycletta sp., Cyrtocapsella tetrapera, C. japonica from the hemipelagic mudstones at several localities in the Nabae complex, including the base of the mélange (Osozawa, 1993). In contrast to the Cretaceous mélange in the northern Shimanto zone (Taira et al., 1980), overlying coherent turbidites are structurally lacking owing to thrust disruption (Fig. 1).
The Tsuro assemblage consists of variously colored (mostly black, gray, white, green, and red) mudstone and silicic tuff overlain by turbidites. The turbidites stratigraphically lie to the northwest of the mudstone and tuff, and, judging from wellpreserved sedimentary structures such as grading, are interpreted as stratigraphic top and facing is NW. Such hemipelagic and terrigenous sequences are repeated by SE-directed thrusts, and thus NW-facing tops are consistent with large-scale duplication along thrust systems (Osozawa, 1993). A thrust sheet of the Tsuro assemblage intercalates the ~10-m-thick mélange as a thin, stratigraphically disrupted, sedimentary layer between hemipelagic mudstone and turbidite (Fig. 1, at locality 5). Hibbard et al. (1992) recognized this unit and referred to it as a contact mélange (see Hibbard et al., 1992, for his definition). This mélange also contains exotic blocks of chert and basalt, including a large chert block reported by Hibbard et al. (1992), which is also described
Figure 2. (A, B) Green chert block in the Hioki mélange at locality 1; A after Osozawa (2006, figure 4A). (C, D) Basalt fragment in the Hioki mélange at locality 1. Rough, spaced, and anastomosing D1 cleavages are visible. Thin section after Osozawa (2006, figure 4B). Open and crossed nicols, respectively. Scale bar is 1 mm.
Sedimentary block-in-matrix fabric affected by tectonic shear in this chapter. Conglomerate and coarse-grained sandstone overlie turbidites in the northwesternmost thrust sheet, constituting a coarsening upward sequence and providing further support that stratigraphic tops face NW. The Misaki assemblage consists of turbidites, showing much folding including sheath folding (Hibbard and Karig, 1987), whereas black mudstone and silicic tuff occupy the base of the thrust sheet (Fig. 1; Osozawa, 1993). The Maruyama intrusive suite comprises gabbro sills and basalt dikes, the latter intruded into the mélange. At two localities, including Maruyama, the dikes show two distinct trends orthogonal to each other (Fig. 1). Unaltered glass-chilled rinds are exceptionally preserved at locality 6 (Fig. 1), although large amounts of secondary calcite and chlorite have formed at most localities. The Shijujiyama Formation consists mostly of mildly deformed sandstone, although its base is made up of volcanic breccia and pillow basalt, mostly found as float blocks. The breccia contains secondary minerals of chlorite and prehnite. GEOLOGIC STRUCTURE In the Nabae complex we recognize five generations of structures: D minus 1 (D–1), D0, D1, D2, and D3 deformation phases, in contrast to the early, intermediate, and late deformation phases defined by Hibbert et al. (1992) and Osozawa (1992). D–1 and D1 cleavages are associated with recrystallization of white micas and correspond to what we refer to here as M minus 1 (M–1) and M1 metamorphic phases. D1 pressure-solution cleavages are the most conspicuous and are widespread, thus they serve as a general guide for correlating deformation stages described in previous work (Hibbard et al., 1992; Osozawa, 1993). The D1 cleav-
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age generally strikes northeast-southwest, and dips southeast at a high angle (Fig. 3). Evidence for the earlier two deformation phases D–1 (M–1) and D0 are important for addressing mélange genesis problems. D–1 Cleavage within a Chert Block and Mud Injection A chert block up to 15 m in length occurs in the mélange unit within the Tsuro assemblage (Hibbard et al., 1992). This chert is folded about a north-northeast–trending axial plane dipping eastsoutheastward at a high angle (Fig. 4A). D1 cleavage planes in the folded chert and muddy matrix have the same strike and dip, and are axially planar to D1 fold structures (Fig. 4A), indicating that both were involved in D1 folding. However, other cleavages, folded by D1 folds, are visible in the chert, indicating an earlier deformation referred to here as D–1 pressure solution and associated metamorphism that generated M–1 sericites (Figs. 4C, 4D). Another chert block at locality 1 (Fig. 4B) shows the same relationships, providing further support for an older D–1 event. A muddy clastic dike, labeled D0, is intruded into chert, orthogonal to bedding, with a width of ~5 cm (Figs. 5A, 5B). Microscopic study clearly shows that D–1 pressure solution cleavages and flattened radiolarians in chert are cut by the dike (Figs. 5C, 5D). Although D1 cleavages within the dike are weakly developed parallel to the D–1 cleavages in adjacent cherts, flow structures formed by grain boundary sliding are predominantly parallel to the contact and orthogonal to the D–1 cleavages (Figs. 5C, 5D). An important implication is that D–1 cleavage and related M–1 mica crystallization, probably linked to Eocene subduction that formed the Sakihama mélange, had penetrated the Paleogene
Figure 3. Stereograms of cleavage, fold axis, and bedding of the Nabae complex. Some bedding dips SE at high angles, and the beds are overturned. Stratigraphic top is, however, NW, which is a consequence of D3 indentation. The diversity of fold axes indicates sheath folding.
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Chert
Chert
D1 D 1 axial axi xial a p plane la ane ne cleavage cclle ea avva ag ge e Folded Fo old ded ed D-1 D D-1 ccl cleavage le ea avva age ge
D1 D 1 axial axxiia all p plane la ane n clea cl cleavage eava eava ea vage ge
D1 ffault D1 au a ultt
Fo F Folded old ded d D 1 ccleavage Dle ea avvag va ag ge ge D-1
Rad o Ra Radiolaria ollar aria a
Radi Ra Radiolaria diio d ollar ar a
Figure 4. (A) D1 fold, involving both mudstone matrix and a chert block, in contact mélange of the Tsuro assemblage at locality 5. Axial plane cleavages are visible in mudstone. Filled rectangle: Sample point of thin section C and D. (B) D1 fold in chert block in the Hioki mélange, at locality 1. Filled rectangle: Sample point of thin section E and F. (C, D) Fold accompanied by D1 axial plane cleavage and represented by folding of D–1 cleavage and chert bedding. Folded D–1 cleavages are associated with M–1 sericites, but M1 sericite in this figure is unclear. Spots are statically recrystallized calcite. Thin section after Osozawa (2006, figure 8F). Open and crossed nicols, respectively. (E, F) D1 folded D–1 cleavages and M–1 sericites. M1 sericites are also along D1 cleavages and fault. Brecciation and faulting, left side of axial plane, is D1 effect, and a radiolarian test is affected by D1 pressure solution. Open and crossed nicols, respectively. Scale bar in C–F is 1 mm.
Sedimentary block-in-matrix fabric affected by tectonic shear chert before its incorporation in the matrix. Thus the apparent sequence in the chert block is that the D0 mud dike was injected after the D–1 cleavage development and consolidation of the chert, suggesting that the chert block was exhumed and then mixed into the unconsolidated matrix, and then subjected to the main tectonic deformation of D1. The simplest mechanism for this mixing is a sedimentary process, although mud diapirism cannot be discounted. D–1 Cleavage within Exotic Sandstone in Debris Flow In the Hioki mélange at Maruyama, a 1.5-m-thick conglomerate layer is intercalated within the mélange (Taira et al., 1980). Cropping out at a wave-cut terrace, subrounded sandstone blocks up to 30 cm in diameter are dispersed within the sandy matrix
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(Fig. 6A). Larger blocks are concentrated close to the top and bottom parts of the layer, the largest being at the top. Their long axes tend to have been rotated clockwise relative to the layer, indicating a history of dextral shear. In thin section, the sandstone fragments show significant diversity in terms of their contents of quartz, feldspar, and other minerals; detrital grain size; variations of strained quartz; and cleavage development (Figs. 6B–6D), and they clearly reflect multiple sources. Radiolarians and foraminiferans are also present in two of the sampled fragments (Fig. 6E). Other than sandstone, fragments of chert, basalt, granitic rock, and metamorphic rock are also present (Fig. 6F). A pressure solution cleavage (Figs. 6C, 6D) is present in many of these fragments, and this cleavage is randomly oriented in the blocks, indicating that this cleavage is a pre-sedimentary fabric. All fragments in the matrix
Figure 5. (A, B) Mud injection in chert block, same as shown in Figure 4A. B is after Osozawa (2006, figure 8D). (C, D) D–1 cleavages crosscut by injection. M–1 sericites and deformed radiolarians are aligned along the D–1 cleavages in the block, and detrital white micas are aligned along the flow structures in the matrix. D1 cleavage and M1 sericite are mildly developed parallel to the D–1 cleavages in this case. Thin section after Osozawa (2006, figure 8E). Open and crossed nicols, respectively. Scale bar in C and D is 1 mm.
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Figure 6. (A) Sandstone debris-flow layer in the Hioki mélange at locality 1. After Osozawa (2006, figure 4C). (B) Floated sandstone fragment. Crossed nicols. (C, D) D–1 cleavage within a sandstone fragment. D1 cleavages with sericite are almost parallel to the base of figures, but are anastomosing and also around fragments. Open and crossed nicols, respectively. (E) Radiolarians and foraminiferans within muddy sandstone fragments. Crossed nicols. (F) Granite fragment. Crossed nicols. B–F: Thin sections after Osozawa (2006, figure 4D). Scale bar in B, D, E, and F is 1 mm.
Sedimentary block-in-matrix fabric affected by tectonic shear are subrounded (Figs. 6B–6F), and where in contact with each other they show D1 pressure solution effects (Figs. 6B–6F). D1 cleavage, thus defined, was developed parallel to the conglomerate layer and oblique to the randomly oriented D–1 cleavage within the fragments, clearly suggesting that these two cleavages are distinct and not a consequence of progressive rotation of a cleavage. Although these clasts are polygenetic, they are important in the context of this study in that they show an exhuming source with a well-developed pressure solution cleavage. Thus we infer that this source is the same as the chert blocks and assign this pre-sedimentary fabric to D–1. D0 Normal Faults D0 normal faulting occurs in both the mélanges and coherent units, although it is more easily recognized in the coherent units (Fig. 7A; Osozawa, 1993). The fault planes are typically marked by calcite and quartz veining, and are clearly marked by lowangle striations. Some sandstone clastic dikes are observed along the fault planes (Fig. 7B; also Osozawa, 1993, Fig. 5), showing that normal faulting was either the result of a surficial, soft sediment deformation or that it occurred in the shallow part of the subduction interface under high fluid pressures and underconsolidation due to elevated fluid pressures. In addition, the crosscutting of faults by D1 cleavage is a clear indication that the faulting was D0, predating D1 (Figs. 7C–7F). When faults are restored to their original orientation (Osozawa, 1993), it is apparent that the intermediate stress axis was northwest-southeast, which is interpreted to represent the trend of a subduction but still near an actively spreading mid-oceanic ridge (Osozawa, 1993). We interpret these faults as either the product of surficial extension on the trench slope prior to incorporation of the material into the accretionary complex, or early structures formed near the deformation front of the subduction zone under elevated fluid pressures. In the latter case, the extension could represent the kinematic effects of ridge subduction (Osozawa, 1993). Although variable phases of bed-parallel extension that is seen in many other mélanges (e.g., Cowan, 1985; Needham, 1987); these are included in D1 asymmetric Riedel shear. D0 Gabbroic and Basaltic Dikes Gabbroic and basaltic dikes occur at several localities within the complex and are clearly igneous intrusions into the blockin-matrix fabric. As for the normal faults described above, these basaltic dikes are clearly crosscut and overprinted by the D1 cleavage (Fig. 8), and accordingly designated D0 here. An important observation is that cleavages are not cross-cut by the injections, indicating no pre-injection cleavage and no progressive sequence for D0 and D1. In addition, mud injections observed within (characteristically unaltered) glass chill margins (Figs. 8A, 8C, 8E) indicate that the dike was injected into mud rather than into mudstone, before D1 pressure solution acted on the mudstone. This soft sediment deformation during D0 is consistent with the obser-
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vations along D0 normal faults and clastic dikes described above. The basaltic dike trends are similar to those of the normal faults, also indicating a NW-SE–directed intermediate stress axis consistent with that of a subducting ocean ridge (Osozawa, 1993). Additionally, the gabbro intrusion illustrates that the amount of shearing associated with D1 is very small, as shown in the outcrop photos and the aerial photo in Figure 9, where the dike trends orthogonal to the D1 cleavage, axial planes, and shear fabric, but its margin has not been displaced. Notice that in the areas shown in Figure 9, the mélange contains many exotic Paleogene basalt and chert blocks, and significant shear would be expected along the block margins if the block-in-matrix relationships were formed by tectonic strain. For example, the thickness of the chert in the Sakihama mélange is on the order of tens of meters (Osozawa, 2006), which represents a minimum required displacement for incorporating chert into the mélange by shearing. D1 Cleavage and Shear Structure The most pervasive structure in the Nabae complex is D1 pressure solution cleavage, which, as pointed out by Hibbard and Karig (1987), is rough, spaced, and anastomosing. The latter is indicated by alignment of metamorphic white micas (e.g., Figs. 2D, 6E). Pressure solution of microfossil tests and lithic and mineral fragments is also evident (Figs. 10A, 10B). Fabric asymmetries indicative of shear are also observed in the mélanges, whereas contrasting shear structures are not observed in the coherent Tsuro assemblage. Asymmetric pressure shadows (Figs. 10C, 10D) and fringes, and Riedel shear planes, are often observed in the mélanges. Their asymmetry indicates a southeast-directed orthogonal shear and a northeast-directed right-lateral component. The pressure shadows and fringes consist of sericites and chlorites directly related to the pressure solution cleavage, indicating that both are syn-D1. The contact between the Tsuro assemblage and the underlying Sakamoto mélange is bounded by a thrust fault marked by calcite and quartz veining. The thrust crosscuts the D1 cleavage (Figs. 10E, 10F) and therefore represents a brittle, post-D1 deformation. However, the thrust does show evidence of cleavage development along the fault (Figs. 10E, 10F), and the thrusting is considered to be a late feature of D1 deformation. Thrust faults bounding the imbricated sheets of the Tsuro assemblage (Fig. 1) are all of this type of D1 vein faults (Fig. 11). Such imbrication can be explained as a duplex in map scale (Fig. 1), but it should be noted that the Tsuro assemblage is a coherent unit. Focused D1 shear may have formed these D1 thrust faults, but distributed shear may have formed broader shear zones. Folds in the Nabae complex, both in coherent units and mélanges, exhibit axial plane cleavage (Fig. 12; Fig. 3), which corresponds to D1 pressure solution. The folding thus basically reflects D1 deformation, which, although showing SEdirected asymmetry, differs in style between the coherent units and mélanges. Cylindrical folding is observed in the coherent units, and non-cylindrical folding in the mélanges. However, the
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Figure 7. (A) Conjugate normal faults occur in an intercalated sequence of silicic tuff and siliceous mudstone: the Tsuro assemblage at locality 4. Plan view. (B) Clastic dikes intrude into SE-dipping silicic tuff and siliceous mudstone: the Tsuro assemblage, just southeast of locality 5. Dikes are aligned parallel to the normal fault. (C–F) Normal fault with quartz and calcite veins penetrated by D1 cleavage, in the Tsuro assemblage at locality 4. Fault separation of contact between silicic tuff and mudstone visible at upper right side of images is 5 cm, but D1 cleavages are not separated by the veined fault; hence the cleavages postdate the fault. Thin section after Osozawa (1993, figure 6) and Osozawa (2006, figure 7B). Scale bar in B–F is 1 mm.
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Figure 8. (A–D) Chilled margin of basaltic intrusive, the Maruyama intrusive suite, in the Sakamoto mélange, at locality 2, collected from outcrop shown in Figure 9A, showing the basalt as not exotic. Unconsolidated mud is injected into the chilling glass, showing the intrusion precleavage. Such an excellent example of mud injection into basaltic glass is reported in the Miocene Setogawa complex, a correlative part of the Shimanto zone (Osozawa et al., 1990). Thin section after Osozawa (1993, figure 7). (A, C) Open nicols. (B, D) Crossed nicols. Unaltered glass is preserved, and its chemical composition determined by inductively coupled-plasma mass spectroscopy, showing island-arc basalt affinity. (E) Pressure-solution cleavage on glass, plagioclase, and secondary calcite in basalt. Open nicols. (F) Pressure-solution cleavage on mudstone. Open nicols. In crossed nicol images, sericite illustrates anastomosing cleavage. Scale bar in A–F is 1 mm.
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Figure 9. (A) Basaltic dike intruded into the Sakamoto mélange at locality 2. The intrusion and cleavages have orthogonal trends at this locality. The mélange also contains chert and basalt fragments. Thin section images are shown in Figure 8. After Osozawa (2006, figure 5A). (B) Basalt dike at locality 2. The intrusion is gently folded in contrast to the tightly folded mudstone, reflecting the lithologic contrast. (C) Aerial photograph of a gabbro dike and sill, Hioki mélange at Maruyama (locality 1). The gabbro dike is not dislocated by the mélange fabric but is crosscut by it. After Osozawa (2006, figure 4E). V marks gabbro dike.
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Figure 10. (A, B) D1 pressure solution cleavage, Tsuro assemblage at locality 4. A radiolarian is truncated by cleavage. Open and crossed nicols, respectively. After Osozawa (2006, figure 7C). (C, D) D1 asymmetric pressure shadow with dextral sense of shear, contact mélange of the Tsuro assemblage, at locality 5. Open and crossed nicols, respectively. After Osozawa (2006, figure 8B). (E, F) D1 thrust between the Tsuro assemblage and Sakamoto mélange, just southeast of locality 4. Thrusting was accompanied by quartz and calcite veining. Open and crossed nicols, respectively. After Osozawa (2006, figure 7D). Scale bar in A–F is 1 mm.
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Figure 11. D1 thrusts on outcrops. Hanging wall is green mudstone, and footwall is turbidite. Stratigraphically, mudstone below and turbidite above, and such successions, are also observed. (A) Tsuro assemblage at locality 6. (B) Tsuro assemblage at locality 7.
coherent Misaki assemblage exhibits sheath folding (Fig. 12D; Hibbard and Karig, 1987) as also seen in the mélanges. Hibbard and Karig (1987) and Hibbard et al. (1992) described the D1 fold systems as landward (NW) vergent, whereas Osozawa (1992) described the structures as SE vergent. We are uncertain of the origin of this discrepancy. We believe that the stratigraphic top or facing of the Tsuro unit is NW, duplexing and thrusting is SE directed, and asymmetric shear in mélanges is top to the SE, as previously mentioned. In addition, the gravitational top of the fractionated gabbroic sill at Cape Muroto near locality 3 is NW (pointed out since Yajima, 1972), and the fold vergence is SE (Figs. 12A–12C). The nature of exotic clast contacts observed in thin section gives insight into the relationship of D1 deformation and the incorporation of the exotic clasts. The contacts between exotic clasts and the surrounding matrix are commonly pressure solution cleavages (Fig. 13), although in some cases at least one of the contacts is surely a Riedel shear surface. This shows that Riedel shears alone cannot explain the incorporation of isolated exotic fragments into matrix. D2 and D3 Structures The Shiina-Narashi fault (Hibbard and Karig, 1987) (Fig. 1) marks the contact between the Miocene Nabae complex and the Eocene Muroto complex. This feature reflects brittle D2 faulting, and is associated with several centimeters of gouge. Along with D2 folding, defined by warping of D1 cleavage, D2 structures are developed only locally in the Nabae complex. Whereas most of the Shimanto belt trends ENE-WSW, the NE-SW orientation of the Nabae zone is referred to as the Muroto flexure (Hibbard and Karig, 1987; Fig. 1). The Muroto flexure has effectively bent the D2 Shiina-Narashi fault—also overturning the Nabae complex—and is therefore labeled D3.
DISCUSSION Dominance of Sedimentary Processes in the Formation of Block-in-Matrix Fabrics The principal observation from this study is that several lines of evidence suggest that blocks incorporated into the Nabae complex mélanges carry evidence of ductile deformation prior to their incorporation into the mélange matrix. This early fabric, referred to here as D–1, restricted to blocks, was subsequently overprinted by a soft sediment deformation (D0) representing either a surficial process or early deformation of underconsolidated sediments at the deformation front of a subduction zone. These soft sediment structures were then overprinted by a penetrative, pressure solution cleavage and fold systems of the main ductile deformation phase (D1) and two more localized fold complexities (D2 and D3) that are largely irrelevant to this discussion. Like most mélanges, the D0 deformation of the Nabae complex poses a problem in distinguishing the importance of surficial
Figure 12. (A) Thrust associated with SE-directed asymmetric fold, in an intercalated sequence of silicic tuff and siliceous mudstone: the Tsuro assemblage just southeast of locality 5. D1 axial plane cleavage is parallel to the thrust. (B) SE vergent fold, with secondary NW vergent folds in overturned limb, in contact mélange of the Tsuro assemblage, at locality 5. Sandstone layer defines fold, but basalt block is included. (C) SE vergent, asymmetric fold in turbidites, the Misaki assemblage at locality 3. (D) Sheath fold in turbidites, the Misaki assemblage at locality 3. Same outcrop photo is in Hibbard and Karig (1987, figure 6a). Sheath axis plunges NE, but axial plane cleavage can be defined. Bedding-cleavage relation indicates that the sheath lies on the overturned limb of first-order SE vergent asymmetric fold. (E, F) D1 fold with axial plane cleavages in thin section scale: Hioki mélange at locality 1. Open and crossed nicols, respectively. Cleavages in F are anastomosing. Scale bar in A, E, and F is 1 mm.
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processes versus processes associated with early deformation of poorly consolidated sediments as they are carried into the mouth of the subduction zone. Nonetheless, in the Nabae complex, evidence points strongly to a dominance of surface processes in, at the least, the mixing process, and by association, much of the stratal disruption. Perhaps the most compelling observation supporting this conclusion is the occurrence of the debris flow that incorporates clasts that carry an earlier pressure solution fabric that is randomly dispersed in an enveloping rock that carries a homogeneous D1 cleavage. This observation leaves no doubt that part of the sedimentary source for the mélange contained an exhuming, older source terrane with rocks that contained a pressure solution cleavage. Similar observations of older fabrics in chert blocks away from this debris flow suggest that they too were derived from this exhuming source terrane. Their provenance would be lost in the complexity of D0–D1 overprints. In this case, however, the presence of mafic dikes that were themselves only weakly deformed prior to development of the D1 cleavage, but are orthogonal to the mélange fabrics, suggests that, at the least, the mélange had not been carried far into the subduction interface before it was stratally disrupted. Although the depth of emplacement is uncertain, and it is possible that a few kilometers of slip along the subduction interface could have produced the mélange by tectonic processes, the simpler interpretation, in light of other observations, is that the stratal disruption in this complex occurred prior to accretion as an olistostrome complex. Implications for General Problems of Mélange Formation Pressure solution cleavage is not an expected feature of ocean-ridge spreading systems. It has been proposed that in some cases exotic blocks may have been derived from the outer trench slope—perhaps from subducting seamount, normal fault,
or transform fault scarps (e.g., Taira et al., 1982; Xenophontos and Osozawa, 2004)—although this possibility has not been actively pursued. The most plausible setting for cleavage to develop is in an accretionary prism, at the inner trench slope of a subduction zone. In our study the evidence suggests that older Paleogene basalt, chert, hemipelagic rocks, and sandstone were accreted in an older prism setting, in which D–1 cleavage could be formed. The deformed and metamorphosed accretionary products would then be exhumed by detachment faulting and thrusting (Schoonover and Osozawa, 2004; Osozawa and Pavlis, 2007), followed by gravitational collapse and debris flow into the evolving (i.e., Miocene) trench axis (Fig. 14). These D0 deposits would have undergone further D1 deformation with attendant cleavage and shear. The story is more complex, because the Paleocene chert in the Eocene Sakihama mélange has a cleavage that predates mélange deposition, and fragments of this chert are present in the Miocene Nabae complex (Osozawa, 2006). The problem we pose in this chapter for the Nabae mélange is not unique to this area. Indeed, the evidence we use in this analysis is analogous to the approach used years ago in analyzing the “knocker problem” in the Franciscan Complex (e.g., Cowan and Page, 1975). In fact, new evidence for a sedimentary origin of block-in-matrix fabrics in Franciscan mélanges is presented in this volume (Erickson, this volume; Wakabayashi and Dilek, this volume). In many respects, however, the problem in the Nabae complex is less complex, because, unlike the Franciscan, where the “exotic blocks” contain high-pressure metamorphic assemblages and ultramafic mantle rocks, the degree of exhumation needed to recycle the blocks found in the Nabae complex is much smaller: at most a few kilometers of exhumation. Thus, relatively wellknown processes, such as out-of-sequence thrusting, wedge extrusion, or forearc extension, could easily produce sufficient exhumation to produce the observed recycling. Indeed, the presence
Figure 13. Debris flow in thin section scale, in contact mélange of the Tsuro assemblage at locality 5. Gneissose granite and microcline fragments are included. Finer grained flow is associated. D1 pressure solution cleavages are mildly developed parallel to the flow. Open and crossed nicols, respectively. Scale bar in A and B is 1 mm.
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Figure 14. Recycling model. Accreted terrigeneous and oceanic materials, having undergone D–1 deformation and exhumation, were redeposited as debris flows by mass wasting, whose products underwent further accretion and D1 deformation.
of D2 thrust and normal fault systems may represent structures that are at least partially responsible for the exhumation. It was proposed that whereas subducted and subductioneroded material (including igneous oceanic crust) is mostly recycled to the mantle, this is balanced by the addition of magmatic arc components to the continental crust (von Huene and Scholl, 1991). Results of the present study moreover demonstrate that oceanic and terrigenous material represents a significant portion of material recycled within subduction zones, albeit at shallower levels and on smaller scales (Fig. 14), as suggested originally by Cowan and Page (1975). The Nabae mélanges show different accommodation of D1 deformation, the mélange zone corresponding as a whole to a single but wide shear zone, in which the shear strain was concentrated. In contrast, the coherent Tsuro assemblage was internally affected only by slaty cleavage, and therefore lacks an asymmetric shear fabric. However, D1 shear fabric in the coherent Misaki assemblage is represented by sheath folding, as also observed in the Nago complex, Okinawa (Schoonover and Osozawa, 2004). Further in contrast, the Yuwan and Ashio mélanges exhibit blockin-matrix fabric but completely lack asymmetric shear fabric, showing only slaty cleavage (Osozawa and Yoshida, 1997; Osozawa et al., 2008, 2009). Together, these diverse examples indicate that the nature of fabric development in a subduction complex is independent of whether the affected unit is a mélange or a coherent unit. Although it is difficult to clearly extend the results of this study to the general problems of tectonic versus sedimentary origins of mélange systems, it is important to note that the block-inmatrix mélange style of the Nabae complex is typical of many mélanges, particularly where exotic blocks are incorporated into the mélange. Thus we suggest that more studies should look carefully at the details of block-matrix relationships in mélanges to further address the importance of sedimentary versus tectonic processes in their formation. Many systems will clearly be more ambiguous than the case we describe here, because the unusual association of the mélange with a ridge subduction event pro-
vided a key part of the history for this study. Nonetheless, other techniques, particularly detrital thermochronology combined with U-Pb zircon geochronology, may provide new tests to examine this long-lived problem. CONCLUSIONS 1. The block-in-matrix fabric of mélanges in the Nabae accretionary complex shows the D0 sedimentary structure of a debris flow, which includes exotic blocks with an earlier D–1 fabric, overprinted locally by D1 shearing. The latter, however, does not significantly disturb the previous D0 normal faults and basaltic dikes, because shear displacement as a whole is rather small. Collectively these relationships show that sedimentary processes formed the block-in-matrix fabric. 2. Recycling of oceanic and terrigenous materials is therefore to be expected, such that growth of the accretionary prism involves both sedimentary and tectonic recycling. 3. Recognition of distinct stages of deformation is possible because the deformation is not progressive, as expected for tectonic mélange, so a better understanding of the structural development of an accretionary prism through time is possible. ACKNOWLEDGMENTS The official constructive review was given by John Wakabayashi, coeditor of the present volume, Jim Hibbard, and Dan Orange. Darrel Cowan kindly reviewed an earlier version of the manuscript. Tetsuya Mizutani offered his unpublished photo data. REFERENCES CITED Brandon, M.T., 1989, Deformational styles in a sequence of olistostromal mélanges, Pacific Rim Complex, western Vancouver Island, Canada: Geological Society of America Bulletin, v. 101, p. 1520–1542, doi:10.1130/0016-7606(1989)101<1520:DSIASO>2.3.CO;2.
