Tectonosomes and Olistostromes in the Argille Scagliose of the Northern Apennines, Italy (GSA Special Paper 335)

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Copyright© 1999, The Geological Society of America, Inc. (GSA). All rights reserved. GSA grants permis ion to individual scientists to make unlimited photocopies of one or more items from this volume for noncommercial purpo es advancing science or education, including classroom use. Permission is granted to individual to make photocopies of any item in this volume for other noncommercial, nonprofit purposes provided that the appropriate fee ($0.25 per page) is paid directly to the Copyright Clearance Center, 222 Rosewood Drive, Danvers, MA 01923, USA, phone (978) 750-8400, http://www.copyright.com (include title and ISBN when paying). Written permission is required from GSA for all other forms of capture or reproduction of any item in th.e volume including, but not limited to, all types of electronic or digital scanning or other digital or manual transformation of articles or any portion thereof, uch as abstracts, into computer-readable and/or tran mittable form for personal or corporate u e, either noncommercial or commercial, for-profit or otherwise. Send permission requests to GSA Copyrights. Copyright is not claimed on any material prepared wholly by government employees within the scope of their employment. Publi bed by The Geological Society of America, Inc. 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301 Printed in U.S.A. GSA Books Science Editor Abhijit Basu Library of Congress Cataloging-in-Publication Data Pini, Gian Andrea, 1957Tectono omes and olistostromes in the argille scagliose of the Northern Apennines, Italy I Gian Andrea Pini. p.

em. -- (Special paper ; 335)

.Includes bibliographical references (p. ISBN 0-8137-2335-3 l . Rocks, Sedimentary--Italy--Apennines. -Apennin.es.

3. Geology--Italy--Apennines.

2. Olistostromes--ltaly-

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Special papers (Geological Society of America) ; 335. QE47 l .P528 1999 552'.5'0945--dc21

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Cover: Top--Typical landscape of the Bologna Apennines; the badlands are developed on Lower Cretaceous Ligurian tectonosomes. The whitish bodies

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marly turbiditic beds,

which have been boudinaged and strongly deformed in brittle simple shear condition. Bologna Apenniites, Sillaro VaHey. Bottom-Stratal disruption and stacking of a strati­ graphic level (the thick limestone bed) in th.e Lower Cretaceous Ligurian tectonosomes. Bologna Apennines, ldice Valley. For more information see Figs. 7B, SA, and 25A.

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Geological Society of America Special Paper 335 1999

Tectonosomes and olistostromes in the argille scagliose of the Northern Apennines, Italy Gian Andrea Pini Dipartimento di Scienze della Terra e Geologico-ambientali, Università di Bologna, via Zamboni 67, I-40127 Bologna, Italy

ABSTRACT Rocks characterized by an apparently chaotic block in matrix fabric and known as argille scagliose are widespread in the Apennines of Italy. Most of the argille scagliose can be differentiated into two kinds of units, tectonosomes and olistostromes, which have different sedimentological and structural characteristics and implications. The tectonosomes derived from polyphase deformation of lithostratigraphic units of an oceanic cover succession. Stratal disruption resulted from (1) folding and boudinage of nonconsolidated sediments, compatible with shallow tectonic deformation or gravitational spreading and slumping, and (2) prevalent brittle, simple-shear deformation, related to boudinage, transposition of bedding, stacking of blocks, and thrusting. A pervasive scaly fabric enabled the tectonic mobility of the matrix that accompanied the simple shear deformation. Olistostromes are chaotic sedimentary bodies. They have been interpreted as syntectonic deposits related to episodes of slope failure at the front of thrust systems in thrust-bordered, piggyback basins. The inferred depositional mechanisms are debris avalanches that become debris flows downslope. All of these bodies, whether tectonosomes or olistostromes, have, at first observation, common characteristics, such as block-in-matrix fabric and the seemingly chaotic disposition of blocks. The criteria for differentiating tectonosomes from olistostromes may be useful to distinguish tectonically disrupted lithostratigraphic units from submarine debris flow and slide deposits in melanges in other orogenic belts. The basic criteria for field distinction are composition and age of component rocks, different order-disorder ratio in the distribution of blocks, and the presence of scaly fabric or of a matrix hosting millimeter to submillimeter clasts (brecciated matrix).

INTRODUCTION

subsequent usage of diverse terms such as Ligurian melange, Chaotic Complex, Undifferentiated Complex, complesso caotico eterogeneo (heterogeneous chaotic complex), or caotico eterogeneo (heterogeneous chaos), to indicate the same rocks throughout the Apennine chain (Bailey and McCallien, 1963; Abbate and Sagri, 1970; Boccaletti and Coli, 1982; Boccaletti et al., 1987). Some of the argille scagliose have been recognized as sedimentary bodies inside layered sequences of normal marine strata. These bodies were defined as olistostromes by Flores (1955, 1956) (see discussion and references in Abbate et al., 1970a; Elter

The terms argille scagliose and olistostrome were coined in the Apennine chain and in Sicily, respectively (Fig. 1). The term argille scagliose has been widely utilized in the Italian geologic literature to indicate structural units, tectono-stratigraphic units, or sedimentary bodies that are characterized by a chaotic appearance in outcrop, by a prevailing block-in-matrix fabric, and by a certain degree of rock mixing (e.g., see Merla, 1952; Maxwell, 1959; Hsü, 1965; Page, 1963, 1978). These attributes explain the

Pini, G. A., 1999, Tectonosomes and olistostromes in the argille scagliose of the Northern Apennines, Italy, in Pini, G. A., Tectonosomes and olistostromes in the argille scagliose of the Northern Apennines, Italy: Boulder, Colorado, Geological Society of America Special Paper 335.

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Figure 1. Index map showing the geographic position of the Northern Apennines chain.

and Trevisan, 1973; Hoedemaeker, 1973). The term olistostrome (or olisthostrome) is generally utilized to indicate the products of mass-gravitational transport, mostly related to the mechanism of muddy debris flows, in accretionary wedges, orogenic belts, and passive-margin sequences (see Hoedemaeker, 1973; Moore and Karig, 1976; Abbate et al., 1981a; Raymond, 1984; Cowan, 1985; Brandon, 1989; Chanier and Ferrière, 1991; Taira et al., 1992a; Jeanbourquin, 1994; Orange and Underwood, 1995; Steen and Andresen, 1997). The term olistostromal processes refers to mechanisms of stratal disruption and rock mixing related to gravitational mass processes (Steen and Andresen, 1997). After some of the chaotic bodies in the Apennines were differentiated as olistostromes, the terms argille scagliose, Chaotic Complex, or heterogeneous chaos were still applied to a significant part of a large allochthonous nappe in the Apennines, the Ligurian nappe. Even though the identification of olistostromes is supported in many places by the field observation of sedimentary bottom and top contacts within well-studied sedimentary sequences, the relationships among the argille scagliose and the other units of the Ligurian nappe are typically obscure, and consequently their interpretation has changed through time. Nevertheless, the argille scagliose were always considered to be large masses of rocks derived

from a distant paleogeographic setting, and the disruption of these masses was related to gravitational tectonics or to massgravitational processes (Merla, 1952; Hsü, 1965, 1967; Abbate and Sagri, 1970; Page, 1978; De Jager, 1979; Pieri and Mattavelli, 1986). Thus, a conceptual linkage between olistostromes and the argille scagliose has always been maintained, or at least never been excluded. Moreover, the textural characteristics of olistostromes have frequently been confused with some of the peculiar characteristics of argille scagliose, such as their scaly fabric. In spite of the unresolved or uncertain origin of many of the bodies of argille scagliose in the Apennines, the terms olistostrome and argille scagliose have been widely applied to melanges in other orogenic chains. The terms carry with them the implication that the block in matrix fabrics, stratal disruption, and other attributes of these melanges might have originated by gravitational tectonics, or by sedimentary mechanisms such as mass flows (e.g., Rigo de Righi and Cortesini, 1964; Hsü, 1965; Maxwell, 1974; Cowan, 1978; Page, 1978; Page and Suppe, 1981; Swarbick and Naylor, 1980; Abbate et al., 1981a; Aalto, 1982; Cowan, 1985). The argille scagliose of the Northern Apennines can be subdivided into (1) strongly deformed formations, or tectonosomes; (2) sedimentary bodies emplaced by debris flows or avalanches, or olistostromes; and (3) normally bedded sedimentary units. The term tectonosome here describes deformed formations that prevailingly have a block in matrix fabric, that resulted from intense stratal disruption. Due to this fabric, the outcrop aspects of the tectonosomes are apparently similar, at a first observation, to those of the olistostromes. The criteria for distinguishing tectonosomes and olistostromes in a block in matrix rock assemblage derive from Pini (1987a, 1987b) and Castellarin and Pini (1989); they are similar to those in works by Bettelli and Panini (1987, 1989, 1992), based on work in the Northern Apennines between Modena and Parma. These basic criteria, which were published in Italian journals, are described here because they are not easily accessible to the international scientific community. These criteria may be useful for identifying and differentiating tectonosomes and olistostromes in melanges in orogenic belts elsewhere. The paper has the following objectives. 1. Herein I review the setting of the Ligurian nappe and the evolution of the concepts and nomenclature regarding chaotic units and argille scagliose. 2. I present the characteristics, including fabric, composition, and disposition, that enable the differentiation of tectonosomes and olistostromes in the apparently similar disrupted rocks of the argille scagliose. 3. I describe the mechanism of stratal disruption and the structural evolution of the tectonosomes of the Apenninic border, south of the town of Bologna. 4. I define the mechanism of olistostrome genesis, based on present-day examples of slope failure at thrust fronts and mud diapirism.

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy STRUCTURAL SETTING OF THE LIGURIAN UNITS IN THE NORTHERN APENNINES The geology of the Northern Apennines has been discussed in several papers. This chapter makes reference to summary and comprehensive papers, and to maps by Abbate et al. (1970b, 1988); Boccaletti et al. (1971, 1982, 1990a, 1990b); Elter (1973), Dallan Nardi and Nardi (1974), Carmignani et al. (1978, 1994), Funiciello et al. (1979), Reutter and Groshurt (1978), Reutter (1981), Boccaletti and Coli (1982), Treves (1985), Castellarin and Vai (1986), Ricci Lucchi (1986), Sestini et al. (1988); Bettelli et al. (1989a); Lavecchia et al. (1989), Patacca and Scandone (1989), Carmignani and Kligfield (1990), Patacca et al. (1990), Bigi et al. (1991, 1992), Castellarin et al. (1992, 1994, 1995), Baldacci et al. (1994), Castellarin (1994), and Conti and Gelmini (1994). The Northern Apennine chain, the geographic position of which is shown in Figure 1, is characterized by predominantly Cenozoic thrust faults and folds that in general verge toward the northeast. The structurally higher thrust sheets or nappes came

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from progressively farther to the west-southwest and were emplaced onto the Adria microplate. The sedimentary cover and the basement of the western sector of the Adria microplate have been diffusely incorporated into the chain. The chain has a buried continuation under the Quaternary cover of the Po Plain and the Adriatic Sea. The buried Apennine front, shown in Figure 2, has been located on seismic reflection profiles (see Pieri and Groppi, 1981; Pieri, 1983; Castellarin et al., 1985). In the first part of this section, the disposition of structures and rocks in the exposed chain is summarized with reference to Figures 2 and 3. Simplified paleogeographic models, showing where the units came from and when they were emplaced onto the Adria plate, are in Figure 4. In paleogeographic and structural terms, units that originated in more westerly positions are conventionally called internal, and those that were situated in more easterly areas, external. The geology of the Apennines is complicated by the continuous interaction between tectonics and sedimentation. This interaction, and the proliferation of formal and informal stratigraphic

Figure 2. Structural map of the northwestern part of the Apennines. Adapted from Funiciello et al. (1979); Boccaletti and Colli (1982); Bigi et al. (1992).

Figure 3. Major structural units, lithostratigraphic units, and paleogeographic domains of the Northern Apennines. Block letters indicate the stratigraphic units and the paleogeographic domains; structural units are in italics. The superposition of the structural units occurred progressively from southwest (internal units) to northeast (external). Abbreviations: Sd—Santerno Pliocene-Pleistocene deposits; Co—Colombacci Formation; Gs—Gessoso-solfifera Formation; Ma—Marnoso-arenacea Formation; cc—Carbonate-clastic succession; cs—Carbonate succession; ev—Upper Triassic evaporites and carbonates; Hb—Hercynian basement; CG—Castel Guerrino Sandstone; CF—Monte Cervarola and Monte Falterona Sandstones; Ps— Polychromous shales; Po—Porretta Formation; Pm—Ponte della Venturina melange; PM—Metamorphic “Pseudo Macigno”; VT—Continental deposits of the Tuscan “Verrucano”; Mc—“Macigno” Sandstone; Mo—Monte Modino Sandstone; FM—Fiumalbo and Marmoreto Marlstones; Mm—Monte Modino melange; Pi—Pievepelago Formation; Ca—Canetolo Complex; Ef—Eocene flysches; Pf—Paleocene-Eocene flysches and marly covers; Hf—Helminthoid flysches; bc—Ligurian “basal complexes”; op—Ligurian ophiolites.

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Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy

Figure 4. Paleogeographic sketch maps of the Northern Apennines at different stages. Sketch A: The Alpine Neo-Tethys, Late Jurassic–Early Cretaceous. (Adapted from Castellarin, 1982b; Abbate and Sagri, 1984; Abbate et al., 1988.) B: The paleogeography at the end of the Mesoalpine tectonic phase, in the latest part of the middle Eocene. After Castellarin (1994). C: The earliest Neogene tectonic phase, early-middle Miocene (after Castellarin, 1994).

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units, make it difficult to understand the stratigraphic and tectonic relationships among the rocks constituting the principal structural units. In this paper, the following conventions have been adopted. 1. Formations have been used as the fundamental rock-stratigraphic unit. The usage of the informal names has been limited as much as possible, and only the widely diffused and recognized informal terms have been adopted. In addition, the incredible number of formation names has been simplified; similar formations are grouped under the name of the most widespread and well known. The term complex has been utilized in the Apennines as either a formal or informal special stratigraphic unit following the correct usage defined on the International Stratigraphic Code, as in the case of the Canetolo Complex, Chaotic Complex, Casanova Complex, complesso caotico eterogeneo. However, the same terms have been utilized to indicate the same rocks, but as structural units. In this paper, the term has been utilized only with its stratigraphic meaning. 3. Formal and informal supergroups are rock-stratigraphic units from higher in the hierarchy. These comprise formations of the Ligurian domains (Fig. 5) that are in stratigraphic continuity, and which were deposited in the same basin (Fig. 4A). 4. Domains are the fundamental units of the paleogeography; they are indicated in Figures 3 and 4A. Domains include one or more supergroups, which have been deposited in the same basin or in nearby basins (Fig. 4A). 5. Fault-bounded structural units are the fundamental map units in the Northern Apennines. There are different relationships between stratigraphic units and structural units. A structural unit may correspond to a single formally or informally defined formation. More typically, a structural unit includes two or more formations that are part of the same supergroup. However, certain distinctive formations and informal stratigraphic units, e.g., the Palombini Shales, are present in more than one structural unit. Several structural units, such as the Caio, Cassio, Monghidoro, and Canetolo structural units and the Modino and Cervarola structural units in Plate 1, take their names from one of the constituting stratigraphic units: the Caio, Cassio, and Monghidoro flysches, the Canetolo Complex, and the Monte Modino and Monte Cervarola Sandstones. These names may be confusing, but are adopted here because they are widely used in the geological literature of the Northern Apennines. Major units and structures The highest and the most far-traveled structural unit is the Ligurian nappe. It came from the most westerly position, was thrust over all the other structural units, and currently extends from the Tyrrhenian Sea to the southwestern margin of the Po Plain (Figs. 2, 3, and Plate 1). In the southwestern part of the Northern Apennines, which extend from the crest of the chain to the Tyrrhenian Sea, the Ligurian nappe is structurally above the diverse Tuscan structural units. The higher and more internal of these units, the Tuscan nappe, was stripped off the Tuscan base-

Figure 5. Schematic stratigraphic diagram of the Ligurian Supergroups. Only the major formal and informal names of the lithostratigraphic units are shown (E, M, L—Early, Middle, Late). Modified from Vai and Castellarin (1993).

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy ment that constitutes the westernmost part of Adria. The Tuscan nappe and the associated Modino and Pievepelago units were thrust over the Apuane Alps metamorphic unit. The Sestola-Vidiciatico, Cervarola, and Castel Guerrino units are thrust sheets considered to be more external than the Tuscan nappe, but the link among these units, the Tuscan nappe, the Apuane Alps, and the Adria basement, are still unresolved. In the northeastern side of the chain, facing the Po Plain and the Adriatic Sea, the Ligurian nappe was emplaced onto the Romagna-Umbria and the Marche-Adriatic fold and thrust belts, which were developed largely in the stratified cover of the Adria plate. The original structural relationships among the major units have been complicated by younger out of sequence thrusting, the most evident effects of which were the emplacement of the Tuscan units onto the Ligurian nappe and the further translation of Cervarola and Castel Guerrino thrust sheets over the RomagnaUmbria fold and thrust belt (Figs. 2, 3, and Plate 1). Furthermore, the Tyrrhenian side of the Apennines, to the Apennine crest, was subsequently affected by extensional tectonics related to the opening of the Tyrrhenian Sea. Extension resulted in the formation of the Tuscan grabens and the exhumation of the Apuane Alps metamorphic core complex (Carmignani and Kligfield, 1990). Ligurian nappe. The Ligurian nappe comprises many different fault-bounded structural units, some of which are differentiated in Plate 1. The units include ophiolites and weakly or nonmetamorphosed sedimentary rocks. The sedimentary rocks have been subdivided into several formal and informal stratigraphic units, some of which are named in Figure 5. Formations that are typically in stratigraphic continuity have been collected into groups and supergroups (see Abbate and Sagri, 1970; Vai and Castellarin, 1993), as shown in Figure 5. The stratigraphic units of different groups and supergroups that have common sedimentological and compositional characteristics may in turn be grouped into some distinctive stratigraphic intervals. These intervals generally are marked by a difference in composition and style of deposition, and are separated by angular unconformities or locally by decollements (Vai and Castellarin, 1993). The most widespread intervals consist, from younger to older (Fig. 5), of the following. The Paleocene–lower Eocene flysches and the upper Campanian to Maastrichtian helminthoid flysches are calcareous or arenaceous turbidites and are underlain by basal complexes. The basal complexes include Cenomanian to Campanian calcareous-arenaceous-shaly turbidite successions (Ostia and Scabiazza Sandstones, Val Lavagna and Ronco Formations); Neocomian to Cenomanian shaly-calcareous deposits (Palombini Shales, Varicolored Shales), which are typically associated with Upper Jurassic sedimentary rocks (Calpionella Limestones, diaspri, or radiolarites, ophiolitic breccia, and ophiolitic sandstones) directly overlying the ophiolites (serpentinized peridotite, gabbros, basalts, pillow lava, ophicalcites). The age of the basal complexes extends to earliest Eocene, where they crop out with lower Eocene flysches, e.g., the Calvana Supergroup (Fig. 5).

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The basal complex and helminthoid flysch are informal, but widely recognized, names commonly used in Apennine geology. These terms have been adopted here to simplify the stratigraphy in the tectono-stratigraphic diagram (Fig. 3), and to summarize the stratigraphic sketch (Fig. 5). The Upper Cretaceous flysches were called helminthoid flysches before their distinction in several formations (Fig. 5). This definition derives from the presence of trace fossils (Helminthoidea labirinthica) on the top of several beds in the successions. The basal complex indicates the set of formations, and of chaotic rocks, that is the stratigraphic base of a helminthoid flysch. The supergroups correspond to different basins inside the Mesozoic Ligurian paleo-ocean, a part of the Neo-Tethys, which separated the European plate to the north and west from the Adria microplate to the south and east (see Channel et al., 1979; Abbate et al., 1988) after the break-up of Pangea (Fig. 4A). These basins are usually grouped in two different subdomains, the internal and external Ligurian domains (Abbate and Sagri, 1970; Sestini et al., 1988; Vai and Castellarin, 1993; Baldacci et al., 1994)(Fig. 5). Epiligurian successions. The Epiligurian successions (see Ricci Lucchi and Ori, 1985; Ricci Lucchi, 1986; Bettelli et al., 1989b) are middle-upper Eocene to upper Pliocene strata. These sediments were deposited unconformably on more strongly deformed units of the Ligurian nappe, as recognized by Merla (1952) and confirmed subsequently (e.g., Maxwell, 1959; Sestini, 1970; Ricci Lucchi, 1986; Vai and Castellarin, 1993; Mutti et al., 1995). All these authors pointed out the lesser degree of allochthony of the Epiligurian rocks with respect to the Ligurian units, as well the progressive decrease in tectonic translation upward in the successions (semiallochthonous units; Selli, 1962; Ricci Lucchi, 1987). These attributes imply that the Epiligurian successions were deposited on the Ligurian nappe during its translation (e.g., Merla, 1952; Maxwell, 1959). Other structural units. The other structural units and nappes (Tuscan nappe, Cervarola-Falterona, Sestola-Vidiciatico, and Castel Guerrino units, Romagna-Umbria and Marche-Adriatic fold and thrust belt) derive from the stacking of the sedimentary successions deposited in the Tuscan and Umbrian domains of the Adria microplate (Figs. 3 and 4). The sedimentary successions have a common typical stratigraphy (Fig. 3). The late Oligocene marks the beginning of the deposition of various siliciclastic turbiditic sequences, such as the Macigno, Monte Cervarola, Monte Falterona, and Castel Guerrino Sandstones, and the Marnoso-arenacea Formation (Fig. 3). The turbiditic sequences, which become progressively younger from southwest to northeast, were deposited in a migrating foredeep, at the front of the Ligurian nappe (see Boccaletti and Guazzone, 1970; Bortolotti et al., 1970; Reutter, 1981; Ricci Lucchi, 1986; Vai, 1987, 1988; Lavecchia et al., 1989; Boccaletti et al. 1990a; Conti and Gelmini, 1994). These Neogene turbiditic sequences commonly include large chaotic sedimentary bodies, derived from the Ligurian nappe, that were emplaced by gravitational mass movements into the foredeep (olistostromes of Tuscan and Umbrian successions, see Abbate et al., 1970a, 1981a; Elter and Trevisan, 1973; Ricci

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Lucchi, 1975; De Jager, 1979). These bodies are commonly interpreted as the sedimentary precursors of the submarine emplacement of the Ligurian nappe onto the foredeep sequences (precursory olistostromes; Elter and Trevisan, 1973). Certain structural units identified as Subligurian in Figure 3 and Plate 1 are considered to have been deposited in an intermediate paleogeographic position between the Tuscan and Ligurian domains (Abbate and Sagri, 1970, 1984; Sestini et al., 1988; Vai and Castellarin, 1993). They may derive from a thinned Adria margin, west of the Tuscan domain. They crop out as structural units at the base of the Ligurian nappe. Compressional history The age of the structural arrangement of the Ligurian units has been long discussed and is generally encompassed by two models: (1) all the Ligurian units have been arranged mostly by Late Cretaceous to early-middle Eocene tectonic phases (Merla, 1952; Abbate et al. 1970b, 1988; Papani, 1971; Elter, 1973; Zanzucchi, 1978, 1980; Plesi et al.; 1993; Vai and Castellarin, 1993; Carmignani et al., 1994; Castellarin, 1994); (2) only the internal Ligurian units were arranged by the Cretaceous-Eocene phases; the external Ligurian units were deformed mostly by a subsequent Oligocene to early Miocene tectonics (Boccaletti et al., 1971; Reutter and Groshurt, 1978; Reutter, 1981; Marroni and Meccheri, 1993). In either case, this structural arrangement is related to the closing of the Ligurian paleo-ocean. These tectonic phases caused the obduction of the Ligurian units onto the western margin of the Adria microplate. The deformation of the Adria margin occurred from Oligocene to Quaternary time as the Ligurian nappe was translated on the continental margin (Boccaletti et al., 1971; Reutter and Groshurt, 1978; Reutter, 1981; Castellarin et al., 1985, 1994, 1995; Conti and Gelmini, 1994). Compressional tectonics is still active along the Apennines–Po Plain border and in the buried chain (Bartolini et al., 1983; Castellarin et al., 1985; Castellarin and Vai, 1986; Patacca and Scandone, 1989). The Neogene northeastward migration of the contractional fronts (Boccaletti and Guazzone, 1970; Reutter et al., 1978) is associated with the synchronous diffusion of extensional tectonics in the Tyrrhenian area (Elter et al., 1975; Boccaletti et al., 1982, 1990a; Vai, 1987, 1988; Sartori et al., 1989; Lavecchia and Stoppa, 1989; Carmignani and Kligfield, 1990; Patacca et al., 1990; Castellarin, 1994). The mechanism of Ligurian nappe translation was long considered as passive, involving gravitational sliding from the stacked chain and into the foredeep (see Migliorini, 1948; Merla, 1952; Abbate et al., 1970b; Elter and Trevisan, 1973; De Jager, 1979; Pieri and Mattaveli, 1986; De Feyter, 1992; Farabegoli et al., 1992), or related to gravitational spreading (see Van Wamel and Zwart, 1990). However, the geometrical identity and the direct kinematic link between the structures in the Ligurian nappe and the compressional thrusts and folds in its tectonic substratum, as well as their common age of activity, suggest that the Ligurian

nappe was emplaced as a result of compressional shortening in an orogenic wedge (Castellarin et al., 1985; Castellarin and Pini, 1989; Conti, 1987, 1994; Conti and Gelmini, 1994; Conti and Tosatti, 1996). Marine sedimentation continued on top of the Ligurian nappe as it advanced northeastward. The resulting Epiligurian successions have been progressively carried as satellite basins (Ricci Lucchi, 1986) on top of the Ligurian nappe during its movement over the margin of Adria (Figs. 2 and 4C). Epiligurian successions represent a posttectonic deposit on Ligurian structural units that had been deformed by the Late Cretaceous and Eocene tectonic phases. However, the Epiligurian successions and the Ligurian units were deformed together by late Paleogene–Neogene faulting and thrusting that affected the entire Ligurian nappe (Conti, 1987, 1994; Pini, 1987b; Castellarin and Pini, 1989; Conti and Gelmini, 1994). The sedimentation of the Epiligurian units occurred concurrently with the evolution of these tectonic structures, so that the shape of the basins and the sedimentation have been controlled by thrusts (piggy-back basins; Ori and Friend, 1984). CONTROVERSY REGARDING THE ARGILLE SCAGLIOSE AND OLISTOSTROMES The Ligurian nappe contains three principal components that differ in their styles and degrees of deformation. Most abundant are well-bedded stratigraphic successions with perfectly continuous strata, more or less affected by faults and folds (e.g., the Caio, Cassio, and Monghidoro units, Fig. 6). A second, less-abundant component consists of deformed stratigraphic units, in which the continuity of beds is generally preserved even though folds and boudinage were pervasively developed (e.g., the basal complexes in Plate 1). These bedded sequences are associated with more thoroughly disrupted units, which are the main topic of this paper. These apparently chaotic and variably disrupted assemblages consist of blocks of different dimensions, ranging from centimeters to hundreds of meters, that are dispersed in a clayey or shaly matrix. The blocks are fragments of single calcareous and siliciclastic turbidite beds, packages of turbidite beds with internal stratal continuity (herein called bed packages), or slabs of ophiolites and ophiolitic breccias. These disrupted units have been described by a variety of terms. For example, these rocks have been called argille scagliose, or more recently Chaotic Complex and heterogeneous chaos. Some of these rocks have been identified as olistostromes. The complicated evolution of the nomenclature is reviewed in the following, because the terms argille scagliose and olistostrome have been widely applied to or compared with disrupted rocks in other orogenic belts (see Hsü, 1965; Hoedemaeker, 1973; Maxwell, 1974; Page, 1978; Aalto, 1981, 1982; Hisada, 1983; Raymond, 1984, and references cited therein; Cowan, 1985, and references cited therein). The term argille scagliose was first introduced by Bianconi (1840) to describe a peculiar characteristic of the matrix of cer-

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy tain rocks in the Bologna Apennines. These rocks can be progressively split into small, smooth and lustrous or polished scales. Argille scagliose literally means scaly clays in Italian: this term was only descriptive and was applied at the centimeter or decimeter scale. With time, however, the scale at which it was applied expanded from centimeter to outcrop and map scale. Accordingly, the term has been extended to include the entire rock assemblage: the matrix, fragments of single beds, and bed packages. With this expansion in meaning, the term was also extended geographically to similar rock assemblages in the Ligurian and Subligurian units throughout the Apennine chain (e.g., Scarabelli, 1848; Bombicci, 1882; Bonarelli, 1902; Anelli, 1927, 1938; Migliorini, 1933, 1948; Ogniben, 1953, 1956; Beneo, 1955, 1956a, 1956b; Rigo de Righi, 1956; Ruggieri, 1956, 1958; Signorini, 1956; Trevisan, 1956). The term was extended further by Merla (1952). His use of Argille Scagliose implied that all the large masses—from hundreds of meters to several tens of kilometers—of nondisrupted, well-bedded sequences, like the Ligurian flysches and the basal complexes, and the large masses of ophiolites, were huge slabs included in the chaotic disrupted units, which serve as a matrix at regional scale (see also Page, 1963; Hsü, 1967). In this view, only the Epiligurian successions—the Loiano succession, in the original Merla definition—are not considered to be argille scagliose, because they are recognized as less allochthonous, having been deposited on the just-deformed Ligurian units. This extreme extension of the term was soon abandoned by Merla (1956), who proposed that (1) the term argille scagliose should only refer to the chaotic assemblages of blocks in clayey matrix, and (2) the entire Ligurian nappe, including the argille scagliose and the coherent units, should be named “terreni alloctoni,” which may be translated as allochthonous masses or assemblages. In any case, even ignoring the extension of the term, it passed from having an originally descriptive meaning to indicating a certain genesis. The argille scagliose came to be considered as the product of submarine landslides (Beneo, 1956b; Ruggieri, 1956; Trevisan, 1956; Gorler and Reutter, 1968), the product of pseudodiapirism due to tangential tectonics related to the gravitational emplacement of a nappe (Ogniben, 1953), or the result of stratal disruption of the mechanically weak part of a gravitational nappe during gravity sliding (Merla, 1952; Signorini, 1956; Maxwell, 1959; Page, 1963; Hsü, 1967; Abbate and Sagri, 1970). Moreover, regardless of the terms used or the local mechanism invoked to explain the stratal disruption and chaotic fabrics, the disrupted rocks came to be almost unanimously interpreted as having resulted from the gravitational emplacement of the Ligurian nappe. The history of argille scagliose intersects that of another well-used and often abused term, olistostrome (Flores, 1955, 1956). Olistostromes are bodies characterized by a chaotic assemblage of blocks in a mudstone-rich matrix. The term was originally applied to bodies in the Tertiary succession of Sicily (Flores, 1955, 1956). The bodies were distinguished at the beginning from the argille scagliose because of the different block in matrix fabric: the matrix consists of clays that include millimeter

9

to centimeter argillaceous clasts that have angular shapes (argille brecciate, or brecciated clays) or rounded shape (argille puddingoidi, or clayey puddingstone) (Beneo, 1955, 1956a; Rigo de Righi, 1956). Beneo and Rigo de Righi considered the fabric to reflect the different structural and stratigraphic position of the rocks: argille brecciate and argille puddingoidi are prominent characteristics of bodies included in the Miocene-Pliocene normal-bedded sediments (the olistostromes); the argille scagliose are typical of the largest allochthonous masses constituting the substratum of the Miocene-Pliocene successions (Ligurian and Subligurian gravitational nappes). All three types, however, were considered a common component of the olistostromes (Beneo, 1955, 1956a, 1956b; Flores, 1955, 1956). A more genetic distinction was proposed by Ogniben (1953) in Sicily: he suggested that a block in brecciated matrix texture may be linked to submarine mass transport (i.e., olistostromes are submarine muddy debris flows); a block in scaly matrix character may due to in situ deformations of sequences, linked to pseudo-diapirism of the argille scagliose during orogenic sliding. The term olistostrome has been extended to the bodies of argille scagliose that are included inside normal-bedded deposits in the Tuscan and Umbrian successions of the Northern Apennines (Abbate et al., 1970a, 1981a; Elter and Trevisan, 1973; Ricci Lucchi, 1975; Labaume, 1992). The stratigraphic positions of some of those olistostromes are shown in Figure 3. They are considered to be bodies deposited by debris flows and submarine slides derived from the front of the Ligurian nappe (Abbate et al., 1970a; Elter and Trevisan, 1973; Ricci Lucchi, 1975, 1986; Labaume, 1992; Berti et al., 1994). Similar sedimentary bodies have also been recognized inside the nondisrupted, coherent units of the Ligurian nappe (e.g., Abbate et al., 1970a; Naylor, 1981, 1982). Moreover, some well-confined sedimentary bodies (olistostromes) and blocks (olistoliths) of argille scagliose were described by Selli (1943-44), Fazzini and Tacoli (1963), Papani (1963, 1971), and Elter and Trevisan (1973) as being inside the well-bedded Epiligurian successions in the Emilia Apennines. These olistostromes comprise mostly Ligurian or Subligurian rocks and have a prevailing block in matrix aspect. They include either single blocks or large slabs of bed packages. The matrix has a brecciated texture, and part of the matrix has a scaly fabric (Sestini, 1968; Elter and Trevisan, 1973; Abbate et al., 1981a). These features are considered typical of an origin from large submarine slides and debris flows. The fissility or scaly parting was considered by Abbate et al. (1981a) to be related to compaction and postcompaction tectonics. At the beginning of the 1960s, the nondisrupted Ligurian successions, such as flysches, part of the basal complexes, and the largest masses of ophiolites were interpreted as thrust sheets, characterized by internal folds and thrusts and bounded by tectonic contacts (Elter, 1960, 1973; Zanzucchi, 1972, 1978; Reutter and Groshurt, 1978). The entire Ligurian nappe was again considered as gravitational, but as a gravitational thrust sheet it was able to develop internal tectonic structures. At this stage, the terms Chaotic Complex and Undifferentiated Complex were sug-

10

G. A. Pini

gested as an alternative to argille scagliose (see Abbate and Sagri, 1970), because the latter was considered to have been applied too extensively. The term Chaotic Complex pointed out again the chaotic character of some of these rocks, whereas Undifferentiated Complex indicated that some of the argille scagliose “with the progress of stratigraphic and structural studies...will be split into various units” (Abbate and Sagri, 1970, p. 314). But, at the scale of regional mapping, the argille scagliose were again interpreted as a unique, chaotic assemblage (Chaotic Complex in Boccaletti and Coli, 1982). Some papers published since the second half of the 1980s confirm that the argille scagliose can be split into units of different nature and origin. Chaotic sedimentary bodies, derived from debris flows, and related to the Epiligurian successions have been widely recognized inside the argille scagliose in the ModenaParma Apennines by Bettelli and Panini (1985). These bodies, defined as sedimentary melanges, overlie other sedimentary or tectonic melanges, the internal organization and meaning of which were not investigated. The sedimentary melanges are characterized by a block in matrix texture and the matrix is of brecciated clays. The significant impact of the paper by Bettelli and Panini (1985) is that some parts of the previously undifferentiated argille scagliose of the Ligurian units were recognized as true sedimentary rocks, belonging to the Epiligurian successions, on the basis of their textural and compositional characteristics. Castellarin et al. (1986) also proposed to differentiate units inside the argille scagliose south and southwest of Bologna. This paper presented a radically different suggestion from a structural and regional point of view: that the argille scagliose, or parts of them, may be the product of tectonic deformation. In fact, the ordered distribution of internal features, the presence of tectonic lamination, and the constant lithologic content of some units, enabled the differentiation of several structural units inside the argille scagliose at the map scale. The macroscopic to mesoscopic characteristics were related to the penetrative development of scaly fabric in the shaly matrix, at mesoscopic to microscopic scales. Castellarin et al. (1986) suggested that these fabrics and structures derived from pervasive simple-shear deformation of mud-dominated Ligurian successions. Pini (1987a) recognized two fundamental kinds of textures and fabrics in the argille scagliose of the Bologna Apennines. On geologic maps, he subdivided the argille scagliose into deformed Ligurian formations, called tectonosomes in this paper, and olistostromes (Pini, 1987b). Very similar observations led to the same distinction in the Po Plain facing the Apennines south of Modena and Parma, and in the Romagna region (Bettelli and Panini, 1987; Conti, 1987). Starting from those papers, it became easier to distinguish and map deformed Ligurian formations, olistostromes, and normal Epiligurian deposits inside the argille scagliose (see Bettelli and Panini, 1989, 1992; Bettelli et al., 1989a, 1994a, 1994b, 1996b; Castellarin and Pini, 1989; De Nardo, 1992, 1994; Pini, 1993; Capitani, 1993; Conti, 1994; De Libero, 1994a, 1994b).

The reconstruction of the thrust systems that arranged the Ligurian units in the region led to the reinterpretation of the entire Ligurian nappe as thrust sheets that were active in contractional tectonics (Castellarin et al., 1985; Castellarin and Pini, 1989; Conti, 1987, 1994; Conti and Gelmini, 1994; Conti and Tosatti, 1996). Some workers, however, still interpret the nappe to be a gravitationally emplaced sheet and its internal structure as having resulted from gravitational gliding (De Feyter, 1992). CRITERIA FOR FIELD- TO MAP-SCALE DISTINCTION OF TECTONOSOMES AND OLISTOSTROMES IN THE ARGILLE SCAGLIOSE OF THE BOLOGNA APENNINES The argille scagliose are highly disrupted rocks, characterized by the lack of continuity of beds, and have a prevailing fabric of blocks dispersed in a matrix (block in matrix fabric). Blocks consist of marls, siltstones, sandstones, or limestones, and the matrix consists of shales. These rocks have been considered as the results of either chaotic mixing of, or extremely complicated interposition of, different stratigraphic units. Similar rocks are one of the component of melanges in orogenic belts and accretionary prisms (e.g., Maxwell, 1974; Connelly, 1978; Cowan, 1978, 1985; Aalto, 1981; Bachman, 1982; Pettinga, 1982; Raymond, 1984; Sarwar and De Jong, 1984; Yilmaz and Maxwell, 1984; Silberling et al., 1988; Brandon, 1989; Taira et al., 1992a; Orange et al., 1993; Orange and Underwood, 1995; Miller and Gray, 1996). Some of the major problems in the interpretation of these rocks are defining the effective degree and nature of the mixing, describing their textural aspects, and defining the contribution of tectonic versus sedimentary processes to the stratal disruption (e.g., see Bachman, 1982; Raymond, 1984; Underwood, 1984; Cowan, 1985; Orange and Underwood, 1995). In view of the fact that the argille scagliose have been frequently compared with melanges in other chains (see Hsü, 1965; Page, 1978; Raymond, 1984; Cowan, 1985), and that it is now possible to distinguish units of different nature and origin, the argille scagliose may be an interesting study case for the other highly disrupted units and rocks in other chains and accretionary prisms. The homogeneous and chaotic argille scagliose shown in Plate 1, as it was considered by earlier workers, can be subdivided into three kinds of mappable rock bodies: belts and bodies of strongly deformed Ligurian stratigraphic units, or tectonosomes; bodies of muddy debris-flow deposits, or olistostromes; and middle-upper Eocene to Oligocene intervals of normal pelagic sedimentation. This chapter describes the characteristics that enable the field and map distinction of tectonosomes and olistostromes as recognized in the Bologna Apennines. Figure 6 describes the basic characteristics of olistostromes and tectonosomes; Plate 2A shows the distribution of tectonosome and olistostrome bodies in the Bologna Apennines.