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Cloos, M., 1982, Flow mélanges: Numerical modeling and geologic constraints on their origin in the Franciscan subduction complex: Geological Society of America Bulletin, v. 93, p. 330–345, doi:10.1130/0016 -7606(1982)93<330:FMNMAG>2.0.CO;2. Cowan, D.S., 1985, Structural styles in Mesozoic and Cenozoic mélanges in the western Cordillera of North America: Geological Society of America Bulletin, v. 96, p. 451–462, doi:10.1130/0016-7606(1985)96<451: SSIMAC>2.0.CO;2. Cowan, D.S., and Page, B.M., 1975, Recycled Franciscan material in Franciscan mélange west of Paso Robles, California: Geological Society of America Bulletin, v. 86, p. 1089–1095. Dela Pierre, F., Festa, A., and Irace, A., 2007, Interaction of tectonic, sedimentary, and diapiric processes in the origin of chaotic sediments: An example from the Messinian of Torino Hill (Tertiary Piedmont Basin, northwestern Italy): Geological Society of America Bulletin, v. 119, p. 1107–1119, doi:10.1130/B26072.1. DiTullio, L., and Byrne, T., 1990, Deformation paths in the shallow levels of an accretionary prism: The Eocene Shimanto belt of southwest Japan: Geological Society of America Bulletin, v. 102, p. 1420–1438, doi:10.1130/0016-7606(1990)102<1420:DPITSL>2.3.CO;2. Fisher, D., and Byrne, T., 1987, Structural evolution of underthrusted sediments, Kodiak Islands, Alaska: Tectonics, v. 6, p. 775–793, doi:10.1029/ TC006i006p00775. Hashimoto, Y., and Kimura, G., 1999, Underplating process from mélange formation to duplexing: Example from the Cretaceous Shimanto Belt, Kii Peninsula, southwest Japan: Tectonics, v. 18, p. 92–107, doi:10.1029/1998TC900014. Hibbard, J., and Karig, D., 1987, Sheath-like folds and progressive fold deformation in Tertiary sedimentary rocks of the Shimanto accretionary complex, Japan: Journal of Structural Geology, v. 9, p. 845–857, doi:10.1016/0191-8141(87)90085-X. Hibbard, J., Karig, D., and Taira, A., 1992, Anomalous structural evolution of the Shimanto accretionary prism at Murotomisaki, Shikoku Island, Japan: Island Arc, v. 1, p. 133–147, doi:10.1111/j.1440-1738.1992.tb00065.x. Hsü, K.J., 1968, Principles of mélanges and their bearing on the FranciscanKnoxville paradox: Geological Society of America Bulletin, v. 79, p. 1063– 1074, doi:10.1130/0016-7606(1968)79[1063:POMATB]2.0.CO;2. Ikesawa, E., Kimura, G., Sato, K., Ikehara-Ohmori, K., Kitamura, Y., Yamaguchi, A., Ujiie, K., and Hashimoto, Y., 2005, Tectonic incorporation of the upper part of oceanic crust to overriding plate of a convergent margin: An example from the Cretaceous–early Tertiary Mugi Mélange, the Shimanto Belt, Japan: Tectonophysics, v. 401, p. 217–230, doi:10.1016/j .tecto.2005.01.005. Kimura, G., and Mukai, A., 1991, Underplated units in an accretionary complex: Mélange of the Shimanto Belt of eastern Shikoku, Southwest Japan: Tectonics, v. 10, p. 31–50, doi:10.1029/90TC00799. Needham, D.T., 1987, Asymmetric extensional structures and their implications for the generation of mélanges: Geological Magazine, v. 124, p. 311–318, doi:10.1017/S0016756800016642. Needham, D.T., and Mackenzie, J.S., 1988, Structural evolution of the Shimanto Belt accretionary complex in the area of the Gokase River, Kyushu, SW Japan: Journal of the Geological Society [London], v. 145, p. 85–94, doi:10.1144/gsjgs.145.1.0085. Onishi, C.T., and Kimura, G., 1995, Change in fabric of mélange in the Shimanto Belt, Japan: Change in relative convergence?: Tectonics, v. 14, p. 1273–1289, doi:10.1029/95TC01929. Osozawa, S., 1984, Geology of Amami Oshima, Central Ryukyu Islands, with special reference to effect of gravity transportation on geologic structure: Japan, Science Report of Tohoku University, Sendai, 2nd Ser. (Geology), no. 54, p. 165–189. Osozawa, S., 1992, Double ridge subduction recorded in the Shimanto accretionary complex, Japan, and plate reconstruction: Geology, v. 20, p. 939– 942, doi:10.1130/0091-7613(1992)020<0939:DRSRIT>2.3.CO;2. Osozawa, S., 1993, Normal faults in accretionary complex formed at trench-trench-ridge triple junction, as an indicator of angle between
the trench and subducted ridge: Island Arc, v. 2, p. 142–151, doi:10.1111/j.1440-1738.1993.tb00082.x. Osozawa, S., 2006, Mélanges and dikes of the Miocene Nabae complex, Muroto Peninsula, Shikoku, Japan: Journal of the Geological Society of Japan, v. 112, Supplement, Excursion Guide Book, Annual Meeting of the Geological Society of Japan, 113th, Kochi, p. 41–53. Osozawa, S., and Pavlis, T., 2007, The high P/T Sambagawa extrusion wedge, Japan: Journal of Structural Geology, v. 29, p. 1131–1147, doi:10.1016/j .jsg.2007.03.014. Osozawa, S., and Yoshida, T., 1997, Arc-type and intra-plate type ridge basalts formed at trench-trench-ridge triple junction, implication for the extensive sub-ridge mantle heterogeneity: Island Arc, v. 6, p. 197–212, doi:10.1111/j.1440-1738.1997.tb00170.x. Osozawa, S., Sakai, T., and Naito, T., 1990, Miocene subduction of an active mid-ocean ridge and origin of the Setogawa ophiolite, central Japan: Journal of Geology, v. 98, p. 763–771, doi:10.1086/629439. Osozawa, S., Koitabashi, T., Katsube, S., and Flower, M., 2008, Medium P/T metamorphism in a subduction zone: A new type of regional metamorphism in Japanese accretionary complexes, inferred from b cell dimension of potassic white mica, in Columbus, F., ed., Structural Geology Research: New York, Nova Science Publishers, p. 1–19. Osozawa, S., Morimoto, J., and Flower, M., 2009, Block-in-matrix fabrics that lack shearing but possess composite cleavage planes: A sedimentary mélange origin for the Yuwan accretionary complex in the Ryukyu island arc, Japan: Geological Society of America Bulletin, v. 121, p. 1190–1203, doi:10.1130/B26038.1. Sakai, H., 1987, Storm barnacle beds and their deformation in the Murotomisaki olistostrome and mélange Complex, Shikoku: Journal of the Geological Society of Japan, v. 93, p. 617–620. Schoonover, M., and Osozawa, S., 2004, Exhumation process of the Nago subduction related metamorphic rocks, Okinawa, Ryukyu island arc: Tectonophysics, v. 393, p. 221–240, doi:10.1016/j.tecto.2004.07.036. Şengor, A.M.C., 2003, The repeated rediscovery of melanges and its implications for the possibility and the role of objective evidence in the scientific enterprise, in Dilek, Y., and Newcomb, S., eds., Ophiolite Concept and the Evolution of Geologic Thought: Geological Society of America Special Paper 373, p. 385–445. Shreve, R.L., and Cloos, M., 1986, Dynamics of sediment subduction, melange formation, and prism accretion: Journal of Geophysical Research, v. 91, p. 10,229 –10,245. Taira, A., Tashiro, M., Okamura, M., and Katto, J., 1980, The geology of the Shimanto belt in Kochi prefecture, Shikoku, Japan, in Taira, A., and Tashiro, M., eds., Geology and Paleontology of the Shimanto Belt, Selected Papers in Honor of Prof. Jiro Katto: Kochi, Japan, Rinyakosaikai Press, p. 319–389. Taira, A., Okada, H., Whitaker, J.H.McD., and Smith, A.J., 1982, The Shimanto Belt of Japan: Cretaceous–lower Miocene active-margin sedimentation, in Leggett, J.K., ed., Trench-Forearc Geology: Geological Society [London] Special Publication 10, p. 5–26. Taira, A., Byrne, T., and Ashi, J., 1992, Photographic Atlas of an Accretionary Prism: Geologic Structures of the Shimanto Belt: Tokyo, University of Tokyo Press, 124 p. von Huene, R., and Scholl, D.W., 1991, Observations at convergent margins concerning sediment subduction, subduction erosion, and the growth of continental crust: Reviews of Geophysics, v. 29, p. 279–316, doi:10.1029/91RG00969. Xenophontos, C., and Osozawa, S., 2004, Travel time of accreted igneous assemblages in Western Pacific orogenic belts and their associated sedimentary rocks: Tectonophysics, v. 393, p. 241–261, doi:10.1016/j .tecto.2004.07.037. Yajima, T., 1972, Petrology of the Murotomisaki gabbroic complex: Journal of Japanese Association of Mineralogy, Petrology, and Economic Geology, v. 67, p. 218–241. Manuscript Accepted by the Society 21 December 2010
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The Geological Society of America Special Paper 480 2011
Numerical estimation of duplex thickening in a deep-level accretionary prism: A proposal for network duplex Hikaru Ueno* Tobishima Corporation, 2, Sanbancho, Chiyodaku, Tokyo 102-8332, Japan Ken-ichiro Hisada† Yujiro Ogawa§ Doctoral Program in Earth Evolution Sciences, University of Tsukuba, Tsukuba, Ibaraki 305-8572, Japan
ABSTRACT Two types of thrust duplex structures were identified in excellent exposures of the deep level of the Jurassic to Cretaceous accretionary complex in the Kanto Mountains, central Japan, and the thickening ratio and shortening ratio were calculated. Simple (S type) and composite (C type) duplexes are mapped in an excavation site 100 × 40 m in extent. The beds of the S type and C type duplexes were thickened by factors of 5.8 and 6.0, respectively; however, the C type duplex includes four orders of smaller duplexes within it that underwent their own shortening. Thus the total thickening factor may attain at least 6–13, indicating a comparable degree of thickening at the level of greenschist facies conditions (approximately10 km or more in depth) in the accretionary prism.
INTRODUCTION The internal structure of active accretionary complexes has been investigated by ocean drilling and seismic profiling. Studies of the Nankai Trough, the Barbados Ridge, and other sites have revealed aspects of the accretionary prism such as offscraping by in-sequence thrusts, underplating by duplexes, and out-of-sequence thrusts (e.g., Kagami et al., 1983; Moore, 1989). In shallow levels of accretionary complexes, duplexes are important structures for
thickening and shortening sediments (Hirono and Ogawa, 1998; Yamamoto et al., 2005). However, structures at deeper levels in accretionary complexes are poorly understood because of their extreme stratal complexity and lack of index fossils, and the mechanisms for thickening and shortening are uncertain. Accretionary complexes of Mesozoic to Cenozoic age are well exposed in the Japanese Islands, and their internal structures and evolution, from shallow to deeper levels, can be observed in various complexes from Kyushu through Shikoku, and from
*Current address: Japan Railway Construction, Transport and Technology Agency, Honmachi 6-50-1, Yokohama 231-8315, Japan;
[email protected]. †
[email protected]. § Corresponding author, current address: Century Tsukuba-Miraidaira C-740, Yokodai 1-127-2, Tsukubamirai-shi 300-2358, Japan;
[email protected].
Ueno, H., Hisada, K.-I., and Ogawa, Y., 2011, Numerical estimation of duplex thickening in a deep-level accretionary prism: A proposal for network duplex, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 207–213, doi:10.1130/2011.2480(09). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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Kii to the Kanto Mountains. At deeper levels of the accretionary complex, duplexing can be a key process for underplating. Our recent field studies documented the thickening and hence the shortening of sediments associated with thrust duplex formation along bedding planes in the Jurassic to Early Cretaceous accretionary complex of the Chichibu Belt in the Kanto Mountains. In this chapter we propose the concept of the network duplex from detailed observations at large, fully exposed outcrops, and we discuss thickening processes of the Jurassic accretionary complex, together with an estimation of the thickening ratio. GEOLOGIC SETTING The Chichibu Belt extends ~1000 km along the Pacific side of the Japanese Islands (Fig. 1A). In the Kanto Mountains, 100 km west of Tokyo, the Chichibu Belt (Jurassic to Early Cretaceous accretionary complex), here divided into northern and southern belts, lies between the Sambagawa Belt (Sambagawa high pressure–low temperature [HP-LT] type metamorphic rocks) on the north and the Shimanto Belt (Late Cretaceous to
Cenozoic accretionary complex) on the south (Fig. 1B). The southern Chichibu Belt is fault bounded on its northern side with the Cretaceous Sanchu Group, which consists mainly of siltstone, sandstone, and conglomerate representing forearc basin deposits (e.g., Takei, 1963), and on its southern side it is in fault contact with the Shimanto Belt, composed mainly of Cretaceous to Miocene accretionary complexes. The southern Chichibu Belt in the study area consists mainly of Early Jurassic to Early Cretaceous formations, some parts of which are complexly deformed into mélange-like bodies. These formations are composed of sheared shale beds with blocks of chert, limestone, basaltic rocks, sandstone, and siliceous shale. Some units are lithologically identified as the Otchizawa Formation (Okubo and Horiguchi, 1969) and the Hamadaira Group (Hisada and Kishida, 1986) (Fig. 1C). For this chapter we subdivided the Hamadaira Group on the basis of lithofacies and age into four units, the Nogurizawa, Ryokami, Ogamata, and Kawakami Formations (Iwasaki et al., 1989; Ueno et al., 1990). In detailed mapping of an excellent bedrock exposure of 100 × 40 m extent at a dam construction site, we found that the so-called mélange-like bodies are identical to systematically
Figure 1. Index map of the Hamadaira Group, Nagano Prefecture. Pref.—Prefecture.
Numerical estimation of duplex thickening in a deep-level accretionary prism deformed stratal formations and should be regarded not as mélange but as duplex structures. This locality, called the Mikawa excavation site, is in the upper reaches of the Mikawa River, Minamiaiki Village, Nagano Prefecture (Fig. 1C). NETWORK DUPLEX The strata at the Mikawa excavation site belong to the Early Jurassic to Early Cretaceous Ryokami Formation of the Hamadaira Group, and consist of chert, shale, sandstone, and alternating beds of sandstone and shale in ascending (apparently stratigraphic) order. The bedding planes generally strike east-west, from N60° W to N70° E, and the strata dip northward at 40° to 80° (Fig. 2A). The chert is bedded and is commonly black to dark gray and partly reddish brown, and is in fault contact with the shale and sandstone. This fault strikes NW and dips almost vertically. The shale is black and partly dark gray where siliceous and is intercalated with fine-grained sandstone and tuffaceous shale layers several centimeters in thickness. The shale is also in fault contact with the sandstone. The bedding generally strikes WNW and dips northward at 60° to 70°. The sandstone is generally fine grained and includes abundant shale clasts and shale intercalations that contain very fine grained sandstone laminations. The sandstone contacts the alternating beds of sandstone and shale
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on low-angle faults. These faults strike WNW or WSW and dip northward at <30°. The Mikawa excavation site was mapped in detail, and several thrust duplexes were recognized (Ueno and Hisada, 2006) (Fig. 2B). The duplexes are of two types: antiformal stack and hinterland dipping. Foreland-dipping duplexes were not observed. The hinterland-dipping duplex (Fig. 3A) is ~10 m in length and is composed of four horses ranging from 1.5 to 5 m in length. (In measuring the length of each horse in this study we simply measured between both ends as shown in Fig. 4, not attempting to account for elongation or shortening during deformation. This is discussed further in the section on Thickening and Shortening by Duplex, below). The third horse consists of three smaller horses of sandstone, labeled as first-order subhorses (sub1-horse 1, 2, and 3 in Fig. 3B). These form an antiformal stack. Moreover, one of the first-order subhorses is composed of three secondorder subhorses (sub2-horses 1, 2, and 3 in Fig. 3B) between 30 and 50 cm in length. The three first-order subhorses also form a hinterland-dipping duplex. These two levels of hinterland-dipping duplexes show the same sense of dip-slip, indicating that they formed at the same time during the same thrust movement. In this case, the duplex is composed of several smaller duplexes. Duplexes of a larger scale, on the order of 10 m, are also recognized. As shown in Figure 2A, thrust faults are developed
Figure 2. Network duplexes of Mikawa excavation site. (B) Shown in A on map of larger region. Blue lines mark network thrusts, and dashed pink lines mark larger duplex boundaries.
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parallel or subparallel to each other, branching off and rejoining. They bound sigmoidal blocks of sandstone and shale. These larger duplexes have geometries similar to those of the smaller hinterland-dipping duplexes and antiformal stacks (red circles in Fig. 2B), and they are all developed along the thrusts. Striations on the fault planes support a displacement sense mostly of tops to the south or southeast, roughly perpendicular to the trend of thrust faults mapped to the east, as shown in Figure 1 (Ueno and Hisada, 2006). This observation suggests that the thrusts are responsible for duplication of the strata in this outcrop. Therefore, the sigmoidal blocks correspond to horses, and the thrusts correspond to the roof and floor thrusts. The thrusts are composite, made up of a network of several orders of thrust sets. We call the resulting set of nested duplexes a network duplex, and we call the thrusts network thrusts (blue lines in Fig. 2B). The network duplex is characterized by network thrusts with several horses between them. We recognized
A
>20 network duplexes (Fig. 2B) among network thrusts at this excavation site. Also at this excavation site we recognized a larger scale of duplexes, labeled duplexes 1 and 2 in Figure 2A (pink dashed lines). In duplex 1, we counted more than 70 horses of massive sandstone, several meters in width. The geometry of this duplex corresponds to an antiformal stack. In duplex 2, we identified several horses of massive sandstone, wider than those of duplex 1. TWO TYPES OF DUPLEX The three main types of duplex that have been proposed are the hinterland-dipping duplex, the antiformal stack, and the foreland-dipping duplex (Boyer and Elliott, 1982; Mitra and Boyer, 1986). In this study we recognized two types of duplex at this excavation site, a simple type (S type) duplex and a
Roof thrust
2nd horse 4th horse
B
1st horse
3rd horse
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B 2nd horse
Sub1-horse 1 Sub1-horse 2 1 -horse 3 2 Sub Sub2-horse 2 Sub -horse 3
Sub2-horse 1
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Figure 3. Photographs and illustrations of a hinterland-dipping duplex. The location is within the rectangle enclosed by black dashed lines in Figure 2B. These structures appear superficially like broken formation, with local development of sandstone blocks surrounded by shale.
Numerical estimation of duplex thickening in a deep-level accretionary prism combination type (C type) duplex (Fig. 4A). The identification of two types of duplex is based on their geometries in relation to the thickness of horses. The S type duplex is characterized by a combination of several tens of horses of similar width (Fig. 4B). The geometry of this duplex corresponds to the hinterland-dipping duplex. The C type duplex is characterized by a combination of several tens of horses and subhorses of various widths. (We denote orders of subhorses using a superscript, with a larger number indicating a smaller horse.) At the Mikawa excavation site, six horses (a to f) were identified in the C type duplex (Fig. 4A). Each of these horses is itself a C type duplex, because it is a combination of subhorses of various widths. In the case of horse a (Fig. 4C), there are 13 sub1-horses. In the case of horse c (Fig. 4C), there are 15 sub1-horses, of which six (subhorses 22–27 in Fig. 4C) constitute a hinterland-dipping duplex, and another three (subhorses 28–30) constitute an antiformal stack as shown in Figure 3. Accordingly, the C type duplex can be treated as a combination of S type duplexes. In the map of the whole outcrop (Fig. 2A), duplex 1 consists of many horses with similar thickness. Thus, duplex 1 corresponds to an S type duplex. Duplex 2 corresponds to a C type
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duplex, because it is a combination of duplexes of various widths, that is, a combination of subn-horses (n = 1–4). THICKENING AND SHORTENING BY DUPLEX For ancient accretionary complexes, it is known that the extensive underplating occurred through duplex formation at deeper levels, such as in the Kodiak Island, Shimanto, and Franciscan complexes (e.g., Moore and Byrne, 1987; Sample and Moore, 1987; Murata, 1991; Isozaki, 1996; Kimura et al., 1996; Hashimoto and Kimura, 1999). Hirono and Ogawa (1998) described duplexes of the Emi Group in the Boso Peninsula, central Japan. The typical duplex is several tens of centimeters in height and several meters in length, with several horses on the scale of tens of centimeters. They reported that the duplex array is 120 m thick after restoration of subsequent faulting, splaying in the lower to the upper zone in the direction of duplex development. They also concluded that the thickening factor of duplexes—duplicated thickness divided by original thickness—is 3.5. The Ryokami Formation in this study is characterized by network duplexes on the outcrop scale (Figs. 2–4). At this locality
Figure 4. Thickening estimates from the representative duplexes highlighted in map A, showing the parameters used to estimate thickening: H for thickness and L for length. Inset B shows the thickening estimate for the S type duplex, and inset C shows the thickening estimate for the C type duplex. For discussion on the measurement of bed length and its influence on shortening factors, see text.
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the true thickening factor is uncertain because there is no key bed, but most horses have similar lithology and appear to be of the same bed. We estimated the thickening and shortening factors of the S and C type duplexes at the Mikawa excavation site (Fig. 4) as follows. The S type duplex shown in Figure 4B is composed of 25 horses and measures 30 m in length (L1) and 7.5 m in width (H1) at the outcrop. This measurement takes no account of stratal deformation from the original beds. The average thickness of these horses (H2) is ~1.3 m, and their total length (L2) is 136 m. The apparent thickening factor of the S type duplex is H1/H2 = 7.5/1.3 = 5.8. The total length of the horses decreases as L1/L2 = 30/136, which means that the apparent shortening factor is ~1/4.5 (to 22%) by duplication within the S type duplex, given that little volume decrease is thought to have occurred during deformation. The C type duplex shown in Figure 4C is composed of six horses, formed by 38 subhorses. The mean thickness of the subhorses is ~1.2 m (H3), whereas the present thickness is 7.2 m (H4); and their total length is 158 m (L3), whereas the present length is 49 m (L4). The apparent thickening factor of the C type duplex is H4/H3 = 7.2/1.2 = 6.0. The total length of the horses decreases to 31% by duplication, meaning that the apparent shortening ratio is L4/L3 = 49/158, ~1/3.2. These calculations indicate that the thickening and shortening ratio of S and C type duplexes are similar in order, in spite of their different geometries. If all the assumptions for calculation, including no volume change, are correct, the thickness and shortening factors are of inverse value. However, there are many unclear factors for this estimation, and we need further controls for more accurate results. Turning to the larger duplexes at the Mikawa excavation site (Fig. 2A), duplex 1 consists of 94 horses. The average thickness of each horse (H5) is 4.1 m, and the maximum thickness of duplex 1 (H6) is 42.2 m. Thus the thickening factor is H6/H5 = 42.2/4.1 = 10.3. In duplex 2, we counted 56 horses. The average thickness of each horse (H7) is 5.2 m, the maximum thickness of duplex 2 (H8) is 63.2 m, and the thickening factor of duplex 2 is H8/H7 = 63.2/5.2 = 12.3. In this case too, the order of the factors is similar to those of the previously documented smaller scales. Given that a single horse in duplex 2 is an aggregation of several nested levels of subhorses, the total thickening factor can be expected to be larger. Sandstones are predominant in the Ogamata and Ryokami Formations. This may be due to duplication. An estimate of the true thickening and shortening might be approached under the assumption that the strata are shortened during the early to late stages of deformation by layerparallel shortening. At the same time the beds may be stretched by layer-parallel shearing. Also, in many cases of duplexing, unfolding might occur. We have no way of knowing the ratio of those deformations. We assumed little volume change during duplexing, because we found no strong mylonitic texture within the sandstone beds with only some slight preferred orientation of grains.
However, measuring the present length of the layer introduces considerable errors, because we did not take into account the lenticular shape of the sand layers, and, moreover, each horse has had its ends stretched while being shaped into a lenticular body. This factor may increase the length measurements by several tens of percentages. Therefore, in measuring the length, both elongation and shortening effects are involved. Thus our present estimation of thickening and shortening by duplexing has several elements of uncertainty. More information on shortening or lengthening of each horse is needed for better calculations of the shortening ratio by duplex formation. CONCLUSIONS Numerous duplexes were mapped at the Mikawa excavation site. On the outcrop scale, these form complex assemblages of nested structures, here termed network duplexes. We propose the existence of S type and C type duplexes on the basis of the geometries. In some young accretionary complexes, underplating is recognized at shallow levels (Hirono and Ogawa, 1998; Yamamoto et al., 2000, 2005). Our observations indicate that duplexes may be developed at deep as well as shallow levels. Duplexes have previously been identified on the outcrop scale of several meters (Hirono and Ogawa, 1998) and a few kilometers (e.g., Murata, 1998). In this study, it is clear that duplexes also range in scale from several centimeters to several meters, and from several tens to several hundreds of meters. The factors of thickening and shortening of beds by duplexing were estimated, and the apparent thickening factors attain 5.8 for the S type, and 6.0 for the C type, and for larger scale duplexing these factors range from 10.3 to 12.3. There are many other elements to be considered, and the total thickening factors may be even greater. ACKNOWLEDGMENTS We thank Darrel S. Cowan (University of Washington) for valuable comments and suggestions on an early draft of this chapter, and Hidetoshi Hara (Geological Survey of Japan, AIST) for constructive discussions and valuable comments throughout our work. We appreciate critical reviews by Richard Sedlock, Juan Luis Alonso, and John Wakabayashi. REFERENCES CITED Boyer, S.E., and Elliott, D., 1982, Thrust systems: American Association of Petroleum Geologists Bulletin, v. 66, p. 1196–1230. Hashimoto, Y., and Kimura, G., 1999, Underplating process from mélange formation to duplexing: Example from the Cretaceous Shimanto Belt, Kii Peninsula, southwest Japan: Tectonics, v. 18, p. 92–107, doi:10.1029/1998TC900014. Hirono, T., and Ogawa, Y., 1998, Duplex arrays and thickening of accretionary prisms: An example from Boso Peninsula, Japan: Geology, v. 26, p. 779– 782, doi:10.1130/0091-7613(1998)026<0779:DAATOA>2.3.CO;2. Hisada, K., and Kishida, Y., 1986, The Hamadaira Group in the western Kanto Mountains, central Japan—The developmental processes of the
Numerical estimation of duplex thickening in a deep-level accretionary prism Jurassic–Lower Cretaceous accretionary prism: Journal of the Geological Society of Japan, v. 92, p. 569–590. Isozaki, Y., 1996, Anatomy and genesis of a subduction-related orogen: A new view of geotectonic subdivision and evolution of the Japanese Islands: Island Arc, v. 5, p. 289–320, doi:10.1111/j.1440-1738.1996.tb00033.x. Iwasaki, T., Sashida, K., and Igo, H., 1989, Mesozoic strata of the KitaaikiKawakami area in Minamisaku Country, Nagano Prefecture, northwest Kanto Mountains, central Japan: Journal of the Geological Society of Japan, v. 95, p. 733–753. Kagami, H., Shiono, S., and Taira, A., 1983, Subduction of plate at the Nankai Trough and formation of accretionary prism: Science, v. 53, p. 429–438. Kimura, G., Maruyama, S., Isozaki, Y., and Terabayashi, M., 1996, Wellpreserved underplating structure of the jadeitized Franciscan complex, Pacheco Pass, California: Geology, v. 24, p. 75–78, doi:10.1130/0091 -7613(1996)024<0075:WPUSOT>2.3.CO;2. Mitra, G., and Boyer, S.E., 1986, Energy balance and deformation mechanisms of duplexes: Journal of Structural Geology, v. 8, p. 291–304, doi:10.1016/0191-8141(86)90050-7. Moore, J.C., 1989, Tectonics and hydrogeology of accretionary prisms: Role of the decollement zone: Journal of Structural Geology, v. 11, p. 95–106. Moore, J.C., and Byrne, T., 1987, Thickening of fault zones: A mechanism of mélange formation in accreting sediments: Geology, v. 15, p. 1040–1043, doi:10.1130/0091-7613(1987)15<1040:TOFZAM>2.0.CO;2. Murata, A., 1991, Duplex structures of the Uchinohae Formation in the Shimanto Terrane, Kyushu, Southwest Japan: Journal of the Geological Society of Japan, v. 97, p. 39–52. Murata, A., 1998, Duplexes and low-angle nappe structures of the Shimanto terrane, southwest Japan: Memoirs of the Geological Society of Japan, no. 50, p. 147–158 (in Japanese with English abstract).
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Okubo, M., and Horiguchi, M., 1969, Geology of the Manba District: Tokyo, Geological Survey of Japan, Quadrangle Ser., scale 1:50,000, v. 26, p. 66 p. Sample, J.C., and Moore, J.C., 1987, Structural style and kinematics of an underplated slate belt, Kodiak and adjacent islands, Alaska: Geological Society of America Bulletin, v. 99, p. 7–20, doi:10.1130/0016 -7606(1987)99<7:SSAKOA>2.0.CO;2. Takei, K., 1963, Stratigraphy and geological structure of the Cretaceous System in the eastern part of the Sanchu Graben, Kanto Mountainland, Japan: Journal of the Geological Society of Japan, v. 69, p. 130–146. Ueno, H., and Hisada, K., 2006, Slip sense in the southern Chichibu Belt in the Kanto Mountains, Japan: An indicator of plate subduction direction: Journal of the Geological Society of Thailand, v. 1, p. 83–90. Ueno, H., Hisada, K., and Igo, H., 1990, Imbricate structure and exotic mass observed in the Ryokami area, Kanto Mountains: Japan, Annual Report of the Institute of Geoscience, University of Tsukuba, no. 16, p. 46–49. Yamamoto, Y., Ohta, Y., and Ogawa, Y., 2000, Implication for the two stage layer-parallel faults in the context of the Izu forearc collision zone: Examples from the Miura accretionary prism, central Japan: Tectonophysics, v. 325, p. 133–144, doi:10.1016/S0040-1951(00)00134-7. Yamamoto, Y., Mukoyoshi, H., and Ogawa, Y., 2005, Large and small structural characteristics in shallow burial accretionary prism on land: Rapidly uplifted Neogene accreted sediments in the Miura-Boso Peninsula: Tectonics, v. 24, TC5008, doi:10.1029/2005TC001823.
MANUSCRIPT ACCEPTED BY THE SOCIETY 21 DECEMBER 2010
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The Geological Society of America Special Paper 480 2011
Tectonic, sedimentary, and diapiric formation of the Messinian mélange: Tertiary Piedmont Basin (northwestern Italy) Andrea Festa* Dipartimento di Scienze della Terra, Università di Torino, Via Valperga Caluso, 35, 10125 Torino, Italy
ABSTRACT The Messinian mélange of the Tertiary Piedmont Basin is the product of different but interrelated processes (tectonic, gravitational, and diapiric) that operated sequentially over a short time span (intra-Messinian time) and in a geodynamic environment (episutural basin) for which mélanges have so far been poorly described. It is composed of different mappable bodies of (non-metamorphic) mixed rocks characterized by a strong facies convergence. Their geometric and stratigraphic position, the internal organization, and the nature of the bounding surfaces allow the defining of some criteria to distinguish different units of mixed rocks (tectonically disrupted unit, gravity-driven sedimentary unit, and diapiric disrupted unit), in each of which the role of a different prevailing mélange-forming process can be inferred. None of these processes operated in isolation. They were linked by complex and intimate mutual interactions and triggered by intra-Messinian tectonics. The latter produced self-generating processes of mélange formation in which gravitational and diapiric processes triggered and affected each other. Different pulses of overpressured fluids (often rich in methane) strongly governed sediment deformation and also played a crucial role in influencing the time relationships and causative links between the different mélange-forming processes. Faulting may have triggered gas hydrate dissociation, promoting the upward rise of overpressured fluids. These fluids reduced the shear strength of the overlying sediments, promoting large-scale gravity-driven phenomena. Loading provided by rapid emplacement of the gravity-driven sedimentary bodies could have, in turn, developed new overpressured conditions necessary to promote the upward rise of poorly consolidated sediments and shale diapirism. INTRODUCTION The term mélange has a long history of academic debates regarding its origin. First coined by Edward Greenly (1919) for a unit of mixed rocks in Anglesey (North Wales), which he interpreted as the product of tectonic shearing, the term mélange has assumed with time different meanings both descriptive and genetic.