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy

OLISTOSTROMES

TECTONOSOMES

ORIGINAL STRATIGRAPHY MAPPABLE CHARACTERISTIC

UCE units (upper Pelagic shales with scarce thinbedded turbidites of various Campanian to lower/middle composition (siliciclastic, Eocene) carbonate and mixed)

LC units (Neocomian to Albian)

Pelagic shales with scarce interbedded carbonate and siliciclastic turbidites beds

Tens to hundreds of meters bodies, hosted by the LC units with tectonic contacts. Depending on the scale of mapping, the UCE may be single deformed stratigraphic units or part of a tectono-stratigraphic unit with LC units

MESOSCOPIC CHARACTERISTIC

High structural order by parallel orientation of blocks and matrix features (i.e., pseudobedding) and bedding of boudinaged bed packages

Kilometer belts, separated by tectonic contacts, often including some UCE bodies. Depending on the scale of mapping, the LC may be single deformed stratigraphic units or part of a tectono-stratigraphic unit

Indeterminable

Sedimentary chaotic bodies, often interbedded with normal sediments Chaotic block in matrix including decameter slabs characterized by a proper fabric, such as slabs of tectonosomes and poorly deformed stratigraphic units

Type B

MICROSCOPIC TO MESOSCOPIC CHARACTERISTIC

Normal to poorly consolidated clays, with only scarce scaly fabric

Highly consolidated shales with pervasive and penetrative scaly fabric.

Chaotic block in matrix (blocks from centimeters to 1 m)

Type A

11

Normal to poorly consolidated clays including millimeter clasts (brecciated texture). Fluidal features and open honeycomb texture of clays.

Figure 6. Comparative chart of the descriptive characteristics differentiating tectonosomes and olistostromes. LC—Lower Cretaceous.

Mesoscopic characteristics The first step in mapping and distinguishing the bodies is observing outcrops and hand samples in the field. At this scale, tectonosomes and olistostromes are both characterized by an apparently prevailing block in matrix texture, but the block shapes, the distribution of the blocks, and the aspect of the matrix enable the two type of bodies to be distinguished. Tectonosomes. The first difference evident at outcrop scale is an order versus disorder contrast: tectonosomes are ordered structurally, whereas olistostromes are mostly characterized by a chaotic distribution of textural features. The fabric of tectonosomes has the following attributes. 1. Competent beds—limestones, sandstones, siltstones, and marlstones—are fragmented into blocks, which have definite shapes, sharp outlines, and an evident triaxial geometry where the plane of major extension coincides with the original bedding plane (Figs. 7 and 8, A and B). 2. The blocks show a common trend and a strong parallel preferred orientation of the bedding plane (Figs. 7 and 8, A and B, and 9).

3. The shales are subdivided into well-defined, elongate bodies of different color (Fig. 8, C and D) that are disposed parallel to the blocks (Fig. 7B). The fabric components are arranged between two end members: boudinaged bed package and blocks in a shaly matrix (Figs. 7 and 9). In boudinaged bed packages the continuity of a single competent bed is still recognizable for meters. The different shaly beds, highlighted by color contrasts, are flattened and elongated in a pinch and swell style. The original bedding is preserved, and this overall body may be described as a boudinaged sequence. The structural order is defined by fragments and boudins of what were obviously originally continuous beds. In blocks in shaly matrix the blocks have a strictly common orientation of longer axes, which coincide with the bedding plane, but the original bedding is completely disrupted. Blocks from the same bed are not continuous and are dispersed in a shaly matrix. The matrix is subdivided into thin and very elongated (elongation/thickness ratio 50/1) bodies, highlighted by different colors. These elongated bodies, evident in Figure 8C, are here defined as bands, following Castellarin and Pini (1989). Matrix bands and the blocks both have a strong com-

12

G. A. Pini

mon planar orientation that simulates a bedding, here defined as pseudobedding. The orientation of blocks and matrix bands in the blocks in shaly matrix end member and the orientation of bedding in the boudinaged bed packages are strictly parallel at the outcrop scale. True bedding makes a gradual transition to pseudobedding, even in the same outcrop (Figs.7B and 9). Moreover, the blocks in shaly matrix fabric have the same composition and almost the same shape as the bed fragments in boudinaged bed packages, and the matrix bands in the blocks in shaly matrix derive from the shaly beds in boudinaged bed packages (Fig. 7B). The common structural order and origin of these end members suggest that the tectonosomes are

composed of the strongly deformed (complete stratal disruption) and weakly deformed (merely boudinaged) parts of the same stratigraphic units. Where complete stratal disruption prevails, the general aspect of these rocks is that of blocks in matrix. Olistostromes. Disorder is the major impressive characteristic of all the olistostromes (Fig. 10, A and B). 1. Blocks of several different lithotypes—limestone, siltstone, sandstone, marlstone, shale, or clays—are randomly present in the matrix. 2. Blocks have either definite shapes and sharp outlines, or, in the case of marlstone, shale, and clay blocks, irregular shapes and locally diffuse outlines (Fig. 10B).

Figure 7. LC1 tectonosome (valley of the Idice River). General view (A) and close up (B), showing the spatial relationships between boudinaged bed packages (BBP) and blocks in shaly matrix (BSM). B: The typical angular and prismatic shape of the blocks of thick calcareous and thin siliciclastic turbidites. The shaly beds are highly boudinaged and pass to the matrix bands of the BSM fabric. The limestone beds have a basal laminated silty interval, and most of the beds are composed of massive micritic limestones. The complicated calcite-filled vein framework affecting the limestone beds is only schematic. Similar veins affect also the fine sandstone, and the siltstone beds. (Modified from Amorosi et al. (1992).

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy

13

Figure 8. Photographs of tectonosomes. The Neocomian LC1 tectonosomes (A) are characterized by thick fine-grained calcareous turbidites (meter blocks) and thin siliciclastic turbidites (white arrow). See Fig. 7b for further details. The Aptian-Albian LC2 tectonosomes (B) are characterized by the presence of fine grained (fine sandstone to siltstone) siliciclastic turbidite beds. (Hammer length = ~30 cm.) A line drawing B is in Figure 27A. Varicolored bands in the shaly matrix are particularly evident in the LC2a tectonosomes (C) and derive from the boudinage and stacking of shaly beds. The Paleocene to lower–middle Eocene UCE2 tectonosomes (D) are characterized by thick beds of red, gray and whitish clays and marly clays, with some thin levels of whitish fine sandstone (arrow) and of “manganesifere” siltstone (see also Figs. 19 and 20). Location of A,B and D in Plate 2B. Location of C is between the Reno and Samoggia Rivers (Plate 1).

3. Blocks tend to have an isodiametric shape, instead of a triaxial geometry. 4. Blocks are randomly distributed in the shaly matrix; only a weak orientation of roughly triaxial blocks is locally observable where it is developed parallel to the basal contact of olistostromes. 5. At the scale of the outcrop, the shaly matrix is homogeneous (Fig. 10A), rather than characterized by the varicolored bodies as in the tectonosomes; only a weak banding (Fig. 10B) may have developed parallel to the basal contact. The dimensions of blocks range from small fragments, to a single bed, and to packages of beds, from a few centimeters to tens of meters (Figs. 10, A and B, and 11). The distribution of the dimension of blocks changes, also in the same olistostrome, and may be grouped into two types (Fig. 11).

1. Type A olistostromes display a block in matrix texture. Blocks are meter to centimeter sized fragments of single beds of limestones, sandstones, marlstones, and shales, floating in a matrix. 2. Type B olistostromes display large blocks (meters to tens of meters) of Ligurian units (Fig. 11), which have certain textures and fabric: boudinaged bed packages and blocks in shaly matrix in the blocks from the tectonosome units, and weakly deformed bed packages in less-deformed units. These large blocks may be in contact with a matrix that fills the interstices, and in this case the olistostromes are considered as clast supported. Alternatively, the blocks are supported by the matrix. In both cases, the matrix has a block in matrix fabric, similar to the fabric that characterizes the type A olistostromes. The bedding and pseudobedding in the largest blocks do not have a common trend, but are instead

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G. A. Pini

Figure 9. Schematic illustration of an outcrop of LC2 tectonosome in the Zena Valley. A: Drawing showing the block dispersal pattern. B: The key for distribution of the two different mesoscopic fabrics in the outcrop: boudinaged bed packages and blocks in shaly matrix.

randomly oriented. Some of these type B olistostromes are composed of large blocks of tectonosomes, in places clast supported, and the clasts are so abundant that only very careful observation enables one to identify the single blocks and to distinguish these olistostromes from the tectonosomes. Microscopic to mesoscopic characteristics The other characterizing feature is the fabric or texture of the matrix at a mesoscopic to microscopic scale. The thick weathered carapaces and the cover of recent mudflows tend to homogenize the features of the tectonosomes and olistostromes and lead to confusion at this scale of observation. Where the weathered surface materials have been removed, more accurate observations result in the following characterizations. The fundamental distinction is between the brecciated matrix of olistostromes and the scaly matrix of tectonosomes (Fig. 12). The typical scaly fabric of the tectonosomes is assumed by the penetrative and pervasive development of curved and anastomosing cleavage planes (scaly cleavage), which subdivide the rocks in millimeter- and centimeter-sized splinters (Figs. 12A and 13A). The brecciated matrix of olistostromes is composed of clays containing millimeter to submillimeter sized clasts (Figs. 12, B and C, and 13B). (A complete description of the different kinds of matrix is presented later.) The distribution of scales is as important as the presence of the scaly fabric. For example, several olistostromes have been tectonized and assume a centimeter to decimeter scaly fabric,

which is usually less pervasive than the one in tectonosomes. This is not a general rule, however, because the younger tectonosomes may not develop a scaly fabric. I consider the brecciated texture to be the major discriminating characteristic: its generalized presence is indicative of olistostromes, and its absence is indicative of tectonosomes, or more simply of horizons of normal sedimentation. The presence of brecciated texture also enables the recognition of strongly deformed olistostromes, because the smaller clasts are always evident despite a pervasive development of the scaly fabric (Fig. 12B), in agreement with G. Bettelli and F. Panini (1989; and 1986, personal commun.). These criteria must be applied with care. The olistostromes commonly include blocks of tectonites as large sedimentary clasts (olistoliths), which show all the characters of the tectonosomes, as well as the lack of brecciated texture. Mappable characteristics, original stratigraphy, present stratigraphy A fundamental distinction between tectonosomes and olistostromes is evident from their map characteristics and the regional distributions of ages and lithotypes. Tectonosomes. The tectonosomes may be subdivided into several typical rock units, each of which is defined on the basis of its overall character and age. These units crop out as distinct, mappable bodies and belts (Plate 2A). Each is in tectonic contact with other tectonosomes or with olistostromes, and may be considered as an individual structural unit, whether it is a large, map

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy scale unit, or a small, meter- to decimeter-scale body (Plate 2A and Fig. 14). Each single structural unit is of the same age throughout. Furthermore, belts and bodies of the same age everywhere display the same kind of rock association. Thus, even if each unit displays different degrees of stratal disruption, it does not include exotic blocks that were foreign to the parent stratigraphic unit. Each single structural units derives from the deformation of a single rock-stratigraphic unit that was deformed, partly disrupted, and organized in bodies and belts, without losing its stratigraphic and chronological integrity. In Figure 15, the label tectonosomes indicates that the informal rock-stratigraphic units LC1, LC1a, LC2, LC2a, UCE1, and UCE2 were strongly and heterogenously deformed, but never lost their lithologic and chronological coherence. The term LC indicates Lower Creta-

15

ceous rock-stratigraphic units and the tectonosomes derived from them, whereas UCE indicates Upper Cretaceous–lower Eocene units and tectonosomes. The LC tectonosomes cover large areas (Plate 2A) below the Epiligurian deposits. These tectonosomes crop out as continuous, widespread units, or as single fault-bounded, laterally continuous and elongated, narrow belts with pinched-out edges (Plate 2A and Fig. 14). The continuous units extend to tens of square kilometers in area, whereas each wedge extends to some kilometers in length and is hundreds of meters (to 1km) wide. The UCE tectonosomes everywhere crop out as isolated lensshaped bodies (Castellarin et al., 1986) along the contact between different belts of LC or inside a single LC unit (Plate 2A and Fig. 14). The length of the bodies ranges from hundreds of meters to few meters; the thickness changes from tens of meters to meters.

Figure 11. Photograph of an Oligocene(?)–lower Miocene(?) olistostrome (EO1-2, see text). The top contact with the Bismantova Formation (B) is marked. The olistostrome lies on the LC2 tectonosome, labeled as LC2, with a sedimentary contact. The olistostrome displays two different facies: the type A is characterized by a block in matrix texture, with blocks that not exceed 2 m in length. Differently colored matrix constitutes some elongated, darker bodies. The Type B is characterized by large, from meters to tens of meters, olistoliths of the LC2 tectonosomes (the darkest rocks), partly separated by a block in matrix texture (the lightest masses). The LC2 tectonosome is characterized by meter-scale bodies of white marlstones. Figure 10. A: The block in matrix texture, labeled as bm, including large LC2 tectonosome olistoliths, labeled as t, of an Oligocene(?)–lower Miocene(?) olistostrome (see EO1-2 in text, and Plate 2A and Fig. 15). The top contacts with Bismantova Formation (B) and Pliocene IntraApenninic deposits (P) are marked. B: The distribution of the blocks in the matrix of an Oligocene(?)–lower Miocene(?) olistostrome (EO1-2). The black arrow indicates a marlstone block from the marlstones of the LC2 tectonosome, whose border has been fragmented and dispersed in the matrix.

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Figure 12. The schematic block diagrams, adapted from Pini (1987a), summarize the most evident differences between the matrix in tectonosomes (A) and olistostromes (B, C). The tectonosome matrix (A) is generally characterized by a pervasive and penetrative scaly fabric, due to the superposition of a centimetric scaly cleavage on a millimeter scaly cleavage. A brecciated texture (C) is everywhere evident in olistostromes, even if strongly deformed by development of a pervasive centimetric scaly fabric due to postdepositional deformation (B).

Figure 13. Examples of tectonosome (A) and olistostrome (B) matrix.

Even the largest bodies, however, crop out as isolated masses, rather than as continuous bodies as in the case of the LC belts. Olistostromes. In contrast, the olistostromes display a mixture of rocks of notably different ages and provenance at every scale, from microscopic to map scale. Thus, olistostromes are genetically related to the mixing of rocks with diverse origins. It is not possible to reconstruct from the components a single, original stratigraphic unit, and even the attribution of these components to a single parent succession is generally impossible. The widely distributed olistostromes are well-defined sedimentary bodies with top and bottom sedimentary contacts (Plate 2A and Figs. 10 and 11). These olistostromes are associated with the Epiligurian deposits, and are here called Epiligurian olistostromes, or EO in Plate 2A and Figure 15, according to the successions in which they are interbedded. Their association with normal sediments enables the age of the deposition of the olistostromes to be dated (Plate 2A and Figure 15). The ages of the events of olistostrome emplacement are discussed in the following. Some less-widely distributed olistostromes are related to the Ligurian units and are here defined as Ligurian olistostromes, or LO in Plate 2A and Figure 15. The nature of the contacts is the only characteristic that differentiates the Ligurian olistostromes

from the Epiligurian olistostromes. Ligurian olistostromes are everywhere in tectonic contact with the tectonosomes; no preserved sedimentary contacts have been observed. Etymology and usage of the terms Tectonosomes. Tectonosomes are here defined as the block in matrix rocks deriving from an original in situ deformation of Ligurian stratigraphic units. The term tectonosome is here utilized instead of a term such as broken formation, disrupted formation, or dismembered sequence (Raymond, 1984), to emphasize that: (1) tectonosomes consist of both merely boudinaged and completely disrupted parts of the same stratigraphic unit, and both types are typically present in the same outcrop and grade into one another; (2) the high structural order evident at the outcrop to map scale coincides with a structural fabric; (3) the deformation structures, which are responsible for the development of the fabric, are easily recognizable from outcrop to microscopic scale; and (4) it is generally possible to reconstruct the history of deformation of these rocks. The term tectonite, which I utilized in other works to indicate these units, has been abandoned here. The term tectonite typically

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy

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Figure 14. Map and outcrop relationships between the tectonosomes in a system of tectonic wedges (Idice-Quaderna tectonic wedge system, Fig. 16B. The geologic map (A) emphasizes the distribution of the tectonosomes in belts and bodies. LC tectonosomes crop out as elongated belts, separated by tectonic contacts. Some bodies of LC and UCE are hosted by LC belts. The northeastward panoramic view (B) shows the contacts and disposition of some LC tectonosome belts that steeply dip toward the northnortheast. The location of the view point is in Fig. 14a.

refers to a rock (Turner and Weiss, 1963; Sander, 1970; Raymond, 1975) “whose fabric reflects the history of its deformation” (Bates and Jackson, 1980, p. 641). The term is also used to indicate a discrete zone of concentrated deformation, such as a mylonitic zone. I need here a term to generically indicate a rock body, not simply a rock, which may crop out for large distances of several kilometers. Furthermore, the bodies of deformed units contain rocks that, because of their complete disruption and fabric, may be considered as tectonites, but these rocks do not constitute an entire tectonosome. In addition, the term tectonite is commonly used for metamorphic rocks, in which the fabric was produced by crystalplastic flow in the solid state (see Turner and Weiss, 1963). The term has been extended to low-temperature metamorphic rocks showing a fracture or scaly cleavage (see the S-F tectonites of Raymond, 1975) or to rocks with some discrete planes of discontinuity (S-C tectonites, Platt and Vissers, 1980). On the contrary, the rocks I am describing are characterized only by shallow-level deformation related to fractures, cleavage, and veins, or to independent particulate flow of nonconsolidated sediments.

The term tectonosome seems to describe better the deformed Ligurian stratigraphic units, which are the subject of this paper. The term somai means body in Greek. The Greek term tektonikos is an adjective related to the activity of carpenters and architects (Mercier and Vergely, 1992), and implies the concept of internal organization and of being well built. This seems to me to point out the strong structural order that is typical of tectonosomes. Tectonosomes will designate, in general, the entire system of structural units and bodies of the Bologna Apennines that derive from deformed Ligurian stratigraphic units. Tectonosome will be utilized here to indicate the single structural unit and body, of any size, in an outcrop or on a geologic map. The definition of tectonosomes is complicated by the scale of mapping: a single belt or a widespread unit of the LC tectonosomes may include bodies of the UCE tectonosome units (Plate 2A and Fig. 14). This situation has also been observed in the LC group of units: some small lens-shaped bodies are hosted by the rocks of other LC units (see Fig. 14A). Thus, the definition of the the term depends on the scale of mapping (cf. Plate 2A and Fig. 14).

Figure 15. Schematic stratigraphic diagram of the Bologna Apennines. The sources of the stratigraphic data are Pini (1987b, 1993), and data of Negri, Pini, and Borsetti. (E, M, L—Early, Middle, Late.)