Mélanges are the product of different processes (tectonic, sedimentary, and diapiric) of fragmentation and mixing that occur in various geodynamic environments (e.g., subduction zones and accretionary prisms, Cowan, 1978, 1985; Aalto, 1981; Cloos, 1982; Barber et al., 1986; Brown and Westbrook, 1988; Cloos and Shreve, 1988; Onishi and Kimura, 1995; forearc and backarc basin, Page and Suppe, 1981; transform and overthrust fault
*
[email protected] Festa, A., 2011, Tectonic, sedimentary, and diapiric formation of the Messinian mélange: Tertiary Piedmont Basin (northwestern Italy), in Wakabayashi, J., and Dilek,Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 215–232, doi:10.1130/2011.2480(10). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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zones, Vollmer and Bosworth, 1984; Schultz, 1981; transform faults, Saleeby, 1984; continental margin sedimentary wedges, Jacobi, 1984; continental interior sedimentary wedges, Krieger, 1977; episutural basins, Dela Pierre et al., 2007, and elsewhere; see also Camerlenghi and Pini, 2009, and Festa et al., 2010a, 2010b, for a complete review), and they interact in a complex way. Tectonic, sedimentary (gravitational), and diapiric processes are in fact often linked by intimate and mutual interactions, and none of these processes operated in isolation (e.g., Lash, 1987; Maltman, 1994). For example, gravitational movements can be prompted by tectonic stress, faulting, and earthquakes able to induce the disruption of the coherent stratigraphic succession and/or local slope over-steepening necessary to trigger sediment failure (e.g., Lash, 1987; Clennell, 1992; Maltman, 1994). Mobilization of slope sediments may also occur in response to diapiric movements of overpressured sediments whose upward rise can be encouraged by the occurrence of mechanical discontinuities that work as preferential pathways (e.g., Barber et al., 1986; Lash, 1987; Barber and Brown., 1988; Brown and Westbrook, 1988; Clennell, 1992; Maltman, 1994; Kopf, 2002). The upward migration of overpressured sediments can be, in turn, caused by the rapid emplacement of gravity-driven deposits that are able to develop the necessary overpressure to promote shale diapirism and/or mud volcanoes (e.g., Barber et al., 1986; Lash, 1987; Collison, 1994; Maltman, 1994). Moreover, tectonic disruption and the mixing of originally coherent sequences, gravitational submarine downslope movements, and shale diapirism may occur over a short time span in a given geodynamic environment (e.g., Lash, 1987; Pini, 1999; Dela Pierre et al., 2007; Festa et al., 2010a, 2010b). Their products, which commonly show a strong facies convergence, are likely to be overprinted by later deformation and metamorphism that can obscure the prevailing formation processes. As a consequence, especially in ancient examples, the same mélange studied by different authors gives rise to conflicting interpretations (see, e.g., Raymond, 1984b cum biblio). In the light of these problems, Raymond (1975, 1984a) and the Penrose Conference on Mélanges in 1978 (see Silver and Beutner, 1980) provided evidence that the only agreement about mélanges is that they are “mappable” “bodies of mixed rocks.” For these reasons, he suggests using the term mélange in a prudent, descriptive way and not for describing processes of formation. In this context, different studies have proposed valuable criteria for distinguishing mélanges formed in different ways and in different environments (e.g., Naylor, 1982; Byrne, 1984; Cowan, 1985; Orange, 1990; Clennell, 1992) with special attention to polygenic mélanges (e.g., Raymond, 1984a; Saleeby, 1984; Yilmaz and Maxwell, 1984; Barber et al., 1986; Lash, 1987; Orange, 1990; Pini, 1999; Cowan and Pini, 2001; Dela Pierre et al., 2007). Further complications occur when unconsolidated sediments are considered. In this case, different factors (e.g., rate of consolidation, permeability of the deforming materials, effective stress, mineral dehydratation, hydrocarbon generation, gas hydrate dissociation, etc.) can strongly modify the diagnostic features (e.g., internal organization, nature of the contacts, geometric and strati-
graphic position, etc.) useful for identifying different formation processes. In particular, fluid pressures in excess of the equilibrium hydrostatic pressure strongly govern sediment deformation (Brown and Orange, 1993) and in turn exert a fundamental influence on the mechanical behavior of sediments (Maltman, 1994). This is the case for the Messinian mélange of the Tertiary Piedmont Basin (northwestern Italy), an episutural basin formed after the main Alpine collisional event (Figs. 1A and 1C). Previous works (Irace, 2004; Festa et al., 2005a; Dela Pierre et al., 2007) attributed its origin to the interaction of tectonic, sedimentary, and diapiric processes occurring sequentially in partially lithified sediments and in a short time span (intra-Messinian time). In this chapter we present an update, at regional scale, of previous research. New data have come from the northern sector of the Tertiary Piedmont Basin (Torino Hill and Monferrato; Fig. 1A) in which the lack of later deformation and metamorphism has preserved excellent examples and a complete case history for understanding the time relationships and causative links between the main mélange-forming processes. REGIONAL GEOLOGICAL SETTING The Tertiary Piedmont Basin (Fig. 1A), in the internal sector of the Western Alps, is a large episutural basin (retro-foreland basin sensu d’Atri et al., 2002) developed since the late Eocene. The sedimentary evolution of this basin has been strongly controlled by the thick-skinned thrusting related to progressive migration of the orogenic belt-foredeep pair toward the inner domains of the Adria plate. The sediments of the Tertiary Piedmont Basin have in fact been deposited unconformably, after the meso-alpine collisional event, on the junction between the Alps and Apennines (Fig. 1C). This junction corresponds to a complex jigsaw of buried tectono-stratigraphic domains constituting the metamorphic Western Alpine units, the SE-vergent front of the South Alpine thrust system, and the northeast-vergent external front of the non-metamorphic (External Ligurian) Apennine chain (e.g., Bonsignore et al., 1969; Biella et al., 1988, 1997; Gelati and Gnaccolini, 1988; Castellarin, 1994; Mutti et al., 1995; Roure et al., 1996; Piana, 2000; Mosca, 2006). Since the Oligocene onward, deposition has been strongly influenced by the opening of the Balearic Sea and the building of the Apenninic thrust belt (Boccaletti et al., 1990). As a result, different tectono-sedimentary domains, characterized by only partially comparable structural and stratigraphic evolution, developed (Piana and Polino, 1995; Biella et al., 1997). They are the Torino Hill and Monferrato to the north, and the Langhe, Alto Monferrato, and Borbera Grue to the south (Fig. 1A). The data discussed in this chapter come from the northern part of the Tertiary Piedmont Basin. The Torino Hill succession rests unconformably on a metamorphic basement, buried at a depth of 2–3 km (Bonsignore et al., 1969; Miletto and Polino, 1992; Polino et al., 1992), interpreted as the South Alpine basement (Mosca, 2006). The structural setting of the Torino Hill consists of a northeast-southwest transpressive flower structure formed mainly
Tectonic, sedimentary, and diapiric formation of the Messinian mélange before the late Burdigalian, which evolved in Langhian to Tortonian times as an asymmetric open anticline (Festa et al., 2005b, 2009b). The Rio Freddo Deformation Zone (sensu Piana and Polino, 1995), a regional, long-lived, northwest-striking fault zone, separates the Torino Hill and Monferrato successions (Figs. 1A and
NW
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2A). It consists of steeply dipping strike-slip faults, associated with minor northeast- to east-directed mesoscale thrusts and contractional duplexes (Piana, 2000). At the regional scale the Rio Freddo Deformation Zone has been understood as the surface expression of a deep-seated shear zone (Piana and Polino, 1995; Piana, 2000)
SE
Figure 1. (A) Schematic structural map of the Tertiary Piedmont Basin (TPB; modified from Bigi et al., 1990). (B) Location of Figure 1A. (C) Geological cross section (modified from Biella et al., 1988; Roure et al., 1990; Bello and Fantoni, 2002; Carminati et al., 2004; Festa et al., 2005b, 2009b). The trace of the section is shown in Figure 1A.
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active mainly since Rupelian to Burdigalian times (Festa et al., 2005b, 2009b; Fig. 1C). The Monferrato succession unconformably overlies the Cretaceous to lower Eocene non-metamorphic External Ligurian units (Bonsignore et al., 1969; Dela Pierre et al., 2003a), and its magnetic basement is buried at a depth of 8–10 km (Cassano et al., 1986; Miletto and Polino, 1992; Polino et al., 1992). The structural setting consists of a mosaic of northwest-striking rock slices, bounded by transpressive faults (e.g., Rio Freddo Deformation Zone and Villadeati Fault, Fig. 2A) and mainly made up of pre-Langhian sediments, alternating with open synclines and/or monoclines made up of Langhian to Tortonian sediments (Piana, 2000; Dela Pierre et al., 2003a).
The Torino Hill and Monferrato successions were deformed by contractional faulting stages of Rupelian, Burdigalian, and Serravallian age that show different styles of deformation in the two tectono-stratigraphic domains. Regional north-south shortenings were prevalent during intraMessinian time and are consistent with the northwest and northward movement of the Padane Thrust Front, which, to the north, overthrusted the Torino Hill and Monferrato onto the Po Plain foredeep (Fig. 1A). These shortenings caused, at the surface, only minor displacement along the preexisting faults (Rio Freddo Deformation Zone and Villadeati Fault), whereas the tectono-stratigraphic units have essentially undergone a gentle southward tilting responsible for the emplacement of the Messinian mélange.
Figure 2. (A) Structural map of the Torino Hill and Monferrato successions (modified from Dela Pierre et al., 2003b; Festa et al., 2009a). (B) Stratigraphic sketch of the Torino Hill and Monferrato successions.
Tectonic, sedimentary, and diapiric formation of the Messinian mélange THE MESSINIAN MÉLANGE The Messinian mélange of the Tertiary Piedmont Basin is a non-metamorphic sedimentary body of mixed rocks resulting from the dismemberment of the originally coherent stratigraphic succession (Dela Pierre et al., 2002, 2007; Irace, 2004; Festa et al., 2005a, 2009b; Irace et al., 2005) made up of pre-evaporitic Tortonian–lower Messinian hemipelagic marls (Sant’Agata Fossili marls) and by shallow-water Messinian gypsum (Gessoso Solfifera Formation) deposited in response to the Mediterranean salinity crisis (Hsü et al., 1973; Cita et al., 1978; Clauzon et al., 1996; Krijgsman et al., 1999). This mélange is attributed to the lower part of the post-evaporitic interval, and it crops out over an area of several hundreds of square kilometers (Fig. 1A) with the local exception of the Alba succession, where the “normal” Tortonian to Messinian succession is preserved (Sturani, 1973, 1978). The Messinian mélange consists of a prevalent fine-grained unconsolidated matrix that envelops hard blocks ranging in size from meters to several hundreds of meters. The matrix commonly consists of mud breccias composed of millimeter- to centimeter-sized angular clasts of unconsolidated Tortonian to lower Messinian sediments. The blocks, all of them referable to the Messinian, consist of primary selenitic gypsum (~0.1 m to >100 m in size) and a wide range of evaporitic, bioclastic, and methane-derived carbonates (~0.1 to 10 m in size). The latter include similar blocks of thinly laminated peloidal carbonates, clast supported breccias, and cemented mud breccias containing the remains of chemo-symbiotic communities (Lucina bivalves and tube worms). These blocks formed as a result of carbonate precipitation driven by the bacterial degradation of methane, mainly on the basis of their strongly δ13C-depleted isotopic signature (−25‰ to −50‰ Peedee belemnite) (Clari et al., 1988, 1994, 2004; Cavagna et al., 1999; Clari and Martire, 2000; Dela Pierre et al., 2002). Moreover, the positive δ18O value (from 2‰ to 8‰ Peedee belemnite) suggests that fluids could be sourced from gas hydrate dissociation. The same authors interpreted these blocks as authigenic carbonates, related to the rise of methane-rich fluids through the sedimentary column. Previous studies (Dela Pierre et al., 2002, 2003a; Irace, 2004; Irace et al., 2005) described the Messinian mélange as the Valle Versa chaotic complex (sensu Dela Pierre et al., 2002). The latter is correlated with the lower part of the post-evaporitic succession of the Apennine foredeep (p-ev1 Synthem of Roveri et al., 2003) and the Messinian chaotic deposits of Sicily and Spain (e.g., Roveri et al., 2003; Artoni et al., 2004; Lucente et al., 2005; Manzi et al., 2005), and its origin has been related to gravitydriven phenomena triggered by intra-Messinian tectonics (Dela Pierre et al., 2002, 2003a; Irace, 2004; Irace et al., 2005). The Valle Versa chaotic complex (Figs. 2A and 2B) forms a lenticular sedimentary body (~200 m of maximum thickness) bounded between well-bedded facies through two discontinuity surfaces (Dela Pierre et al., 2002, 2003a; Irace, 2004; Festa et al., 2005a, 2009b; Irace et al., 2005). The lower one is an erosional surface cutting into older marine sediments (late Oligocene to early Mes-
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sinian) or locally into Messinian primary evaporites (Gessoso Solfifera Formation). The upper one separates the Valle Versa chaotic complex from post-evaporitic Lago Mare sediments or lower Pliocene marine deposits (Argille Azzurre Formation). Recently, on the basis of a structural and stratigraphic study carried out in a gypsum quarry in the northern sector of the Tertiary Piedmont Basin (Moncucco quarry, Torino Hill), Dela Pierre et al. (2007) showed that the Valle Versa chaotic complex fully corresponds to gravity-driven sedimentary deposits, and that it is only a part of the Messinian mélange. The latter shows, in fact, a more complex stratigraphic and structural arrangement, and it is a composite chaotic unit formed of different bodies of mixed rocks (Fig. 2B) originated by different processes (tectonic, gravitational, and diapiric). In each of these the role of a prevailing mélange-forming process can be inferred. FIELD DATA: BODIES OF MIXED ROCKS Geologic mapping and structural and stratigraphic analyses allow us to define some criteria to distinguish the different bodies of mixed rocks that compose the Messinian mélange of the Tertiary Piedmont Basin. Based on the geometric and stratigraphic position, the internal organization, and the nature of the bounding surfaces of the different chaotic products, three units consisting of mappable bodies of mixed rocks are defined. Tectonically Disrupted Unit of Mixed Rocks The lower part of the Messinian mélange (Figs. 2B and 3) is bounded at the top by the angular unconformity corresponding to the base of the gravity-driven sedimentary unit (see below). The tectonically disrupted unit rarely crops out (Moncucco, Banengo, and Codana quarries); it has a maximum thickness of ~150–200 m, and in many cases it is partially reworked by subsequent gravity-driven phenomena. The imprint of tectonic deformation in the genesis of this unit of mixed rocks is inferred on the basis of the following criteria (Clennell, 1992; Pini, 1999; Cowan and Pini, 2001; Dela Pierre et al., 2007): 1. Repetition at different scales of structural associations (that are) all coherent with the intra-Messinian regional stress field (Festa et al., 2005a; Dela Pierre et al., 2007). The latter, northsouth directed (Fig. 4E), is related to the northward migration of the Padane thrust front (Festa et al., 2005b), and it is represented by north-south to northwest-southeast dextral transpressive fault systems with a left-handed en échelon arrangement (Fig. 2A). These faults are generally sealed by the gravity-driven sedimentary unit, as is clearly observable in the Moncucco quarry (Festa et al., 2005b; Dela Pierre et al., 2007) and evidenced by geologic mapping data in the Codana gypsum quarry (Figs. 3A and 3C). In the Moncucco quarry (Fig. 3A), the faults juxtaposed the largest Messinian gypsum blocks (up to several hundred meters in size) against the Sant’Agata Fossili marls (Tortonian–lower Messinian). This defines map-scale S-C dextral-transpressive
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Figure 3. (A) Geological map of the Moncucco quarry (modified from Dela Pierre et al., 2007). (B) Geological map of the Banengo quarry (modified from Bissacca, 2006). (C) Geological map of the Codana quarry (adapted from Bosticco et al., 1992). Note that in Figures 3A and 3B the tectonically disrupted unit (TDU) consists of tectonic blocks of both the Gessoso Solfifera Formation (GSF) and the Sant’Agata Fossili marls (SAF).
Figure 4. Tectonically disrupted unit: (A) Panoramic view of the Banengo quarry. The tectonically disrupted unit (TDU), consisting of tectonic blocks of both GSF and SAF, is bounded at the top by the angular unconformity (dotted line) corresponding to the base of the gravity-driven sedimentary unit (GDU). VVC—Valle Versa chaotic complex (upper Messinian); GSF—Gessoso Solfifera Formation (Messinian). Black dashed lines and V-shaped symbols indicate the bedding of GSF. (B) Detail of Figure 4A: Elongated block of gypsum is tectonically enclosed in the Sant’Agata Fossili marls (SAF), of Tortonian—early Messinian age, and aligned to the shear zone of the thrusts (black lines). (C) Detail of Figure 4B: reverse mesoscale S-C shear zone in the SAF marls. Arrows indicate the sense of shear. (D) Schematic cross section of the Codana quarry (trace in Fig. 3C). Dextral strike-slip faults depict a flower structure (adapted from Bosticco et al., 1992). Symbols as in A. (E) Mesoscale intra-Messinian faults and relative slip vectors from the Moncucco, Banengo, and Codana quarries (Schmidt net, lower hemisphere). The first two nets are referred to the studied quarries (Moncucco, Banego, and Codana), whereas the third net relates to the total studied area. Gray and white circles indicate, respectively, long-axes lineation of elongated blocks ranging in size from 10 cm to 100 cm and from 1 m to >100 m. Note the parallelism between the long axes of elongated blocks and the main faults. (F) Strike-slip flower structure in the Codana quarry (location in Fig. 4D): Elongated gypsum blocks are aligned with the shear zones. Symbols as in A. (G) Lozenge-shaped block of gypsum tectonically enclosed in the SAF marls and aligned with the main faults (white lines) of the Moncucco quarry. White dashed lines indicate the bedding of the SAF marls. (H) Detail of the shear zone of Figure 4G: Intersection of R and P shears defines centimeter-spaced L-shear lenses in the SAF marls. Arrows indicate the sense of shear.
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shear zones, north-northwest striking, consisting of subparallel bounding faults linked by northwest-southeast synthetic oblique ones. In the Codana gypsum quarry (Fig. 3C), steeply dipping northwest-southeast dextral faults juxtaposed huge slices of gypsum (up to several hundred meters wide) and depict, in section, a positive flower structure (Figs. 4D and 4F). In the Banengo gypsum quarry (Fig. 3B) the main evidence of the north-south regional shortening is a roughly east-northeast–striking thrust surface that superposed the Sant’Agata Fossili marls onto a huge block (several hundreds of meters) of gypsum (Fig. 4A). Similar structural associations are observable at the mesoscale mainly in the fine-grained sediments (Sant’Agata Fossili marls). They are represented by mesoscale shear zones and a pervasive scaly fabric. Both right-lateral and reverse mesoscale shear zones consist of a well-developed S-C fabric, defined by centimeterto decimeter-sized lithons (Fig. 4C). The latter often acquire a sigmoidal to anastomosed geometry according to north-south regional shortening. The scaly fabric is defined by millimeterto centimeter-sized shiny and striated foliation and consists of pervasive shear lenses (L sensu Naylor et al., 1986) defined by the intersection of R and P shears (Fig. 4H). This is related to the increase of shear deformation close to the main fault surfaces. 2. Structurally ordered blocks-in-matrix fabric and the elongated and lozenge shape of the blocks. According to Pini (1999), these are distinctive features of tectonically disrupted bodies. Close to the main regional fault systems (Rio Freddo Deformation Zone and Villadeati Fault; Fig. 2A), the gypsum blocks, from one to several hundred meters in size, have elongate and lozenge shapes and are tectonically enclosed within the Sant’Agata Fossili marls (Moncucco, Codana, and Banengo quarries; Figs. 4A, 4B, 4E, 4F, 4G). These blocks depict a regular and structurally ordered left-handed en échelon arrangement that is consistent with that of the above-described faults. 3. Alignment of the blocks to the shear zones. The long axes of elongated gypsum blocks are, at all scales, parallel to the main shear zones (Fig. 4E). At the map scale (Figs. 2A and 3A) this is clearly observable close to the Rio Freddo Deformation Zone and the Villadeati Fault. At the mesoscale, the elongated gypsum blocks, one to several meters in size, show a steeply dipping bedding strongly aligned with the S-C shear zones and the minor faults (Moncucco and Codana quarries; Figs. 4B, 4C, 4F, 4G). This is consistent with the transpressive movements that in turn are able to reorient the planar surfaces to their alignment with the main shear zones. Also in the Banengo quarry an elongated block of gypsum, tens of meters wide, is aligned with the main thrust surface (Figs. 4A and 4B). The data collected suggest that at all scales the gypsum blocks are tectonic slices wrenched from the previously coherent evaporitic succession (Gessoso Solfifera Formation) and juxtaposed with the Sant’Agata Fossili marls during the intra-Messinian tectonic episode. 4. Decrease of stratal disruption away from the faults. At map scale (Fig. 2A) this feature can be easily observed away from the main regional faults (Rio Freddo Deformation Zone and Villadeati Fault). Here, the huge gypsum blocks show irregular
shapes and are randomly distributed within the matrix (Figs. 5B and 5C). At the mesoscale the decrease of stratal disruption away from the faults is more easily observable (Moncucco, Banengo, and Codana quarries); the degree of deformation gradually passes from a pervasive scaly fabric, close to the fault, to an S-C fabric, and finally to a still recognizable bedding (Fig. 4G). In these cases, the relay zones between minor faults are linked by small-scale folds with an oblique trend of the axes with respect to the same faults (Moncucco and Banengo quarries). Gravity-Driven Sedimentary Unit of Mixed Rocks The gravity-driven sedimentary unit, corresponding to the Valle Versa chaotic complex of Dela Pierre et al. (2002) and Unit 2 of Dela Pierre et al. (2007), overlies the tectonically disrupted unit (Figs. 3 and 5A). It shows a maximum thickness of ~200 m. The imprint of gravity-driven sedimentary processes in the formation of the Messinian mélange is inferred on the basis of the following criteria (Elter and Trevisan, 1973; Abbate et al., 1981; Pini, 1999; Cowan and Pini, 2001; Dela Pierre et al., 2007): 1. Occurrence of lower and upper depositional contacts as discontinuity surfaces. The lower depositional contact consists of an irregular erosional surface defining a lenticular shape of the gravity-driven sedimentary unit. As it is clearly observable both at map (Figs. 2A, 3, 5A, 5B, 5C) and outcrop scales (Fig. 5D), it cuts into sediments of different age generally represented by the underlying Messinian evaporitic sediments (Gessoso Solfifera Formation) and Tortonian–lower Messinian hemipelagic Sant’Agata Fossili marls, but also by Oligocene (close to the Villadeati Fault) to Serravallian (east of the Moncucco quarry) marine sediments. Locally, as in the Moncucco quarry (Fig. 5A), the lower depositional contact is a karstic surface attributed to subaerial exposure of the gypsum during the intra-Messinian tectonic episode and superposed by Pliocene to Quaternary karst phenomena (Fioraso and Boano, 2002; Fioraso et al., 2004).
Figure 5. Gravity-driven sedimentary unit: (A) Cross section of the Moncucco quarry (trace in Fig. 3A), showing the structural and stratigraphic relationships between the tectonically disrupted unit, the gravity-driven sedimentary unit, and the diapiric disrupted unit. Note that the gravity-driven sedimentary unit is bounded between two discontinuity surfaces and that the tectonically disrupted unit (TDU) consists of tectonic blocks of both the Gessoso Solfifera Formation (GSF) and the Sant’Agata Fossili marls (SAF). (B) Geological map of the area to the west of Moncucco quarry. Note the random distribution of mesoscale blocks within the gravity-driven sedimentary unit. (C) Geological map of the area west of the Banengo quarry (modified after Dela Pierre et al., 2003b; Bissacca, 2006). Note the random distribution of blocks within the gravity-driven sedimentary unit. (D) Lower depositional contact of the gravity-driven sedimentary unit in the Banengo quarry. (E) Highly disordered arrangement of the blocks-inmatrix fabric: angular to rounded carbonate and gypsum blocks float with a random distribution in a fine-grained matrix (west of Banengo quarry). (F) Isotropic texture of the brecciated matrix. Angular to rounded clasts of light stained marls and black arenite are observable (Moncucco quarry).
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The upper depositional contact is an undulating discontinuity surface separating the gravity-driven sedimentary unit from the well-bedded lower Pliocene marine sediments of the Argille Azzurre Formation or locally (Moncucco quarry, Festa et al., 2005a; Dela Pierre et al., 2007; Banengo quarry) from upper Messinian Lago Mare brackish-water sediments (Figs. 2, 3A, 5A). 2. Highly disordered arrangement of the blocks-in-matrix fabric. This feature strongly contrasts with the structural ordered blocks-in-matrix fabric of the tectonically disrupted unit. The blocks (Messinian selenitic gypsum, gypsum-rudites, various carbonate rocks, and methane-derived carbonates), ranging in size from meters to several hundreds of meters, float with a random distribution in a fine-grained matrix. This structure, easily observed at the outcrop scale (Fig. 5E), is also shown at the map scale (Figs. 2A, 5B, 5C) where the huge blocks of gypsum are randomly distributed far away from the main faults. 3. Composition and isotropic texture of the matrix. The matrix enclosing the blocks consists of beige to gray, weakly consolidated mud breccias (Fig. 5F). It is characterized by angular to rounded clasts, ranging in size from millimeters to centimeters and consisting of dark mudstone and less commonly of light stained marls and black arenite. Detailed micropaleontological analyses (Dela Pierre et al., 2002, 2003a, 2007; Irace, 2004; Festa et al., 2009b) of the mud breccias have revealed the occurrence of prevalent reworked late Tortonian to early Messinian planktonic foraminiferans and locally (north of Montiglio, Dela Pierre et al., 2002) those of Eocene to early–middle Miocene age. The lack of any preferred orientation of the clasts defines the isotropic texture well (Figs. 5E and 5F). This allows us to compare these mud breccias to the “argille brecciate” (brecciated clays) described in many classical olistostromes of the Mediterranean area (e.g., Beneo, 1956; Rigo de Righi, 1956; Abbate et al., 1970; Elter and Trevisan, 1973; Pini, 1999; Cowan and Pini, 2001).
Moncucco quarry, for example (Fig. 6A), a tens-of-meters-wide diapiric body of lower Messinian marls, bounded by subvertical north-northeast–south-southwest transpressive faults, pierces the previously coupled gravity-driven sedimentary unit and the tectonically disrupted one. In the Codana and Banengo quarries some evidence can be observed only at the mesoscale (Fig. 7A). In all the observed cases, the faults, northwest-southeast to northeastsouthwest striking, are characterized by a short lateral continuity along the strike (hundreds of meters) and only minor displacements (tens of meters) that strongly contrast with the high intensity of deformation and the high degree of structural complexities recorded (Festa et al., 2005a). These are scaly cleavages, veins, flow fabric, and swirling and stretching of the layers (Fig. 7F), consistent with extrusion mechanisms of poorly consolidated and overpressured fine-grained sediments (sensu Higgings and Saunders, 1967; Barber et al., 1986). Generally, these faults, which are sealed by upper Messinian brackish-water sediments (Lago Mare marls) and lower Pliocene marine muds (Argille Azzurre Formation), displace those responsible for the origin of the tectonically disrupted unit or represent their reactivation. 2. Opposite sense of shear on opposite margins of the structure. This is a diagnostic feature already described in other ancient diapirs (e.g., Orange, 1990). In the example of the Moncucco quarry, the mesoscale structural associations (e.g., scaly cleavages and S-C fabric) characterizing the margins of the diapir (see below, “transitional zone”) indicate oblique-dextral (Fig. 6B) and oblique-sinistral (Fig. 6C) movements on the northwestern and southeastern margins, respectively (Festa et al., 2005a; Dela Pierre et al., 2007; Fig. 6A). In the Banengo and Codana quarries, 1-m-wide fluidal features, squeezed from the main thrust surface, show S-C fabric and centimeter-scale scaly cleavages, indicating the opposite sense of shear on the opposite margins (Fig. 7A). 3. Threefold zonation of deformation and the blocks-inmatrix arrangement inside the diapir (Festa et al., 2005a; Dela
Diapiric Disrupted Unit of Mixed Rocks This consists of meters (Codana and Banengo quarries) to tens of meters (Moncucco quarry) of strongly deformed Tortonian to lower Messinian muds (Sant’Agata Fossili marls). The most convincing scenario crops out in the Moncucco gypsum quarry (Festa et al., 2005a; Dela Pierre et al., 2007; Festa et al., 2010b), where the diapiric unit pierces the gravity-driven sedimentary unit, thus postdating it. The imprint of shale diapirism in the genesis of the Messinian mélange is inferred on the basis of the following criteria (Higgings and Saunders, 1967; Bishop, 1978; Knipe, 1986; Komar, 1972; Barber et al., 1986; Orange, 1990; Festa et al., 2005a; Dela Pierre et al., 2007): 1. Presence of intrusive contacts. At all scales (from meters to several tens of meters), the bounding contacts between the diapiric body and the host gypsum rocks are sharp and consist of subvertical transpressive faults (Figs. 6A, 6D, 6E). Along these faults the Tortonian to lower Messinian muddy sediments (Sant’Agata Fossili marls) pierce the younger Messinian gypsum rocks. In the
Figure 6. Diapiric disrupted unit of the Moncucco quarry: (A) Drawing of the diapiric disrupted unit (modified after Dela Pierre et al., 2007), showing the threefold zonation of deformation and the blocksin-matrix arrangement. (B) Mesoscale data from the northwestern margin of the diapiric body (Schmidt net, lower hemisphere). Comparing nets of Figures 6B and 6C, slickenlines show an opposite sense of shear on the opposite margins of the structure. Note the parallelism between the long axes of elongated blocks (gray circles) and the main faults. (C) Mesoscale data from the southeastern margin of the diapiric body (Schmidt net, lower hemisphere). Legend in B. (D) Intrusive contact between the diapiric body and the host rocks of the tectonically disrupted unit. Note the threefold zonation of deformation and the different blocks-in-matrix arrangement of the transitional and marginal zones. (E) Core zone: strongly asymmetric folds with irregular axial trends and steeply dipping, plunging axes. (F) Fluidal features of the marginal zone separate the S-C fabric of the transitional zone from the tectonically disrupted unit. (G) Detail of the marginal zone: anisotropic mud breccias showing fluidal features evidenced by the alignment of whitish elongated clasts. (H) Detail of the transitional zone: S-C fabric envelops hard blocks spread along the shear zone. Arrows indicate the sense of shear.
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Pierre et al., 2007). This is consistent with the extrusion mechanisms of poorly consolidated and overpressured fine-grained sediments. The clearest example is exposed in the Moncucco gypsum quarry where, on the basis of the increased amount of deformation toward the margin, three zones (Fig. 6A) have been defined (Festa et al., 2005a; Dela Pierre et al., 2007): (1) the core zone (Figs. 6D and 6E), ~10 m wide, where angular and loosely clustered blocks of gypsum and carbonates up to several tens of meters in size are randomly distributed within the muddy matrix (Sant’Agata Fossili marls); in the matrix the bedding is still well recognizable, and it is folded by strongly asymmetric cylindrical folds with irregular axial trends and steeply plunging axes (Fig. 6E); the scaly cleavage is
not present here; (2) the transitional zone (Figs. 6D, 6F, 6H), up to 1 m wide; the muddy sediments are strongly deformed here by a scaly cleavage and an S-C fabric parallel to the margins (Fig. 6H) that clearly indicate the opposing sense of shear on the opposite margins of the structure (Figs. 6B and 6C); decimeter sized carbonate and gypsum blocks, elongated parallel to the scaly cleavage and floating in the Sant’Agata Fossili marls, can be seen here (Figs. 6D and 6H); millimeter to centimeter broken disaggregated hard clasts are spread along the shear zones and depict wisps and tail features; (3) the marginal zone (Figs. 6F, 6G, 6H), which separates the transitional zone from the host rocks. It is made of a thin collar (a few decimeters wide) of anisotropic mud breccias (Fig. 6G) that, in
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Figure 7. Diapiric disrupted unit: (A) Opposite sense of shear on the opposite margins in a diapiric body of the Banengo quarry. (B) Mud injection at the tip of mesoscale faults in the Moncucco quarry. (C) Drawing of B, showing the fluidal features (dashed lines) and the blocks-inmatrix arrangement. (D) Hydraulic fractures filled by injection of muddy sediments in a block of gypsum-rudites wrenched from the margins of the diapiric body of the Moncucco quarry. (E) Drawing of C, showing the alignment of the scaly cleavage to the margins of the fracture. Arrows indicate the sense of shear. (F) Detail of fluidal features in the muddy sediments (Codana quarry). A sort of sigma-type hard block with asymmetric wings can be seen.
Tectonic, sedimentary, and diapiric formation of the Messinian mélange contrast to those observed in the gravity-driven sedimentary unit, are characterized by fluidal features evidenced by the alignment of elongated clasts to the external boundaries of the diapir. The mud breccias are composed of sky-blue to brown clays that envelop angular clasts, millimeter to centimeter in size, of whitish marls and black sands. 4. Occurrence of hydraulic fractures within the hard blocks extruded by the diapir. Decimeter- to meter-sized, long-axes gypsum blocks, wrenched from the host rocks, commonly show hydraulic fractures (up to 5 cm wide) filled by injections of strongly deformed muddy sediments (e.g., Moncucco quarry; Fig. 7D). The latter are pervaded by a millimeter- to centimeter-sized scaly cleavage aligned to the margins of the fractures (Fig. 7E). Other examples of hydraulic fractures are represented by mud injections (Figs. 7B and 7F), up to some meters wide, developed at the tip of mesoscale faults (Banengo, Codana, and Moncucco quarries). These consist of subvertical flame structures with anastomosing arrays of polished curved surfaces and S-C fabric that cuts both the muddy sediments (Sant’Agata Fossili marls) and the gypsum. 5. Occurrence of methane-derived carbonates (Clari et al., 1988, 1994, 2004; Cavagna et al., 1999; Clari and Martire, 2000; Dela Pierre et al., 2002). These blocks (see the major section, above, on The Messinian Mélange) are interpreted as authigenic carbonates, related to the rise of methane-rich fluids through the sedimentary column (Clari et al., 1988, 1994, 2004; Cavagna et al., 1999; Clari and Martire, 2000; Dela Pierre et al., 2002). In the Torino Hill and Monferrato tectono-sedimentary domains the blocks of methane-derived carbonates are commonly and randomly enveloped in the weakly consolidated matrix of the gravity-driven sedimentary unit. Direct evidence for the involvement of methane-rich fluids in the emplacement of diapirs through the gravity-driven sedimentary unit is lacking, except in the northern sector of the Monferrato (Verrua Savoia), where Messinian mud volcanoes, formed by the upward rise of methane-rich fluids, are preserved (Clari et al., 2004). DISCUSSION: TIME RELATIONSHIPS AND CAUSATIVE LINKS BETWEEN DIFFERENT PROCESSES OF MÉLANGE FORMATION Tectonic, gravitational, and diapiric processes worked together in the genesis of the Messinian mélange of the Tertiary Piedmont Basin and are linked by a common triggering mechanism: the regional intra-Messinian tectonic deformation responsible for the northward migration of the Padane Thrust Front. In this context, valuable criteria (geometric and stratigraphic position, internal organization, nature of the bounding surfaces, etc.) allow us to distinguish the role played by different formation processes and to define their time relationships (Fig. 8) and causative links (Fig. 9). Tectonic versus Gravitational Processes Tectonic processes are the oldest, and their products (tectonically disrupted unit of mixed rocks) are unconformably overlain
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by the gravity-driven sedimentary unit. They took place during intra-Messinian time, after deposition of the primary evaporites (deposited in response to the salinity crisis) but prior to deposition of the upper Messinian bodies of mixed rocks (Fig. 8). The role played by tectonic processes can be observed at all scales. On a regional scale the northward migration of the Padane Thrust Front during intra-Messinian time produced a southward tilting of the basin, inducing the slope toward over-steepening necessary to trigger sediment failure (Festa et al., 2005b). As assessed in the literature (e.g., Martinsen, 1994), a very low angle slope can produce various kinds of mass gravitational movements in poorly consolidated sediments. In the Torino Hill and Monferrato tectonosedimentary domains these processes completely disrupted the previously coherent Messinian succession (Fig. 8), mixing it with non-consolidated and younger sediments. As a consequence, the structureless and highly disordered blocks-in-matrix arrangement of the gravity-driven sedimentary unit was formed. On a local scale, northeast-southwest to northwest-southeast transpressive faults, linked by approximately east-west thrusts and reverse faults, also played a prominent twofold role in the genesis and emplacement of the gravity-driven sedimentary unit (Figs. 8 and 9). These faults were responsible for (1) the in situ disruption of the previously coherent evaporitic and pre-evaporitic succession, resulting mainly in the juxtaposition of huge blocks of gypsum (up to several hundred meters in size) and Tortonian muddy sediments (e.g., Codana and Moncucco quarries); locally (Banengo quarry) the superposition of Tortonian sediments onto huge blocks of gypsum also occurred; and (2) the creation of mechanical discontinuities, which, fracturing the impermeable layers of the evaporites, worked as conduits for the upward rise of fluids. A fluid pressure increase within a sediment can reduce the strength of the mass such that it is no longer stable so that deformation might ensue (Maltman, 1994). In the studied cases, further pore-pressure increments caused by the upward rise of fluids along faults could have approached the shear strength of the layers, which lowered the effective stress, activating the failure of mechanically weakened sediments (Fig. 9). The emplacement of the gravity-driven sedimentary unit took place (Figs. 8 and 9). However, the local upward rise of fluids along faults, often overpressured and in some cases rich in methane, could have in turn produced shale diapirism (see below). Tectonic versus Diapiric Processes In many sectors of the Torino Hill and Monferrato tectonosedimentary domains, diapiric processes seem to have occurred prior to and/or during emplacement of the gravity-driven sedimentary unit (Fig. 8). Locally (north of Montiglio), Eocene to middle Miocene planktonic foraminiferans, mixed with Tortonian to lower Messinian ones in the brecciated matrix of the gravity-driven sedimentary unit, were in fact interpreted by Dela Pierre et al. (2002) as the product of diapiric processes reworked by mass-wasting phenomena. The upward rise of overpressured muds could have
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sampled parts of the middle Eocene to Messinian host succession (Dela Pierre et al., 2002). The age of the brecciated matrix is, in fact, older than the age of the hard blocks dismembered by the previous coherent Messinian succession. Further evidence is the occurrence of several reworked blocks, from meters to tens of meters wide, of methane-derived carbonates within the gravity-driven sedimentary unit. Part of these are considered to be the remnants of shale diapirs and mud volcanoes induced by the upward rising of gas-rich fluids (mainly methane) sourced by gas hydrate destabilization (e.g., Clari et al., 1988, 1994, 2004; Cavagna et al., 1999; Clari and Martire, 2000; Dela Pierre et al., 2002; Irace, 2004). The upward rising of methane-rich fluids could have favored overpressured conditions (Fig. 9) within the fine-grained sediments trapped below the low-permeability evaporitic layers. Evidence that in low-permeability sediments methane generation can lead to overpressure and diapirism, especially in shallow-water depths, is described worldwide (Brown and Westbrook, 1988; Hovland and Judd, 1988; Brown, 1990; Reed
et al., 1990; Moore and Vrolijk, 1992; Brown and Orange, 1993). In the case studied the dismemberment of the low-permeability evaporitic layers by faults and other mechanical discontinuities (fractures, joints, bedding surfaces, etc.) could have controlled the upward rise of these fluids (Fig. 9). They could have, in turn, caused the sediments to deviate from a critical failure state, promoting local gravity-driven phenomena able to conceal the traces of shale diapirism (Fig. 9). Studies of present-day accretionary complexes have revealed, in fact, that shale diapirism is an effective mechanism in generating huge volumes of mélanges (Barber et al., 1986; Lash, 1987; Barber and Brown, 1988; Clennell, 1992; Camerlenghi et al., 1995; Harris et al., 1998; Kopf, 2002). Gravity-Driven Sedimentary versus Diapiric Processes Different examples (Moncucco, Codana, and Banengo quarries) show that shale diapirism also represents the final formation process (Fig. 8). Diapiric bodies pierce gravity-driven and
Superposition of mechanisms, processes, and products Figure 8. Time relationships between tectonic, gravitational, and diapiric processes in the origin of the Messinan mélange (not in scale). TDU—tectonically disrupted unit; GDU—gravity-driven sedimentary unit; DDU—diapiric disrupted unit; LPL—Argille Azzurre Formation (lower Pliocene); LM—“Lago-Mare” deposits (upper Messinian); VVC—Valle Versa chaotic complex (upper Messinian); GSF—Gessoso Solfifera Formation (Messinian); SAF—Sant’Agata Fossili Marls (upper Tortonian–lower Messinian); ESS—middle Eocene to Serravallian marine sediments.
Tectonic, sedimentary, and diapiric formation of the Messinian mélange tectonically disrupted units, causing their partial reorganization (Fig. 8). In the Moncucco quarry, it is evidenced by the rotation of the hard blocks floating in the matrix of the gravity-driven sedimentary unit up to their alignment with the margins of the diapiric bodies (Fig. 6A). The rise of the diapirs could have been caused by the combined effects of sedimentary loading, provided by the deposition of gravity-driven chaotic sediments, the presence of lowpermeability layers (primary evaporites), and faulting. In this context, gravity and local-to-regional tectonics have worked together to promote local shale diapirism. Where the fine-grained sediments of the tectonically disrupted unit were progressively but rapidly buried by the emplacement of the gravity-driven sedimentary unit, they could have developed the necessary overpressure to promote shale diapirism (Fig. 9). Rapid burial could have caused pore-fluid dissipation and, as a consequence, promoted the upward rise of overpressured sediments exceeding hydrostatic pressure (Collison, 1994; Maltman, 1994; Dela Pierre et al., 2007). This is well evidenced in the Moncucco quarry, where the main diapiric body is characterized by a threefold zonation of
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deformation, hydraulic fractures within the blocks, and the presence of a narrow zone of fluidal mud breccias that envelop the diapir all around the margins. This evidence is consistent with the extrusion mechanisms of poorly consolidated and overpressured fine-grained sediments under metastable conditions (e.g., Barber et al., 1986; Dela Pierre et al., 2007). In particular, this is confirmed by the unusual occurrence in the Moncucco diapiric body of a thin collar of fluidal mud breccias separating the diapiric body from the host rocks (Festa et al., 2005a; Dela Pierre et al., 2007). These low-viscous and quasi-fluid materials encouraged the rise of the more viscous diapir body, evidencing overpressured conditions. Fault surfaces probably acted as mechanical conduits for the rise of overpressured sediments produced by the emplacement of the gravity-driven sedimentary unit. This probably occurred by the reactivation of the previously formed faults (Fig. 8), responsible for the tectonic dismemberment of the lowpermeable evaporitic layers, which therefore acquired peculiar characteristics. Their short lateral continuity along the strike, and only minor displacements, contrast with the high intensity of deformation of the mesoscale structural associations (scaly cleavages, fluidal features, anastomosed arrays, etc.) produced by the upward rise of poorly consolidated and overpressured sediments. These faults, which represent the preferential pathway for the upward rise of overpressured sediments, displace the gravitydriven sedimentary unit. The gravity-driven sedimentary unit and the diapiric disrupted unit are unconformably overlain by well-bedded postchaotic upper Messinian (Lago Mare deposits) and lower Pliocene (Argille Azzurre Formation) sediments (Fig. 8). This constrains to intra-Messinian times the mutual relationships between gravitational and diapiric processes and indicates that all the complex and interrelated processes responsible for the genesis of the Messinian mélange faded out in late Messinian times (Fig. 8). CONCLUSIONS
Figure 9. Causative links between tectonic, gravitational, and diapiric processes in the formation of the Messinian mélange. Arrows indicate linkages and/or cause-effect relationships between processes of mélange formation and products. See text (Discussion) for a detailed explanation.
An update of previous research (Festa et al., 2005a; Dela Pierre et al., 2007) allows us to evaluate better the role of tectonic, gravitational, and diapiric processes in the genesis of the Messinian mélange of the Tertiary Piedmont Basin, their time relationships, and the causative links. Here, the strong facies convergence of their products required, at first, the definition of valuable criteria to distinguish between the different mélangeforming processes. The repetition, at different scales, of structural associations coherent with the regional north-south stress field, the structurally ordered blocks-in-matrix fabric and the elongated shape of the blocks, their alignment with the shear zones, and the decrease of stratal disruption far away from the faults are key diagnostic features of mélanges of tectonic origin (tectonically disrupted unit). The occurrence of lower and upper depositional contacts as discontinuity surfaces, the highly disordered arrangement of the blocks-in-matrix fabric (strongly contrasting with the structurally ordered arrangement of the tectonically disrupted unit), and the
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composition and isotropic texture of the brecciated matrix identify a mélange of gravity-driven origin (gravity-driven sedimentary unit). The presence of intrusive contacts, opposing shear directions on the opposite margins of the structure, the threefold zonation of deformation, the blocks-in-matrix arrangement, and the occurrence of hydraulic fractures within the hard blocks are diagnostic features of mélanges originated by diapiric processes (diapiric disrupted unit). The mutual relationships and complex causative links between tectonic, gravitational, and diapiric processes reflect a complex deformational history (Fig. 9) that occurred in an unusual geodynamic context (episutural basin) where mélanges have so far been poorly described (see also Festa et al., 2010a). Here, faulting, overpressure, the presence of low-permeability layers (primary evaporites), and the circulation of methane-rich fluids in the sedimentary column represent the main factors that controlled the various mélange-forming processes that occurred over a very short time span (intra-Messinian time). None of these processes operated in isolation; rather, they were linked by complex and intimate mutual interactions (Fig. 9), and they were triggered by a common mechanism (intraMessinian regional tectonics). This promoted self-generating processes of mélange formation in which gravitational and diapiric processes, subjected to different pulses of overpressured fluids, affected and triggered each other. Overpressured fluids strongly govern sediment deformation (e.g., Brown and Orange, 1993; Maltman, 1994) and also play a crucial role in influencing the mechanical behavior of the Messinian sediments of the Tertiary Piedmont Basin. Faulting may have triggered the gas hydrate dissociation, changing pressure conditions inside the sedimentary column and/ or opening pathways to the upward rise of deep-seated, warmer overpressured fluids (Festa et al., 2007). This first pulse of fluids reduced the shear strength of the overlying sediments, triggering the large-scale gravity-driven phenomena. The emplacement of the gravity-driven sedimentary unit took place, concealing a large part of the previous deformational history (blocks of different types of methane-derived carbonates are the only remnants; see Clari et al., 2004, for major details). Loading provided by the rapid emplacement of the gravitydriven sedimentary unit could, in turn, have developed new overpressured conditions in the tectonically disrupted sedimentary succession able to promote the upward rise of poorly consolidated sediments and shale diapirism. The reactivation of previously formed faults could also, in this case, have represented the preferential pathway for the diapiric intrusions that pierced the gravity-driven sedimentary unit. ACKNOWLEDGMENTS I would like to thank F. Dela Pierre, G. Fioraso, and A. Irace for fruitful discussions and for their support during the fieldwork in the Moncucco quarry, and S. Cavagna, F. Mondino, and R. Reville for critical reading of the manuscript. Special
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MANUSCRIPT ACCEPTED BY THE SOCIETY 21 DECEMBER 2010 Printed in the USA
The Geological Society of America Special Paper 480 2011
Recognition of a trench-fill type accretionary prism: Thrust-anticlines, duplexes, and chaotic deposits of the PliocenePleistocene Chikura Group, Boso Peninsula, Japan Satoru Muraoka* Master’s Program in Geosciences, University of Tsukuba, Tsukuba 305-8572, Japan Yujiro Ogawa† Doctoral Program in Earth Evolution Sciences, University of Tsukuba, Tsukuba 305-8572, Japan
ABSTRACT The Pliocene-Pleistocene Chikura Group, southern tip of the Boso Peninsula, central Japan, occurs northeast of the present Sagami trough of the Philippine Sea plate subduction boundary. This group has many bedding-parallel shortening structures, including thrust-anticlines, duplexes, and small-scale conjugate sets of thrusts in addition to various kinds of chaotic deposits. The group forms one large synclinorium with smaller scale folds, but its relationship to accretionary prism evolution has not been explained. On the basis of geological structures examined on uplifted coastal benches, we propose that the lower half of the group was deposited on the subduction plate boundary as trench fill. When the trench was filled, the frontal thrust jumped seaward, causing landward tilting of the earlier trench fill deposits, after which the upper part of the group was deposited in a slope basin setting. The key observation to unravel the sedimentation and deformation is the recognition of the chaotic deposits, specifically whether they have a methane-bearing, fluid-supported chemosynthetic biocommunity (Calyptogena) and calcite-cemented sediments (chimneys or pipes). The chaotic deposits that bear such methane-related materials suggest that the deposition has occurred on the thrust at the landward slope foot, and that the emplacement or depositional mechanism is either as a debris flow or an injection (diapir). As a result, it is concluded that at least the lower half of the Chikura Group is a kind of accretionary prism of the trench-fill type, similar to the Sagami Basin at the present time. We conclude that the lower half of the Chikura Group records accretionary prism development in a trench-fill environment, similar to the present day Sagami Basin.
*Current address: Atmosphere and Ocean Research Institute, University of Tokyo, 5-1-5 Kashiwanoha Kashiwa 277-8564, Japan; e-mail:
[email protected]. † Current address: Yokodai 1-127-2, Century Minamidaira C-740, Tsukubamirai-Shi 300-2358, Japan; e-mail:
[email protected]. Muraoka, S., and Ogawa, Y., 2011, Recognition of a trench-fill type accretionary prism: Thrust-anticlines, duplexes, and chaotic deposits of the Pliocene-Pleistocene Chikura Group, Boso Peninsula, Japan, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 233–246, doi:10.1130/2011.2480(11). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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Muraoka and Ogawa accreted to the Honshu side (Ogawa and Taniguchi, 1988; Ogawa et al., 1989; Hanamura and Ogawa, 1993; Ogawa et al., 2008; Saito, 1992). Recent ocean research on the Sagami trough, including the France-Japan KAIKO projects (Nakamura et al., 1987; Pautot et al., 1987; Ogawa et al., 1989), has verified that there are at least three domains of sedimentary basins and accretionary prisms (Ogawa et al., 2008). On the landward side of the Sagami trough the best examples of subaerial accretionary prisms–forearc basins developed since the Miocene are exposed on the Miura and Boso Peninsulas. South of the Boso Peninsula is a unique tectonic setting of a trench-trench-trench triple junction, the Boso triple junction (Ogawa et al., 1989; Seno et al., 1989) (Fig. 1). Good exposures of Neogene accretionary prism deposits and structures lie along the Pacific coastlines in the southern part of the two peninsulas (owing to the uplift of the erosional bench by occasional subduction-related large earthquakes (Shishikura, 2003; Kawakami and Shishikura, 2006). One of the best examples of such Neogene deposits is the Pliocene-Pleistocene Chikura Group at the southern tip of the Boso Peninsula (Figs. 1 and 2). This chapter focuses on the critical observation and structural mapping of the Chikura Group and presents observations that demonstrate how the strata were horizontally shortened, how such structures are related to chaotic deposits, and whether
INTRODUCTION Japan is one of the type localities for convergent plate margins, where many accretionary prisms have been developed since the Paleozoic; some are ancient on-land examples, some are submarine, and some are recent. The best documented ones for the latter are the Nankai and Sagami trough prisms (Taira et al., 1989; Ogawa and Taniguchi, 1988; Hanamura and Ogawa, 1993; Ogawa et al., 2008). The Nankai prism, along the northern subduction boundary of the Philippine Sea plate under the Eurasian plate, has been studied on many ocean drilling legs by the Deep Sea Drilling Project, Ocean Drilling Program, and Integrated Ocean Drilling Program (DSDP, ODP, IODP) Legs 31, 56, 87, 131, 169, 196, Exp 314, 315, 316, etc.) and related studies (e.g., the France-Japan KAIKO project; Kobayashi, 2002; Kawamura et al., 2009), but no drilling has ever been done in the Sagami trough. Although the Nankai and Sagami troughs are parts of the northeast subduction boundary on the Philippine Sea plate, they are distinct. The Nankai trough is characterized by nearly orthogonal convergence, whereas the Sagami trough is oblique, and it is the place where the frontal part of the Izu volcanic arc is subducting beneath the Honshu continental arc (Fig. 1). Therefore, the Sagami trough has developed mostly from the Izu forearc sediments obliquely
North American plate 0.9 cm/yr
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Figure 1. Index map for the study area, showing subduction boundaries in Japan. Adopted from Seno et al. (1989).
Recognition of a trench-fill type accretionary prism, Boso Peninsula, Japan or not these structures were formed by accretionary processes. Among the coastal exposures, several representative areas were selected (Fig. 2). The structures are shown in the geologic map of Kawakami and Shishikura (2006). GEOLOGICAL SETTING The Chikura Group exposed on the Boso Peninsula forms one large synclinorium. At the northern boundary the group unconformably overlies the middle Miocene to lower Pliocene Miura accretionary prism (Yamamoto and Kawakami, 2005; Kawakami and Shishikura, 2006) (Fig. 2). In contrast, at the southern boundary, the Chikura Group may have been thrust over the present accretionary prisms off the Boso Peninsula (Sato et al., 2005). The Chikura Group is divided into the Shirahama, Shiramazu, Mera, and Hata Formations, in ascending stratigraphic order. Strata are composed mainly of volcaniclastic sandstone and siltstone of deep marine environments with numerous tuff
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and tuff breccia intercalations and chaotic beds in various zones. In particular, the lowermost Shirahama Formation comprises reddish-brown, coarse tuffaceous (andesitic) sandstone or breccia of pyroclastic-flow or debris-flow origin. The Chikura Group ranges in age from 3.68 to 0.78 Ma (Kawakami and Shishikura, 2006, citing data and analysis of Kanie et al., 1997; Kameo and Sato, 1999; Kameo et al., 2003; and Saito, 1999). Within the Chikura Group the Shiramazu Formation is estimated to have been deposited from 3.68 to 3.31 Ma, the Mera Formation from 3.31 to 1.21 Ma, and the Hata Formation from 1.95 to 0.85 Ma. Thus the two formations for which we describe detailed structures here, the Shirahama and Shiramazu Formations, were deposited between 3.68 and 3.31 Ma. Because the age of the Shirahama Formation is still unknown, the exact stratigraphic relationship between the Shirahama and Shiramazu is not known. Although the Shiramazu Formation generally overlies the Shirahama Formation, it locally interfingers with the latter.
35°00'N Tateyama Nishizaki Chikura Kotto Kawaguchi
Pacific Ocean
Okawa
Mera
5 km 34°53'N 139°45'E
Shioura
Nemoto Cape Nojimazaki 139°52'E
Minami-Boso Group Kagamigaura Fm. around 3.58 Ma
Hedate Fm. 5.6 ~ 3.75 Ma
Chikura Group
Nishizaki Fm. 9.9 ~ 4.18/5.20 Ma
Toyofusa Group
Nagaogawa ss Mem. (Hata Fm.)
Takigawa Fm.
Hata Fm. 1.95 ~ 0.85 Ma
Higashinagata Fm.
Rendaiji Congl. Mem. (Mera Fm.) Mera Fm. 3.31 ~ 1.21 Ma
Miura Group
139°59'E
≤0.78 Ma
Kamo Fm.
1.21 ~ 0.78 Ma
Shiramazu Fm. 3.68 ~ 3.31 Ma Nojimazaki Congl. Mem. (Shirahama Fm.) Shirahama Fm.
Fault
Syncline, anticline
Figure 2. Geologic map of the southern tip of the Boso Peninsula, showing the mapped areas along the southern coastlines. Adopted from Kawakami and Shishikura (2006).
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STRATAL DISTRIBUTION AND DEFORMATION STRUCTURES Five coastal exposures, mostly of the lower parts of the Chikura Group, were examined in detail (Fig. 2). At the Nojimazaki area the Shirahama and Shiramazu Formations are exposed; at the Shioura, Okawa, and Kawaguchi-Kotto areas the Shiramazu Formation is exposed, and at the Nemoto-Kido-Mera areas the Mera Formation is exposed. Nojimazaki Area At the Nojimazaki area the Shirahama Formation is distributed, forming gentle folds of some hundred meters halfwavelength (Iitsuka, 1999; Futakami et al., 2001) (Fig. 3). Most of the Shirahama Formation is composed of 100% volcanic arc detritus, such as basaltic-andesitic tuff, and it interfingers with siltstone of the Shiramazu Formation. The Izu arc is the likely source for the volcanic detritus, because offshore of Boso, such volcanic debris is derived only from the Izu arc. The Nojimazaki Conglomerate Member, which comprises abundant andesitic or basaltic cobbles and pebbles with minor sandstone and chert pebbles, is a zone within the Shirahama Formation at the Cape Nojimazaki lighthouse. Clasts of sandstone and chert indicate a source in the Kanto Mountains on mainland Honshu, Japan. The presence of Honshu arc– and Izu arc–derived debris requires that the conglomerate was deposited where it received coarse cobbles from both sources. Such an environment would be the proto-Sagami trough, because this would be a submarine trough between the mainland and the Izu arc. Two kinds of chaotic deposits are developed in this area (Fig. 3). One type, exposed at the northeast corner of Cape Nojimazaki, is composed of clasts of folded siltstone of the Shiramazu Formation that were injected into the Shiramazu-Shirahama formation boundary. The second type of chaotic deposit is exposed west of Odo harbor, to the east of Cape Nojimazaki. These deposits lack siltstone blocks and contain many fragments of a chemosynthetic clam fossil (Calyptogena) and calcareous sandstone blocks (calcite-cemented chimneys or pipes) (Iitsuka, 1999; Futakami et al., 2001) (Fig. 4). In both cases, matrices are coarse volcanic fragments. Shioura Area At Shioura, a thrust-anticline on a 100 m scale was developed in the lower Shiramazu Formation with many south-verging small thrusts (Figs. 5 and 6). The displacement of the main thrust is measured at a maximum of 17 m by means of tracing layers dragged across the fault. This thrust is slightly above the contact with the underlying Shirahama Formation, in which a bedding subparallel thrust forms a southwestward thrust anticline (Fig. 5B). The bedding subparallel thrusts may be common large and small structures in this area.
Okawa Area In the Okawa area, long coastal exposures display various structures that accommodate layer-parallel shortening of the lowermost zone of the Shiramazu Formation, predominantly as duplexes, and also as thrust-anticlines (Figs. 7 and 8). Two blocks of the subjacent Shirahama Formation are overthrust from the north (Figs. 7 and 8A). This may be coincident with many small thrusts in the Shirahama Formation, in which thrusts come up from the subhorizontal sliding plane (Fig. 8B). Examples of thrust-anticlines, all verging south, are shown in Figures 9 and 10. Many conjugate sets of thrust faults are developed in front of the thrust-anticlines, indicating 300% horizontal shortening (Figs. 7 and 8C). A large-scale conjugate set contains small-scale sets. Duplex structures along bedding-parallel thrusts are commonly observed in various zones (Fig. 8D). These structures are shown in Figure 7B. Kawaguchi-Kido-Kotto Area Middle to upper strata of the Shiramazu Formation appear to have fairly uniform, steeply dipping orientations along the north-south–trending coast, extending from Kawaguchi to Kotto. Detailed mapping, however, reveals complicated deformation structures (Aung, 2007) (Fig. 11). In the Kawaguchi area, chaotic deposits, which include fragments of Calyptogena and vein-structure–bearing siltstone blocks, are intercalated within alternating sandstone and siltstone. Because the vein-structure– bearing siltstone is diagnostic in the subjacent Miura Group, which is exposed just north, such blocks must have been derived from there. We interpret some chaotic beds as debris flow beds because their clasts were derived from subjacent layers, whereas we believe that some chaotic beds without Calyptogena formed as a result of injection or liquefaction. This is because their components are derived from coeval, stratigraphically equivalent zones. At least eight sets of thrust-anticlines and many duplexes are developed in this area (Aung, 2007) (Fig. 11). Layer-parallel faults and low-angle thrusts are abundantly developed in the Chikura Group. South of this area a low-angle thrust as a detachment fault steps up from the lower beds. In one example (TAK1, 2) at the toe of this thrust a single layer is repeated seven times, forming an antiformal duplex stack, and this was folded further at a later stage, forming complicated thrust-duplexes (Fig. 12A). This structure also exhibits back thrusts verging northeast, but most of the thrust structures verge southwest. To the north in the Kido area, a large thrust-anticline (TAK3, 4) is developed, in an L shape, with tripling of the same beds (Fig. 12B). Many folds of approximately a few meters halfwavelength are developed at Kido harbor. Farther north, in the Kotto area, two large thrust-anticlines are developed (TAK6, 8) (Figs. 12C and 12D). Many duplexes are interspersed with overthrusts. Within the large-scale thrust-anticlines and the smaller scale thrust-anticlines, a
Recognition of a trench-fill type accretionary prism, Boso Peninsula, Japan
Conglomerate and sandstone Volcanic sandstone Tuffaceous sandstone and siltstone Sandstone and siltstone Chaotic deposit 1 Chaotic deposit 2 Strike trace Fault Anticline Syncline
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Road
Concrete
18 15 22
18 Concrete 25 20
12
22
Odo harbor
45 43 22 32 32
Nojimazaki lighthouse 17
28
24
23 35 44
30
N
34°54'
Sea
139°53'
100 m
Cape Nojimazaki
Figure 3. Detailed geologic map of the Cape Nojimazaki area, after Iitsuka (1999). Two types of chaotic deposits are explained in the text.
A
B
Figure 4. Outcrop photographs of chaotic deposits: (A) with Calyptogena fragments and calcite-cemented sandstone blocks to the west of Odo harbor, and (B) with folded siltstone layers, without Calyptogena fragments, at Cape Nojimazaki. In both cases the matrix is coarse sandstone and pebbles, showing evidence of liquefaction and injection.
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A
A main thrust
12
Younging of bed
15
Reverse fault Strike trace
22
road
Shioura 20 Fig. 6A
Sea
Fig. 6B
18 80 20
B
N 100 m
B N 100 m
Figure 5. (A) Sketch map of the Shioura area, showing a thrust anticline of south vergence. (B) Schematic plan section from above at Shioura, showing a thrust anticline, the representative structure of the Shiramazu Formation.