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Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy 1. A tectonosome corresponds to a single deformed lithostratigraphic unit, if no guest bodies of other units are present (e.g., most of the LC1 and LC1a units) or if the clayey bodies are so large that they can be mapped at the scale in question. 2. A tectonosome is instead a tectono-stratigraphic unit if guest bodies of other units have been recognized, but cannot be mapped because of their small dimensions compared to the map scale. At a general map scale (1:10,000 to 1:50,000; see Pini, 1987b, 1993) (Plate 2A), each LC belt has to be generally defined as a tectono-stratigraphic unit, because the smaller guest bodies of unit UCE and of other LC units are too small to be mapped. Similarly, the mappable UCE bodies are commonly single, deformed stratigraphic units; some of the largest bodies that include both the UCE1 and UCE2 units may be defined as tectono-stratigraphic units (see later text examples of UCE bodies). Olistostromes. The term olistostrome has been used here, following the accepted definition of Abbate et al. (1981a), mostly because of the priority of the term, its worldwide usage, and the strictly genetic and morphological connections with the olistostromes of the Tuscan and Romagna-Umbria successions. The term derives from the Greek words stroma and olistomai that mean accumulation and to slide, respectively (Flores, 1955, 1956), and implies the concept of sediments having accumulated by sliding. In this paper, the spelling of the term olistostrome, is used instead of olistostrome, suggested by Hoedemaeker (1973), because of the priority (Flores, 1956) and its wider usage. The terms type A and type B describe two different sedimentary facies of the olistostromes, related to the distribution of the dimensions of the clasts. The genetic implications and the meaning of this distinction are discussed later. Considering both type A and type B olistostromes, the dimensions of blocks range from less than 1 mm to tens of meters. The presence of such a variety of clasts and block dimensions implies some difficulties about how to define a dimensional limit between matrix and blocks, independent of the scale of observations. I propose the following textural distinction based on the internal texture or fabric of blocks. All the blocks that contain bed packages are named slabs, as usual in the Italian literature (see Abbate et al., 1981a), or olistoliths, as defined by Flores (1955, 1956) and commonly used in the international literature. Each olistolith can be characterized, depending on its internal fabric, as constituting of either blocks in shaly matrix, boudinaged bed packages, or nondeformed bed packages. The pieces of a single bed are named block. The dimension of 1 mm can be assumed as the limit between the blocks and the clasts belonging to the matrix; the latter are here defined as microclasts, following Swarbick and Naylor (1980) and Abbate et al. (1981a). “Clast” is here utilized to indicate all the lithic objects in the homogeneous clays, including the slabs and olistoliths, the blocks, and the microclasts. Matrix here includes the homogeneous shales and the microclasts that have dimensions of less than 1 mm.

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Because of the previous indiscriminate usage of the term olistostrome, and the presence of different textures, some specific limitations are to be imposed. I suggest the term olistostromes for all the disrupted units that are sedimentary bodies, following Castellarin and Pini (1989), with these restricted meanings: (1) olistostrome refers to an entire sedimentary body, regardless of its fabric (type A and type B); (2) block in matrix is reserved for the texture that characterizes all the type A olistostromes and surrounds the slabs in the type B bodies; (3) the matrix may be termed brecciated matrix, as defined by Abbate et al. (1981a) and in general agreement with the previous authors (see argille brecciate, or brecciated clays, in Ogniben, 1953; Beneo, 1955, 1956a, 1956b; Rigo de Righi, 1956). Because of the peculiar characteristics of the matrix, I propose to define this texture of the olistostromes as block in brecciated matrix texture. STRATIGRAPHY AND SETTING OF TECTONOSOMES, OLISTOSTROMES, AND RELATED EPILIGURIAN DEPOSITS IN THE BOLOGNA APENNINES After the argille scagliose are subdivided into tectonosomes and olistostromes, the distribution of the main structural units in the Bologna Apennines appears completely changed. This change is evident on geologic maps when the argille scagliose complex in Plate 1 is compared with the Ligurian structural units in Plate 2A. The areal distribution of Ligurian units, which include both tectonosomes and strongly deformed olistostromes, is greatly reduced, whereas the Epiligurian successions have been extended to include a great part of what was previously considered to be argille scagliose. After subdivision, the Epiligurian contains numerous olistostromes, associated with thin intervals of normal, hemipelagic deposits and thin-bedded turbidites of the Monte Piano and Antognola Formations. The stratigraphic relationships among the different units are summarized in Figure 15. Stratigraphy of tectonosomes and Ligurian olistostromes A brief description of the lithotypes is reported here, because the changes in composition strongly influence the tectonic behavior of the rock assemblages. Therefore, some of the peculiar characteristics of the Apennine tectonosomes, compared to those of disrupted units and rocks of other orogenic chains, might be due to compositional differences. LC units of tectonosomes. Unit LC1 is Neocomian. The age of unit LC1a is Aptian-Albian. They are composed of 30–100 cmthick limestone beds and siliciclastic thin-bedded turbidites included in prevailing gray shales (unit LC1, Figs. 7, A and B, and 8A), or in varicolored shales (LC1a). Limestones are fine-grained, calcareous turbidites, mostly composed of pelagic mud, with a very thin interval of fine, laminated calcareous or hybrid sandstone at the base. These ages have been revealed by the concordance among several dates of blocks of limestones and sandstones

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(foraminifers), and some dates from the matrix (nannoplankton dating by A. Negri, 1994–1995, personal communs.). Units LC2 and LC2a are mostly composed of shales characterized by bands and ribbons of different colors: black, dark gray, green, gray, red, light red (Fig. 8C). The shales host 5-10-cmthick siltstone and fine sandstone turbidite beds (thin-bedded turbidites) (Fig. 8B). The LC2 unit is characterized by 70–200-cm-thick beds of greenish-white marlstones (Fig. 11), which are sometimes recognizable as turbidites mostly composed of hybrid calcareous mud with a fine-sandstone interval. Unit LC2a includes common 2–5-cm-thick dark brown (manganesifere, following the traditional definition) siltstones and scarce 20–30-cm-thick) calcarenite beds. Unit LC2 may be referred to Aptian-Albian on the basis of several dates from limestone and marlstone, as well as dating of shales. The other varicolored shales (LC2a) have an uncertain age because only a few blocks of calcarenite and limestone have been dated by foraminifers. They have Aptian-Albian ages. This age has been confirmed by more dating of nannoplankton associations in the matrix (A. Negri dating, 1994–1995, personal communs.). The original stratigraphic relationships between these units are no longer observable. They presumably constitute part of a single stratigraphic succession, characterized by the progressive upward decrease of the calcareous turbidite input and contemporaneous increase of the fine-grained, siliciclastic component. UCE unit of tectonosomes. The associations UCE1 and UCE2 have late Campanian to Maastrichtian and Paleocene to early–middle Eocene ages, respectively. The typical aspects of unit UCE1 are dark gray shales alternating with gray clayey sands and turbidite quartzarenites, and they contain mixed calcareous-siliciclastic turbidite beds. Mixed turbidites (Fontana et al., 1994) are characterized by a siliciclastic coarser interval followed by a calcareous lutitic interval. The siliciclastic interval typically shows parallel laminations. Unit UCE2 is characterized by light green and/or gray clays and light red silty clays, which include thin layers of black-brown (manganesifere) turbiditic fine sandstones and siltstones and bodies of white marly clays (Fig. 8D). The age of the rocks in UCE2 is still an open problem, because of their presence inside deformed belts and their typical sterility. Only ~20% of the outcrops have been dated (A. M. Borsetti, 1984–1987, and A. Negri, 1994–1995, personal communs.), so that, in spite of the coherence of dates and compositions, a common age of these bodies cannot be completely confirmed. Tectonosomes as a tectono-stratigraphic unit from a single Ligurian sequence. The lithostratigraphic components of the tectonosomes may correlate at a regional scale with lithostratigraphic units known to be parts of other, nondisrupted Ligurian successions. For example, units LC1 and LC1a show some lithologic relationships with the Palombini Shales of the Baganza and Sambro Supergroups (Fig. 5); units LC2 and LC2a may be correlated with the Varicolored Shales of Baganza Supergroup or to the Variegate Shales of the Sambro Supergroup. Units UCE1 and UCE2 have some similarity to the varicolored shales of the basal

complex of the Monte Morello flysch (Bettelli and Panini, 1992). The argille scagliose of the Apennines west of Bologna toward Parma and in the Ligurian nappe in Romagna (Fig. 2) have been interpreted as consisting of either Epiligurian olistostromes, tectonic melanges, or variably disrupted Ligurian units, derived from “basal complexes” (Bettelli and Panini, 1989, 1992; Conti, 1987, 1994; Bettelli et al., 1994b; De Nardo, 1994). The basal complexes I and II have been correlated with the basal complex of the Baganza Supergroup (Fig. 5), and the basal complex III has been considered as a continuous sequence extending from the Lower Cretaceous to the lower Eocene (Sillaro varicolored shales, Lavino Supergroup) (see Bettelli and Panini, 1992; De Nardo, 1994; Bettelli et al., 1994b, 1996a). These latter units represent the direct, physical continuation of the tectonosomes of the Bologna Apennines. Ligurian olistostromes. These olistostromes are arranged, together with the tectonosomes, in systems of tectonic wedges that are unconformably covered by Epiligurian deposits of different ages. Thus, the age of the Ligurian olistostromes ranges from the youngest age of included blocks, Turonian, to the oldest age of the Epiligurian cover, the middle–late Eocene. Stratigraphy of the Epiligurian deposits Epiligurian deposits extend to the late Pliocene stage. The age of the base of these successions, the stratigraphic base of the Monte Piano Formation, is still discussed. Papani (1971), and more recently Cerrina Feroni et al. (1990, 1992) proposed a late Eocene age; instead, Bettelli et al. (1989b) and De Nardo et al. (1992) gave an age of the latest part of the middle Eocene. New data from the Bologna Apennines (A. Negri, 1995, personal commun.) indicate a late Eocene age, but a latest middle Eocene age cannot be excluded. The stratigraphic scheme of the Epiligurian successions in the Bologna Apennines is shown in Figure 15. The Epiligurian deposits can be subdivided into different successions on the basis of the presence of olistostromes. 1. Olistostromes are small bodies interbedded in Epiligurian successions of normal pelagic deposition. Their dimensions are at most some hundreds of meters in length and to 20 m thick. 2. Thick and continuous elongated bodies of olistostromes constitute the largest part of the lower Epiligurian succession, below the Bismantova Formation. They are separated by thin, but continuous horizons of the Antognola Formation (Plate 2A). The dimensions of olistostromes are tens of meters in thickness and kilometers in length. 3. A single olistostrome, here defined as EO1-2, replaces all the Epiligurian succession below the Bismantova Formation, overlying thin and discontinuous levels of Monte Piano Formation and unconformably overlying the tectonosomes (Plate 2A). In this case, the olistostrome dimensions are several kilometers long and to 100 m thick. The olistostrome covers an area of tens of square kilometers, even where it is dislocated by continuous thrust fronts (Plate 2A).

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy Stratigraphy of the Epiligurian olistostromes. The sedimentary contacts at the base of the olistostromes and the association of the olistostromes with horizons of normal sedimentation enable the reconstruction of a stratigraphy of the bodies (Fig. 15). The bodies may be grouped in clusters at different stratigraphic levels. 1. The EO4 cluster corresponds to late Tortonian–early Messinian olistostromes along the Sillaro line, associated with preevaporitic marls. 2. The EO3 cluster of olistostromes is interbedded in the Serravallian–early Tortonian deposits of the Bismantova Formation. 3. The EO2 cluster groups together the olistostromes in the Antognola Formation (Upper Oligocene–lower Miocene). 4. EO1-2 indicates a single body, which overlies the tectonosomes and conformably covers discontinuous lenses of the Monte Piano Formation. It is apparently covered by late Oligocene to middle Miocene deposits (Antognola and Bismantova Formations), and typically includes bodies of the Monte Piano Formation. A possible age is late Oligocene to early Miocene. 5. The EO1 cluster includes the olistostromes in the Monte Piano Formation (middle Eocene–lower Oligocene). Normal stratigraphic units associated with the olistostromes. The normal stratigraphic units are hemipelagic and thin turbidite deposits associated with the largest Epiligurian olistostromes. These deposits, together with the tectonosomes and the olistostromes, were commonly tectonized by Neogene faults. Two different formations constitute these units. 1. The Monte Piano Formation displays different facies: light green and light red silty clays with intercalations of white quartz sands and thin layers of manganesifere siltstones; white coarse quartzarenite with interbedded dark gray clays; light gray siliciclastic turbidites and gray hemipelagites. The formation is present as lens-shaped bodies at the base of EO1-2 olistostrome and as blocks and slabs inside most of the Epiligurian olistostromes. In the first case, it directly overlies the tectonosomes along a welldefined unconformity; the formation locally includes small EO1 olistostromes (see the area around the Sillaro River, Plate 2A). This unit has a late Eocene–early Oligocene age (from both foraminifer and nannoplankton dating, performed by A. M. Borsetti, 1984–1987, and A. Negri, 1995, personal communs.), but an age of the latest part of the middle Eocene cannot yet be excluded. 2. The prevailing gray-green shales and marls of the Antognola Formation are interbedded as thin horizons among many bodies of olistostromes along the Sillaro line (Plate 2A and Fig. 15). The Antognola Formation is also present as small, discontinuous bodies in the EO1-2 olistostrome. Similar rocks of the same formations crop out in continuity with the other Epiligurian stratigraphic intervals, where the olistostromes are not so well developed (Plate 2A and Fig. 15). Structural setting of the Bologna Apennines The differentiation of the former argille scagliose into tectonosomes and olistostromes has enabled a better definition of

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the pattern of tectonic structures in the Bologna Apennines (Fig. 16, A and B). The Bologna Apennines can be subdivided into two distinct structural belts, which are characterized by a different development of regional-scale tectonic structures. External structural belt. The structural belt closer to the Po Plain displays elongated (2 to 10 km long) and thin (from hundreds of meters to 1 km thick) slices and wedges of tectonosomes, olistostromes, and Epiligurian deposits. These wedges, named here tectonic wedges (Fig. 16B), are bounded by tectonic contacts, mainly dipping northward at high angles and striking parallel to the Apennine–Po Plain border (Plate 2A and Figs. 14 and 16B). The tectonic wedges can be grouped into two distinct systems, the areal distribution of which is shown by Figure 16B. The Idice–Quaderna wedge system extends between the Idice and Quaderna Rivers. This wedge system includes tectonosomes and LO olistostromes only and it is overlain and closed by the middle–upper Eocene base of the local Epiligurian succession. The age of structural arrangement is necessarily older than latest middle Eocene. The Idice–Savena wedge system extends between the valleys of the Idice and Savena Rivers. The wedge system involves the EO1-2 olistostrome and the Monte Piano and Antognola Formations. Some tectonic contacts are unconformably covered by the upper Oligocene part of the Antognola Formation: the age of deformation can be older than late Oligocene, but younger than the top of the Monte Piano Formation (Oligocene).Therefore, these structures may be Oligocene in age. The rest of the wedges cut also the Antognola Formation, and the other Epiligurian deposits, from the Serravallian Bismantova Formation upward, unconformably overlie this system. This last group of structures could have developed during a Burdigalian tectonic phase. Internal structural belt. The other structural belt is south of the external structural belt. The systems of tectonic wedges are here substituted by continuous outcrops of Tertiary Epiligurian successions, with an areally large exposure of Tertiary olistostromes, EO1-2, and tectonosomes. The southwesternmost part of this structural belt coincides with the northward, and seaward, border of the Pliocene Intra-Apenninic basin (Ricci Lucchi et al., 1981), one of the major and younger Epiligurian basins, which is shown in Plate 1. This basin and the underlying older Epiligurian deposits separate the Bologna Apennines from the High Bologna Apennines, which are characterized by the presence of other, nondisrupted, Ligurian structural units, such as the Monghidoro unit (Plate 1). The internal and external structural belts are separated by one of the arcuate thrusts of the Sillaro arc (Castellarin and Pini, 1989) (Fig. 16A). This arcuate system of frontal, oblique, and lateral ramps has been related to the prevailing dextral, transpressive activity of the Sillaro line. Transpression corresponds to the emplacement of the Ligurian units at the Po Plain–Apennine border. Some backthrusts in the frontal area, close to the Po Plain (Plate 2A and Fig. 16A), are connected to this thrust system. The thrusts and backthrusts of the Sillaro arc deform tectonosomes, olistostromes, and all the Epiligurian deposits as young as Pleis-

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Figure 16. Structural sketch maps of the Bologna Apennines. A: Major thrust faults, normal faults, and strike-slip faults of the Sillaro arc tectonic system. Modified from Castellarin and Pini (1989). Slip vectors are averages computed from many individual measurements made in fault rocks. B: Location and age of the tectonic wedges of tectonosomes and olistostromes. Adapted from Pini (1993).

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy tocene. The thrusts were active during Serravallian-Tortonian (?), late Messinian, Pliocene, and early Pleistocene tectonic phases (Pieri and Groppi, 1981; Pieri, 1983; Castellarin et al., 1985; Castellarin and Pini, 1989). TECTONOSOMES The various units of tectonosomes include the LC units, which are present as belts and widespread units at regional scale, and the UCE units, which are present as bodies included in the LC units. The smallest bodies include only one rock association, whereas the largest bodies instead include both associations UCE1 and UCE2. The difference in areal distribution also reflects a different structural style: the shape and the distribution of the blocks and the charaters of the matrix are quite different in LC and UCE units. The common characteristics of all the tectonosomes are the strong internal order and the presence of different degrees of stratal disruption in the same outcrop, even if the prevailing aspect is block in matrix texture. Structural order is depicted by the common, planar orientation of blocks, bodies, and matrix features evident at every scale, as shown by Figures 7–9, 14, and 17. At a larger scale, the scale of an entire outcrop, the boudinage pseudobedding is in most cases parallel to the main contacts between the units (Figs. 14A and 17). This relationship defines a planar attitude that is parallel to the general setting of the tectonosome wedges and to the major tectonic structures (Fig. 14A). The strike of this planar attitude is generally northwest-southeast parallel to the Apennine border. It is progressively rotated to a northeast-southwest direction as it approaches the Sillaro Valley. Here, the effects of the lateral ramp system of the Sillaro arc are more evident (Pini, 1987b; Castellarin and Pini, 1989). The dip is either to the northeast or southwest (Figs. 7A and 14) and generally follows the dip of the main contact between the units. Local variability of the dip direction is related to largescale, tens of meters to 100 m, folds (Figs. 7A and 14), and to Neogene thrusts associated with the Sillaro arc system (Fig. 16A). These thrusts rearranged the planar attitude and forced it parallel to the thrust front (Pini and Castellarin, 1989). Stratal disruption in UCE units UCE2 Unit. These units are characterized by the prevailence of boudinaged bed packages and blocks in shaly matrix. The blocks in shaly matrix developed as subordinate, discontinuously present and narrow belts (tens of centimeters) along the contacts with the LC units. Boudinage occured by pinch and swell of the siltstone and fine sandstone beds (Figs. 18 and 19C). The deformation of the clays and marlstone beds gave roughly lens-shaped bodies with long and thinned tails (Figs. 8D, 18, 19, A and B, and 20). Fine sandstone and siltstone beds were deformed by pinch and swell without the development of any evident structures and discontinuities. The marlstones and clays formed boudinage and were folded by continuous deformation, without the development

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of any cleavages and fabric. Mesoscopically ductile here defines a deformation that, at the mesoscopic to macroscopic scale, does not produce mechanical discontinuities. A three-dimensional study of the boudinage style has been possible on the compact sandstones-siltstone beds only, because of their smaller dimensions (Fig. 21). The blocks have an irregular oblate ellipsoidal shape, with X = Y > Z strain ellipsoid axes, and the shape of boudins is always symmetric in cross section. Neither the asymmetrical lens shapes nor rotations of bodies, indicative of a simple shear component (Cowan, 1990; Kano et al., 1991), have been observed. This boudinage may imply an almost equal extension in all the directions of the X-Y plane, which coincides with the bedding plane, so that the extension is axially symmetric with respect to the normal to the bedding plane (axially symmetric extension, Cowan, 1982). This shape of blocks also occurs where the thinning rate of the sandy beds and the separation of boudins are greatest. In this case, the separation of boudins seems to develop preferentially in a single direction, ranging from east-northeast–west-southwest to northeast-southwest, instead of in all directions in the X-Y plane. Some of the boudin edges are deformed in elongated tails that developed in the same direction as the greatest separation of the smaller blocks (Figs. 19A and 20). The increase in deformation corresponds to the passage from axially symmetric extension—graphically represented by an imaginary, flattening-type (Ramsay, 1967) strain ellipsoid—to a certain component of monodirectional boudinage, which implies an extension in a single direction in the plane of bedding, represented by a constriction-type (Ramsay, 1967) strain ellipsoid (Fig. 21). The continuity of both clayey and sandy beds is limited to only a few meters (Figs. 8D and 17). This limit may be an original attribute of the sedimentary beds, enhanced by the layer-parallel, symmetric extension. Only in some places is the oblique truncation of the sequences evident in the clayey beds, inducing zones of concentrated displacement without the development of any discontinuities such as scaly fabric or any kind of cleavage (ductile shear zones in clays). The same ductile shear zones are responsible for the decimeter- to meter-scale interfingering and indentation of the different clayey beds that produced small (from 40 cm to about 1 m in length) bodies hosted in the other beds (Figs. 18, 19A, and 20). These guest bodies, as they are termed here, are not to be considered exotic, because their roots are still visible and developed inside the same unit. UCE1 Unit. The UCE1 tectonosome crops out in very few outcrops. The largest is associated with UCE2 and is shown in Figs. 22 and 23. In UCE1, boudinaged bed packages are still prevailing, but blocks in shaly matrix characterize the border zone with the LC and UCE2 units and are also present as narrow— from tens of centimeters to few meters—and discrete belts inside the bodies (Fig. 23). In the boudinaged bed packages, the boudinage is marked by pinching and swelling of the sandstone beds and is associated with isoclinal folding (Fig. 23). In three dimensions, the boudinage produced boudins with an irregular oblate ellipsoidal shape, with X = Y > Z axes (Fig. 24A). The separation

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Figure 17. Lens-shaped body of an UCE2 tectonosome (lower Eocene), in an LC2 tectonosome of Aptian-Albian age, just beneath the basal sedimentary contact of an EO1-2 olistostrome. The photograph (A) and the explanatory drawing (B) show the disposition of the western half of the UCE lens-shape body and its internal structures, in relation to the structures in LC unit.

of boudins, which seems to have developed in all directions in the plane of major extension (X-Y), suggests the idea of axially symmetric extension (see plane A on Fig. 24A). The isoclinal folds are symmetric and disharmonic and have a common orientation of axes and axial planes. Thickening and thinning of sandstone beds (see left parts of Fig. 23, A and B) are enhanced in both the hinges and limbs of the folds. Boudinage seems not to be related to folding, but rather to predate it. The mixed turbidites show a disharmonic boudinage: the micrite limestone intervals do not follow the strongly elongated,

pinch-and-swell shape of the siliciclastic sandy base (Fig. 24B). The shape of limestone blocks is everywhere almost prismatic, with rounded corners, or roughly ellipsoidal. The different shape of the boudins reflects a difference in deformation. The sandstone beds show a mesoscopically ductile deformation. The limestone blocks, associated with a mesoscopically ductile deformation, are brecciated and contain subangular to rounded fragments of limestone of different sizes dispersed in a matrix of the same limestones, which plastically accommodated the deformation without developing evident structures.