Figure 6. Outcrop photograph of the thrust anticline at Shioura: (A) looking east; (B) looking west.
Recognition of a trench-fill type accretionary prism, Boso Peninsula, Japan horizontally duplicated duplex structure, a kind of antiformal stack-type duplex, is involved. Such duplexed layers are duplicated again by later thrusts, so there were at least two stages of thrust-duplex development, subparallel to the bedding plane (Fig. 11B).
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This formation, to the west of Cape Nojimazaki, displays repetition of several folds of one hundred meters half-wavelength (Figs. 13C-1 and 13C-2). Many duplications of a single bed are further duplicated, and are finally isoclinally folded, at Nemoto (Fig. 14). As shown in this figure, at least three stages of shortening are observed.
Nemoto-Mera Area Other Areas: A Large-Scale Submarine Slide Body The Mera Formation, above the Shiramazu, is characterized by convolute folds with northeastward vergence and smallscale thrusts with southwestward vergence (Figs. 13A and 13B).
A
Younging of bed Reverse fault Strike trace
DISCUSSION
Shirahama Formation Conjugate sets of thrust anticline
The outcrop of a large-scale submarine slide body, first recognized by us in the Hata Formation, north of Shirahama, was described by Yamamoto et al. (2007) (Fig. 15). The unit consists of large blocks in a sand matrix that appears to originally have been mobilized by liquefaction, all of which came from the same formation. These slide bodies appear to have been derived from the south, based on features in another small slide body south of this large-scale slide body (Ito and Sugiyama, 1989). Similar chaotic submarine slide bodies are found in the same zones of the Hata Formation, corresponding to a depositional age of ca. 2.0 Ma (Kawakami and Shishikura, 2006). These slide bodies show that northward dipping of the previously horizontal or south-tilting slope began during deposition of the Hata Formation.
As described above, the Chikura Group displays many kinds of structures, showing bedding-parallel shortening, mostly of southward vergence, but also with some northward vergence. We discuss their development below, starting with deposition and ending with deformation. The concluding interpretation is illustrated in Figure 16.
Fig. 10 Fig. 9
Sea
Where Was the Chikura Group Deposited? N 50 m
B
50 m
N
Figure 7. (A) Sketch map of the Okawa area, showing thrust anticlines and conjugate sets of thrust faults in front. Notice that two slices are thrust over, toward the south, from the subjacent Shirahama Formation. (B) Schematic cross section from above at Okawa, showing a thrust anticline as the basic structure of the Shiramazu Formation.
Most researchers have concluded that the Chikura Group was deposited in a trench-slope basin or forearc basin. Among them, Kotake (1988) noted that the sedimentation and deformation occurred on the trench or trench slope. However, looking at his reconstruction of the sedimentation, we view this as being incompatible with accretion on the Honshu arc side. Kawakami and Shishikura (2006) describe very detailed stratigraphic correlation of the total Chikura Group, based on precise tracing of many key tuff beds, but they only depict stratigraphic positions with no real discussion of sedimentary environments. Aung (2007) found many bedding-parallel shortening structures for the first time, but he thought they were all due to gravitational sliding, in spite of an absence of known normal faults. Those researchers did not account for the origin and depositional environment of large amounts of Izu arc–derived volcanic rocks, nor did they explain the mechanism of layer-parallel shortening. Considering that the Shirahama and Shiramazu Formations are interfingered with each other, and that the former is dominantly composed of the direct deposition of the coarse Izu
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volcanic-arc materials, it is reasonable to think that at least the lower, southern part of the Chikura Group was deposited in front of the incoming Izu arc on the lower plate of the subduction zone (Fig. 16, stage 1). The Nojimazaki Conglomerate Member, locally intercalated between the Shirahama and Shiramazu, is composed mainly of basaltic and andesitic cobbles and pebbles from the Izu arc with subordinate ones of sandstone and chert from the Honshu arc. This suggests that the sedimentary basin was on the lower plate of the subduction boundary (trench or paleo-Sagami trough) between the Philippine Sea plate (for the Izu arc) and the Eurasia plate (for the Honshu arc) (Fig. 16, stage 2). One more important component is that most of the beds are generally dipping to the north, with some rare southward dips originating by means of folds and faults. However, in the
northernmost part, the Chikura Group unconformably overlies the Miura Group, producing a northward-onlapping relationship between the two units. If the sedimentary environment was in a trench along the subduction boundary, the basin must have been inclined to the north during the final stage and overlapping to the north, above the already accreted Miura accretionary prism, when the sedimentation that preceded the Shiramazu Formation (and its uppermost part) occurred (Fig. 16, stage 3). Most of the bedding parallel shortening may have occurred during that stage of sedimentation. How Were the Chaotic Deposits Formed? The chaotic deposits of the Chikura Group hold the key to reconstructing the depositional environment. Other than the
A
B
C
D
Figure 8. Representative outcrop photographs of faults and folds of the Shirahama and Shiramazu Formations. (A) The Shirahama Formation thrusts over the Shiramazu southward at Okawa. (B) Subhorizontal thrust fault cuts the Shirahama Formation southward, east of Okawa. (C) Conjugate thrust faults in the Shiramazu Formation at Okawa. (D) Duplex structures in the Shiramazu Formation in the Kotto area.
Recognition of a trench-fill type accretionary prism, Boso Peninsula, Japan debris flow or slump type deposits, many chaotic deposits on the Boso Peninsula are of diapir or injection origin. Two types of diapiric chaotic units occur in the study area: One contains chemosynthetic clam fragments and calcareous sandstone blocks and pipes, whereas the other contains folded siltstone blocks. The matrices in both types are coarse sands and pebbles of volcanic origin. Both types intrude the Shirahama Formation as dikes or sills. The injection process of these chaotic deposits may be due to rising of the pore-fluid pressure of the lower Shiramazu Formation. The mud diapir, which was caused by high fluid pressure, brought up siltstone blocks with Calyptogena fossil fragments, probably related to the melting of methane hydrate, along with related clams and calcareous beds (Iitsuka, 1999; Futakami et al., 2001).
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A
A N 5m Reverse fault Younging of bed N Hammer
5m
B
Bag for scale Reverse fault Younging of bed
B
B Hammer
Looking west
Looking northeast
Figure 9. Sketch map and outcrop photographs of folding, faulting, and repetition of key beds at Okawa, west of the area of Figure 10.
Figure 10. Sketch map and outcrop photograph of folding, faulting, and repetition of key beds at Okawa, east of the area of Figure 9.
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A
TAK8 TAK7
Kotto
TAK6
Younging of bed Fault Strike trace
TAK5
Kido TAK4
TAK3
Sea
In the north, some chaotic deposits contain vein structure–bearing siltstone, which is common in the subjacent Miura Group, strongly suggesting that the Miura accretionary prism supplied material to the chaotic deposits by debris flows or by diapirs from underneath. As known from the KAIKO project, particularly by the research using the submersible, Calyptogena colonies and calcareous-cement sediments dominantly occur on the seabed in the trench’s landward slope, particularly on the deformation front or large thrust trace, where there is seepage of methanebearing fluid (Kobayashi, 2002). As a result, sedimentation with methane-bearing fluid seepage is concentrated at the trench and along the trench slope close to large thrusts (Fig. 16). As for the large-scale submarine slide bodies in the Hata Formation, we propose the following model. Before formation of these slide bodies, the basin (a trench filling basin) was wide with some thrusts cutting the trench slope, supplying methane-bearing fluid along the thrusts; but after the southern part became uplifted and inclined to the north, the situation changed. This change included a jump of the deformation front to the south, and at the same time a jump of the sedimentary basin to the north (Figs. 16, stages 5 and 6). The lack of strong deformation of the Hata Formation and superjacent strata suggests that the depositional environment became like a forearc basin or trench-slope basin, with northward dipping and scant horizontal shortening. What Does the Difference in Structure between the Shiramazu-Mera and Hata Mean?
Kawaguchi TAK2
139°58'
TAK1 100 m
B
34°56'
200 m
N
The occurrence of complicated folds and thrusts in the Chikura Group, mostly in the Shiramazu and Mera Formations, is significant for the tectonic reconstruction. The duplicated layers, or duplex structures, were further duplicated in the second stage thrust-anticlines. The former stage commonly resulted in three to five times, or even more, repetition of a single layer, which was further folded or thrusted in a later stage. Such structures are not developed in the superjacent Hata Formation, suggesting that folding and thrusting occurred before deposition of the Hata Formation. SUMMARY AND CONCLUSIONS
Figure 11. (A) Simplified sketch map of the Shiramazu Formation in the Kawaguchi-Kotto area, showing major thrust anticlines, TAK1 through TAK8. Largely revised after Aung (2007). (B) Schematic cross section from above at Kawaguchi-Kido-Kotto, showing thrust-anticlines, the representative structures of the Shiramazu Formation.
Based on the description above, the following stages of development of sedimentary succession and deformation are summarized for the history of the Chikura Group. During the first stage the lower Shirahama Formation filled the trough between the Philippine Sea plate subduction boundary below the Eurasia plate. The lower Shiramazu Formation from the Izu arc side interfingered with the Shirahama Formation (Fig. 16, stage 1). Detachment faults stepped up through these formations from the subduction boundary to the surface of the Shiramazu Formation. Methane-bearing fluid seeped through the detachment
Recognition of a trench-fill type accretionary prism, Boso Peninsula, Japan faults, and where these faults daylighted, colonies of Calyptogena were developed (Fig. 16, stage 2). During the second stage the upper Shirahama Formation was deposited on these formations, and conglomerate composed mostly of volcanic material including pebbles from the Honshu arc were deposited in the canyon between the two arcs. Again, the upper Shiramazu Formation covered the upper Shirahama Formation. Subsequently, mud diapirs rose from the lower Shiramazu Formation to the upper Shiramazu. In addition, at the surface, debris flows from the collapsed slope sediments of the Miura Group formed the chaotic deposits that contain fragments of Calyptogena and vein structure–bearing siltstone (Fig. 16, stage 2). In the third stage the Shirahama and Shiramazu Formations were deposited on the previous formations, resulting in interfingering relationships. Then the detachment faults further stepped up, developing thrust-anticlines and duplexes. After
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that, deposition of the Shirahama and Shiramazu Formations ended (Fig. 16, stage 3). During the fourth stage the Mera Formation was deposited on the Shiramazu Formation. When deposition of the Mera ended, the existing detachment faults stepped up through the sequence, causing development of further thrust-anticlines and duplexes. In the upper Mera Formation, folds of some hundred meters half-wavelength were developed (Fig. 16, stage 4). During the fifth stage the subduction front jumped southward, accreting the Shirahama, Shiramazu, and Mera Formations to the Miura Group side (Fig. 16, stage 5). In the sixth and final stage the Hata Formation covered the slope basin on top of the Chikura Group. During Hata Formation deposition the Chikura Group began to be inclined northward. At the same time the chaotic sedimentary body of the Hata Formation was developed from south to north (Fig. 16, stage 6).
A
B
C
D
Figure 12. Representative outcrop photographs of thrust anticlines and associated duplex structures. (A) Kawaguchi area; (B) Kido area; and (C, D) Kotto area.
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A
B
C-1
C-2
Figure 13. Representative outcrop photographs of the Mera Formation. (A) Northward inclined convolution; (B) southward thrust structure, both at Nemoto; and (C) folds at Mera. C-1 is an anticline and C-2 is a syncline.
B
A
B
Thrust
Hammer
Antic
line
40 38
Thrust
Thr
ust
Synclin
e
35 70
N 65 3m
Reverse fault Younging of bed
Figure 14. Sketch map (A) and outcrop photograph (B) of the three-stage deformation of the folded duplex structure at Nemoto.
Recognition of a trench-fill type accretionary prism, Boso Peninsula, Japan
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Figure 15. Large-scale submarine slide from the right side (south) in the Hata Formation, north of Shirahama.
S
Paleo-Sagami trough
N
ca. 10 km
1
Izu arc Philippine Sea plate
2 3 Legend
4
Chikura Group Hata Fm. Mera Fm. Shiramazu Fm.
5
Nojimazaki Congl. Mem. (Shirahama Fm.) Shirahama Fm.
Miura Group Nishizaki Fm.
Sagami trough
6
Debris flow Calyptogena fossil Mud diapir
Figure 16. Cartoon showing progressive development of the Chikura Group from stages 1 through 6. Notice that deformation occurs synchronously with sedimentation in the trench-fill deposits between the Izu and Honshu arcs. Calyptogena (methane-bearing fluid supported) clam colonies (marked by X) are present at the thrust on the seabed, showing southward prograding of the thrust front. At stage 6 a jump of the large thrust causes uplift of the southern edge to provide a trigger for large-scale submarine slide bodies from the Hata Formation to the north. See text for further explanation.
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ACKNOWLEDGMENTS We acknowledge John Wakabayashi and Yildirim Dilek for inviting us to contribute this chapter. Field data are adopted from earlier work by Takeshi Kuwano, Akinori Iitsuka, and Than Tin Aung, whom we thank. Field assistance by Yoko Michiguchi and other Press laboratory members, University of Tsukuba, is also greatly appreciated. The English revision, aided by helpful discussions with Richard Walker, is likewise greatly appreciated. REFERENCES CITED Aung, T.T., 2007, Tectonically induced gravity sliding: Recognition and identification of examples from Pliocene Chikura Group, southernmost Boso Peninsula, Japan: Earth Evolution Sciences, v. 1, p. 15–20. Futakami, M., Ito, M., and Matsukawa, M., 2001, Conglomerates with Vesicomyid Bivalvia (Calyptogena) of the Shiramazu Formation of the Chikura Group in Shirahama, Chiba—An example of mud diapirs: Journal of the Geological Society of Japan, v. 107, p. 611–619. Hanamura, Y., and Ogawa, Y., 1993, Layer-parallel faults, duplexes, imbricate thrust and vein structures of the Miura Group: Key to understanding the Izu fore-arc sediment accretion to Honsyu forearc: Island Arc, v. 3, p. 126–141. Iitsuka, A., 1999, Occurrence and origin of breccia-like deposits in the Shiramazu Formation, southernmost part of the Boso Peninsula [B.S. thesis]: Tsukuba, Japan, University of Tsukuba. Ito, T., and Sugiyama, S., 1989, Basal structures of the Pleistocene Chikura submarine sliding sheet in the southernmost Boso Peninsula, central Japan, in Taira, A., and Masuda, F., eds., Sedimentary Facies in the Active Plate Margin: Tokyo, Terra Publications, p. 511–528. Kameo, K., and Sato, T., 1999, Recent development of calcareous nannofossil biostratigraphy and its application: Neogene and Quaternary stratigraphy of offshore Joban based on calcareous nannofossils: Journal of the Japanese Association for Petroleum Technology, v. 64, p. 16–27. Kameo, K., Saito, K., Kotake, N., and Okada, M., 2003, Late Pliocene sea surface environments in the Pacific side of central Japan based on calcareous nannofossils from the lower part of the Chikura Group, southernmost part of the Boso Peninsula: Journal of the Geological Society of Japan, v. 109, p. 478–488. Kanie, Y., Hattori, M., Kuramochi, T., Okada, H., Ohba, T., and Honma, C., 1997, Two Vesicomyid Bivalvia in the Shiramazu Formation of the Chikura Group in the southern-most part of the Boso Peninsula: Journal of the Geological Society of Japan, v. 103, p. 794–797. Kawakami, S., and Shishikura, M., 2006, Geology of Tateyama District: Geological Survey of Japan, Quadrangle Ser., scale 1:50,000, AIST, 82 p. Kawamura, K., Ogawa, Y., Anma, R., Yokoyama, S., Kawakami, S., Dilek, Y., Moore, G.F., Hirano, S., Yamaguchi, A., Sasaki, T., YK05-08 Leg 2, and YK06-02 Shipboard Scientific Parties, 2009, Structural architecture and active deformation of the Nankai Accretionary Prism, Japan: Submersible survey results from the Tenryu Submarine Canyon: Geological Society of America Bulletin, v. 121, p. 1629–1646, doi:10.1130/B26219.1. Kobayashi, K., 2002, Tectonic significance of the cold seepage zones in the eastern Nankai accretionary wedge—An outcome of the 15 years’ KAIKO projects: Marine Geology, v. 187, p. 3–30, doi:10.1016/S0025 -3227(02)00242-6.
Kotake, N., 1988, Upper Cenozoic marine sediments in southern part of the Boso Peninsula, Japan: Journal of the Geological Society of Japan, v. 94, p. 187–206. Nakamura, K., Renard, V., Angelier, J., Azema, J., Bourgois, J., Deplus, C., Fujioka, K., Hamano, Y., Huchon, P., Kinoshita, H., Labaume, P., Ogawa, Y., Seno, T., Takeuchi, A., Tanahashi, M., Uchiyama, A., and Vigneresse, J.L., 1987, Oblique and near collision subduction, Sagami and Suruga Troughs: Preliminary results of the French-Japanese 1984 Kaiko cruise, Leg 2: Earth and Planetary Science Letters, v. 83, p. 229–242, doi:10.1016/0012-821X(87)90068-9. Ogawa, Y., and Taniguchi, H., 1988, Geology and tectonics of the Miura-Boso Peninsulas and the adjacent area: Modern Geology, v. 12, p. 147–168. Ogawa, Y., Seno, T., Tokuyama, H., Akiyoshi, H., Fujioka, K., and Taniguchi, H., 1989, Structure and development of the Sagami Trough and off-Boso triple junction: Tectonophysics, v. 160, p. 135–150, doi:10.1016/0040 -1951(89)90388-0. Ogawa, Y., Takami, Y., and Takazawa, S., 2008, Oblique subduction in an island arc collision setting: Unique sedimentation, accretion and deformation processes in the Boso TTT-type triple junction area, NW Pacific, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 155–170, doi:10.1130/2008.2436(07). Pautot, G., Nakamura, K., Huchon, P., Angelier, J., Bourgois, J., Fujioka, K., Kanazawa, K., Nakamura, Y., Ogawa, Y., Séguret, M., and Takeuchi, A., 1987, Deep-sea submersible survey in the Suruga, Sagami and Japan trenches, preliminary results of the 1985 Kaiko cruise, Leg 2: Earth and Planetary Science Letters, v. 83, p. 300–312, doi:10.1016/0012 -821X(87)90073-2. Saito, S., 1992, Stratigraphy of Cenozoic strata in the southern terminus area of the Institute of Geology and Paleontology: Sendai, Japan, Tohoku University, v. 93, p. 1–37. Saito, T., 1999, Revision of Cenozoic magnetostratigraphy and the calibration of planktonic microfossil biostratigraphy of Japan against this new time scale: Journal of the Japanese Association for Petroleum Technology, v. 64, p. 2–15. Sato, H., Hirata, N., Koketsu, K., Okaya, D., Abe, S., Kobayashi, R., Matsubara, M., Iwasaki, T., Ito, T., Ikawa, T., Kawanaka, T., Kasahara, K., and Harder, S., 2005, Earthquake source fault beneath Tokyo: Science, v. 309, p. 462–464, doi:10.1126/science.1110489. Seno, T., Ogawa, Y., Tokuyama, H., Nishiyama, E., and Taira, A., 1989, Tectonic evolution of the triple junction off central Honshu for the past 1 million years: Tectonophysics, v. 160, p. 91–116, doi:10.1016/0040-1951 (89)90386-7. Shishikura, M., 2003, Cycle of interpolate earthquake along the Sagami Trough, deduced from tectonic geomorphology: Bulletin of the Earthquake Research Institute, University of Tokyo, v. 78, p. 245–254. Taira, A., Tokuyama, H., and Soh, W., 1989, Accretion tectonics and evolution of Japan, in Ben-Avraham, Z., ed., The Evolution of the Pacific Ocean Margins: New York, Oxford University Press, p. 100–132. Yamamoto, Y., and Kawakami, S., 2005, Rapid tectonics of the Late Miocene Boso accretionary prism related to the Izu-Bonin arc collision: Island Arc, v. 14, p. 178–198, doi:10.1111/j.1440-1738.2005.00463.x. Yamamoto, Y., Ogawa, Y., Uchino, T., Muraoka, S., and Chiba, T., 2007, Largescale chaotically mixed sedimentary body within the Late Pliocene to Pleistocene Chikura Group, Central Japan: Island Arc, v. 16, p. 505–507, doi:10.1111/j.1440-1738.2007.00587.x.
MANUSCRIPT ACCEPTED BY THE SOCIETY 21 DECEMBER 2010
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The Geological Society of America Special Paper 480 2011
Implication of dark bands in Miocene–Pliocene accretionary prism, Boso Peninsula, central Japan Yoko Michiguchi* Yujiro Ogawa† Doctoral Program in Earth Evolution Sciences, University of Tsukuba, Tsukuba 305-8572, Japan
ABSTRACT Thin, planar, dark, lamination-like bands are found in host siltstones in the Miocene-Pliocene metamorphosed Miura-Boso accretionary prism, southern Boso Peninsula, Japan. We classified the bands into four types on the basis of distribution, crosscutting relations, and internal textures. Type 1-1 dark bands are developed parallel to the bedding plane and do not include crushed or deformed grains within the band. Type 1-2 bands are also developed parallel to the bedding plane, but grain alignment within the band cuts obliquely across that in the host rock. Type 2 bands include ductilely deformed grains similar to an S-C′ structure, whereas type 3 bands have cataclastic grains. All the dark bands except type 1-1 (being an open fracture with little displacement) are shear bands or slip planes formed from sedimentation to accretion, although the formation mechanisms between the four types are different. These deformation bands are affected by the state of consolidation and magnitude of stress during formation, reflecting the deformation processes. Type 1-1 bands show evidence of independent particulate flow from excess pore-fluid-pressure generation, which occurs just after sedimentation. Type 1-2 bands are flexural-slip faults formed during formation of folds; type 2 bands are sliding planes formed from submarine landslides, whereas type 3 bands are thrust faults formed during accretion.
INTRODUCTION When sediments are brought into the frontal zone of an accretionary prism, they have already undergone various stress conditions such as gravitational stress, compressive stress from tectonics, and pore-fluid pressure. The changes in stress conditions influence physical properties of the sediments such as permeability, strength, controlled deformation, and controlled fracturing. Therefore, the understanding of stress history soon after
sedimentation, in terms of these factors, is extremely important in interpreting the deformation processes. We investigated thin (a few to tens of millimeters thick), planar, dark, lamination-like structures developed in the siltstones at Nishikawana in the southern Boso Peninsula in central Japan, because of the insight they might give into the stress and strain history of sediments that were newly incorporated into an accretionary prism. The sedimentary strata in this area are known as the Miura-Boso accretionary prism, which is of Miocene–Pliocene
*Current address: Japan Nuclear Energy Safety Organization, Toranomon 4-3-20, Tokyo 105-0001, Japan;
[email protected]. † Current address: Century Tsukuba-Miraidaira C-740, Obari, Yokodai 1-127-2, Tsukubamirai-shi 300-2358, Japan. Michiguchi, Y., and Ogawa, Y., 2011, Implication of dark bands in Miocene–Pliocene accretionary prism, Boso Peninsula, central Japan, in Wakabayashi, J., and Dilek,Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 247–260, doi:10.1130/2011.2480(12). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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age and unmetamorphosed (Ogawa and Taniguchi, 1988; Ogawa et al., 1989; Yamamoto, 2006) (Fig. 1). These strata afford an excellent opportunity to study deformational structures from the pre-accretionary stage to the post-accretionary stage. Such lamination-like structures are developed at various places and are important for understanding the deformational history before, during, and after formation of the accretionary prism. We refer to such structures as dark bands (Fig. 2). Dark bands have been interpreted as healed faults in the lower Miocene Emi Group (Ishimaru and Miyata, 1991) and as layer-parallel faults during accretion in the Miura Group of the Miura-Boso accretionary prism (Yamamoto et al., 2000). They are thought to be formed during the early stage of accretionary prism development. In convergent plate margins such as Nankai, Chile, and Costa Rica, similar dark bands have been interpreted as deformation bands or compactive shear bands, which are oblique to the bedding plane (Lundberg and Moore, 1986; Maltman et al., 1993; Ujiie et al., 2004). They appear as kink-band, shear-zone, or fault-like features. The structures that form the toe of the Nankai accretionary prism are also thought to be earlystage structures in an accretionary prism. This is because they are reported only from present convergent plate margins (Ujiie et al., 2004) and not from on-land accretionary prisms (except for the Miura-Boso prism). As a result of our detailed fieldwork at Nishikawana, we classified the dark bands into four major types on the basis of distribution, crosscutting relations, and internal structures (Fig. 3). Dark bands develop at various places and in relation to various structures, such as parallel to a bedding plane, along the sliding plane of inferred submarine landslide deposits, along the thrust fault in an accretion stage, etc. (Figs. 2 and 3). Therefore, dark bands are inferred to form in various stages such as before and during accretion. Deformation structures in porous sandstone are known as deformation bands (Fossen et al., 2007). They are lowdisplacement deformation zones that are millimeters to centimeters thick, and are similar to dark bands at a glance. They have
PAC
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Figure 1. Plate boundaries of the Izu-Bonin island arc, a plate collision zone, and the only trench-trench-trench (TTT)–type Boso triple junction (southeast of the Boso peninsula). Arrow on top indicates the study area, Nishikawana. EUR—Eurasian plate, NAM—North American plate, PAC—Pacific Ocean plate, PHS—Philippine Sea plate.
been classified on the basis of internal structures into subtypes reflecting deformation mechanisms, granular flow, cataclasis, phyllosilicate smearing, dissolution, and cementation. These types are dependent on the internal and external conditions during deformation, such as mineralogy, grain size, cementation, porosity, stress state, etc. (Hirono, 2005; Fossen et al., 2007). Deformation bands in siltstone on land have scarcely been studied. This study attempts to interpret the formation mechanisms of dark bands developed in siltstones on the basis of fieldwork and microscopic observation. Fundamental questions we address in this chapter include the following: How are dark bands formed? What are the stress conditions, consolidated conditions? How are dark bands related to accretionary prism development? GEOLOGICAL SETTING OF THE STUDY AREA The southern Boso Peninsula lies in central Japan, east of the Izu arc collision zone, where the north-south–trending Izu-Bonin arc collides with the Honshu arc (Ogawa et al., 2007; Fig. 1). The southern Boso Peninsula is composed mainly of a shallow accretionary prism (Miura-Boso accretionary prism) including the upper Miocene–lower Pliocene Miura Groups, deposited in the deep marine environment of the Izu-Bonin forearc (Ogawa and Taniguchi, 1988). The sediments of the Miura-Boso accretionary prism did not undergo severe metamorphism because the sediments were quickly incorporated into the accretionary prism during arc-arc collision and then uplifted to the Honshu forearc area without deep burial (Ogawa and Taniguchi, 1988; Ogawa et al., 1989; Yamamoto et al., 2005; Yamamoto, 2006). Therefore, the strata preserve several deformational structures under pre-lithified conditions from pre- to post-accretion, which are (1) paleo-liquefied layers and vein structures such as microfractures developed oblique to the bedding planes that may have formed from shearing and/or earthquake motion soon after sedimentation, (2) submarine landslide deposits that may have formed from increased slope instability, and (3) thrust faults, folds, and other structures directly associated with subduction (Hanamura and Ogawa, 1993; Ogawa and Taniguchi, 1988; Ogawa et al., 1989; Yamamoto, 2006; Ohsumi and Ogawa, 2008). The main study area is Nishikawana, southwest of Tateyama on the southern Boso Peninsula (Fig. 1), where the Nishizaki Formation of the Miura Group is distributed (Kawakami and Shishikura, 2006). On the basis of a maximum estimated paleotemperature of <50 °C and a high porosity (30%–50%), Yamamoto et al. (2005) suggested that these rocks had not been buried deeper than 1000 m. Strata are composed mainly of alternating volcaniclastic sandstones and diatomaceous siltstones of deep-sea depositional environments. CHARACTERISTICS OF DARK BANDS We classify the dark bands found in the rocks of the field area into four major types: type 1-1, type 1-2, type 2, and type 3.
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Volcaniclastic sandstone
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Figure 2. Outcrop photos of dark bands at Nishikawana. Dark bands are thin (a few to tens of millimeters thick), planar, dark, and laminationlike structures. They are classified into four major types: type 1-1, type 1-2, type 2, and type 3. As the occurrences of types 1-1 and 1-2 are similar, we showed them as type 1 bands in this figure. (A–C) Type 1: dark band developed parallel to the bedding plane. (C) Some type 1 bands cut vein structures, and others do not. (D) Type 2: dark band on submarine sliding plane. (E) Type 3: dark band developed along the thrust during accretion.
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Types 1-1 and 1-2 bands are parallel to the bedding, type 2 bands are developed along the basal slide plane of submarine landslide deposits, and type 3 bands are thrust faults that appear to have formed during subduction-accretion (Figs. 2 and 3). Each type of dark band was formed during a different stage or stages of the evolution of these deposits, and by different events. The field appearance and relationships of types 1-1 and 1-2 are similar, so we distinguish them on the basis of their microscopic textures. We observed internal structures of dark bands in thin section under an optical microscope, and cut-surface textures under a scanning electron microscope (SEM). Compositions of minerals in the dark bands and their host siltstones were studied by X-ray diffractometer (XRD) analysis. All the collected samples were friable, so the samples were impregnated to make thin sections with Quetoll-651, a water-miscible epoxy resin, to prevent the collapse and dilation of the texture. We polished the rocks with oil instead of water to prevent collapse and dilatation of the texture. SEM observation was done using a Hitachi S-3000N microscope. The samples were ion-spattered by an Au-Pd target. XRD analysis for major minerals and clay minerals was conducted to determine the composition of the clay minerals concentrated along the dark bands. For clay mineral detection, fractions <2 mm were concentrated by elutriation. In this study, “grains” mean framework mineral grains such as quartz, feldspar, etc., and
“clay minerals” are flakes of clay minerals in the matrix. Porosity measurements for some samples were carried out by a mercury intrusion method, using a porosimeter (Micromeritics Autopore 9520), at the National Institute of Advanced Industrial Science and Technology of Japan. Type 1 Some dark bands vary from parallel to oblique bedding and similarly from thrust to normal faults; but the ones with mostly layer-parallel orientations were classified as type 1. Type 1 bands are developed only in siltstones and parallel to the bedding plane in alternating sandstones and siltstones (Figs. 2A–2C). Some dark bands form a single plane and others anastomose. Type 1 bands cut and dislocate vein structures, and vice versa, although the amount of displacement is not significant (Figs. 2C and 3). Thus we believe that type 1 bands are a type of fault or fracture. We also believe that some type 1 bands occurred almost synchronously with the vein structure formation: i.e., just after sedimentation. The thin sections and SEM images show that the grain size inside type 1 bands is almost the same as that of the host siltstone (Figs. 4 and 5). There are no deformed and/or crushed grains inside type 1 bands and on the boundaries between type 1 bands
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Figure 3. Cutting relations of dark bands in the Nishikawana area. As the occurrences of types 1-1 and 1-2 are similar, we showed them as type 1 in this figure, but they were formed during different stages. Type 1: primary accretion; type 2: submarine landslide; type 3: thrust movement during accretion. We believe that thrust movement could cause a submarine landslide at any time.