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy Along the passage from boudinaged bed packages to the belts of blocks in shaly matrix, the dispersion of sandstone blocks was caused by further boudinage of beds as well as by the increasing complexity in the shapes of the blocks (Fig. 24A, plane B). The increasing complexity of shapes is related to centimeter-decimeter polyclinal and disharmonic folds. In the zone of bed fragmentation (plane B on Fig. 24A), the boudins and blocks have a preferential direction of development in spite of their complex shapes. In three dimensions, the boudins range between an almost spherical shape and a threeaxes ellipsoidal shape. The X axes of the ellipsoidal boudins are all oriented almost east-west. Also, the separation of blocks seemed to have occurred preferentially in an east-west direction. This phase of boudinage and block separation can be related to a certain component of monodirectional boudinage along an east-west direction. The mesoscopically ductile, large-scale folding (Fig. 23, see fold axes in Fig. 24A) seems to be associated genetically with this phase of blocks in shaly matrix development; this interpretation is suggested by the parallelism between the axial planes and the borders of the blocks in shaly matrix.

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Mechanisms of mesoscopically ductile deformation. All these structures are characterized by ductile behavior of the rocks. In fine sandstones and siltstones and in the micritic limestones of the mixed turbidites, the mesoscopically ductile deformation is not related to intracrystalline deformation, such as crystal plasticity (Schmid, 1983; Knipe, 1986b, and references therein). Cataclastic flow related to microscale fragmentation of grains (Knipe, 1986b, and references therein) can also be excluded on the basis of direct observations in thin section and with the scanning electron microscope (SEM). Excluding crystal plasticity and cataclastic flow, the continuous deformation may instead be related to intergranular particulate flow in granular rocks such as sandstones and siltstones, due dominantly to grain-boundary sliding (independent particulate flow; see Borradaile, 1981; Maltman, 1984; Knipe, 1986b). This kind of deformation can be developed in nonconsolidated sediments, and it is aided by a high content and pressure of fluids and a low effective confining pressure (e.g., Maltman, 1984; Knipe, 1986b; Lucas and Moore, 1986). This kind of deformational mechanism may also occur in poorly consolidated sandstones, as proposed by Spörli (1989) and, generally, in partly lithified sediments (Knipe, 1986b). In this case it implies a continuous process of

Figure 18. Photograph and the explanatory line drawing illustrating a section of the body in Figure 17, showing boudinage inside UCE2 body and the interaction between LC and UCE along the borders of the body.

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Figure 19. Details of the UCE2 tectonosome body in Figure 17. A, B: Extremely extended boudins made from a bed of the “manganesifere” siltstones. These pictures are explained by the line drawings in Figure 20. C: A less-extended, lens-shaped boudin resulting from the boudinage of a “manganesifere” siltstone bed.

Figure 20. A, B: Illustration of the photographs of Figure 19 showing details of mesoscopic ductile style of boudinage of a “manganesifere” siltstone bed of the UCE2 tectonosome.

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Figure 21. Schematic diagrams of mesoscopically ductile deformation of UCE2.

Figure 22. A quarry showing the relationships between the UCE and LC tectonosomes. The location of the outcrop is in Figure 14B. B is modified from Amorosi et al. (1992).

breaking and crystallization of the cement—possibly associated with pressure solution—grain rotations and translations, and a certain component of intragranular deformation (grain breakage dependent particulate flow; see Maltman, 1984, 1994; Knipe, 1986b). No direct evidence of this kind of process has been found in thin section. The deformation of clays occurred without the devel-

opment of scaly fabric and therefore may indicate disagreggation and intergranular flow of the clay grains in a nonconsolidated, water-rich sediment (Knipe, 1986a). The rounded shape of the blocks and the dispersion of sand clasts inside the clays around the blocks (Fig. 19B) confirm the weak consolidation of all the materials, including sandstones and clayey rocks, during deformation.

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Figure 23. The floor of the quarry shown in Figure 22. The line drawing shows the contact between UCE1 and UCE2 tectonosomes and the internal structures of the two units. The mesoscopic fabric of the UCE1 consists of boudinaged bed packages and passes to blocks in shaly matrix.

Brittle-ductile brecciation in limestones indicates different degrees of cementation of limestones. The matrix derives from nonconsolidated carbonatic mud, whereas the clasts are presumably derived from consolidated part of the same bed (Labaume, 1987; De Libero, 1994b, 1993, personal commun.).

Stratal disruption in LC units The LC tectonosomes have both boudinaged bed packages (BBP) and blocks in shaly matrix; the former are prevalent and the bed packages are dispersed and included in the prevailing

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Figure 24. Schematic diagrams of (A) mesoscopically ductile deformation of UCE1 tectonosomes. B: The differential deformation of the siliciclastic and the calcareous intervals of a mixed turbidite bed of unit UCE1.

blocks in shaly matrix. The common planar attitude of pseudobedding and bedding is the sum of the interaction of different mesoscale to macroscale structures. Macroscale to mesoscale observations. Large-scale duplications of part of the sequence are evident where thick beds are a marker or where it is possible to recognize the polarity of bedding (Figs. 7B and 25). The duplications are evidenced by the overturning of a large section of the sequence, some tens of meters in thickness over some tens of meters in length (Fig. 25A). These structures may be reconstructed as an isoclinal fold, although there is no direct evidence for such a fold because of the lack of continuity of the beds and the lack of a hinge. A reconstruction of the folded(?) stratigraphic units using the thick beds shows that a great part of the sequence has been elided in the fold nucleus. The right-side-up and upside-down fragments of a single original thick bed are very close to one another, a few meters to centimeters, (Figs. 7B and 25A), and include strongly deformed blocks in shaly matrix, which may accommodate the elision of part of the sequence. There is no way to reconstruct the sense of vergence of the overturning, or the direction of a possible fold axis. Both the right-side-up and the upside-down limbs were affected by mesoscopic thrust faults (Figs. 25, A and B). These mesoscopic thrusts are evidenced by the repetitions of one, two, or three boudinaged beds, which are inferred from field and photo analysis. The thrust plane can be only inferred on the basis of the displacement of the boudinaged sequences, because the plane does not correspond to a true, discrete basal slip surface. The thrust movement seems to have been accommodated by a

2–3-m-thick zone of further deformation of boudinaged bed packages, or by a zone of blocks in shaly matrix. The LC units are also characterized by small, meter-scale folds, shown as mesoscopic folds in Figure 25b. These open to close, disharmonic folds directly deform the shaly beds and matrix bands and passively reorient the blocks at the scale of meters (Fig. 26B). The folds accomplished the overturning of parts of the sequence that are about 2–3 m in thick. Mesoscale observations. The main characteristic of the LC tectonosomes is the sharp, angular, and prismatic shape of the blocks (Figs. 7, 8, A and B, 26, and 27, A–C). Only the degree of dispersion, but not the shapes, of blocks changes from boudinaged bed packages to blocks in shaly matrix (Figs. 7B, 8, and 9). The origin of blocks is related to a fracture boudinage in the limestone and siltstone-sandstone beds. More sets of calcite-filled veins that are almost orthogonal to the bedding are present in the blocks (Figs. 7B and 27A). These sets commonly predetermined the surfaces along which the blocks separated, giving the classic sharp, angular, and prismatic shape. The largest—centimeters to tens of centimeters—veins typically border the blocks in boudinaged bed packages and blocks in shaly matrix, and are typically asymmetric, with oblique fibrous calcite filling (Figs. 26A and 27B). The same kinds of veins connect two blocks that were going to be separated in boudinaged bed packages (Figs. 26B and 27B). All these veins, in spite of their orientation and thickness, when observed in an optical microscope, do not include fragments of clays from the matrix and grains of the host rocks (clean veins; Byrne, 1994, or clear veins; Orange et al., 1993). More-

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Figure 25. Sketches of the observed and interpreted relationships between mesoscopic deformations and macroscopic structures in LC1 tectonosomes (A) and between mesoscopic deformations in unit LC2 (B). The different lithotype associations of the parent lithostratigraphic units force the contrasting, different scale of the interpretation in A and B. The calcareous turbidites of the LC1 tectonosomes are markers proportionate enough to reconstruct outcrop- to map-scale structures. A and B are from the outcrop shown in Figures 7, 8A, and 9.

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Figure 26. Mesoscopic scale structures in UCE2 tectonosomes. Boudinage and stacking of sandstone-siltstone beds are outlined in A. The shaly beds with planar, parallel cleavage did not develop scaly fabric. Figure A’ is a closeup of the stacking of blocks. The photograph of this outcrop is in Figure 8B. B: Line drawing of a mesoscale, disharmonic fold, which passively rotates the blocks. This fold contains a smaller scale rootless fold of a thin siltstone bed, formed during a previous tectonic phase. This rootless fold traps nonscaly and noncleaved clays in the hinge zone (Figs. 27D and 29D).

over, all the observed veins have sharp and clear-cut borders with the host rocks. The marly beds mostly have a broadly asymmetric lens shape, with sharp borders (Fig. 28, A and B). The shape of these bodies is associated with a pressure-solution cleavage that subdivides the rocks into centimeter-scale lithons that have been deformed into sigmoidal shapes by faults (Fig. 28, B and C). This association passes to a noncohesive fault breccia (Sibson, 1977) in the more thinned parts of the bodies (Fig. 28, B and C). The zones of boudinaged bed packages show that the fracture boudinage of beds was accompanied by low-angle stacking of blocks (Fig. 26A). Low-angle stacking has counterparts even in the blocks in shaly matrix: stacking of meter to decimeter parts of the pseudobedding is evidenced in several cases by the lowangle interlacing of matrix bands (see Figs. 7 and 8C). Even though discontinuity-dominated deformation is the major characteristic of these rocks, there are some relicts of continuous deformations, defined before as mesoscopically ductile deformation, that crop out as single blocks—relict blocks—or as entire bed packages. Some blocks are the relicts of pinch and swell boudinage of limestones (Figs. 27F and 29A) and sandstones-siltstones. The boudinage, commonly associated with fault troughs (Hoedemaeker, 1973), occurred by plastic accommodation of rocks without observable mesoscopic or microscopic discontinuities.

Some sandstone-siltstone blocks are the relicts of rootless, noncylindrical folds, with marked thickening of the hinge (Figs. 27G and 29B) and curved axial planes. Folding occurred by the plastic accommodation of the sandstones and siltstones, and is associated with discontinuous, curving, and anastomosing veins filled by the darkest and finest siltstone. These veins preferably develop in the hinge zone, either parallel or at various angles with respect to the sedimentary lamination. All the calcite-filled veins cut across the fold hinges. These folds, here defined as type 1 relict folds, are rarely observed in the Aptian-Albian units. Other limestone blocks show a rough pinch and swell structure due to belts of brecciation that resulted in rounded and subangular fragments inside a continuous, plastically accommodated, homogeneous matrix of the same rocks (Figs. 27E and 29C) Some blocks are the remnants of rootless, noncylindrical, isoclinal folds of sandstone-siltstone beds (Figs. 26B, 27D, and 29D), with parallel limbs and a moderate thickening of the hinge. Paired folds are preserved in blocks from thin beds of laminated siltstones, and display an evident asymmetry (Fig. 27D). The curvature of the hinge lines is either within the axial surface and therefore almost plane (plane noncylindrical folds, Hibbard and Karig, 1987) or implies a slight warping of the axial surface. The evident expulsion of the clays from the nucleus enabled the parallelism of the limbs (Figs. 27 and 29D). These folds have different geometrical relationships with the calcite-filled veins (Fig.

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27D): some sets of veins have been passively rotated in the hinge zone; fan-shaped veins, striking grossly parallel to the fold axis, developed in the hinge; and one or two sets of veins cut across the hinge. The calcite-filled veins parallel to the fold axis accommodate an incipient transposition of the hinges in the asymmetric paired folds. These folds, defined here as type 2 relict folds, due

to the evident difference in style with type 1, are present in a relatively high concentration (up to 10%–20% of the blocks) in certain parts of the LC, both of Neocomian-Barremian and Aptian-Albian age. Marlstone bodies have also been deformed into complex, very thinned and folded boudins (Fig. 30). The continuity at the

Figure 27. Structural anatomy of blocks in the LC tectonosomes. Deformation mostly resulted in pervasive development of mechanical discontinuities, such as veins, fractures, and faults (discontinuitydominated deformation, A–C). Some mesoscopic ductile, pinch-and swell-boudins (F) and rootless folds (G) are still evident as “relict structures”, overprinted by the discontinuity-dominated deformations. Some other blocks show an intermediate deformation (D–E), which may be interpreted as the transition between consolidated and nonconsolidated sediments.

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy

Figure 28. Marly bodies with discontinuity-dominated deformation. A: The distribution of white marly bodies in a block in shaly matrix fabric. The sharp, well-defined asymmetric lens shape and the sharp borders contrast with the style of the bodies in Figure 30, in spite of the same age and lithotype. B and C: The internal structures in a marly body, interpreted as a fragment of a turbidite bed. The general deformation of the body is associated to pressure-solution cleavage, faults, and cataclastic breccia.

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outcrop scale, as well the presence of beds with a slightly different composition, define a boudinaged and folded sequence (boudinaged bed packages). These structures, which were associated with the continuous deformation of rocks, formed without pressure-solution cleavage, faults, or cataclastic flow. The boudins have fringed borders, with interfingering and dispersion of rounded pieces of marlstones in the shales, that suggest an enhanced ductile behavior. Microscale observations: The scaly fabric of the matrix. The greatest part of the shaly rocks is subdivided into small lensshaped or prismatic-shaped scales (Fig. 31). The scales define the textural aspect of the rocks from outcrop to microscopic scale and define a scaly fabric, as defined for the UCE2 tectonosomes. In the LC units, the scaly fabric is characterized by two different classes of scales (Figs. 31B and 32A) that are easily distinguishable in the field with the occasional help of a hand lens, and that can be completely studied using the SEM (Fig. 33). Centimetric, prismatic to lens-shaped scales, here defined as scaly lithons, result from the interlacing of two or more sets of

subparallel discontinuous cleavage (Figs. 31, A and B, 33A, and 34), the spacing of which varies from 1–6 cm. This scaly cleavage is here defined as centimeter scaly cleavage. The cleavage surfaces are everywhere very well striated (Fig. 33B), polished, and typically covered by films of fibrous, striated calcite. All these features indicate slip along discrete planes (microfaults). The shales are pervasively subdivided into submillimeter scales, here defined as scaly microlithons, by closely spaced cleavages (Figs. 31B and 32, A and B), whose surfaces are not striated, but everywhere polished. The substantial differences in scale and morphological aspect suggest that a different name is appropriate, so I call it here millimeter scaly cleavage. The spacing is too pervasive and close to allow an appreciation of the shape of scales during field observation: using the SEM, the spacing of cleavage is about 30 µm and the scales show a lens shape (Fig. 33, C and D), defined by the interlacing of two or more sets of anastomosing, narrow (less then 1 µm thick) belts. This kind of scaly cleavage may be defined as a continuous and anastomosing, but not disjunctive, cleavage (De

Figure 29. Examples of mesoscopic ductile deformation in LC tectonosomes. A: A lens-shaped boudin from a calcareous turbidite bed of unit LC1. B: The ductile-style folding (type 1 folds) inside a lensshaped boudin in unit LC2a. C: A brittle-ductile in situ breccia in a block of unit LC1. D: A rootless fold hinge (type 2 folds) that constitutes a block in the boudinaged bed package of unit LC2 in Figure 26B (see also Fig. 27d).

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Figure 30. Ductile-style deformation of two marly beds of the LC2 tectonosomes at the outcrop scale. The two marly beds—interpreted as turbidite beds—are distinguishable by their different color, as well as by the presence of a fine sandstone interval associated with the whitest one. They define a strongly boudinaged and folded bed package, lacking discontinuity-dominated deformations. All the deformation occurred by plastic accommodation in the marlstones.

Libero, 1994a, 1994b). Belts are zones in which clay minerals have a strong parallel preferred orientation; the belts appear in relief with respect to the cleavage domains (Fig. 33D), because of the strong synkinematic recrystallization and lattice compaction of the clay minerals. The systematic reorientation is locally associated with an opposite, symmetric drag of the clay platelets around the cleavage. This feature may indicate that these planes are microscopic shear zones with a limited amount of slip.

The scaly fabric envelope (Figs. 31, B and D, and 34) is defined as the mean direction of the scales. This is useful for routine measurement, but it takes into account only the most evident system of scaly fabric. Therefore, in the case of a superposition of centimeter or millimeter scaly cleavage, the scaly fabric envelope indicates only the mean direction of the centimetric scales. Detailed measurements of the geometrical relationship between millimeter and centimeter scaly cleavage and of the distribution

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Figure 31. Scaly cleavages and scaly fabric in the LC tectonosomes. A and B: Relationships between competent blocks of sandstone and siltstone, bands of the matrix, and a centimetric scaly clevage (CSC) in an LC2 tectonosome. Adapted from Pini (1992b). The scaly fabric envelope is defined as the mean direction of the scales. The relationships between scaly fabric and a siltstone block are evident in the center of C and the explanatory diagram D. In this picture a millimeter scaly cleavage (MSC) and a CSC developed almost parallel to one another. CSC planes are the most prominent and continuous (some centimeters) discontinuities in the shaly matrix, and the MSC are the shorter discontinuities, better evident to the left of the block (LC2a tectonosomes).

of each single set of planes have been done at several stations in the field (Pini, 1987b; Pini and Castellarin, 1989). Scaly fabric is definitively pervasive in LC units, and therefore the shales may be termed scaly clays or scaly shales. The LC units involved in the tectonic wedge systems developed a pervasive scaly fabric, which originated from the centimeter cleavage

that cuts, bends and sigmoidally deforms the millimeter cleavage, giving what is here defined as scale-in-scale structure (Figs. 32A, 33C, and 34). Where the LC units crop out as large units that are not involved in close repetition (see Plate 2A Sillaro River area), they have a prevailing millimeter scaly cleavage. Belts, from 1 m to tens of meters thick, with a scale-in-scale structure (centimeter

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy

Figure 32. Scaly fabric, scaly cleavages, and relict parallel cleavage. A: Single centimetric scale, or lithon, produced by centimeter scaly cleavage, that includes several submillimeter scales, or microlithons, produced by millimeter scaly cleavage (MSC). This structure has been defined as scale-in-scale structure (LC1 tectonosome). A “relict” parallel-planar cleavage resulting from compaction, evident on the right side of the B, makes a gradual transition to MSC on the left. MSC overprinted the parallel-planar cleavage, which developed parallel to lamination of a siltstone bed (lower right) (LC2 tectonosome).

and millimeter scaly cleavage), are commonly observed. This fabric characterizes the Neogene thrusts or surrounds the lens shaped bodies of the UCE tectonosomes. Parts of the non scaly clays are present in roughly lensshaped bodies, commonly associated with some blocks (Figs. 26, A and B, and 32B). These bodies show a parallel, planar cleavage that has been overprinted by the scaly fabric at a low angle (Fig. 32B). Other clays have been trapped in the hinges of type 2 folds (Figs. 26B, 27D, and 29D) and have been preserved by a scaly fabric that completely surrounds the block. Furthermore, these trapped clays do not show the planar, parallel cleavage. Interpretation of microscale to mesoscale structures in the LC units. The interpretations are developed on the basis of geometry and relationships among blocks, bodies, matrix bands, and scaly fabric, and on the analysis of the distributions of the structures in blocks and bodies, as observed and measured in field (Castellarin et al., 1986; Pini, 1987b and 1994–1996, personal commun., unpublished new data). New evidence presented here relates to the deformational style and the mechanism of deformation of the blocks, and to the recognition of the relict structures. The geometrical relationship among the relicts of mesoscopically ductile deformation and the prevailing discontinuity-dominated

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deformations enable the progressive deformation of LC tectonosomes to be reconstructed. Structures related to mesoscopically ductile deformation. The shape of the ductile pinch-and-swell boudinage (Figs. 27F and 29A) seems to have developed in a monodirectional sense and is commonly associated with a single set of conjugate fault troughs and ridges. These structures indicate an extension that occurred in a single, preferential direction parallel to the bedding. The only remnants of an entirely ductile folding, predating the discontinuity-dominated deformation, are the rootless hinges of the type 1 relict folds (Fig. 27G). The geometry of the fold hinges suggests a non-cylindrical, isoclinal folding, with axial surface folding. The marked plastic bending of the sedimentary lamination in the fold hinges and the pinch and swell structure of the sandstone-siltstone beds were accommodated only by independent particulate flow of clasts, without grain breakage, so that the shallow-level character of this deformation is clearly indicated (Byrne, 1994). The presence of veins filled by the same sediment confirms the shallow-level character, and, together with the extreme mobility of the rocks, indicates that the deformation occurred in nonconsolidated or very poorly consolidated sediments, characterized by the high content and pressure of intergranular fluids (see discussion on vein structure in Lundberg and Moore, 1986; Hanamura and Ogawa, 1993). In micrite limestones, the deformation probably would also have occurred in uncemented or unconsolidated sediments. A similar, high mobility of nonconsolidated sediments is suggested by the ductilestyle, boudinaged and folded marlstone beds. The ductile-brittle breccia in limestones (Figs. 27E and 29C) may be related, as in the examples from the UCE1 units, to the increasing consolidation of sedimentary materials and its variability in individual limestone beds (Labaume, 1987; De Libero, 1994b, and 1993, personal commun.). The belts of breccia, which were developed in a single system, indicate a monodirectional boudinage for a single block. Structures related to discontinuity-dominated deformation. This kind of deformation dominates the LC tectonosomes. The evolution of the structures is partitioned into three different incremental steps, summarized in Figure 35. The stages of deformation indicated in the figure are described in the following. Stage I. The fracture boudinage of limestones and sandstones occurs, and determines the calcite-filled vein systems. A study of the distribution and compositional-textural characters of the veins is in progress and at this time it is not possible to anticipate the mechanism and the geometries of boudinage. However, the boudinage seems to have developed parallel to the bedding, in different steps related to both hydraulic fracturing and distinct, subsequent events of monodirectional boudinage. The type 2 relict folds may develop during this stage. Folds are kinematically related to the calcite-filled veins striking parallel to the fold hinge; folds rotated some of the systems of veins, and seem to be cut across by some subsequent vein systems. All the calcite-filled veins are related to deformation of at least partly lithified rocks, as indicated by their sharp and straight

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Figure 33. Scanning electron microscope photograph of the scaly fabric of tectonites. A: A lithon of a centimeter scaly clevage (CSC) in the matrix of the LC2 tectonosome from the outcrop shown in Figure 26 (100x). Striation and “lunulle de traction” (Petit et al., 1983) on a plane of CSC are evident in B, at the magnification of 200x. The sample was taken in the matrix of an LC2 tectonosome, west side of the Idice River valley. Planes of CSC and microlithons of a millimeter scaly cleavage (MSC) are the subjects of photograph C (500x). This sample came from an LC1 tectonosome, from the same outcrop shown in Figure 32A. A closeup of the MSC planes is in D, at the magnification of 2000x. The sample was collected from an LC1 tectonosome, on the east side of the Idice River.

border with the host rocks and by the lack of inclusions (Byrne, 1994). The deformation of the sandstone in the hinge of the type 2 folds is related to independent particulate flow, but the bending of the sedimentary lamination was partly accommodated by calcite veins, so these structures may be related to a ductile-brittle deformation in incompletely lithified, but not unlithified, sandstones (Byrne, 1994). Accommodation by vein development is particularly evident in the partly transposed asymmetric folds (Fig. 27D). During this step the planar, parallel cleavage that is present as a relict structure of the shales (Figs. 26, A and B, and 32B) may have developed. This cleavage developed parallel to the bedding plane and may be related to pure compaction of sediments. Part of the millimeter scaly cleavage might have developed in this transitional stage of consolidation. Scaly fabric may facilitate the escape of water and cause the consolidation of the clayey rocks (Karig, 1986; Carson and Berglund, 1986; Moore et al., 1986). A certain degree of clay consolidation may explain the

fact that all the observed calcite-filled veins in the sandstones and limestones are clear, and do not host clay grains from the matrix (Byrne, 1994). However, the clays were expelled from the type 2 fold hinges without developing any kind of cleavage (Fig. 29D): these features indicate independent granular flow of the clays, related to their wet, poorly consolidated condition. The lack of clay inclusions in calcite-filled veins may be also related to high pore pressure during vein development. Stages II and III. The second stage (II) is the phase of block separation. The largest, asymmetric calcite-filled veins in limestone and sandstone, shown in Figure 27B record a change in the deformation: the beginning of separation of boudins and blocks. The asymmetric shape and the oblique disposition of the calcite fibers indicate that the largest veins represent the first moment of block rotation with respect to the boudinage plane (Fig. 35). The scaly-fabric envelope and the bands of matrix and shales almost everywhere make a small angle, from 5° to 10°, with the boudi-

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy

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Figure 34. Schematic diagrams of the scaly fabric. A: The observed relationships between the centimetric scaly cleavage (CSC), the lithons (l), and the scaly-fabric envelope. The schematic block diagram in B shows the distribution of the millimetric scaly cleavage (MSC) and microlithons (ml) inside a single CSC lithon (scale in scale structure). C: Relative sense of shear of the planes that constitute CSC lithons and MSC microlithons, in an ideal case of scale in scale structure.

nage plane in the boudinaged bed packages. The angle with the single blocks increases locally because of the rotation of the blocks. The smaller, less elongated blocks rotated the most (see Cowan, 1990), and caused the evident perturbation of the matrix features (Figs. 31, A and B). The last stage (III), the passage from boudinaged bed packages to blocks in shaly matrix, implies the complete disruption of the sequences. In the field, this passage is gradational; it is defined by the further separation and by the stacking of boudins, to the definitive interruption of the continuity of beds. Nevertheless, blocks and bodies maintain the common parallelism of the plane of longer axes (bedding plane). This stage is represented in the boudinaged bed packages by the stacking of blocks that occurred at low angles to the bedding (Figs. 25, 26A, 31A, and 35). Other common structures are the normal faults that displace the bedding in boudinaged bed packages (Fig. 25B). Both stacking and normal faults are evident in north-northeast–south-southwest to northeast-southwest sections and vanish or become confusing in the other sections. Normal faults and stacking of blocks do not correspond to discrete fault planes in the shales, but they are accommodated by the reorientation of the scaly fabric

along the same direction as the blocks and, only rarely, by a decrease in the spacing of scaly fabric. The general mesoscopic planar attitude of LC, which defines the structural order evident in the field, has been generated by stages II and III of the structural evolution. These stages were not necessarily separated in time: they might have worked together. In several outcrops of boudinaged bed packages, the same scaly fabric accompanies both separated and stacked blocks (e.g., see Fig. 31, A and B). This association of modes became more effective at the border of the boudinaged bed packages and led to the pseudobedding in blocks in shaly matrix and the disruption of the beds. Blocks in blocks in shaly matrix fabric everywhere have an angle of 10° to 20° with respect to the matrix features (e.g., see, Fig. 31, A and B); this characteristic is common with the boudinaged bed packages and may demonstrate that blocks in shaly matrix is the result of both boudinage and stacking from the packages. In addition, the two stages seems to correspond to a coherent structural system that may be described as related to simple shear. The rocks behaved in different ways, depending on their composition and competence (Castellarin et al., 1986).

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Figure 35. Schematic diagrams of discontinuity-dominated deformation of LC tectonosomes, characterized by pervasive development of veins, fractures, faults and scaly fabric. The diagrams show a general interpretation of the three different deformational stages recognized in the field by structural analysis.

The sandstone and limestone blocks behave as objects in a tectonic matrix (see Passchier and Simpson, 1986; Bjornerud, 1989), and only the asymmetric veins played an active role in the separation and stacking of the blocks. These blocks might have different behaviors (Fig. 36). 1. The blocks might follow the direction of the boudinage plane, or the smaller blocks might have been rotated in the same direction as the rotational component of simple shear, deforming the scaly fabric. The flow inside the matrix is low to moderate and the object-matrix velocity gradient is low (see Bjornerud, 1989). 2. The blocks might have participated in a zone of major tectonic transport, and have been stacked and laid out in the direction of scaly fabric (block stacking in Fig. 36). The smaller

blocks might have been rotated in the opposite sense, in the same direction as the rotational component of the simple shear (Fig. 31, A and B). The flow of the matrix is moderate to high and the object-matrix velocity gradient is low (see Bjornerud, 1989). 3. The blocks might have been more highly rotated in the direction opposite to the simple shear, and the scaly fabric rotated in the same direction, following the blocks. These cases may be interpreted as zones of matrix stacking, as indicated in Figure 36, due to a mechanism of thickening similar to the one induced by the P foliation in shear zones and fault rocks (Rutter et al., 1986; Cowan and Brandon, 1994). The marly bodies developed the pressure-solution cleavage fault system and cataclastic brecciation (Fig. 28, B and C). The

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy association of pressure-solution cleavage and faults is very similar to the associations of pressure-solution cleavage and R-shears (Fig. 37), which are typical of pressure-solution fault rocks in marlstones (Koopman, 1983; Groshong, 1988, Pini, 1992a, 1992c). This is a mechanism conceptually similar to that which formed SC-tectonites (see Platt and Vissers, 1984), and it implies an extension related to simple shear. Cataclastic deformation accommodated larger strain in the thinned part of the bodies. This passage from pressure-solution–dominated rocks to a cataclasite is considered as a natural evolution in fault rocks in marlstones under progressively larger strain (Ghisetti, 1987). The shales behaved as a very mobile tectonic matrix that absorbed the greatest part of the strain, but was able to transmit stress to the blocks, as in the case of the marly bodies (Castellarin et al., 1986). The scaly fabric is everywhere, in the observed cases of both boudinaged bed packages and blocks in shaly matrix, oriented at a small angle with respect to the boudinage planes. The angle varies from 5° to 30°, depending on the different structural situations. This angle has been consistently observed, regardless of whether the scaly fabric is an millimeter scaly cleavage, or a centimeter overprinting an millimeter scaly cleavage. At a scale of centimeters, all the contacts between bodies and bands in the shales and matrix, even in the zones not affected by stacking and faulting, are due to the pervasive development of scaly fabric, which cut and transposed the original

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stratigraphic contacts. The matrix bands are therefore the product of a mechanism of oblique transposition instead of merely boudinage. This inference suggests that the scaly fabric is related to a simple-shear regime. All the deformations attributed to stages II and III seem to have occurred in completely consolidated materials. The angular shape of the blocks and their passive behavior, the pressure-solution cleavage–R fault association, and, mostly, the cataclastic flow in the marlstones are elements that suggest a complete consolidation of the competent beds. The scaly fabric was rotated together with the blocks during stage II, and participated in the block stacking during stage III. Thus, the largest part of the scaly fabric may have developed during stage I. Moreover, the presence of nonscaly clays preserved in the hinge of the type 2 relict folds suggests that the scaly fabric postdates the phase of the relict folds, and developed in the last phase of stage I, or contemporaneously with the block separation of stage II. This process might have produced part of the millimeter scaly cleavage, because some of the cleavage developed in just-consolidated shales and overprinted a planar, parallel cleavage that may be interpreted as a compaction cleavage. This late millimeter scaly cleavage is oblique to the bedding and to the compaction cleavage, so that it does not seem compatible with stage I of deformation, but more likely was associated with simple-shear–related deformation in stages II and III. This history also may explain the transposition

Figure 36. Relationships between the scaly fabric envelope and the blocks during stage III (block stacking).

Figure 37. General schematic diagram (left, A) at the mesoscopic scale showing the distribution and relative geometry of the structures related to the discontinuity-dominated deformation of a marly bed. From Castellarin et al. (1986). The geometrical relationships between pressure-solution cleavage, or PSC, and fault systems in pressure-solution fault rocks are summarized on the right (B). R: synthetic Riedel shear; R’: antithetic Riedel shear; P: thrust shear; X: antithetic extensional shear (Groshong, 1988; Pini, 1992c).

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of the sedimentary contacts of the clayey beds by the millimeter scaly cleavage. The centimeter scaly cleavage overprinted the millimeter cleavage and also developed in just-consolidated shales. The last stage (stage III) also produced a local decrease in the spacing of the scaly fabric, due to the nucleation of new cleavage planes. Thus, stratal disruption is the result of the combined effect of boudinage and stacking of blocks, which can be described as a microscale to mesoscale transposition that occurred in brittle conditions. The origin of this brittle transposition is related to the simple-shear flow of the matrix, which induced the boudinage and the rotation of blocks, and the stacking and faulting of blocks. Thus, the structural order of tectonosomes may be defined as a planar, tectonic foliation, which extends from outcrop scale to microscale. Relationships between the microscale to mesoscale and the mesoscale to macroscale structures in the LC units. The general genetic interpretation of the entire system of the discontinuitydominated deformations derives from the reconstruction of the relationship between the microscale to mesoscale and the mesoscale to macroscale structures. The tectonic setting of most of the LC tectonosomes seems to have been characterized by a system of mesoscale to macroscale thrusts and folds that affected almost-consolidated materials and that produced similar effects at different scales; for example (1) the thickening of sequences at smaller scales, associated with the mesoscale thrusting and the stacking of blocks; (2) the thickening of the sequences inside a single unit, associated with the overturning of extensive parts of the units caused by folds or thrusts; and (3) the generation of the belts and the smaller bodies of tectonosomes from a single stratigraphic succession at the macroscale. These deformations were accompanied by the separation of blocks, the stacking of blocks, the development of scaly fabric, and the boudinage-veining at the microscale to mesoscale. All of these processes were normally associated with mesoscale thrusting (Figs. 25, A and B, and 36). The plane of boudinage and of block separation is generally parallel to the thrust plane. The boudinage veining of the sequences (stage I), related to symmetric, layer-parallel extension, might have taken place before the mesoscale thrusting. The characteristic block separation (stage II), however, occurred in an environment of simple shear. The rotation of blocks was in the sense expected to be induced by the rotational component of the simple shear associated with the mesoscale thrusts (Fig. 36). The stacking of blocks developed in both the hanging walls and the footwalls of the mesoscale thrusts (Figs. 25, A and B). Block stacking is typically at a low angle, from 10° to 20°, with respect to the mesoscale thrust plane, and has the same vergence. In this prevailing case, the block stacking may be interpreted as diffuse simple-shear deformation, geometrically and kinetically similar to P shears (Skempton, 1966; Morgenstern and Tchalenko, 1967; Tchalenko, 1968, 1970). Further localized tectonic thickening of the system of matrix and blocks produced the stacking of matrix, in the same way that P foliation results in thickening in fault rocks (Cowan and Brandon, 1994).