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Y R Figure 4. Photomicrographs of dark bands. (A) Type 1-1 band: Grain size inside is almost the same as that of the host rock. There are no deformed or crushed grains inside, but there are rotated grains. Alignment of flakes of clay minerals on each boundary between the dark band and the host rock is symmetrical with respect to the center line of the band. (B) Type 1-2 band: Flakes of clay minerals inside type 1-2 band are distributed pervasively in the dark bands and aligned in two preferred orientations. The dark band of type 1-2 seems to have been formed by a simple shear. (C) Type 2 band: The grains are ductile and deformed along the shear plane. Grains and clay minerals are also arranged in agreement with the directions of the shear plane. (D) Type 3 band: The grain size is smaller than that in the host rock, and clay minerals are mainly arranged in two directions, one parallel and another oblique to the direction of the type 3 band.
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and the host rock. In type 1 bands, two different alignments of flakes of clay minerals were observed, thus further dividing them into type 1-1 bands (Fig. 4A) and type 1-2 bands (Fig. 4B), based on the internal textures. In the case of type 1-1 bands, flakes of clay minerals are concentrated and rearranged to form the preferred orientation,
A Type 1-1
Inside dark band ban
in particular along the boundary between type 1-1 bands and the host rock. Flakes of clay minerals are aligned in two directions between the interior of the band and the band boundary. The obliquity of the two planes indicates shearing along the band boundary, which in this case represents a Riedel shear (Kano and Murata, 1998). Alignment of flakes of clay minerals on each
B Type 1-2
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Figure 5. SEM images. (A) Type 1-1 band: Scaly arrangement of clays is observed at the boundary between type 1–1 and the host rock. (B) Type 1-2 band: Grains array in preferred directions. (C, D) Type 2 band: Deformed grains in two directions are observed inside the sliding zone, similar to those in the S-C′ structure. (E) Type 3 band: Grains are strongly sheared and crushed. (F) Grains in the host rock are not deformed.
Implication of dark bands in Miocene–Pliocene accretionary prism boundary is symmetrical with respect to the dark band center line, suggesting an opposing sense of shear at each boundary. This in turn suggests the intrusion of band material with respect to its boundaries. On the other hand, flakes of clay minerals inside type 1-2 bands are distributed pervasively in the dark bands and aligned in two preferred orientations. The direction of grains within type 1-1 bands is parallel to the alignment of grains in the host rock, but that of type 1-2 cuts obliquely across. The orientation of the clay flakes suggests that type 1-2 dark bands are the product of simple shear.
XRD analysis of types 1-1 and 1-2 bands shows the same mineral components as the host rock, although the clay minerals of the dark bands are richer in smectite than those of the host rock. Using two sets of comparative measurements, we found that type 1-1 bands had 4% and 7% less porosity, respectively, than the host rock (Figs. 6B and 6D), and we consider the uncertainty of the porosity measurement to be less than a few percentages. This suggests that a significant reduction occurred in the type 1-1 bands. The void size distribution of the dark bands has a concentration with the median smaller than that of the host rock, and which has a bimodal or multimodal void space distribution (Figs. 6A–6D).
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Figure 6. Result of porosity measurement by mercury intrusion: The left side shows the result of host rock, and the right side shows that for dark bands. (A–D) Type 1 bands. (E, F) Type 3 bands. Porosities of all samples decreased 4%–7% more than those of the host rock.
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These results suggest that the difference between types 1-1 and the host rock is in the modal arrangement.
DISCUSSION Conditions of Formation
Type 2 Type 2 bands are distributed along the boundaries of the duplicated strata within chaotic blocks and along the contacts of the chaotic blocks and the substratum. We interpret the chaotic blocks as submarine landslide deposits (Michiguchi, 2008). The landslide movement apparently resulted in duplication of strata within the chaotic blocks, based on the scale and nature of folds along the block boundaries. Therefore, we believe that type 2 dark bands form along slide planes developed in and along the basal slip surface of a submarine landslide synchronous with the formation of the chaotic blocks. Type 2 bands include various types of grains (feldspar, quartz, radiolarian tests, etc.) having various sizes (Fig. 4C). Two shear planes were recognized by the arrangement of deformed grains and surrounding smaller grains: one is parallel to and the other is oblique to the boundaries of type 2 bands. In particular, highly deformed large grains form an array parallel to the oblique plane where flakes of clay minerals are concentrated (Fig. 4C). SEM images show that inside the shear planes an S-C′ structure was developed, as defined by the alignment of flakes of clay minerals (Figs. 5C and 5D). The flakes are ductilely deformed. S-C′ orientations indicate a normal sense of displacement within the shear planes (Figs. 5C and 5D). The submarine landslide and its internal structures require a formation that reflects an extensional stress field. Type 3 Type 3 bands are distributed along thrust faults (Figs. 2E and 3) and cut type 1-1 bands. At Nishikawana there are many folds and thrust imbrications that were formed during accretion (Yamamoto and Kawakami, 2005). Type 3 bands are developed along such faults. Petrographic examination shows that the grain sizes of type 3 bands are slightly smaller than those of the host rocks (Fig. 4D). Grain alignments inside type 3 bands cut obliquely to those in the host rock. Grains and flakes of clay minerals array in a preferred orientation, indicating shear planes that are developed in two directions (Fig. 4D). We interpreted them as Y-shear and R-shear planes. The microscopic shear sense observed in type 3 bands corresponds to the macroscopic shear sense of the fault along which type 3 bands are found. Shear planes are developed pervasively in type 3 bands. SEM observation indicated that grains inside type 3 are strongly sheared (Fig. 5E), whereas those in the host rock are not deformed. Thus, type 3 bands are thrust faults. Based on one measurement, the porosity of type 3 bands is ~7% lower than that of the host rocks (Figs. 6E and 6F). Type 3 bands were formed by thrust faulting that accompanied grainsize reduction and texture collapse.
Dark bands exhibit little or no displacement across them. Previous studies concluded that the dark bands formed during subduction-accretion, but our study suggests that some of them also formed before and after accretion, as explained below. Moreover, the characteristics of the dark bands indicate several deformation modes, reflecting different physical properties (permeability, consolidation state, etc.) and stress conditions during deformation. Type 1-1 The arrangement of flakes of clay minerals, composed mainly of smectite, along both boundaries of the bands have a Riedel shear geometry, and this geometry indicates an opposing sense of shear along the two boundaries, symmetrical with respect to the center of the band (Fig. 4A). Therefore, this indicates intrusion of the band into the host rock. The absence of deformed and crushed grains and the presence of rotated grains, in a clay mineral matrix along the boundaries, suggest that the texture of type 1-1 bands was formed by independent particulate flow (Borradaile, 1981). This type of flow represents sliding on grain boundaries without fracturing and deforming the grains, and it occurs by an increase of pore-fluid pressure under a low confining pressure. When pore-fluid pressure increases, sliding occurs easily, because the sliding friction on grain contact decreases proportionally to the effective stress reduction (Borradaile, 1981). Type 1-1 band formation was synchronous with and also postdated the development of veins in these rocks, based on crosscutting relationships. This suggests that type 1-1 bands formed as a consequence of an increase of pore-fluid pressure under un- or semilithified conditions just after sedimentation. The lack of internal shearing within the bands leads us to conclude that type 1-1 bands may have formed as tensile hydrofractures. Type 1-2 Type 1-2 bands internally resemble type 1-1 bands in that they both lack deformed and crushed grains, but they differ from the latter because shearing is much more pervasive within the band (Fig. 4B). Type 1-2 bands are pervasively sheared, as illustrated by their internal structure seen in photomicrographs (Fig. 4B). Grain alignments in the band are not parallel to those in the host rocks. Type 1-2 bands are concentrated in the limbs of synclines developed during accretion. Field investigations of exposures parallel to synclinal fold axes showed that most type 1-2 bands cut and dislocate other faults that obliquely cut the strata and that type 1-2 bands are transitional from layer-parallel faults to
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thrust faults. In fold profile exposures, type 1-2 bands displace strata with a thrust sense of shear. Flexural-slip faults, formed during folding parallel to bedding planes, are well known (Yeats, 1986; Tanner, 1992; Vannucchi et al., 2003). Such flexural-slip faults have a thrust sense of shear due to lateral shortening. The geometry and shear sense of the flexural-slip faults are symmetrical about the axial plane of the fold. The amount of displacement is the largest in the limb of a fold. Such characteristics of flexural-slip faults correspond to those of type 1-2 bands. Thus it is inferred that type 1-2 bands are flexural-slip faults.
are the sliding planes of submarine landslides that occurred under the same unlithified conditions as type 1-1 bands.
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Formation Mechanisms
Thin-section and SEM images show ductilely deformed grains and the preferred orientation of grains (Figs. 4C, 5C, 5D). The displacement of a normal fault sense within type 2 bands and the textures similar to S-C′ structures indicate that type 2 bands were formed during bedding-parallel extension, and that deformation occurred while the beds were subhorizontal. We conclude that type 2 bands
The formation conditions of each type of dark band were inferred on the basis of detailed observations. To evaluate the formation mechanisms of the dark bands we must first review soft sediment deformation associated with burial and compaction, which is also called normal consolidation (Fig. 7, line A–B; Miyata et al., 1993; Jones, 1994; Bolton and Maltman, 1998;
Type 3 Grains inside the bands are crushed and are finer than those of the host rock, suggesting cataclastic deformation by frictional sliding (Figs. 4D and 5E). The Riedel shears within type 3 bands exhibit the same shear sense as the macroscopic shear sense associated with the features at outcrop scale. Thus, type 3 bands are thrust faults.
ine nl al tio e rm lida lin No nso te sta co al itic
Cr
A Deposition
Porosity
Increased pore pressure
D Dewatering
B Drainage condition or rapid loading E
Type 1 Type 2
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F I
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Figure 7. Theoretical stress path for sediments undergoing several phases. Sediments are compacted with appropriate loss of volume during burial (the state on the normal consolidation line, A–B–F). Line B–C is shearing under drained conditions. Lines B–D and F–H denote shearing under overconsolidated conditions by unloading of overburden or development of pore-fluid pressure. Original figure adapted from Bolton and Maltman (1998) and Bolton et al. (1998) with new pass lines from A to I.
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Bolton et al., 1998). Even in the sediments that underwent equal normal consolidation, the style of deformation largely depends on whether sediments are deformed under drained conditions. Excess pore-fluid pressure is readily generated under undrained conditions. An increase in the pore-fluid pressure causes a decrease in the effective stress. Bolton and Maltman (1998) reported from an experimental study that under drained conditions there is no fluid pressure change with progressive compaction, and that the deformation is ductile and strain-hardened with continuous volume loss (Fig. 7, line B–C). In contrast, if excess pore-fluid pressure is generated under undrained conditions, pressurized fluid is introduced into the sediments, inducing a state of overconsolidation (Fig. 7, line B–D), leading to brittle deformation. A shear zone will propagate into nearby areas, and grains will be aligned intensely (Bolton and Maltman, 1998; Bolton et al., 1998). Thus, pore-fluid pressure greatly influences the deformation mode during shearing in unlithified to semilithified sediments. Type 1-1 Two major factors that build up excess pore-fluid pressure in normal consolidated sediments are permeability and sedimentation rate (Jones, 1994; Bolton et al., 1998). It is known that the permeability of mud is generally much lower than the permeability of sand (e.g., Middleton and Wilcock, 1994); therefore, excess porefluid pressure is easily generated along the boundary of different permeability. In addition, when the drainage rate is less than the sedimentation rate of the overlying sediments, excess pore-fluid pressure is also generated. A horizontal layer with higher porefluid pressure is generated in this way to form features such as a water sill (Fyfe et al., 1978) and dish structure (Tsuji and Miyata, 1987). Type 1-1 bands were developed within host siltstones, indicating a critical level of such conditions in the host rock for these structures as well as a layer-parallel zone of dislocation. There may be two types of mechanisms that create a horizontal dark band as a tensile fracture or a shear fracture without significant displacement. In general, a tensile fracture develops parallel to the maximum principal stress (σ1) and perpendicular to the minimum principal stress (σ3; Fyfe et al., 1978). The direction of the maximum stress just after sedimentation is usually perpendicular to the bedding plane, corresponding to the gravitational force. Nevertheless, type 1-1 bands were developed parallel to the bedding plane. One explanation may be that the maximum stress axis is locally parallel to the bedding plane when type 1-1 is formed. This may occur exceptionally in front of the deformation front of an accretionary prism (Mikada et al., 2006). Another potential cause of anomalous stress directions is earthquakegenerated ground motions, which are cyclic and form horizontal shear planes (Ohsumi and Ogawa, 2008) but with minimum resultant displacement. During earthquake movement, liquefaction can occur in places to produce high pore-fluid pressure, and the liquefied materials may inject into the fracture from the host rocks by means of absorption by decreased stress. The preferred
orientation of grains and clay minerals along the boundary, and the reduction of average void space size with respect to those of the host rock, might have formed during these processes (Hirono, 2005). Such phenomena may produce a symmetrical shear sense of orientation inside type 1-1 bands. Type 1-1 bands appear to be the consequence of excess pore-fluid pressure. Thus we suggest that gravitational sliding may have been induced easily if the increase in pore-fluid pressure was large enough, and if the slope was gravitationally semi-stable. The displacements of type 1-1 bands are small, but an earthquake or submarine landslide may have been the cause of the horizontal shear if the effective stress on that plane was nearly zero (Fig. 8A). Type 1-2 Grains inside the bands are brittlely sheared and are arrayed obliquely to the alignment in the host rock. Brittle deformation in soft sediments is known to occur in the state of overconsolidation, which is caused by the decrease in effective stress (Bolton and Maltman, 1998). The folds in this area were formed during the offscraping stage of accretion and later uplift. The uplift and removal of overburden owing to accretion caused the decrease in effective stress, and such folds have shears along the bedding plane with brittle bands known as flexural-slip faults (Fig. 7, line F–H, and Fig. 8C). Accordingly, type 1-2 bands are flexural-slip faults that were associated with fold formation during accretion and subsequent uplift. Type 2 Submarine landslide deposits associated with type 2 bands apparently occurred close to the seabed, because the deposits appear to have been formed under un- or semilithified conditions similar to those associated with type 1-1 bands. However, type 2 bands are characterized by different internal texture than type 1-1: the former is ductile, and the latter is brittle. These textural contrasts may reflect differences in the state of stress during band formation. When rapid loading occurs, such as by a submarine landslide, pore-fluid pressure is generated underneath the landslide body. Then the effective stress is thought to be decreased. However, as suggested by Bolton and Maltman (1998), because the pore-fluid pressure increases in accordance with the increase in total stress to preserve void space, the effective stress does not change. Therefore, the sediments will stay on the normal consolidation line, and they will deform ductilely when shearing occurs (Fig. 7, line B–C). Our field observations appear consistent with this experimental result. Type 3 Thrust faults are formed by an increase in the differential pressure or a decrease in the effective stress owing to uplift and
Implication of dark bands in Miocene–Pliocene accretionary prism removal of overburden. In the case of type 3 bands the grains were deformed brittlely. As above, brittle deformation in soft sediments occurs in the state of overconsolidation (Bolton and Maltman, 1998). In this area, thrust faults occurred during accretion (Yamamoto, 2006). Therefore, it is inferred that this type of band occurs during the offscraping and uplifting stage of accretion, resulting in shearing associated with frictional sliding and cataclasis (Fig. 7, line F–H, and Fig. 8C). The decrease in porosity during formation of type 3 bands may occur by textural collapse, dewatering, or drainage (Fig. 7, line I).
A
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The internal structures of type 3 bands are similar to compactive shear bands, and these bands may in fact be equivalent to such shear bands, which were reported from the Nankai accretionary prism by Maltman et al. (1993) and Ujiie et al. (2004). Implication of Dark Bands All the dark bands in Nishikawana are types of shear planes, with a probable case of tensile fracture and injection such as in type 1-1 bands. However, the texture of each type of band is
σ1
B
σ3
τ
τ σ1
σ3
ղ 3
ԥPf
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σ3* σ * 1
σ*
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ձ 0
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σ3*' σ1*'
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C
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T
0
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Figure 8. Mohr construction of effective normal stress (σ*) against shear stress (τ); σ1* is the effective maximum principal stress, and σ3* is the effective minimum principal stress. T is tensile strength. Adapted from Platt (1990). When the circle is tangent to the failure envelope, failure occurs. If the circle is tangent to the T, the failure mode is the tensile fracture. If the point of tangency is larger than 0 in the σ* axis, it is compressive failure. (A) Type 1-1 band. (B) Type 2 band. (C) Types 1-2 and 3 bands. Pf—pore-fluid pressure.
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1. Sedimentation
2. Slump & slide
3. Fault and fold
4. Thrust movement
Offscraping
Migration pathway of sediment
1. Type 1-1
2. Type 2
Bedding plane
Bedding plane
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~1 mm
~1 mm
3. Type 1-2
4. Type 3 Bedding plane
Horizontal tectonic compression
~1 mm
~1 km Figure 9. Summary of formation stage and internal structure in each type of band. 1: Type 1-1 band, caused by an increase in excess pore pressure from low permeability or liquefaction by earthquake motion, etc., during the sedimentation stage, which is an independent particulate flow. 2: Type 2 band is shearing under normal consolidation during the submarine sliding stage. 3, 4: Type 1-2 band shows a flexural slip fault, and type 3 band shows thrust faulting during accretion with overconsolidation conditions to convert frictional sliding into cataclasis.
Implication of dark bands in Miocene–Pliocene accretionary prism different and was influenced by the stress and consolidated conditions when each dark band was formed (Fig. 9). Type 1-1 bands were formed at the time of generation of pore pressure or shear stress under un- or semilithified conditions just after sedimentation. The generation of excess pore pressure was caused by rapid sedimentation from gravity. An earthquake or submarine landslide may be the cause of the horizontal shear if the effective stress on that plane is nearly zero. Therefore, we consider that the surface just after sedimentation was unstable, possibly caused by an increase in pore pressure and a localized horizontal pressure component. Type 2 bands were the sliding planes of a landslide. The internal structures of type 2 bands suggest that rapid loading caused ductile deformation (Bolton and Maltman, 1998). Thus, type 2 was also formed by gravity. As submarine landslides are thought to have been caused by excess pore pressure, it is inferred that the type 1-1 band was a cause of submarine landslides and that the type 2 band was a result. In contrast, type 1-2 bands were flexural slip faults, and type 3 bands were thrust faults formed as a result of strong lateral compression during accretion. In a subduction zone the compaction bands are formed because of large lateral compression and differential pressures, thus causing tectonic movement. Dark bands are not universally related to accretionary prism formation, for they apparently formed during sedimentation before accretion, during accretion, and after accretion. The stress conditions changed from a tensile field to a compressional field, and from proto-accretion to post-accretion (Fig. 9). Because some type 1-1 bands were caused by gravitational force, acting ubiquitously, they are always apt to occur if the conditions are favorable. Investigating dark bands clarifies the changes in stress conditions from proto-accretion to post-accretion, and the deformation modes responsible for the geological phenomena. What Makes These Bands Black? Physical differences between the dark band and host rock are due to the decrease in porosity and the rearrangement of mineral grains and flakes of clay into preferred orientation because of shearing with escaping fluid, i.e., corresponding to dewatering along the drainage. The color contrast between dark bands and host rock may be associated with a change in the chemicalelectric conditions from oxidation to reduction in order to influence the reduction of several materials, such as organics, oxidized iron, manganese, etc. (Oohashi and Kobayashi, 2008). The origin of the color contrast could be further explored by analyzing the chemical and mineralogical compounds. ACKNOWLEDGMENTS We thank John Wakabayashi and Yildirim Dilek, who invited us to contribute this chapter. Hidetoshi Hara (Advanced Industrial Science and Technology, Japan) explained the preparation of samples for clay mineral analysis. Manabu Takahashi and
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Daisaku Sato (Advanced Industrial Science and Technology, Japan) are also thanked for their cooperation in the pore-size analysis. Helpful discussions with Ryo Anma and the Press Laboratory members of the University of Tsukuba are also gratefully acknowledged. REFERENCES CITED Bolton, A., and Maltman, A., 1998, Fluid-flow pathways in actively deforming sediments: The role of pore fluid pressures and volume change: Marine and Petroleum Geology, v. 15, p. 281–297, doi:10.1016/S0264 -8172(98)00025-7. Bolton, A., Maltman, A., and Clennell, M., 1998, The importance of overpressure timing and permeability evolution in fine-grained sediments undergoing shear: Journal of Structural Geology, v. 20, p. 1013–1022, doi:10.1016/S0191-8141(98)00030-3. Borradaile, G.J., 1981, Particulate flow of rock and the formation of cleavage: Tectonophysics, v. 72, p. 305–321, doi:10.1016/0040-1951(81)90243-2. Fossen, H., Schultz, R.A., Shipton, Z.K., and Mair, K., 2007, Deformation bands in sandstone: A review: Journal of the Geological Society [London], v. 164, p. 755–769, doi:10.1144/0016-76492006-036. Fyfe, W.S., Price, N.J., and Thompson, A.B., 1978, Fluids in the Earth’s Crust: Their Significance in Metamorphic, Tectonic, and Chemical Transport Processes: Amsterdam, Elsevier Scientific, 383 p. Hanamura, Y., and Ogawa, Y., 1993, Layer-parallel faults, duplexes, imbricate thrust and vein structures of the Miura Group: Key to understanding the Izu fore-arc sediment accretion to Honsyu forearc: Island Arc, v. 2, p. 126–141, doi:10.1111/j.1440-1738.1993.tb00081.x. Hirono, T., 2005, The role of dewatering in the progressive deformation of a sandy accretionary wedge: Constraints from direct imagings of fluid flow and void structure: Tectonophysics, v. 397, p. 261–280, doi:10.1016/j .tecto.2004.12.006. Ishimaru, K., and Miyata, Y., 1991, Morphology of shear planes in soft sediments—Effects of compaction load and sediment composition: Journal of the Geological Society of Japan, v. 97, p. 713–727. Jones, M., 1994, Mechanical principles of sediment deformation, in Maltman, A., ed., The Geological Deformation of Sediments: London, Chapman and Hall, p. 37–71. Kano, K., and Murata, A., eds., 1998, The Structure Geology: Asakura Syobo, p. 25–37, 48–56, 129–144, 235–248. Kawakami, S., and Shishikura, M., 2006, Geology of the Tateyama District: Quadrangle Ser., scale 1:50,000: Geological Survey of Japan, AIST, 82 p. (in Japanese with English abstract, 5 p.). Lundberg, N., and Moore, J.C., 1986, Macroscopic structural features in Deep Sea Drilling Project cores from forearc regions, in Moore, J.C., ed., Structural Fabrics in Deep Sea Drilling Project Cores from Forearcs: Geological Society of America Memoir 166, p. 13–44. Maltman, A.J., Byrne, T., Karig, D., and Lallement, S., 1993, Deformation at the toe of an active accretionary prism: Synopsis of results from ODP Leg 131, Nankai, SW Japan: Journal of Structural Geology, v. 15, p. 949–964, doi:10.1016/0191-8141(93)90169-B. Michiguchi, Y., 2008, Formation process and mechanism of chaotic blocks in the Nishizaki Formation, southern Boso Peninsula: Journal of the Geological Society of Japan, v. 114, p. 461–473. Middleton, V.G., and Wilcock, R.P., eds., 1994, Mechanics in the Earth and Environmental Sciences: Cambridge, UK, Cambridge University Press, 459 p. Mikada, H., Ienaga, M., Goto, T., and Kasaya, T., 2006, Current research status and meaning of fluid pressure monitoring at the Nankai Trough: Journal of Geography, v. 115, p. 367–382. Miyata, Y., Kamai, T., Suzuki, K., and Matsuda, S., 1993, Measurement of pore water pressure and volume change during deformation tests of unlithified sediments: Gekkan-Chikyu: Kaiyo-Shuppan, Tokyo, v. 15, p. 615–620. Ogawa, Y., and Taniguchi, H., 1988, Geology and tectonics of the Miura–Boso Peninsulas and the adjacent area: Modern Geology, v. 12, p. 147–168. Ogawa, Y., Seno, T., Tokuyama, H., Akiyoshi, H., Fujioka, K., and Taniguchi, H., 1989, Structure and development of the Sagami Trough and off-Boso triple junction: Tectonophysics, v. 160, p. 135–150, doi:10.1016/0040 -1951(89)90388-0. Ogawa, Y., Takami, Y., and Takazawa, S., 2007, Oblique subduction in an island arc collision setting: Unique sedimentation, accretion, and deformation
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processes in the Boso TTT-type triple junction area, NW Pacific, in Draut, A.E., Clift, P.D., and Scholl, D.W., eds., Formation and Applications of the Sedimentary Record in Arc Collision Zones: Geological Society of America Special Paper 436, p. 155–170. Ohsumi, T., and Ogawa, Y., 2008, Vein structures, like ripple marks, are formed by short-wavelength shear waves: Journal of Structural Geology, v. 30, p. 719–724, doi:10.1016/j.jsg.2008.02.002. Oohashi, K., and Kobayashi, K., 2008, Fault geometry and paleo-movement of the central part of the Ushikubi fault, northern central Japan: Journal of the Geological Society of Japan, v. 114, p. 16–30. Platt, J.P., 1990, Thrust mechanics in highly overpressured accretionary wedges: Journal of Geophysical Research, v. 95, p. 9025–9034, doi:10.1029/ JB095iB06p09025. Tanner, P.W.G., 1992, Morphology and geometry of duplexes formed during flexural-slip folding: Journal of Structural Geology, v. 14, p. 1173–1192, doi:10.1016/0191-8141(92)90068-8. Tsuji, T., and Miyata, Y., 1987, Fluidization and liquefaction of sand beds— Experimental study and examples from Nichinan Group: Journal of the Geological Society of Japan, v. 11, p. 791–808. Ujiie, K., Maltman, A.J., and Sanchez-Gomez, M., 2004, Origin of deformation bands in argillaceous sediments at the toe of the Nankai accretionary prism, southwest Japan: Journal of Structural Geology, v. 26, p. 221–231, doi:10.1016/j.jsg.2003.06.001.
Vannucchi, P., Maltman, A., Bettelli, G., and Clennell, B., 2003, On the nature of scaly fabric and scaly clay: Journal of Structural Geology, v. 25, p. 673–688. Yamamoto, Y., 2006, Systematic variation of shear-induced physical properties and fabrics in the Miura–Boso accretionary prism: The earliest processes during off-scraping: Earth and Planetary Science Letters, v. 244, p. 270– 284, doi:10.1016/j.epsl.2006.01.049. Yamamoto, Y., and Kawakami, S., 2005, Rapid tectonics of the Late Miocene Boso accretionary prism related to the Izu-Bonin arc collision: Island Arc, v. 14, p. 178–198, doi:10.1111/j.1440-1738.2005.00463.x. Yamamoto, Y., Ohta, Y., and Ogawa, Y., 2000, Implication for the two-stage layer-parallel faults in the context of Izu forearc collision zone—Examples from the Miura accretionary prism, central Japan: Tectonophysics, v. 325, p. 133–144, doi:10.1016/S0040-1951(00)00134-7. Yamamoto, Y., Mukouyoshi, H., and Ogawa, Y., 2005, Structural characteristics of shallowly buried accretionary prism: Rapidly uplifted Neogene accreted sediments on the Miura–Boso Peninsula, central Japan: Tectonics, v. 24, doi:10.1029/2005TC001823. Yeats, R.S., 1986, Faults related to folding with examples from New Zealand: Royal Society of New Zealand Bulletin, v. 24, p. 273–292. MANUSCRIPT ACCEPTED BY THE SOCIETY 21 DECEMBER 2010
Printed in the USA
The Geological Society of America Special Paper 480 2011
Geopractitioner approaches to working with antisocial mélanges Edmund W. Medley* Geological Engineer, Belmont, California, USA Dimitrios Zekkos† Department of Civil and Environmental Engineering, 2358 GG Brown Laboratory, 2350 Hayward Street, University of Michigan, Ann Arbor, Michigan 48109-2125, USA
ABSTRACT Although mélanges are exciting, puzzling, and controversial to geologists, it is geopractitioners and contractors who must work with them to engineer the constructed works of Society. Geopractitioners include geotechnical engineers, geological engineers, engineering geologists, and rock engineers. Mélanges are the most intractable bimrocks (block-in-matrix rocks), complex geological mixtures composed of hard blocks of rocks surrounded by weaker matrix, and are famously exemplified by those within the Franciscan Complex of Northern California. Bimrocks also include olistostromes, weathered rocks, fault rocks, and lahars. The conventional characterization, design, and construction procedures used by geopractitioners for well-behaved stratified rocks and soils are not well suited to mélanges. The considerable engineering and construction difficulties related to mélanges burden Society to the extent that they can be considered “antisocial.” Case histories exemplify a recommended systematic procedure for characterization, design, and construction with mélanges. Geopractitioner approaches to characterizing California’s chaotic Franciscan mélanges are applicable to geologists and geopractitioners working in fault zones, weathered rocks, lahars, and other bimrocks, and suggestions are offered for collaborative research between geologists and geopractitioners.
INTRODUCTION Mélanges have been of exciting and controversial concern to geologists since Greenly (1919) first identified “autoclastic mélange.”1 There are thousands of publications on various aspects of mélanges; countless more contributions have been published on other geologically complex rock mixtures of com-
petent blocks of rock embedded within weaker matrix rocks, such as fault rocks, lahars, tillites, and weathered rocks. Whereas geological complexity may be delightful for geologists, it can result in despair for geopractitioners. Geopractitioners specialize in the engineering of earth materials and include geotechnical engineers, geological engineers, engineering geologists, and rock engineers. Geopractitioners are professionals who apply
*
[email protected]. †
[email protected]. 1 I believe that it is unnecessary and affected to use an acute “é” in mélange. However, I have deferred to the publisher’s preference throughout this chapter, on the condition that we shall share café one day.—EWM Medley, E.W., and Zekkos, D., 2011, Geopractitioner approaches to working with antisocial mélanges, in Wakabayashi, J., and Dilek, Y., eds., Mélanges: Processes of Formation and Societal Significance: Geological Society of America Special Paper 480, p. 261–277, doi:10.1130/2011.2480(13). For permission to copy, contact
[email protected]. © 2011 The Geological Society of America. All rights reserved.
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scientific and mathematical principles to the efficient, economical, and safe design, construction and operation of the civil engineering works essential to Society’s built environment: building foundations, port facilities, tunnels, bridges, buried utilities, dams, etc. Geotechnical engineers are civil engineers specializing in soil mechanics but who may have little knowledge of rock mechanics or geological principles (Turner, 2005; Medley, 2009). Geological engineers are trained as engineers with broad training in geology, groundwater, soil mechanics, and rock mechanics. Engineering geologists are applied geologists with some training in the mechanical behaviors of soils, rock, and water but less familiarity with engineering design principles–practices and construction procedures. Rock engineers are often geological engineers or mining engineers with specialized training in rock mechanics. MÉLANGES AND BIMROCKS Mélange bodies are found in more than 60 countries, generally within current and ancient mountain systems (Medley, 1994a). Mélange bodies are famously abundant in the Franciscan Complex (the Franciscan), which covers about one-third of Northern California. Many geologists have described various aspects of Franciscan mélanges, such as Wakabayashi (2008), Berkeland et al. (1972), Fox (1983), Cloos (1990), Blake and Jones (1974), Raymond (1984), Cowan (1985), and Hsü (1985). Franciscan mélanges provide a superb field laboratory for geologists studying convergent margin tectonics: Useful field guides have been prepared by Wahrhaftig (1984), Blake and Harwood (1989), and Wakabayashi (1999). The mélanges of the Franciscan are similar to mélanges elsewhere in the world in appearance, properties, and the problems they present globally to geopractitioners. There are surprisingly scant treatments written about mélanges from the perspective of geopractitioners, despite the myriad of geological publications. Some early papers described engineering experience with Italian olistostromes and argille scagliose (Associazone Geotechnica Italiana, 1977; Aversa et al., 1993; D’Elia et al., 1986). More recent research has been performed by the senior author, and a few others, including Lindquist (1994a), Lindquist and Goodman, 1994), Goodman and Ahlgren (2000), Riedmüller et al. (2001), Sönmez et al. (2004), and Roadifer et al. (2009). Raymond (1984) classified mélanges, as “block-in-matrix rocks,” a term that joins the several aliases for mélanges such as: friction carpets, wildflysch, broken formation, argille scagliose, olistostromes, mega-breccias, sedimentary chaos, varicolored clays, etc., and is one of many geological terms for comminuted, mixed, fragmented, and chaotic rocks as listed by Laznicka (1988) and the Glossary of Geology (Neuendorf et al., 2005). This ample geological language describing various complex geological mixtures is of little significance to engineers, because their interest is not generally on geologic origin but on the mechanical properties of geological materials and their relationships to engineering design and construction. To focus engineers’ attention on the fundamental engineering properties of complex geologi-
cal mixtures for the purposes of design and construction, Medley (1994a) coined the neutral word bimrock from Raymond’s (1984) term block-in-matrix rock. Formally, a bimrock is “a mixture of rocks, composed of geotechnically significant blocks within a bonded matrix of finer texture.” The expression “geotechnically significant blocks” means that there is mechanical contrast between blocks and matrix, and the geometry and proportion of the blocks influence the rock mass properties at the scales of engineering interest, which range between centimeters for laboratory test specimens through tens of kilometers for tunnel lengths. Bimrocks include complex geological mixtures such as olistostromes, weathered rocks, fault rocks, and mélanges. Weathered rocks can include mixtures of decomposed soil surrounding fresher core-stones (Fig. 1). Fault rocks, composed of blocks within gouge and sheared rock (Fig. 2), exist at many scales, with blocks ranging between several tens to hundreds of meters in size to millimeter-sized fragments within gouge (Riedmüller et al., 2001, 2004). Mélanges, the most intractable of bimrocks, contain competent blocks of varied lithologies, commonly embedded in sheared matrices of weaker rock (Fig. 3). Regardless of geological nomenclature and formative processes, simplistically, bimrocks have a similar fabric of relatively hard blocks of rock surrounded by weaker matrix rocks. Yet, despite simplifications, characterization, design, and construction within bimrocks is still challenging because of the considerable spatial, lithological, and mechanical variability, and because geopractitioners often mischaracterize them. The conventional and simplistic expectation in geopractice is that soil and rock masses are orderly and stratified to the extent that the words layers, strata, and formations are frequently used in geopractice. Stratified conditions permit geological, spatial, and mechanical interpolations between relatively scant borings, outcrop observations, and test results. Economic and efficient design and construction can usually follow, to the benefit of designers, contractors, owners, and Society at large. Geopractitioners often
Figure 1. A block-in-matrix rock (bimrock) in the Sierra Nevada of California. Decomposed granite contains hard blocks (core-stones) surrounded by a matrix of dense sandy soil (grus). Photo: E. Medley.
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In general, geopractitioners work with mélanges, rather than try to understand how they came to be. “Work” is used to describe characterization, design, and construction, although in this chapter we mostly address characterization. Geopractitioners strive to solve engineering problems economically, efficiently, and safely, so the geopractice approaches we describe may seem coarse to geologists interested primarily in understanding the scientific details of mélanges and their formation. Some geopractitioners are not aware of the existence of complex geological mixtures of weak matrix and strong blocks. Others regard mélanges as soil-rock mixtures. Geotechnical engineers, with their soils bias, often analyze bimrocks as soil, or assume that the geomechanical properties of the weakest materials in mélanges adequately represent the mixed rock mass, and design on that basis. Some geotechnical engineers and rock engineers may focus on the rock content, and consider the mixture to be stable material. Hence, the words of Prof. Harry Bolton Seed, a principal authority for geopractitioners, are thus appropriate: “The general thrust behind engineering problem solving
is to simplify the problem enough to make it solvable. However, we must check to see if we have oversimplified the problem so much that we get other problems instead of the solution we desire” (Rogers, 2008). So, in making the assumptions described above, geopractitioners invite future problems during construction, because extremely weak shear zones and complex stress conditions may exist that are hidden until the ground is exposed during construction. Civil engineering works require characterization, design, and construction. Geological and geotechnical characterization involves exploration of the ground surface and subsurface by direct and indirect means, sampling and testing of its constituent parts, and assigning to the ground mass mechanical properties that are representative of the material for the purposes of the design of a particular project. The chaotic nature of mélanges has more than usual constraining impact on the design and construction of tunnels, dams, excavations and other civil engineering works. Discussion of the general procedures for the characterization of subsurface conditions is beyond the scope of this paper, but many resources are available (e.g., Hunt, 2005). Specific guidance to geopractice characterization of mélanges is provided in the following section as well as by Medley (2001) and Wakabayashi and Medley (2004). Successful characterizations of mélanges depend on geopractitioner skills at interpreting scant ground clues and anticipating the complexity of the rock mass. Depending on the scope of the work, different aspects of the behavior of the rock mass conditions will be critical, and the characterization may need to be focused on those aspects more than on others. Hence a geopractitioner’s exploration of mélanges is often an economic compromise. Too few borings or scant mapping results in a cheap but inadequate characterization. On the other hand, a comprehensive and expensive program may generate ample observations but still may not provide enough data to accurately model the mélange.
Figure 2. Wall of a quarry located within a major fault zone, California. Sheared rock surrounds hard blocks of relatively intact rock. Blocks range between centimeters and tens of meters in size. Photo: E. Medley–Geosyntec Consultants.
Figure 3. Franciscan Complex mélange, Northern California. Blocks buttress base of slope between landslides in sheared shale matrix. Photo: E. Medley–Exponent, Inc.
also use the expressions “well-behaved” and “nicely behaved” for spatially and mechanically uniform stratified conditions, although even “well-behaved” rock masses have discontinuities (joints, fractures, shears, etc.) that complicate simple characterizations. Accordingly, complex geological mixtures of weak matrix and strong blocks can be considered “badly behaved,” and the worst offenders, mélanges, are so much trouble to geopractitioners and Society that it is irresistible to suggest that “badly behaved” mélanges are “antisocial,” particularly since “deranged” mélanges cost Society considerably more effort, money, and time than well-behaved stratified rocks. GEOPRACTICE WITH MÉLANGES AND SIMILAR BIMROCKS
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Characterizations are also based on laboratory tests of specimens that are inevitably disturbed to some extent by drilling, transportation, and specimen preparation. If characterization is performed correctly, potential construction issues will be recognized early in the project. But the geological-geomechanical model is often only a sketch of the actual geological conditions, and geopractitioners must accommodate uncertainty about the likely behavior of ground masses by qualifying their characterizations and adding generous margins of safety to their estimates of the mechanical properties of the soil or rock masses. The geopractitioner also has a duty to warn the client and the contractor about the uncertainties involved in the characterization and design and to recommend methods to deal with problems anticipated during construction. Design and construction follow characterization, although they proceed together with characterization on large design-build projects. Geopractice design is the commissioned (i.e., compensated) systematic conception, invention, and specification of the appropriate means (geometry, loads, etc.) required for acceptable performance of civil infrastructure as it relates to the geologic environment. In practice, designs are subject to the constraints of ground conditions, costs, time, and labor-equipment resources. Construction is performed by contractors using the plans and specifications prepared by the design engineer. Geopractitioners may work with, and for, contractors, although most often they work for owners, who pay for the work. Flaws in the characterization and design will become evident during construction and may cause ground failures, schedule delays, and contractor claims for additional compensation. Table 1 summarizes some of the construction issues, the design and analyses considerations, and the characterization needs for excavations, slopes, foundations, dams, and underground facilities, some of which are addressed in more detail below. GEOPRACTICE METHODS FOR CHARACTERIZATION OF MÉLANGES A major geopractitioner concern is the evaluation of the strength and deformation properties of soil and rock masses. For mélanges, conventional engineering approaches using simple stratigraphic-type characterizations are of little value. Instead, there should be a focus on volumetric block proportions; block and matrix lithologies; strength and discontinuity fabrics; block sizes, shapes, and orientations; and groundwater regimes (Wakabayashi and Medley, 2004). Characterization typically includes review of aerial photographs and available geological information, geological mapping, drilling and coring, and laboratory testing. Ideally, aerial photos can be used with image processing techniques to identify blocks within mélanges, as tonal contrasts may discriminate large blocks from matrix (Medley, 1994a). More typically, information about the blocks is collected by mapping and drilling. The relative strengths of blocks and matrix of blocks can be evaluated using a geologic hammer and reference to engineering geology
protocols such as that of the Geological Society Engineering Geology Working Party (1995), a Schmidt hammer, or a Point Load Test apparatus. The geometry of blocks and fabric of matrix around the blocks must be evaluated because of possible shears and ellipsoidal or prolate-shaped blocks adversely oriented into proposed excavation slopes or tunnel walls. Observations about large blocks buttressing slopes (such as is evident in Fig. 3), and existing failed rock masses can be used to back-calculate the overall strength of mélange (Kim et al., 2004). It can be difficult to recover good quality core in bimrocks during drilling exploration because of the abrupt variations between blocks and matrix, varying block lithologies, extensive shearing, and the highly fractured nature of small blocks (Fig. 4). Recovery of drilled core tends to be much better in blocks than in matrix because the blocks survive the drilling process, whereas the weak matrix generally does not (Riedmüller et al., 2001). Good core recovery has been reported by using triple-tube core barrels (Roadifer et al., 2009), and the Integral Method (Rocha, 1971). Downhole optical and geophysical techniques (such as borehole televiewers) provide rock mass information of the walls of the drilled holes, and are particularly useful when core recovery is poor. The lengths of blocks in cores, core photographs, or televiewer records should be measured to estimate the linear proportion of blocks versus matrix, and to approximately indicate the range of possible block sizes. Because the matrix of mélanges is prone to slake quickly, it is prudent to promptly protect selected specimens in wax or else protect them within shrink-wrapped plastic film (Waeber, 2008). It would seem that geophysical methods could be useful in estimating the presence and geometry of blocks in bimrocks. However, although successful in discriminating density contrasts, geophysical methods in mélanges are hampered by the myriad of smaller blocks that mask the geophysical response of the larger target blocks.
Figure 4. Typical Franciscan mélange recovered in drilled core. Blocks (light gray) interspersed by weak, sheared matrix. Core from exploration boring BH 103, for Richmond Transport Tunnel, San Francisco, California. Photo: E. Medley.
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TABLE 1. GEOPRACTICE CONSIDERATIONS WHEN WORKING WITH MÉLANGES AND OTHER BIMROCKS Characterization Common to all civil engineering works Identify unusual lithologies Drill rock core using triple tube samplers; shrink wrap core Measure block linear proportions from core and mapping scan lines Beware of optimistic cross sections Perform triaxial strength testing at differing block proportions
Slopes, excavations, and dams Perform direct shear tests on discontinuities and/or matrix Map critical failure surfaces
Tunnels, pipelines, and underground spaces Anticipate that observed ground conditions may provide little basis for accurate subsurface profiles and geological characterization along tunnel alignments
Design/analysis
Construction
Decide soil engineering, rock engineering, or bimrock approaches on basis of volumetric block proportions Evaluate potential destabilizing effects of large blocks, block and matrix discontinuity fabrics, shears, block orientations, and shapes Evaluate variable groundwater conditions in blocks and matrix Predict long-term performance of rock mass Evaluate how block proportions, size distributions, lithology, and individual block and matrix strengths affect construction method and performance
Anticipate possible considerable differences between site conditions as characterized and conditions encountered in construction Review equipment limitations and handling of oversize blocks, fragmentation, rippability, and excavatability of very hard blocks Plan for possible unexpected large pressurized groundwater flows when blocks are penetrated
Assess destabilizing effect of stress and groundwater concentrations Analyze stabilizing/destabilizing effects of block orientations Anticipate through-going shears in matrix and/or blocks
Monitor short-term stability of excavated rock with abruptly variable shear fabric Consider leaving large blocks to protrude at grade or incorporate block stick-ups into the work Monitor potential for block fallouts
Analyze stress and groundwater effects of encountering unexpected large blocks Analyze effect of short stand-up times, squeezing ground in weak matrix Develop detailed geotechnical observation method protocol
Select equipment and construction procedures flexible enough to accommodate possible extreme mixed-face conditions Develop procedures for penetrating large, extremely hard blocks, and handling many intact, small blocks Perform detailed face mapping and realtime analysis of ground deformation data
A common error made by geopractitioners is to map outcrops and draw cross sections (Fig. 5) of mélanges as continuous strata because of the interlayered appearance of blocks and matrix in drill core (Fig. 4). Boring logs in mélanges may include misleading expressions such as “interbedded shales and sandstones,” which implies stratal continuity that does not exist. But blocks may include coherent turbidite or sandstone-shale sequences; for such cases, block boundaries can be discriminated because the shale layers in intact sandstone-siltstone-shales will not contain tiny blocks, whereas sheared matrix will. The upper and lower limits of fragment-free sections of shale-siltstone are the boundaries of intact blocks of actual interlayered rocks. Relatively undisturbed samples can be tested in a geotechnical laboratory to evaluate the density and moisture content of the rock, and depending on the scale of interest, multistage triaxial tests (Bro, 1996, 1997). Ideally, consolidated undrained tests with pore pressure measurements will measure the strength of specimens with various volumetric block proportions (Lindquist, 1994a) as a necessary step in estimating the overall strength of the rock mass as explained further below. Unconfined compression tests are not recommended for mélanges where significant shears exist prior to testing, as these may prejudice the test results. Direct shear testing may be performed to measure the strength of selected discontinuities in blocks or, if the specimen is robust, in matrix.
SIGNIFICANCE OF BLOCK SIZES IN MÉLANGES Geological terms for rock fragments that suggest size, such as boulders, are generally taken to be larger than about 0.3 m in size, with no apparent upper size limit. Such terms should not be used to describe the rock inclusions in mélanges: The neutral word block is preferred. However, it is misleading to refer to block “sizes” or block “diameters” on the basis of drill-core- or outcrop-observed dimensions, because, as a matter of geometrical probability (Fig. 6), block measurements from field mapping or drilling are chords that are almost always shorter than the true “diameter” of a block. (A chord is a line with end points on a curve.) In one dimension, dmod (the maximum observed dimension—Fig. 6) is the chord length formed by the intersection of blocks with sampling lines traversing outcrops (scan lines of Priest, 1993), or linear drill core (Figs. 6 and 7). Franciscan mélanges are self-similar, or scale-independent, having the same general appearances regardless of scale, with a few large blocks and an increasing number of smaller blocks (Cowan, 1985; Medley, 1994a; Medley and Lindquist, 1995). Scale independence has also been observed at a large exposure of an Italian olistostrome (Coli et al., 2008). Figure 8 and its inset are photographs taken at different scales of the same outcrop of Franciscan mélange. Small blocks at one scale of interest (detailed photo in Fig. 8) are part of the matrix at the smaller
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Medley and Zekkos ogy, well graded) or fractal (i.e., conforming to negative power laws), with a few large blocks and increasing numbers of smaller blocks, as shown in Figure 9. Log-histograms, as developed by Bagnold and Barndorff-Nielsen (1980), are useful in presenting the measurements because the histogram size classes embrace several orders of magnitude. In the Franciscan mélanges measured, the range in block sizes exceeds seven orders of magnitude between sand (millimeters) and mountains (tens of kilometers), as illustrated in Figure 10. Despite the considerable difference in scales, the mélanges measured from outcrops and maps shown in Figure 10 indicate individual scale-independent block-size distributions with similar appearances. Consequently, blocks will always be found in Franciscan mélanges regardless of the scale of observation. Figure 5. Bimrocks generally cannot be accurately depicted on cross sections. Borehole contacts should be shown with question marks and not connected by lines interpreted to be stratal boundaries, here drawn between borings. After Wakabayashi and Medley (2004).
scale of the main photo in Figure 8. Many other geological mixtures and comminuted rocks, such as fault gouges (Sammis and Biegel, 1989) and fractured rock masses (Nagahama, 1993), also exhibit self-similar block size distributions. Using geological maps and photographs of outcrop exposures of Franciscan mélanges at many scales, Medley (1994a) measured almost 2000 dmod lengths. The resulting block size distributions were poorly sorted (in soil mechanics terminol-
Figure 6. In two dimensions, such as at an outcrop, a block has an apparent block size of dmod, the maximum observed dimension. In one dimension, such as in a drilled boring, the block size is apparently indicated by the chord, the length of interception between the boring and a block. However, only rarely will a dmod or a chord be equivalent to the actual “diameter” or maximum dimension of a block. After Medley (2002).
DISCRIMINATING BLOCKS FROM MATRIX A sufficient volumetric proportion of blocks in mélanges provides mechanical advantages to a mélange. Since there is a very large range in block sizes in mélanges, the geotechnically significant blocks must be discriminated from matrix. Because of scale independence, any reasonable dimension can be used to scale a mélange rock mass for the problem at hand to bound the range of significant block sizes. Medley (1994a) called such a descriptive length the characteristic engineering dimension (ced), Lc (the ced of Medley, 1994a), or simply, the characteristic dimension. A characteristic dimension is analogous to showing a measuring tape, coin, hand, or spouse in a photograph, without which reference the observer cannot appreciate the scale of the image. The characteristic dimension changes as scales of interest change. Lc may variously be (1) an indicator of the scale of a site, such as √A, where A is the area of the site; (2) the size of the largest mapped or estimated largest block (dmax) at the site; (3) the thickness of a failure zone beneath a landslide; (4) a tunnel
Figure 7. Block-core intersections (chords) rarely indicate true block sizes, or shapes, orientations, and distributions of blocks. Note that shears in the matrix negotiate tortuously around blocks. Sandstone-shale sequence in core is not “interbedded shale and sandstone.” Improbable juxtaposition of rocks (e.g., greenstone and shale) strongly suggest mélange. After Medley (2002).
Geopractitioner approaches to working with antisocial mélanges
Figure 8. Franciscan Complex mélange, Northern California. Note shearing in shale matrix adjacent to the large headland block, in which smaller blocks are oriented subparallel to shearing. Block sizes range between tens of meters and meters. Detail showing matrix in the circled area also has block-in-matrix fabric at the scale of the 3.1-m-long bar. Photo: E. Medley.
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mélange, blocks <0.05Lc in size (Fig. 10) constitute greater than 95% of the total number of blocks but contribute <1% to the total volume of mélange and have a negligible effect on the mechanical behavior of the mélange (Medley, 1994a). For these reasons, the threshold size between blocks and matrix at any scale is taken to be 0.05Lc. In practice, the most conservative block-matrix threshold that can be justified should be selected and used to measure the blocks between these limits. Blocks smaller than 0.05 Lc at one scale (e.g., 1:1000) must be assigned to matrix at a smaller scale (say, 1:10,000), although they may still be of substantial size. If the scale of interest becomes larger, blocks previously demoted into matrix are then promoted to geotechnically significant blocks and must be included in the characterization. Also, large blocks at one scale of interest (e.g., 1:10,000) are not geotechnically significant at a larger scale (e.g., 1:1000) because they are then massive, unmixed blocky rock masses that can be characterized according to conventional rock engineering practice. MECHANICAL PROPERTIES OF MÉLANGES
diameter; (5) a foundation width or length; or (6) the dimension of a laboratory specimen. Scale independence also means that a laboratory scale specimen of mélange is more closely a “model” of the parent rock mass at site scale than is typically encountered in conventional rock engineering. This concept was used by Lindquist (1994) in his pioneering investigation into the strength and deformation properties of mélanges by fabricated physical model mélanges of different volumetric proportions and testing them under multistage triaxial loading to failure. Medley (1994a) and Medley and Lindquist (1995) showed that the largest geotechnically significant block (dmax) within any given volume of Franciscan mélange is ~0.75Lc (Fig. 10). Blocks greater than 0.75Lc result in such a diminished proportion of matrix in a local volume of rock mass that the volume can be considered to be massive, unmixed rock composed mostly of the block. Furthermore, for any given volume of Franciscan
The matrix of Franciscan mélanges is most commonly composed of sheared shale, argillite, siltstone, serpentinite, or sandstone, commonly weakened to the consistency of soil. Intact portions of these lithologies may be competent enough to be blocks. Very weak matrix and the most intense shearing within mélanges may be found in block-poor zones adjacent to blocks exceeding tens of meters in largest dimensions. Indeed, Savina (1982) measured up to 800 shears per meter in a Franciscan mélange. Because of highly sheared matrix, landslides are common in block-poor Franciscan mélanges (Medley and Sanz, 2004), although large blocks appear to add buttress support (Fig. 3). The weakest elements in mélanges are commonly the contacts between blocks and matrix (Fig. 11). In some bimrocks, such as welded tuffs, the contacts may be strong (Sönmez et al., 2004). In outcrops and core, contacts may be marked by a lustrous surface on the blocks and a wafer of sheared material that weathers to a slick film of clay. Matrix shears generally pass around blocks via the block-matrix contacts (Figs. 7 and 11) with
Figure 9. Left plot is a log-histogram of dmods measured from the outcrop of Franciscan mélange shown in Figure 8. The right plot is a log histogram of dmods of blocks measured from a map of Franciscan Complex in Marin County, California (Ellen and Wentworth, 1995). The slopes of the log-log plots are the fractal dimension D. The shapes of the two plots are similar despite the extreme difference in scales. After Medley (1994a).
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the most intense shearing commonly present adjacent to the largest blocks. Blocks within the shears are commonly entrained within and oriented subparallel to shears (Fig. 8). Because shears have a tortuous path through the rock mass, the overall orientation of entrained blocks can also abruptly change from place to place within the rock mass. Relatively modest block-matrix mechanical contrast is necessary for a block-in-matrix rock mass to be considered a bimrock. For physical models of mélange, Lindquist (1994) selected a criterion based on the ratio between block stiffness (E, Young’s modulus) and matrix stiffness to generate triaxially induced shears along block-matrix contacts: (Eblock/Ematrix)~2.0. Sufficient contrast between block and matrix to deflect failure surfaces can be evaluated by a friction angle ratio suggested by Medley (1994a) based on work by Lindquist (1994a), Lindquist and Goodman (1994), and Volpe et al. (1991): (tanφweakest block)/(tanφmatrix)>1.5–2.0.
Figure 10. Normalized block-size distribution curves for 1928 blocks measured from outcrops and geological maps of several Franciscan mélanges ranging over seven orders of magnitude in scale. The sizes of blocks are indicated by dmod, the maximum observed dimension of the blocks in the outcrops and maps. The block sizes are divided by the square root of the area (√A) containing the measured blocks to yield dimensionless dmod/√A. The dimensionless relative frequency for each plot is the number of blocks in any size class divided by the total number of blocks measured for the various data sources. The clustered normalized plots are similar in shape, indicating that the block size distributions are scale independent. The plots peak at ~0.05dmod/√A, which is the block-matrix threshold size at any scale. Blocks smaller than 0.05dmod/√A are assigned to the matrix; they tend to be too small to measure and are undercounted. The largest indicated block size is approximately equivalent to √A (at dmod/√A = 1), but 99% of the blocks are smaller than ~0.75√A, which is defined as the maximum block size (or dmax) at any scale. After Medley (1994).
However, rock strength is a function of both friction angle and cohesion, and the block-matrix contrast will also be influenced by both strength components. Accordingly, as recommended by Prof. Harun Sönmez (2009, personal commun.) the unconfined compressive strengths (UCS) of blocks and matrix could be used to define a strength-based block-matrix contrast criterion: assuming that reliable UCS tests can be performed on weak matrix, UCSblock/UCSmatrix >1.5. For strength ratios or stiffness ratios less than the lower bounds described above, there will be an increased tendency for shears and failure surfaces to pass through blocks rather than around them. The overall strength of a mélange is directly related to the volumetric proportions of the blocks and is largely independent of the strength of the blocks (Lindquist, 1994a). The dependence on volumetric proportion to overall strength is an apparent fundamental property of bimrocks, as indicated by Savely (1990) for the boulder-rich Gila Conglomerate in Arizona; and by Irfan and Tang (1993) for boulder-rich colluvium in Hong Kong. A sufficient block volumetric proportion and range of block sizes can contribute to strength, because failure surfaces are forced tortuously to negotiate around blocks, thus increasing passive resistance to shear failures (Lindquist, 1994a; Medley, 2004; Sönmez, et al., 2006a, 2006b). Where blocks are uniformly sized, failure surfaces tend to have smoother, undulating profiles (figure 4.4 of Medley, 1994a; Medley, 2004), and hence the mixed rock mass has less shear resistance, and a lower overall strength. But there is a dependence on the normal stresses to which the mélange rock mass is subjected: At sufficiently high normal stresses, failure surfaces will penetrate blocks regardless of the mechanical contrast between matrix and blocks. Inherent defects in the blocks will aggravate the effect.
Figure 11. The weakest element in a bimrock is commonly the block-matrix contact. Gwna mélange, Anglesey, Wales. Photo: E. Medley.
Geopractitioner approaches to working with antisocial mélanges
Figure 12. This plot shows that strength of bimrocks increases with volumetric block proportion. The increase is added to the strength of the matrix. After Medley (2001), from data of Lindquist (1994a) and Irfan and Tang (1993).
As shown in Figure 12, Lindquist (1994a) conservatively established that <~25% volumetric block proportion the frictional strength and deformation properties of a physical model mélange was that of the matrix, a conclusion supported by the research of Irfan and Tang (1993). However, Goodman and Ahlgren (2000) and Roadifer et al. (2009) identified increases to overall mélange strength at volumetric block proportions <~25%. Lindquist (1994a) also showed that between ~25% and 75%, the friction angle and modulus of deformation of a physical model mélange proportionally increased, depending on the relative orientation of blocks to applied stresses; and cohesion generally decreased because of increasing weak block-matrix contacts. Cohesion has also been noted to increase with increasing volumetric block proportions in mélanges (Goodman and Ahlgren, 2000; Roadifer et al., 2009) and volcanic agglomer-
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ates (Sönmez et al., 2006b), possibly owing to block-matrix contact strengths being greater than matrix strengths for the materials evaluated. Beyond ~75% block proportion, Lindquist (1994a) showed there was no further increase in bimrock strength, as blocks tended to be in contact, and the rock mass approached the condition of being a jointed rock mass with infilled joints. Franciscan complex mélanges have been assigned strength parameters varying between a block-poor mélange in a landslide: c′ (effective cohesion) of 9.6 kPa to 14.8 kPa (100– 200 pounds per square foot [psf]), and φ′ (effective friction angle) of 25°–35° (Kim et al., 2004); for a block-rich mélange below Scott Dam: c′ of 76 kPa (1590 psf), and φ′ of 39° (Goodman and Ahlgren, 2000); and for mélange below Calaveras Dam: c′ of 6.9 kPa to 39.9 kPa (1000–5740 psf), and φ′ of 25°–45° (Roadifer et al., 2009). Sheared matrix in Franciscan mélange can be extremely weak, with low to zero cohesion and effective friction angles of <10°. Based on the findings summarized above, it is apparent that if geopractitioners follow soil mechanics convention and assume that the mechanical behavior is adequately represented by solely the properties of the weak matrix materials, any mechanical advantage afforded by blocks will be discounted, hence penalizing the rock mass strength. Furthermore, ignoring blocks during characterization increases the risk that they will be neglected during design, only to be rediscovered later by the earthwork or tunneling contractor, with expensive consequences. Because evaluation of the strength of mélanges is not straightforward, the procedural steps of Table 2 may be helpful. ESTIMATING VOLUMETRIC BLOCK PROPORTIONS The direct and relatively simple relationship between strength and volumetric block proportions for mélanges and other bimrocks is appealing to geopractitioners. Nevertheless, to
TABLE 2. FLOWCHART OF STEPS TO CHARACTERIZE MÉLANGE STRENGTH Step 1: Check block-matrix strength contrast using estimates of weakest block and matrix mechanical parameters and following criteria: tan φblock/tanφmatrix ≥1.5–2.0 Eblock/Ematrix ≥2.0 UCSblock/UCSmatrix ≥1.5 Step 2: Select characteristic dimension Lc at scale of engineering interest; e.g.: height of landslide, width of tunnel face; length of foundation, diameter of laboratory specimen, etc. Step 3: Calculate block size thresholds: dmin=0.05*Lc (block-matrix threshold) dmax=0.75*Lc (block–blocky rock mass threshold) Step 4: Drill and core exploration length of (preferably) ≥10*dmax Step 5: Measure total linear block proportion LL Step 6: Assume LL is equivalent to volumetric block proportion Vv and estimate uncertainty range in estimate of Vv (use Fig. 14) Step 7: Establish conservative design values of volumetric block proportion, using lower bound for strength and upper bound for earthwork construction Step 8: Measure lab strengths of specimens with different Vv proportions and construct a plot of specimen strength vs. Vv (such as Fig. 12 or Fig. 18) Step 9: On plot of specimen strength vs. Vv identify critical volumetric block proportions and following three regions: Lower bound strength controls if Vv <15%–25%, use soils or rock engineering analyses Variable strength controls if 15%–25%
65%–75%, use rock engineering analyses
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evaluate the mechanical properties of a bimrock the volumetric block proportion must be estimated, which may not be a simple matter. The volume of blocks in laboratory specimens can be measured by disaggregation, or estimated from specimen densities, if the individual densities of block and matrix materials are known (Lindquist, 1994a). Although the laboratory method does not work at site scales, stereological techniques do. Stereology relates point, linear, and planar observations to volumetric properties and is based on the original work of the mineralogist Rosiwal (1898), who evaluated the mineral content of rock specimens using microscope traverses of thin section slides. A fundamental principle of stereology is that, given enough sampling, linear block proportions and areal block proportions are equivalent to volumetric block proportions (Ross, 1986; Underwood, 1970; Weibel, 1980; Medley, 1994a). Linear block proportions are the ratios of the total lengths of blocks intersected by sample lines to the total length of the sample lines, and depending on the scale of interest, these can be estimated from scan-line measurements at outcrops, measured from drilled cores (Medley, 1994a, 1994b), or from laboratory specimens (Medley, 1994a, 1994b). The areal block proportions can also be measured from outcrops, geological maps, or air photos, using image analysis (Medley, 1994a, 1994b; Coli et al., 2008). As indicated in Table 2, a characteristic dimension (Lc) is selected that describes the scale of the problem at hand (Medley, 1994a, 2004), and the block-matrix threshold selected as 0.05Lc. During mapping, the dmod of the blocks are measured. When logging core, all block-core intersections (chords) greater than 0.05Lc are measured. But in practice, block lengths greater than ~5 cm long could be measured, even if the block-matrix threshold is larger, as this information will be useful for work performed at laboratory scale, where 0.05Lc is 5% of typical 100-mm-diameter core specimens. Blocks in mélanges are not uniformly sized or distributed, so the volumetric block proportion cannot be accurately determined from a few borings. Also the assumption that measured areal or linear block proportions are equivalent to the required volumetric block proportions is only valid if the sampling size is large. However, if the total length of drilling or other liner sampling of at least 10dmax is performed, the linear block proportion reasonably approaches the volumetric block proportion (Medley, 1994a, 1997). But in geopractice exploration, the desirable optimum total length of exploration drilling is frequently not performed because of the expense of drilling, which in mélanges can be higher owing to the troublesome effort required to secure good quality cores from mechanically variable rock. Medley (1997) investigated the potential error in estimates of volumetric block proportion based on the assumption that they are the same as the measured linear block proportions. Four physical models of mélange were fabricated with known block size distributions and volumetric block proportions. The models were explored with hundreds of model boreholes. The extreme variability in linear proportions, even for adjacent bor-
ings, is indicated by Figure 13. Based on thousands of randomized realizations of the model boring data, Medley (1997, 2002) concluded that potentially serious errors result if volumetric block proportions, total block volumes, and block size distributions are assumed, without qualification, to be equivalent to the linear block proportions measured from few borings or outcrops. Medley’s (1997) experiments showed that measured linear block proportions have to be adjusted by an uncertainty factor (actually, the coefficient of variation) to yield an appropriate estimate of the volumetric block proportion (Fig. 14). Uncertainty depends on both the total length of the linear measurements, such as from drilled core, and the linear block proportion itself. In Figure 14 the total length of drilled core is shown as Ndmax, or multiples of the size of the largest expected block, dmax, which can be estimated from field exposures or else assumed to be 0.75Lc at site scale. The volumetric block proportion is within a range between the adjusted lower and upper volumetric block proportions. It is prudent and conservative to apply the uncertainty adjustment downward for the calculated estimates of volumetric block proportions for the purpose of assigning strength parameters for a bimrock. On the other hand, for the purpose of assessments of excavation rippability and construction preparedness in tunneling or earthwork construction, it is prudent and conservative to adjust upward the calculated estimates of volumetric block proportions.