Limited zones show a back-verging stacking, directed in the opposite sense (backstacking) (Figs. 25B and 35). Some normal faults have the same geometry as R shears (Riedel, 1929; Skempton, 1966; Morgenstern and Tchalenko, 1967; Tchalenko, 1968, 1970) with respect to the mesoscale thrust zone (Figs. 25B and 35). These observations suggest a general link between the mesoscale thrusts and the rotation, stacking, and faulting of blocks. The large-scale overturning of the sequences (folds or thrusts?) is still obscure, but the sense of vergence of the minor structures can be inferred (Fig. 27A). The mesoscale thrusts are evident in north-northeast–south-southwest to northeast-southwest sections and disappear in almost west-northwest–eastsoutheast to northwest-southeast sections. More mesoscale thrusts, verging in the same direction, toward either northeast or southwest, characterize separate domains with the same prevailing direction of transport. Some back-verging thrusts (backthrusts?) are present in these domains (Fig. 25B). The vergence of thrusting is independent of the general northeast or soutwes dip of either the boundaries of units or of the foliation. Mesoscale thrusting of the same vergence commonly deformed both the right-side-up and upside-down portions of the deformed successions (Fig. 25A), so the thrusting may be independent and subsequent to the overturning of sequences. Thus, the overturning of sequences developed in an earlier stage of deformation than the mesoscale thrusts, and may be responsible for the fracture boudinage (stage I) or for an early part of the block separation and rotation (stage II). Meter-scale folds have a prevailing northwest-southeast axis, and the axial plane is parallel to the general trend of the foliation. The vergence is either northeast or southwest. The striations on scaly fabric (centimeter scaly cleavage) and on some faults truncating the blocks (Fig. 28C), as well as the fibers of the large asymmetric calcite veins, indicate a northeastsouthwest direction of movement. The same direction is suggested by the pressure-solution cleavage–R fault association and by the striations on the same extensional faults in marly bodies (Castellarin et al., 1986). This similarity may imply that the microscale to mesoscale and the mesoscale structures described here compose a coherent structural system, responsible for stages II and III of the development of the tectonosome fabric. The prevailing direction of tectonic transport suggested by the kinematic indicators is toward the southwest (Castellarin et al., 1986), but this sense of transport is an inference to be confirmed by more analysis. All these structures are responsible for producing the strong, internal organization of the LC units. The common parallel attitude of the bedding and pseudobedding may be defined as a mesoscopic tectonic foliation. The tectonic foliation is generally parallel to the tectonic boundaries between belts of LC units. It is likely that the last stage of microscale to mesoscale deformation (stages II and III) and the mesoscale thrusts were the products of the generation of the belts and the smaller bodies of tectonosomes. These structures, together with the foliation, have been

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy bent and tilted by open, large-scale folds and large-scale thrusts related to the more recent tectonic evolution of the Ligurian nappe (Sillaro arc system). Structural interaction among the tectonosome units The contact between the UCE1 and UCE2 units is affected by small scale, decimeter to meter, asymmetric, open to close folds (Figs. 22, 23). Axes and axial planes are almost vertical because of the general tilting of the entire system of tectonosomes. Axial planes strike either perpendicular or almost parallel to the contact, and in the latter case, the vergence of folds developed in either of the two directions along the contact (Figs. 22 and 23). Small bodies of one unit that are present inside the other along the contact suggest a certain degree of indentation (Fig. 23). True indentation structures, as defined in the following, are not evident in the only section available for inspection. The UCE units are present only in the largest bodies, and only the body of Figures 22 and 23 is exposed well enough to be conveniently observed. The contacts between UCE and LC (Figs. 17, 20, 22, and 23) and between different LC units (Figs. 14) are observable in a wide range of differently oriented sections. Figure 38 shows a general interpretation of the observed situations. The interfingering of the two different units characterizes the contacts and results in elongated and almost triangular-shaped bodies. These structures, herein defined as indentation structures, are present along belts of concentrated deformation placed at low angles, from 10° to 20°, with respect to the contacts. Some slices are completely included inside the host rocks and appear as single, elongated, lozenge-shaped bodies (Figs. 18, 23), defined as guest bodies. In these cases a mixing of different stratigraphic units occurred, but the term “exotic block” should not be used because the “roots” of each single body are still evident. Indentation structure and guest bodies are also present, with the same geometry, inside the UCE and LC, among different shaly and clayey beds. In LC tectonosomes, these structures are evident at the scale of centimeters in the sequences of simply boudinaged bed packages (Fig. 31A). They increase in importance and dimensions, to tens of centimeters, in the zone of block stacking, and have the largest diffusion and the widest dimensional distribution, to 1m, in blocks in shaly matrix (Fig. 8C). They are also present everywhere in the smaller UCE bodies (Figs. 17–20) and are concentrated along the contact with the LC into the largest UCE bodies. All these processes occurred on either side of a body, or a band, with opposite symmetry. These structures are geometrically compatible with a simple-shear deformation, acting either parallel to or slightly oblique to a body or band. A centimeter scaly fabric, as defined for the LC unit, of 4–6 cm spacing, is concentrated along the borders of the UCE bodies, close to the LC tectonosomes, and affected all the mass of the smaller UCE bodies included in the LC tectonosomes, or is associated with the Neogene faults and thrusts of the Sillaro arc. Scaly

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fabric gives rise to, or enhances, the decimeter indentation of different clayey bodies close to the borders (Fig. 18). Scale indentation produces the small-scale (hand lens to microscope) transposition of the contacts between clays of different compositions and colors. Scaly fabric spacing decreases approaching the contact with LC units. Structural evolution of the tectonosomes The fabric and structures described and interpreted in the preceding indicate that various stages of deformation are characterized by different styles and occurred under different conditions of sediment consolidation. The evolution of UCE units is different from that of LC units. LC units. The transition from nonconsolidated sediment to rocks is complete in the older and more-deformed LC, which were presumably deformed close to the sea floor in the first stage and were subsequently buried while undergoing a consolidation related to both burial and tectonic deformation. The structures and the fabric follow this evolution. 1. Most of the relict structures suggest an early stage of mesoscopically ductile deformation of nonconsolidated sediments. Each individual structure likely developed during a monodirectional layer-parallel extension, instead of an axially symmetric extension (Cowan, 1982, 1985). Because only the hinges of the type 1 relict folds are now preserved, and the boudins have been rearranged by the subsequent fracture boudinage, it is difficult to determine the relationships between folding and boudinage. It is also impossible to reconstruct the kinematics of extension for the whole sequence, and, if the structures were related to shallow-level tectonic deformation, to slumping or to gravity induced, layer-parallel spreading. The brittle-ductile brecciation of the limestone beds indicates that carbonatic mud may have had a higher diagenetic potential (De Libero, 1994b) than the sandstone and siltstone. 2. The type 2 relict folds and the calcite-filled veins inside the blocks suggest a transitional stage of deformation as the sediments were consolidating to rocks. A shallow-level tectonic deformation might have pertained instead of deformation by slumping because of the presence of the clear and sharp-bordered calcite-filled veins. The latter indicate that the limestone and sandstone were partly lithified, and possibly were characterized by high pore-fluid pressure; both features are expected to develop under significant burial (see Orange et al., 1993; Byrne, 1994). The asymmetric, incipiently transposed folding of the thinner siltstone beds indicates simple shear and may be compatible with large-scale, thrust-related isoclinal folding (asymmetric intrafolial folds; Onishi and Kimura, 1995). Because of the relative geometrical relationships among macroscopic and mesoscopic structures, this step of mesoscopic deformation might have been genetically related to the large-scale structures (folds, thrusts?) that were responsible for overturning of most of the sequences. 3. The separation and stacking of blocks, which were together responsible for the definitive structural arrangement of

Figure 38. Idealized summary of the geometrical relationships among deformational structures in UCE and LC tectonosomes. A: Relationships among the various discontinuity-dominated structures in LC units, in respect to the disposition of belts and bodies of UCE and LC. A closer observation (B) allows the mesoscale structures in a UCE body to be related to the structures of the hosting LC tectonosome. As generally observed in the field, the mesoscopic and mesoscopic-microscopic structures in the LC tectonosomes (mesoscale thrusts, block stacking, matrix stacking, normal “R” faults, pressure-solution cleavage–R fault association, scaly fabric) are compatible with the same general sense of shear that causes the structures in UCE (indentation structures, asymmetric guest bodies, monodirectional and/or asymmetric boudinage of beds). The closeup diagrams in C and D illustrate the small-scale transposition of the contacts between UCE and LC. The scaly fabric disappears progressively in rocks of UCE away from the contact. D: The trend of the individual scaly-cleavage component of the centimetric scaly clevages (CSC).

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy the LC units, occurred in almost completely consolidated rocks. The stratal disruption was due to the combined effect of simpleshear driven, layer-parallel extension (i.e., block separation) and block stacking. These two kinds of deformation are related to the mesoscopic thrusts. Together, all are responsible for (1) the mesoscopic internal order, expressed by the mesoscopic tectonic foliation, and (2) the outctrop-scale characteristics, such as the guest bodies of the LC units. UCE units. In contrast to the LC units, most of the structures that developed in the UCE are related to mesoscopically ductile deformation. The deformation of UCE, far from the effects of the structural interaction with the LC, occurred mostly in nonconsolidated sediments. The absence of intergranular deformation in sandstones (grain breakage) indicates that the deformation occurred at a very low effective confining stress (Lucas and Moore, 1986). The boudinage, which occurred by axial symmetric exension in UCE1 and UCE2, may be interpreted as the result of a layer-parallel, lateral spreading, without lateral confinement. This may be related to gravity spreading, even if the axially symmetric extension may also be compatible with a shallow-level tectonic deformation (Cowan, 1982). The stratal disruption in unit 1 developed by layer-parallel monodirectional extension along an east-west direction, associated with the isoclinal folding. The limestone beds here show a higher diagenetic potential than do the sandstone-siltstone beds, as indicated by the brittleductile brecciation of the carbonatic part of the mixed turbidites in the UCE1 unit. Relationships of UCE and LC. The contrast between the structural styles of the UCE and LC is more striking where it is present at the passage between the UCE bodies and the hosting LC rocks. Here, the monodirectional boudinage, the duplication of beds, and the transposition of sedimentary contacts in the UCE have a strong geometrical similarity, and, in several cases, geometrical continuity with the features in the LC, but the deformational mechanisms in the two suites of structures differ strongly. The deformation in the UCE is related to ductile shear zones, instead of the concentration of scaly fabric. A pervasive centimeter scaly cleavage developed only close to the border of the UCE as the direct continuation of the belts of concentrated scaly fabric in the LC. This cleavage produced the indentation of the contacts between the UCE and LC and, at smaller scales, the indentation of the sedimentary contacts between different clayey beds in UCE. Far (tens of centimeters to meters) from the contact, the boudinage and the interlacing of clayey beds are related to a mesoscopically ductile deformation of the clays in belts of concentrated displacement. These displacements occurred without the development of any kind of cleavage. The same distribution of structures is evident in the more-competent beds: the sandstone blocks have a prevailing ductile style, pinch-and-swell boudinage, that is overprinted by a fracture boudinage, close to the contacts with the LC units. It is remarkable that, even if the mechanisms of deformation were different, geometrically similar and commonly oriented structures deformed, at every scale, not only the shaly beds and

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bands of both the UCE and LC units, but also the contacts between the different units. This geometrical relationship suggests that the UCE bodies and the LC units both were affected by a similar simple-shear deformation. This deformation is related to the development of tectonic contacts between the UCE and LC units, as well as to the development of the older tectonic wedge system, such as the Idice-Quaderna wedge system. The direction of shear in all the units and for all the contacts is generally northeast-southwest; the sense of shear is either toward the northeast or the southwest. The other deformation in the UCE bodies, which included the first stage of boudinage by axially symmetric extension in both units 1 and 2, the isoclinal folding and the stratal disruption of the same unit 1, and the folding and stacking of the contact between UCE1 and UCE2, seems to have occurred independently from the LC deformation. The basis for this inference is the different geometry of this suite of UCE structures. The monodirectional boudinage inside unit 2, especially where it developed close to the borders with the LC units, might have been associated with the extensional component of the simple shear that deformed the LC and UCE as a whole. The boudinage seems to have acted in the same direction of the general tectonic transport. This latter interpretation, however, needs to be confirmed by further observations. Regional meaning of the tectonosome structures In previous works, the UCE bodies were not recognized as a discrete unit: they were confused with the Aptian-Albian varicolored shales, units 2 and 2a, in the work of Pini (1992b), or with the Monte Piano Formation (Pini, 1987b, Castellarin and Pini, 1989). These erroneous age assignments made it difficult to explain the different observed deformational styles, which were related to different amounts of compaction of clayey rocks of the same age (Castellarin and Pini, 1989; Pini, 1992b; Amorosi et al., 1992). In addition, the failure to recognize the existence and structural arrangement of the two tectonosome units, UCE and LC, resulted in the generic attribution of the tectonosomes to the deformation of a unique stratigraphic succession, the basal complex III, extending in age from early Cretaceous to Early Eocene and with the same degree and style of deformation throughout (Bettelli and Panini, 1989, 1992; De Nardo, 1992, 1994). Associating the different structural styles with a more accurate dating of these rocks (A. M. Borsetti, 1984–1987, personal communs; A. Negri, 1994–1995, personal communs.) may suggest a different view. The marked change in deformational style occurs abruptly along the contact between the UCE and LC units; it suggests a dramatic change in depth of burial and deformation and in the state of consolidation. Moreover, the difference in deformational style between the UCE and LC units corresponds to the different areal distribution of the units. The UCE units have a very limited distribution; they are present as bodies inside the other units and, more commonly, as bodies along the contact between the LC units. These remarkable differences

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between the two groups of tectonosomes are enhanced by the long stratigraphic gap that extends between the youngest LC and the oldest UCE units. This depositional hiatus, or lack of sediment preservation, covers most of late Cretaceous time, from the Cenomanian to Late Campanian, more than 20 m.y. The stratigraphic gap corresponds closely to the age of the Eoalpine phases (the Eo-Alpine events of Trümpy, 1973), which have been invoked by other authors from a regional point of view (Vai and Castellarin, 1993; Mutti et al., 1995). A change in structural setting may be hypothesized to have resulted from tectonic phases separating the deposition of the two different groups of units that were converted to tectonosomes. According to this interpretation, the LC units were deposited until the Albian; afterward, they were strongly deformed during the Late Cretaceous—Cenomanian to Turonian and Santonian to Campanian—to Eoalpine tectonic phases. The sedimentary record, preserved in the UCE units, started again after these tectonic phases and continued from late Campanian to early Eocene time, giving deposits that unconformably overlie the LC tectonosomes. All these units were subsequently deformed together during the middle Eocene (Mesoalpine) phase, as demonstrated by the Quaderna wedge zone. There is no evidence that the stratigraphic gap corresponds to lack of deposition, complete erosion of deposited sediment, or both. A more certain attribution of the tectonic event recorded in the tectonosomes can also be inferred from the interpretation of the LO olistostromes. This interpretation is discussed in the last chapter. Lozenge-shape in tectonosomes: Shear lenses from macroscale to microscale The lens shape—lozenge shape in three dimensions—is an important result of deformation in the tectonosomes described here, even if it is not so striking as the lozenge shape of the blocks and matrix that is a distinctive attribute of other disrupted units in melanges in orogenic chains and accretionary prisms (Moore and Wheeler, 1978; Connelly, 1978; Cowan, 1978, 1982, 1985; Bachman, 1982; Pettinga, 1982; Byrne, 1984, 1994; Nell, 1990; Miller and Gray, 1996). Lens or lozenge shapes are less evident because most of the blocks of the LC units, which are by far the most abundant component of the tectonosomes, have a prismatic, angular shape. Thus, even though it is subordinate, the lens shape is visible at different scales. 1. The termination of belts of units imparts a very elongated lens shape at map scale (Plate 2A and Fig. 14A). 2. The UCE bodies hosted by belts of LC bodies also have a lens shape at a scale that varies from meters to hundreds of meters (Plate 2A and Figs. 14, 17, and 38A); 3. The same shape also characterizes the smaller guest bodies that developed along the contacts between the LC and UCE units, at the scale of tens of centimeters to meters (Figs. 18, 23, 38, A and B). 4. Even the clayey beds inside the single UCE unit assume a lozenge shape, due to boudinage (Figs. 8D, 17–20, 38B). The

boudinage and the stacking of the shaly beds in the LC units produced lens-shape bodies (Figs. 7B, 31A) that pass to the very elongated, lens-shaped bands in the matrix of the blocks in shaly matrix (Figs. 7B, 8C). These structures developed at the scale of tens of centimeters to meters. 5. At a smaller scale, any contacts between different tectonosome units, as well the contact between the different shaly beds or bands in the LC and UCE units, are transposed by scaly fabric or mesoscopically ductile shear zones. The transposition of the contacts produced lens-shape indentation structures and guest bodies, at the scale of tens of centimeters to a few centimeters (Figs. 31A and 38, C and D). 6. The lens shape is typical of the single scales of the scaly fabric, either centimeter or millimeter scaly cleavage (Figs. 32A, 33, A and C, and 34), at a scale of few centimeters to tens of microns. The lozenge-shaped bodies and the indentation structures resulting from the transposition of the contacts between units and beds and bands were generated by the intersection and interlacing of two or three sets of (1) ductile shear zones, (2) belts of concentrated scaly fabric, and (3) cleavage planes. The centimeter and millimeter scaly cleavage fabric also formed in this manner (Fig. 34 B–D). In spite of the different scale and nature, these sets are oriented at a low angle, 10° to 15°, one to the other. The geometry of the sets and the sense of shear along the shear planes or zones are consistent with the disposition of the P, R, and D shears, summarized in Figure 39, that has been observed in experimental models, in fault rocks, in brittle shear zones, and at the regional scale (see Skempton, 1966; Tchalenko, 1968, 1970; Bartlett et al, 1981; Naylor et al., 1986; Hancock, 1985; Groshong, 1988; Pini, 1992b, 1992c; Cowan and Brandon, 1994). In detail, the interlacing systems of P, R, and D planes produce the shear lenses, as defined by Naylor et al. (1986). This association of shear planes and zones is the typical product of an advanced stage of deformation in the progressive experimental deformation of clays (Tchalenko, 1968, 1970). Similar associations are evident at the base of large muddy slides (Tchalenko, 1968, 1970) and are the most typical structures in fault rocks in clays and shales (Groshong, 1988). In these fault rocks, this association is more common as total shear strain increases (Pini, 1992a). This configuration of planes is the one that enabled the maximum flattening of the fault rocks, associated with the maximum tectonic transport of the rocks in the direction of the slip vector of the main structure (Fig. 41). Considering that the clayey rocks behaved as a tectonic matrix, the mobility of which enabled the deformation of entire sequences, the P, R, and D configuration of structures seems to have permitted the contemporaneous boudinage and stacking of the sequences. OLISTOSTROMES The polymict, chaotic sedimentary bodies are the only parts of the argille scagliose that can be considered as similar, from a genetic, textural and compositional point of view, to the olis-

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy

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Figure 39. Distribution of shear planes. A: Shear-plane distribution, following the nomenclature of Skempton (1966). The rotated R’ are geometrically similar to the X planes of Swanson (1988). B: Distribution of the different structural associations of shear planes, which subdivide the fault rocks in clays into lithons (shear lenses of Naylor et al., 1986). The structural association of planes P, R, and D is most stable close to the zone of principal contact, which is the zone of largest shear strain. This change in geometrical configuration of the shear planes is accomplished by the decreasing spacing of the planes. C: Two-dimensional models of evolution, showing the efficiency of the three different shear-plane configurations observed in natural examples (B). The models imply planar, nonwarped shear surfaces and assume that the lithons have not been deformed. Modified from Pini (1992c).

tostromes of the Tuscan successions (see Abbate et al., 1981a) and of the Romagna-Umbria successions. Moreover, the chaotic bodies in the Epiligurian successions—the Epiligurian olistostromes, or EO in Plate 2A and Figure 15—typically have bottom and top sedimentary contacts and are associated with normal sediments (Figs. 10A, 11, 40, A and B, 41B), in the same way as are the Tuscan and Romagna-Umbria olistostromes. Because of these reasons, as well the priority of the name, all these bodies have

been named olistostromes in this work. The chaotic block in matrix rocks associated with the tectonosomes have been interpreted here as olistostromes—the Ligurian olistostromes, or LO, in Plate 2A and Figure 15—because their textural and compositional characteristics are very similar to those of the undoubtedly sedimentary bodies of the Epiligurian and Romagna-Umbria successions. However, their interpretation is not certain because they lack preserved sedimentary contacts. In the following sections the texture and

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Figure 40. Maps of the Bologna Apennines showing the areal distribution of type A and type B olistostromes (A). Distribution and composition of the major olistoliths in some type B olistostromes (B). A is modified from Pini (in Castellarin and Pini, 1989). For location see Plate 2B. (L, M, U—Lower, Middle, Upper).