Figure 13. Plan view of an array of 100 linear block proportions ranging between 0% and 61%, measured for a physical model bimrock with an actual volumetric block proportion of 32%. This data representation clearly shows the spatial variability, the extremes being indicated by the circled values. Coincidentally, the average of the extreme values of liner block proportions (30.5%) closely matches the actual volumetric block proportion of 32%. However, other random pairings of data will generally not yield average linear proportions that match the actual volumetric block proportion. Likewise, in geopractice explorations of mélange, it is unlikely that two borings will yield average linear block proportions that match the actual volumetric block proportion of an explored rock mass. After Medley (1997).
Geopractitioner approaches to working with antisocial mélanges ESTIMATING BLOCK SIZE DISTRIBUTIONS Given the economic ramifications for construction planning, it is of great value to geopractitioners to be able to reasonably estimate the volumes and block size distributions of bimrocks to be excavated or tunneled (Attewell, 1997). Hard blocks in mélanges that are larger than ~0.6 cm in diameter are too large to pass through the apertures of earthmoving scrapers, which are like the gaps in cheese parers and potato peelers. Blocks >~2 m cannot be easily removed by conventional bulldozers. Because blasting is often not permitted during earthwork in urban areas, excavation of large blocks must be performed by fragmenting them with hydraulic rock hammers, resulting in noise, dust, and expensive delays to earthwork grading schedules, although nonexplosive chemical demolition agents are available. As the primary means of obtaining subsurface data is by drilling, it is logical to attempt to estimate block size distributions from chord length distributions. The degree to which 1-D (one dimension) chord length distributions match actual 3-D block size distributions is dependent on the orientation of blocks relative to the boring directions, volumetric block proportion, and total length of drilling. Because observed chord lengths are almost invariably smaller than the actual block diameters (Fig. 15) the frequency of larger block sizes tends to be underesti-
Figure 14. This plot provides the uncertainty in assuming that measured linear block proportions (13% to 55%) are assumed to represent volumetric block proportion. Uncertainty (coefficient of variation) in estimates of volumetric block proportions are a function of the total length of linear measurement, expressed as multiples (N) of the length of the largest block (dmax). The dashed line shows an example for Scott Dam, California, where the total length of drilling performed was equivalent to ~5 times the size of the largest block; i.e., Ndmax was ~5. Based on the drilling, the linear block proportion was ~40%. When assumed to represent the actual volumetric block proportion, there was an uncertainty of 0.2. Hence the adjusted volumetric block proportion was 40% – (0.2*40%), or 32%. After Medley (1997).
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mated, and the frequency of smaller sizes overestimated. Indeed, larger blocks are mischaracterized as smaller blocks to the degree that small block sizes are indicated that may not even be part of the actual 3-D block size distribution (Fig. 15). Estimates of 3-D block size distributions from measurements of 2-D outcrop areas also yield widely erroneous estimates (Haneberg, 2004). It is thus unlikely that drilling and coring into a mélange rock mass will yield actual 3-D block size distributions, so estimating 3-D block size distributions should be attempted with great caution, or better yet, not at all. The practical consequence of underestimating block sizes from exploration drilling is that unpleasant and costly surprises may occur during excavation and tunneling of bimrocks. CONSTRUCTION IN MÉLANGES The size and volumetric proportion of hard blocks, and their individual lithology, strength, and discontinuity fabrics, are critical considerations in the selection of the appropriate method of excavation and the equipment required to safely perform the excavation of weaker matrix material and stronger, commonly very large blocks. Large blocks (greater than tens of meters in size) indicate the possibility of generally weaker sheared matrix adjacent to the block and potentially dangerous ground conditions. Stress distributions in mélanges influence the construction behavior and long-term performance of dam and structure foundations, slopes (Medley and Sanz, 2004), and underground excavations (Button et al., 2003; Moritz et al., 2004; Riedmüller and Schubert, 2002). Stress states depend on the in situ stress and lithologies; size distributions, orientations, and shapes of blocks;
Figure 15. Comparison of true 3-D block size distribution (3D BSD) used in the fabrication of each of four physical model bimrocks. The 1-D chord length distributions (CLDs) were generated from measuring all chord lengths in 100 model borings generated for each model. Despite the 400 borings, the 1-D distributions do not mimic the actual 3-D distribution. Instead, small blocks are indicated that were not incorporated into the actual models, and the proportions of larger block sizes are underestimated. After Medley (2002).
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orientations of matrix shears and groundwater; and the geometry and type of construction. High stresses may be transferred into the sheared matrix surrounding unexpected large blocks as tunnels or excavations under construction approach the blocks. Such rearranged ground stresses, in combination with high in situ tectonic stresses, and/or from topographic effects, may yield problematic to disastrous “squeezing ground” conditions during tunnel excavation (Hoek, 2001; Button et al., 2003). A site evaluation must consider the general hydrogeologic setting and the potential for variable water flow or artesian pressures during and after construction, as groundwater has an important impact on the stress conditions as well as the mechanical properties. Haneberg (1995) demonstrated that in heterogeneous colluvium, variations in the permeability of matrix and blocks perturb fluid flow and influence local factors of safety in mélanges and fault rocks. Unexpected inflows of groundwater may also discharge from large fractured blocks, or stress imbalances may arise from pore water pressure variations around large blocks. In mélanges, the permeabilities may also differ by orders of magnitude over short distances, because intact blocks, fractured blocks, intact matrix, and sheared matrix may have widely varied individual permeabilities. The impact of abrupt changes in permeability, as well as the potential for preferential flow, needs to be considered in the design and may impact construction. The installation of piezometers prior to construction will help evaluate the local hydrogeologic regime. However, the common construction practice of using drilled sub-horizontal drains to lower water levels in mélange rock masses at landslides is only erratically successful. Relatively impermeable matrix drains slowly. Fractured and relatively permeable blocks penetrated by drains may gush water for a while, providing an illusion of success, but discharge often diminishes with time as the blocks empty. GEOTECHNICAL OBSERVATIONAL METHOD The confirmation of design assumptions during construction is important, but based on the generally limited investigations and characterizations of mélanges normally performed, the geopractitioner may be unable to predict the performance of the work during construction. Hence adherence to the geotechnical observational method (Peck, 1969) is essential. This method requires that all possible modes of failure are anticipated and observations and measurements in the field are used to identify critical, potentially unforeseen, conditions. Different rock mass strengths may be assigned for the different failure modes. For example, in an excavation, there may be potential for large-scale slope failures through the bimrock; or else the detachment of a critically oriented block along adversely oriented matrix shears could cause significant damage. Estimates are then made of the likely progression of ground movements as the various failure mechanisms develop, as ground failures are rarely instantaneous. Once failure modes and scenarios are conceived, action criteria based on critical thresholds of ground movement are established,
rock mass and groundwater monitoring programs developed, and consequent remediation procedures prepared. During construction, ground deformations are monitored and ideally analyzed in real time (Moritz et al., 2004). If ground movements approach action thresholds, the construction remediation procedures are implemented to deter the failures. CASE HISTORIES The following four case histories are summaries of experience in geoengineering practice that illustrate many of the points made in this chapter. Case History 1: Mischaracterization of a Landslide As described by Medley and Sanz (2004), there are many factors that should be considered when characterizing and analyzing the slope stability of mélanges and other bimrocks. However, the most fundamental consideration is the recognition that a rock mass is a bimrock. The consequences of nonrecognition can be expensive, as described here. Hillside repairs were proposed to mitigate landslides that intermittently disrupted a main road in California. Based on borings terminated ~2 m into apparent sandstone bedrock, the investigating geopractitioner concluded that each of the landslides was composed of a shallow layer of clay and boulder colluvium sliding on the surface of underlying sandstone bedrock (Fig. 16). He recommended that the most economical repair would be removal of the failed soil down to the solid bedrock and re-grading of the failed slopes. The successful contractor bid for the designed repairs of the several landslides for a fixed price of more than a million dollars. During construction of the first landslide repair, the earthwork contractor encountered pervasively sheared shale containing abundant rock blocks up to several meters in size, which required considerable effort to remove, as blasting was not permitted. Neither solid bedrock, nor even a definite landslide “failure surface” between the landslide and the bedrock was observed; instead there were countless shear surfaces. In an attempt to find continuous bedrock, the excavation was deepened to several tens of meters below the design depth of a few meters. The project was halted when the repair for this first landslide repair had cost almost all the total contract amount set aside for repair of all the landslides. The landslide was actually a deep-seated earth flow in pervasively sheared mélange rather than a shallow soil mass sliding on top of apparent bedrock, as interpreted from exploration drilling. The interpreted bedrock was the graphical artifact of connecting straight lines between the soil-rock contacts intersected by the borings (Figs. 5 and 16). Although the geological chaos was a surprise to the geotechnical engineer, it should not have been, as publically available geological maps showed the locale of the landslide to be within Franciscan mélange, with large blocks protruding prominently from the hillside around the site.
Geopractitioner approaches to working with antisocial mélanges Case History 2: Estimate of Volumetric Block Proportion in an Excavation The Lone Tree Slide was a major landslide in Franciscan mélange that blocked California State Highway 1 near San Francisco (Van Velsor and Walkinshaw, 1993). To stabilize the slope, 950,000 m3 of intact and failed mélange were excavated to an average depth of 37 m. The drilled exploration of the landslide intersected a maximum block chord length of ~8 m. Based on the percentage length of drilled chords through blocks, the project geopractitioners estimated that 5% of the total volume to be excavated would be blocks large enough to be troublesome to excavate and thus eligible for extra compensation as rock. However, the sizes and proportion of the hard, large blocks encountered during excavation were far greater than had been anticipated during design and caused delays and unanticipated expense (Van Velsor and Walkinshaw, 1993). To develop approaches useful for future construction projects, Medley (1994a, 1994b) logged the drilled core and mapped the slopes of the excavation at Lone Tree Slide. The characteristic engineering dimension (Lc) for the landslide was assumed to be the average thickness of the original slide, some 30 m. The block-matrix threshold was selected as 1.5 m (0.05Lc). Several large blocks protruded from the undisturbed
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hillside adjacent to the landslide the exposed parts of which suggested that the intact blocks were at least 30 m or more in size (Medley, 1994a). Accordingly, the size of the largest block (dmax) was assumed to be ~30 m. About 375 m of drilled core from the landslide exploration was reviewed. Block/core chords were measured between 1.0 m and up to 7.9 m in length. The linear block proportion for all the mélange explored was ~10%, which was the weighted average of block-poor mélange within the landslide (about zero percent) and block-rich mélange beneath the slide (28%). The 375 m total length of drilling was equivalent in length to 12.5dmax (where dmax was 30 m). Figure 14 indicates the uncertainty to be at least 0.40 for a measured linear block proportion of 10%, total drilling of 375 m equivalent to 12.5 times the length of the estimated dmax of 30 m (i.e., Ndmax of 12.5), and the assumption that the measured linear block proportion of 10% was equivalent to a volumetric block proportion of 10%. Hence the estimated range of volumetric block proportions ranged between 6% and 14% (10% ± 0.4 times 10%). In a similar situation, for excavation purposes a conservative geopractitioner should adopt the higher bound (in this case, 14%). For an estimate of volumetric block proportion for evaluating the overall rock mass strength, the 6% lower bound was too low to improve the geomechanical properties of the rock mass. Nevertheless, if it had been available at the time, the higher bound estimate of block proportion would have been useful for construction cost purposes. According to the contractor, after the excavation work was completed, the volumetric proportion of excavated blocks was estimated to be between 6% and 11%, which was greater than the original design estimate of 5% but closer to the conservative post-project estimate of 14%. If the estimation had been performed prior to excavation instead of afterward, a suggested range of between 5% and 15% would have been reasonable for estimating the volumetric block proportion as the basis for pre-construction cost estimates to excavate and remove blocks. Case History 3: Estimate of the Strength of a Mélange Underlying a Dam
Figure 16. The upper sketch shows a geopractitioner’s landslide characterization based on shallow borings and an observed outcrop: a shallow soil landslide sliding on an assumed underlying continuous bedrock surface. The lower sketch shows the actual bedrock conditions in which the four borings intersected discrete blocks in a mélange composed of sheared shale matrix and blocks of various dimensions and lithologies. After Medley (2001). BH—borehole.
This case history outlines the approach adopted to characterize the bedrock underlying Scott Dam, which impounds the Eel River at Lake Pillsbury, 160 km north of San Francisco. Built in the 1920s, the dam is a masonry gravity structure ~40 m high. The dam is underlain by a Franciscan mélange (Goodman and Ahlgren, 2000; Hovland et al., 2000). In the 1970s conventional geotechnical analysis indicated that if the strength of the mélange under the dam was the same as the strength of the weak, sheared shale matrix (the thenconventional assumption), the foundation rock would be too weak to resist lateral sliding and the dam would fail by sliding mode along the base contact with mélange. But the dam was still intact more than 50 years after it was built, so the mélange bedrock was clearly stronger, possibly due to the presence of blocks
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in the sheared shale, as suggested by Volpe et al. (1991). A recharacterization of the Franciscan mélange and a reevaluation of the foundation strength was commissioned by the dam owner, an investigation which funded the Ph.D. research of Medley (1994a) and Lindquist (1994a). Characterization required selection of the block-matrix threshold, which in turn required selection of Lc, the characteristic dimension scaling the problem at hand. In general, flexibility is inherent in the selection of Lc for dam foundations. The dam width, dam height, square root of the area (√A) of the dam footprint, or thickness of a critical failure zone could all have been used for Lc. But the likely critical mode of failure was considered to be basal sliding at the contact between Scott Dam and the Franciscan mélange foundation rock. Accordingly, Lc was selected as a postulated 3 m thickness of a basal failure shear zone within the mélange rock at the base of the dam. As described above, geotechnically significant blocks that influence bimrock strength range between ~0.05Lc at the block-matrix threshold, and 0.75 Lc, for the largest block, dmax. For the Scott Dam project, the block-matrix threshold was selected as 0.15 m, or 5% of the postulated 3-m-thick basal shear zone. This criterion was used to discriminate blocks from matrix when reviewing drill core and drill logs. Measurement of block chord lengths in drill logs and photographs (Fig. 17) of Scott Dam core in the assumed critical potential failure zone indicated that the linear block proportion was ~40% for the postulated critical 3-m-thick shear band of failed mélange below the dam. The total length of exploration core required to yield a reasonably accurate estimate of a volumetric block proportion from drilling should be at least 10 times the size of the expected largest block dmax, or 10dmax. At Scott Dam, based on field observations and drilling, the size of the largest block was estimated to be between 30 and 43 m (Medley, 1997), so greater than 300 to 430 m of drilling would have been preferable. In actuality, only ~150 m of core (representing at least 5dmax) was recovered
Figure 17. Typical core from an exploration boring at Scott Dam, California, showing matrix of sheared shale containing blocks that were smaller than 0.05 Lc (0.15 m) and assigned to the matrix, and a measurable block. Photo: Richard E. Goodman.
from the several extensive exploration campaigns performed since the 1970s. Because there were insufficient data, the estimated volumetric block proportion had to be adjusted for uncertainty. As shown in Figure 14, using the procedure described by Medley (1997) the estimated linear block proportion was 40%, the uncertainty was 0.2, and the adjusted estimate was 32% to 48% (40% ± 0.2 times 40%). For purposes of strength estimates, it is prudent to use the lower bound, so a conservatively adjusted volumetric block proportion of 32% was selected for the volumetric block proportion of the Franciscan mélange below the dam. The estimate was subsequently lowered to 31% on the basis of additional exploration drilling (Goodman and Ahlgren, 2000). Because of scale independence, mélange and other bimrocks at the scale of laboratory specimens are closer to being scale models of the parent rock masses than is generally true in geotechnical engineering. Laboratory specimens of Franciscan mélange from the dam foundation (Goodman and Ahlgren, 2000) were tested using multistage triaxial compression tests such as those described by Lindquist (1994), Lindquist and Goodman (1994), Bro (1996, 1997), and Goodman and Ahlgren (2000). Given that the diameter of the laboratory specimens was the characteristic engineering dimension Lc, blocks in the specimens were those intact inclusions that had maximum dimensions between ~5% and 75% of the specimen diameter. The volumetric block proportions of each specimen were measured after specimen disaggregation and wash-sieving to retrieve the blocks. Strength testing of Scott Dam laboratory specimens with different block proportions yielded a plot of effective friction angle as a function of volumetric block proportion (Fig. 18). The overall strength of the foundation rock mass was evaluated using
Figure 18. Plot of effective angle of friction (φ′) as a function of volumetric block proportion, generated from laboratory testing of Franciscan mélange specimens obtained from core drilling at Scott Dam, Northern California. From Medley (2001), after Goodman and Ahlgren (2000).
Geopractitioner approaches to working with antisocial mélanges the adjusted estimate of rock mass volumetric block proportion and the laboratory plot of effective friction angle as a function of volumetric block proportion. For the mélange beneath Scott Dam, the friction angle was estimated to be 39° for the overall 31% volumetric block proportion (Fig. 18). Additional geotechnical engineering analyses confirmed that the strength of the mélange bimrock at the dam foundation was considerably greater than the strength of the matrix alone. On the basis of the geotechnical characterization and analysis of the mélange, both the California Division of Safety of Dams and the Federal Energy Regulatory Commission agreed that Scott Dam was safe and did not require any of the initially proposed structural reinforcement (Goodman and Ahlgren, 2000; Hovland et al., 2000). Essentially the same procedure described here for Scott Dam was used for the evaluation of the shear strength of Franciscan mélange below Calaveras Dam (Roadifer et al., 2009). Case History 4: Evaluation of Permeability of a Mélange In February 2000 a high excavated slope failed at a longestablished residential area of Millbrae, California (Snell and Medley, 2008). The slide was largely within Franciscan Complex mélange. Litigation proceedings followed shortly afterward, during which several allegations were made, one being that the San Andreas Lake water-supply reservoir, some 1500 ft from the landslide, had leaked groundwater through the intervening low broad ridge and caused the landslide. Geological and geotechnical investigations revealed Franciscan mélange within the ridge, consisting of a chaotic mixture of clayey, highly sheared matrix with blocks composed of sandstone, shale, greenstone, and other lithologies. Hydrogeological observations and analyses concluded that the permeability of the mélange ranged between 2.1 × 10−8 and 1.4 × 10−7 cm/s. Detailed hydrogeological analyses showed that the alleged contribution to the landslide of groundwater leaking from the reservoir through the ridge was invalid. Following additional investigations and analyses, it was apparent that other factors had contributed to the landslide. These included long-term degradation of the mélange exposed in the excavated slope, which was aggravated by lack of slope maintenance and the effects of poorly controlled surface drainage. CONCLUSIONS Mélanges, fault rocks, weathered rocks, and similar bimrocks are common and problematic for geopractitioners (geotechnical engineers, geological engineers, engineering geologists, and rock engineers) working in geologically complex areas of the world. Hence it is vital that geopractitioners have at least a conceptual understanding of the existence of bimrocks and are able to recognize them, even if they are unfamiliar with the processes of formation and other geological details. Despite their heterogeneity, mélanges and other bimrocks can be characterized for the purpose of geoengineering design
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and construction, even where there is great uncertainty in the characterization, or where the volumetric proportion of blocks is too small to provide geomechanical benefit. Simple procedures are available to characterize and analyze mélanges, and implementation of these procedures may reduce expensive surprises by focusing the geopractitioners’, owners’, and contractors’ attentions on the difficulties that may be encountered during design and construction. Accordingly, Society will be better served if geopractitioners working with mélanges learn to expect the unexpected. We have written this chapter to increase awareness by geologists of the practical problems faced by engineers working (often heroically) with geologically exciting but chaotic mélanges and other troublesome bimrocks. Geologists have much understanding, which, if shared without unnecessary geological jargon, could assist engineers in recognizing mélanges and other geologically complex rocks with a view to improving the simple approaches described here for their characterization, design, and construction. The writers hope that geologists and geopractitioners are more cautious the next time they use the expressions interbedded or soil with boulders in boring logs or reports, and that they persuade engineers to critically review the cross sections they develop from their field observations of sites in mélanges. Overall, we believe that better geologist-geopractitioner collaboration will reduce the expensive surprises that so often occur in the design and construction of Society’s constructed works. SUGGESTED RESEARCH We suggest the following avenues of cooperative research between geologists and geopractitioners: • Understand better the effects of block-matrix contrast strength on the overall mechanical behavior of bimrocks. • Develop virtual bimrocks using numerical modeling methods to simulate under-controlled conditions of bimrock masses, and investigate geomechanical behaviors of model bimrocks with variations in block-matrix and strength-stiffness contrasts, block shapes and sizes, block size distribution, and spacing. Pan et al. (2008) have made a start on such research. • Survey how various geological disciplines characterize the complexity of the range of geological mixtures, such as ore bodies within waste rock, fault rocks, and discontinuous contaminated soil and groundwater within pristine geology. • Develop statistical, geostatistical, and stereological approaches to understanding and predicting the uncertainties of our estimates of rock block volumes, sizes, shapes, orientations, etc., based on the limited drilling and mapping exploration tools available to geologists and geopractitioners. • Understand better the complex hydrogeological interactions within stressed rock-soil mixtures by learning from structural geology and geomechanics.
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Medley and Zekkos • Develop a database of characterization design and construction experience in bimrocks, such as the recent tunneling experience of Spreng et al. (2008) and Roadifer et al. (2009) in Franciscan mélanges.
ACKNOWLEDGMENTS We are grateful to John Wakabayashi (a structural geologist with rare appreciation of society’s engineering problems) for encouraging us (engineers with rare appreciation of geology) to write this chapter. Invaluable comments were offered in the reviews of Harun Sönmez and Bill Haneberg, both academics/ consultants with rare broad appreciation of both the geology and the engineering issues related to working in mélanges and other bimrocks. REFERENCES CITED Associazone Geotechnica Italiana (AGI), 1977, Proceedings of the International Symposium on the Geotechnics of Structurally Complex Formations: Capri, Italy, Associazone Geotechnica Italiana. Attewell, P.B., 1997, Tunnelling and site investigation: Proceedings of the International Conference on Geotechnical Engineering of Hard Soils–Soft Rocks: Rotterdam, A.A. Balkema, v. 3, p. 1767–1790. Aversa, S., Evangelista, A., Leroueil, S., and Picarelli, L., 1993, Some aspects of the mechanical behaviour of “structured” soils and soft rocks: Proceedings of the International Symposium on Geotechnical Engineering of Hard Soils–Soft Rocks: Rotterdam, A.A. Balkema, v. 1, p. 359–366. Bagnold, R.A., and Barndorff-Nielsen, O., 1980, The pattern of natural size distribution: Sedimentology, v. 27, p. 199–207, doi:10.1111/j.1365-3091.1980 .tb01170.x. Berkland, J.O., Raymond, L.A., Kramer, J.C., Moores, E.M., and O’Day, M., 1972, What is Franciscan?: American Association of Petroleum Geologists Bulletin, v. 56, p. 2295–2302. Blake, M.C., and Harwood, D.S., 1989, Tectonic Evolution of Northern California, Field Trip Guidebook for Trip 108: Washington, D.C., American Geophysical Union. Blake, M.C., and Jones, D.L., 1974, Origin of Franciscan melanges in northern California, in Dott, R.H., and Shaver, R.H., eds., Modern and Ancient Geosynclinal Sedimentation: Tulsa, Society of Economic Paleontologists and Mineralogists, p. 345–357. Bro, A., 1996, A weak rock triaxial cell; Technical Note: International Journal of Rock Mechanics and Mining Sciences, v. 33, p. 71–74, doi:10.1016/0148-9062(95)00047-X. Bro, A., 1997, Analysis of multi-stage triaxial test results for a strain-hardening rock: International Journal of Rock Mechanics and Mining Sciences, v. 34, p. 143–145, doi:10.1016/S1365-1609(97)80040-8. Button, E.A., Schubert, W., Riedmueller, G., Klima, K., and Medley, E.W., 2003, Tunnelling in tectonic mélanges—Accommodating the impacts of geomechanical complexities and anisotropic rock mass fabrics: International Bulletin of Engineering Geology and the Environment. Cloos, M., 1990, Evolution of the geological interpretation of the Franciscan Complex in the San Francisco Bay region: A comparison of crosssections, in Bilodeau, B.J., and Davis, S.O., eds., Geologic Guidebook to the Point Reyes Area, Northern California: Bakersfield, California, American Association of Petroleum Geologists, Pacific Section, p. xxiii–xxxi. Coli, N., Berry, P., Boldini, D., and Castellucci, P., 2008, Analysis of the Block Size Distribution in the Shale-Limestone Chaotic Complex (Tuscany, Italy): San Francisco, Proceedings, Symposium of American Rock Mechanics Association, June 27–July 3, 2008, 7 p. Cowan, D.S., 1985, Structural styles in Mesozoic and Cenozoic melanges in the Western Cordillera of North America: Geological Society of America Bulletin, v. 96, p. 451–462, doi:10.1130/0016-7606(1985)96<451: SSIMAC>2.0.CO;2. D’Elia, B., Distefano, D., Esu, F., and Federico, G., 1986, Slope movements in structurally complex formations, in Tan Tjong Kie, Li Chengxiang, and
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MANUSCRIPT ACCEPTED BY THE SOCIETY 21 DECEMBER 2010 Printed in the USA
Contents Introduction: Characteristics and tectonic settings of melanges, and their significance for societal and engineering problems John Wakabayashi and Yildirim Dilek
Part I. Melange Generation in Oceanic Fracture Zones in Abyssal Settings 1. Serpentinite matrix melange: Implications of mixed provenance for melange formation John W. Shervais, Sung Hi Choi, Warren D. Sharp, Jeffrey Ross, Marchell Zoglman-Schuman, and Samuel B. Mukasa
2. Geochemical mapping of the Kings-Kaweah ophiolite belt, CaliforniaEvidence for progressive melange formation in a large offset transform-subduction initiation environment J. Saleeby
Part II. Melange Formation Associated with Subduction Initiation 3. Constraints on the evolution of the Mesohellenic Ophiolite from subophiolitic metamorphic rocks R. Myhill
4. Role of plutonic and metamorphic block exhumation in a forearc ophiolite melange belt: An example from the Mineoka belt, Japan Ryota Mori, Yujiro Ogawa, Naoto Hirano, Toshiaki Tsunogae, Masanori Kurosawa, and Tae Chiba
Part III. Melange Development in Subduction-Accretion Complexes and in Collisional Settings 5. Melanges of the Franciscan Complex, California: Diverse structural settings, evidence for sedimentary mixing, and their connection to subduction processes John Wakabayashi
6. Tectonic evolution of the Ankara Melange and associated Eldivan ophiolite near Hanfili, central Turkey Anne Dangerfield, Ron Harris, Ender Sanfak10glu, andYildirim Dilek 7. Petrology of a Franciscan olistostrome with a massive sandstone matrix:
The King Ridge Road melange at Cazadero, California Rolfe Erickson
8. Sedimentary block-in-matrix fabric affected by tectonic shear, Miocene Nabae complex, Japan Soichi Osozawa, Terry Pavlis, and Martin F.J. Flower
9. Numerical estimation of duplex thickening in a deep-level accretionary prism: A proposal for network duplex Hikaru Ueno, Ken-ichiro Hisada, and Yujiro Ogawa
Tectonic, sedimentary, and diapiric formation of the Messinian melange: Tertiary Piedmont Basin (northwestern Italy) Andrea Festa
Recognition of a trench-fill type accretionary prism: Thrust-anticlines, duplexes, and chaotic deposits of the Pliocene-Pleistocene Chikura Group, Boso Peninsula, Japan Satoru Muraoka and Yujiro Ogawa
Implication of dark bands in Miocene-Pliocene accretionary prism, Boso Peninsula, central Japan Yoko Michiguchi and Yujiro Ogawa
Part IV. Significance of Melanges for Engineering and Applied Geology Geopractitioner approaches to working with antisocial melanges Edmund W. Medley and Dimitrios Zekkos
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