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy

Figure 41. Details of olistostromes belonging to the Epiligurian successions. A: Block in matrix texture of an Oligocene(?)–lower Miocene(?) olistostrome (EO1-2 olistostrome), with evident depositional banding of the matrix. Some large blocks (arrows) are present in the upper part of the body. B: Block-rich block in matrix texture and the basal contact of an upper Oligocene–lower Miocene olistostrome (EO2) on the clayey deposits of the Antognola Formation (whitest rocks).

composition of the EO, LO, and of some Romagna-Umbria olistostromes (Plate 2A and Figure 15) are discussed and illustrated. Epiligurian olistostromes (EO) Texture. Chaotic block in matrix texture. The most evident textural characteristic, common to all the olistostromes and described previously, is the impression of disorder that derives from observation at the outcrop scale (Figs. 10, A and B, 13B, 41, A and B, and 42, A and B). This impression is related to a chaotic distribution of clasts in the matrix: (1) the dimensions of clasts vary from a few millimeters to 1–2 m (blocks) in the type A and from a few millimeters to tens of meters (olistoliths) in the type B olistostromes; (2) blocks are floating in a shaly matrix and are randomly distributed; (3) blocks and olistoliths of various compositions are randomly assembled in the same outcrop; (4)

49

the blocks have a moderately triaxial to almost isodiametric shape and their dimensions vary from a few centimeters to tens of meters; and (5) the matrix is homogeneous at the mesoscopic scale. In particular, the Epiligurian olistostromes show also the following characteristics. Blocks from the marlstone bodies of the LC2 tectonosomes have irregular shapes and diffuse outlines, as shown in Figure 10B. The blocks make a gradual transition to the matrix; fragmentation and dispersion are localized along their borders. Blocks from Epiligurian clayey deposits have irregular and elongated shapes, with diffuse outlines. Along the border of blocks, the clayey rocks make a gradual transition to the matrix. An inverse grading of the largest blocks is locally observable (Fig. 41A), even if large blocks may also be present close to the base of the type A olistostromes. The orientation of blocks is almost isotropic or weakly anisotropic; there is only a weak orientation of roughly triaxial blocks (Figs. 10, A and B, and 41, A and B). The concentration of blocks in the matrix changes from outcrop to outcrop. It ranges from low, as in the case of Figure 10B and 41A, to moderate (Fig. 10A), to relatively high (Fig. 41B). However, the majority of the volume is matrix, so that the blocks are mostly matrix supported. The matrix hosts some irregular bodies of differently colored matrix (Figs. 10B, 11, 41A). These bodies, and the elongated and irregular blocks of Epiligurian clayey deposits, are concentrated in subparallel, irregular levels that give an impression of decimeter to meter banding, which developed grossly parallel to the basal contact of the olistostromes (Figs. 10B and 41A). Brecciated matrix. Brecciated matrix is a common attribute of all the EO olistostromes. At hand-sample to SEM scale, the clayey matrix is composed of homogeneous clays that contain abundant clasts of different lithotypes, with dimensions ranging from tens of microns to 1–2 cm (Figs. 42, A and B, and 13B). These millimeter-sized clasts have either rounded or angular shapes and may be either irregularly spherical or elongated. These features define the brecciated-matrix texture described in olistostromes by Ogniben (1953), Beneo (1955, 1956a, 1956b), Rigo de Righi (1956), Abbate et al. (1981a, 1981b), and others. This texture is always recognizable at mesoscopic and microscopic scales, even where it is overprinted by scaly cleavage linked to faults or tectonic wedges. At the SEM scale, the homogeneous clays, which include the microclasts, are characterized by an open, honeycomb texture (Fig. 42C). The distribution of millimeter-sized clasts is random. The orientation of clasts ranges from almost isotropic, with no preferred orientation of the most elongated clasts, to anisotropic, with a significant preferred orientation. Figures 13B, 42, A and B, show respectively increasing degrees of common parallel orientation of the elongated clasts. The degree of parallel orientation of clasts varies at the outcrop scale, as well in the same hand specimen. The highest anisotropy is characterized by a higher percentage of clasts that are reoriented. Elongated clasts were reoriented as a function of their dimensions. Generally, the smaller or more elongated clasts are more strongly oriented in the direction of

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B

1 cm

Figure 42. Photographs of brecciated matrix of Epiligurian olistostromes. A and B: Increasing common planar orientation of the microclasts. The arrows in B indicate clasts that are almost entirely constituted by vein-filling calcite. C: Scanning electron microscope photograph showing details of the matrix and enclosed microclasts (200x magnification).

flow; the largest or less elongated are only reoriented in the zones of highest anisotropy. The common orientation of the elongated clasts is, in general, nearly parallel to the basal contacts of the olistostromes. Close to the basal contact of the olistostromes, the anisotropic texture is more strongly developed and the preferred orientation of clasts is more evident (Figs. 42B and 43, A and B). The basal zone of some olistostromes shows an evident banding of the matrix, which is due to the strong common orientation of all the millimeter-sized clasts and to the deformation of some of the clasts (Fig. 43B). Deformed clasts were strongly boudinaged in a ductile style, and some display pinch and swell structures. The deformation presumably occurred when the clasts were not completely consolidated. The general development of the boudinage and the extreme extension of several clasts suggest a flow of the matrix, which produced a fluidal texture. The presence of asymmetric boudinage and sets of extensive ductile faults displacing the clasts, evident in Figure 43A, may indicate that the

flow of the matrix included a simple-shear component. This millimeter banding is accompanied by a millimeter scaly fabric, shown in Figure 43A, that is difficult to observe in the field, except where it has been enhanced by weathering. This scaly fabric, here defined ghost scaly fabric, is a continuous cleavage, morphologically similar to, but less well-developed, than the millimeter scaly cleavage of the LC tectonosomes. Composition and origin of the clasts. At every scale, the EO units show different degrees of mixing of rocks of different types and ages. The olistostromes are polymict at all scales of observation, and for this reason they have been defined as a true, or sedimentary, melange (according to Bettelli and Panini, 1985). The EO units mostly include blocks of LC tectonosomes. Blocks derived from LC tectonosomes are limestones from LC1 and LC1a, sandstones from all the LC units, and marlstones from LC2. All the blocks show the typical brittle structures that have been described previously, such as veins in the limestones and sandstone and siltstones, and the association of pressure-solution

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy

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Figure 43. Matrix with fluidal features of an Oligocene(?)–lower Miocene(?) Epiligurian olistostrome (EO1-2). The photograph of the sample (A) shows a microscopic scaly fabric. The polished cut (B) shows a matrix banding, related to ductile pinch-and-swell boudinage of several microclasts. The white microclasts at the center of the sample show an asymmetric boudinage. Another whitish microclast, immediately above the bottom border, shows a normal ductile fault.

cleavage and R planes in marlstones. Single blocks are composed of pieces of large, asymmetric veins, and some clasts in the matrix are fragments of calcite veins (Fig. 42A). Olistoliths from the LC tectonosomes are abundant, mainly from the LC2 and LC2a, and show the structural association described herein. Thus, all the blocks and olistoliths derived from the LC were included when they were completely consolidated and deformed, or after they had undergone stages I and II of LC deformation. Some Late Cretaceous to Eocene rocks have been recognized by analyzing and dating the blocks. Olistostromes include UCE2 rocks, blocks of Late Cretaceous limestones from some Ligurian successions (Helminthoid flysch[?], UCE1[?]), and blocks and slabs of some Ligurian Eocene flysch (Monte Morello flysch, Fig. 5). The Cretaceous rocks were completely cemented, whereas the rocks of Eocene age were supplied to the older Epiligurian olistostromes when they were not completely lithified. Bodies of partly consolidated pelites of the Epiligurian Monte Piano Formation were also abundantly contributed to the EO rocks. Some olistostrome bodies show a prevalence of olistoliths that were derived from the LC2 and LC2a tectonosomes (Fig. 11, type B). Instead of the abundance of slabs from a single unit, the olistostromes are polymict aggregates, because their matrix includes also blocks of LC1 and LC1a tectonosomes, as well as Upper Cretaceous to Tertiary, Ligurian and Epiligurian rocks. The composition of the millimeter-size clasts of the matrix—previously defined as microclasts—is highly variable: variously colored shales and claystones are the most abundant; microclasts of marlstones, micritic limestones, and siltstone are also common (Figs. 13B, 42, A and B). Fragments of shales and claystones derive from lithotypes that are equivalent to the LC tectonosomes, but the microclasts do not show internal structures and texture, such as millimeter scaly fabric. Thus, the origin of these microclasts cannot be entirely related to the LC tectono-

somes, but may derive from fragmentation of shaly beds of Ligurian units not affected by scaly fabric and compaction-related parallel cleavage. Microclasts of massive and laminated siltstones, whitish marlstones, and micritic limestones may more likely derive from the LC units, even if a provenance from the unit UCE1 or from the other Late Cretaceous to Eocene less-deformed Ligurian units cannot be excluded. All these microclasts have either angular or rounded shapes. Both angular and rounded clasts commonly contain calcite-filled veins (Fig. 42A), which may constitute an entire clast. The angular shape presumably indicates that the rocks were completely consolidated, and brittlely deformed, before they supplied blocks to the olistostrome. The rounded fragments with the veins may derive from erosion of angular fragments; the fragments without veins may derive from the noncomplete dispersion of a mud in water—an origin proposed for clay chips hosted in turbidites. This last mechanism is the possible origin of the white, rounded micritic limestone clasts, probably derived from Monte Morello flysch, in the older olistostromes, as indicated by comparison with the largest blocks. Disposition. These units crop out as well-defined sedimentary bodies. They have a basal sedimentary contact on either tectonosomes (Fig. 11) or horizons of normal sedimentary deposits of the Epiligurian successions (Fig. 41B). The basal contact is commonly well exposed and can be followed for several kilometers, as in the case of the northwestern side of the Sillaro River valley, and along the western flank of the Sellustra River valley. The base of the olistostromes cuts the substratum at a scale of meter; some erosional features result (Fig. 41B). The top is also characterized by a rough topography, which may explain the discontinuous and incomplete presence of Epiligurian sediments on the largest olistostromes (see EO1-2, Plate 2A and Figs. 10A, 11, and 42, A and B).

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The geometry of a single olistostrome can be locally well defined at map scale (Plate 2A and Fig. 40B), as in the case of EO1 and EO2 units inside the Epiligurian sequence at Settefonti, between the Idice and Quaderna Rivers. These bodies have a double lens shape and basal contacts that cut the substratum. In addition, most of the olistostromes display a common orientation of blocks and elongated microclasts and bodies of matrix; this orientation is parallel to the basal contact and may be considered to be bedding. Some examples of this bedding attitude are shown in Figures 10A and 41A and mapped in Figure 40 (A and B). The possibility of recognizing bedding reinforces the identification of olistostromes as sedimentary bodies. The EO units are also present as single tectonic slices associated with the tectonosomes and with other Epiligurian deposits (see the Reno-Idice wedge system, Plate 2A and Fig. 16B) or as a thrust sheet formed by Neogene tectonics (Plate 2A and Fig. 40, A and B). In all cases, olistostromes behave as sedimentary bodies that have been deformed, and the basal sedimentary contact is generally preserved in the thrust-bounded units. Ligurian olistostromes Some rock bodies involved in the pre-Epiligurian tectonic wedge, such as the Quaderna-Idice wedge system, share many of the characteristics of the Epiligurian olistostromes. For this reason they are interpreted as olistostromes, related to the Ligurian unit, and treated in this section as Ligurian olistostromes, here designated as LO. Texture. The aspect at the outcrop is chaotic, without an internal ordered organization. At the outcrop scale, the large-scale banding and the weak planar-preferred orientation of blocks, which are common features in the EO units, have not been observed in these rocks. This difference may be related to the fact that the LO units contain more large slabs, from meters to hundreds of meters, than do the EO units, and that the blocks therein are generally coarser grained than the ones in the EO units. The matrix is typically brecciated and appears very similar to the matrix of the EO units when observed in the field. On closer inspection of polished cuts, the millimeter-size clasts of the brecciated matrix mostly have a rounded shape, which prevails over angular shapes. In detail, only the clasts of limestone and some sandstone-siltstone clasts have an angular shape, while clasts of clays and marlstone, and most of the sandstone-siltstone clasts have a rounded shape. The shaly, angular clasts that are typical of the studied samples of EO units are only very rarely present. Elongated clasts are often present and have a common orientation, which defines an anisotropic texture. The matrix shows different degrees of anisotropy, related to different degrees of planar-preferred orientation of the clasts. The strongest anisotropic texture generally presents a foliated aspect, caused by the pinch-and-swell boudinage of clay, marlstone, and siltstone clasts. This boudinage is asymmetric. It is accomplished on sets of extensive ductile faults displacing the clasts, and by a scaly fabric. Fluidal features and ghost scaly fab-

ric, as defined for the matrix of the EO olistostromes, are also present far from the tectonic borders of the LO belts. The matrix is commonly affected by a centimeter scaly cleavage fabric, developed parallel to the tectonic boundaries of the tectonic wedge, and concentrated close to the tectonic boundaries. The attitude of the scaly fabric is shown in Figure 40B. The angle between the centimeter scaly cleavage and the fluidal features in the matrix is variable from outcrop to outcrop. Composition. The LO units comprise mainly blocks from the LC tectonosomes. Limestone blocks from LC1 and LC1a, or from equivalent formations of the basal complexes (Palombini shales) are the most evident, and they display a well-defined angular shape and sharp outlines. Angular blocks of thin siliciclastic sandstone-siltstone turbidites are also abundant. The olistoliths are either LC tectonosomes or bed packages of Upper Cretaceous turbidite successions. The LC blocks derive mainly from LC1a, LC2 and LC2a, and they display the classic fabric of these units, characterized by blocks in scaly matrix (Figs. 7 and 9) and some boudinaged bed packages. The matrix or shaly beds in these large blocks have a well-developed millimeter spaced scaly cleavage. The Upper Cretaceous successions are either siliciclastic, fine-grained, thin-bedded turbidites, thick beds of coarser siliciclastic turbidites, or thick pelagic calcareous turbidites. Olistoliths of these rocks are shown in Figure 40B. The age of all these rocks is primarily Late Cretaceous; only the thick calcareous turbidites have a more indicative, but not completely confirmed, Cenomanian-Turonian age. The siliciclastic turbidites can be related to the Ostia and Scabiazza Sandstones of the Baganza and Trebbia Supergroups (Fig. 5). Deformation in the Upper Cretaceous successions is limited to folds and boudinage of beds, which only rarely led to separation of the beds. This deformation was mesoscopically ductile and may have developed in incompletely consolidated materials. Other slabs are wellcemented ophiolitic breccias, of undefined age, but probably associated with the Palombini Shales, in accordance with the description of some basal complex sequences, such as that related to the Baganza and Vara Supergroups. The same, or similar, breccias are associated with the EO units, as mapped in Figure 40A. Blocks confidently attributed to the UCE units have never been observed. The submillimeter- to millimeter-sized clasts in the matrix are mostly composed of massive and laminated siltstones, marlstones, micritic limestones, and differently colored shales. The shales are only concentrated in the fine-grained tails of the submillimeter-sized clasts and may derive from the scales of millimeter scaly cleavage in LC tectonosomes. The lack of the varicolored shales in clasts larger than 1 mm characterizes the matrix of these olistostromes and constitutes the major difference with respect to the matrix of the EO units. Disposition and setting. All the LO units are involved in the tectonic wedges and none has preserved sedimentary contacts. The bedding grossly follows the tectonic boundaries, with variable angles. Tectonic overprinting is everywhere very strong: centimeter scaly cleavage pervasively developed parallel to tectonic

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy boundaries. Brecciated texture, however, is always easily recognizable inside the scaly lithons. The texture of these bodies is not so different from that observed in EO and Romagna-Umbria olistostromes that were tectonized along Neogene tectonic lines. All these characteristics suggest that these bodies are olistostromes, but the lack of sedimentary contacts prevents interpreting them with certainty. Romagna-Umbria olistostromes Some of the olistostromes interbedded in the RomagnaUmbria succession are here described because they are classic, undoubted examples of sedimentary bodies deposited by muddy debris flows. Moreover, the source of the sediments that constituted the muddy flows is certain. The rocks that constitute the olistostromes are mostly Ligurian and the source is the Ligurian nappe (as established in the works by Ricci Lucchi, 1975, 1986; De Jager, 1979; Ricci Lucchi et al., 1982; Landuzzi, 1988, 1992, 1994). The olistostromes here studied crop out immediately east of the Sillaro line. These olistostromes represent episodes of large-scale sliding from the Ligurian nappe laterally into the Umbria-Romagna basin (De Jager, 1979; Ricci Lucchi et al., 1982; Ricci Lucchi, 1986; Pini, 1987b; Landuzzi, 1988, 1992). They mark the progressive northeastward movement of the Ligurian nappe onto the Marnoso-arenacea Formation and the Santerno succession. Visignano-Bordignano olistostrome. This olistostrome, shown in Plate 1, crops out across the Santerno Valley (Bruni, 1973; Landuzzi, 1988, 1992, 1994) and constitutes a great olistostrome, about 50 m thick, in the Marnoso-arenacea Formation. The olistostrome is involved, as are all the other stratigraphic horizons of the Marnoso-arenacea Formation, in a large-scale fold-ramp (Landuzzi, 1988, 1992, 1994). In the outcrops closer to the Ligurian nappe, great slabs of LC tectonosomes, mostly unit 2 and 2a, are floating as blocks in matrix. In some of these outcrops, the slabs are so abundant that only careful observation enables one to establish that they are sedimentary clasts in an olistostrome (Pini, 1987b). The thin layers of brecciated matrix that surround these clasts derives from the fragmentation, alteration, and dispersion of the blocks in shaly matrix fabric of LC tectonosomes. Going far from the Sillaro line, toward the Santerno Valley, the slabs progressively become less abundant, and the prevailing aspect progressively becomes one of decimeter blocks dispersed in a brecciated matrix. Most blocks are limestones from the Palombini Shales and sandstone and siltstone from the Varicolored Shales. The dimensions of the blocks progressively decrease farther to the east. These olistostromes also contain large slabs of Osteto marls, interpreted as having been deposited on the Ligurian nappe during its emplacement onto the Marnoso-arenacea basin (see De Jager, 1979; Landuzzi, 1994). The presence of Osteto bodies has the same meaning as the Epiligurian deposits in the EO units. Osteto olistoliths are mainly present close to the Ligurian nappe; they become smaller and less abundant far from the Sillaro line.

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Ca Nova Pagliarizzo and Castello di Fiagnano olistostromes. The location of the outcrops of these olistostromes is shown in Plate 2A, and their stratigraphic position in Figure 15. They were described and mapped by Lucchetti et al. (1963), Bruni (1973), Ricci Lucchi et al. (1982), and Pini (1987b, 1993), and are interbedded respectively inside the upper Pliocene and lower-middle Pliocene deposits of the Santerno succession (Plate 2A). The Ca Nova Pagliarizzo olistostrome displays two different textures. Close to the Sillaro line, or to the Ligurian nappe, meterscale bodies of unit 2 tectonosomes are hosted by a brecciated matrix. The brecciated matrix also includes smaller blocks tens of centimeters in size. Bodies have an irregular isometer shape and fringed margins formed by mixing and amalgamation of lithotypes. Here too, the meter-sized slabs become progressively less abundant away from the Ligurian nappe. The Castello di Fiagnano olistostrome very clearly displays its geometry. It crops out as two separate masses (Plate 2A), the easternmost of which is completely and spectacularly exposed in a single outcrop (Fig. 44A). The body has a biconvex shape and is 60 m long and about 5 m thick. The base of the body fills an erosive channel in the Pliocene clayey deposits. The base is decorated by smaller irregularities, which are very evident in the outcrop as deep cuts in the Pliocene deposits (Fig. 44B). The internal texture of this olistostrome is a typical block in matrix. The blocks are prismatic, but almost isometer, and have sharp outlines. The distribution of the blocks is random; only the more elongated ones have a slight preferred orientation, parallel to the basal contact. Some bodies of matrix of different colors are present, as well as clayey bodies of Pliocene age; these bodies have an irregular lens shape. The contacts with the matrix are irregular, and the Pliocene clays make a gradual transition to the olistostrome matrix as a result of the mixing and amalgamation of the lithotypes. The clayey matrix is typically brecciated. The millimetersized clasts of the brecciated matrix have either rounded or angular, isodiameter, or elongated shapes. The plane of the major axes of the elongated clasts is commonly parallel to the basal contact. This common orientation defines the same moderate anisotropic texture that has been observed in the EO units. The strongest parallel-preferred orientation of the millimeter-sized clasts is concentrated in narrow belts. The anisotropic distribution of clasts becomes a general attribute in zones close to the contact (Fig. 45A). Here, the orientation of clasts is stronger and reaches a texture similar to, but not as enhanced as, the fluidal texture in the LO matrix. The fluidal structure of the matrix is here observable in relation to the direction of the longitudinal axis of the olistostrome, which may in turn represent the direction of flow (Figs. 45, A and B). In polished cuts, the matrix of the basal zone displays strong parallel orientation of all the elongated clasts, independent of their dimensions and axial ratios (Fig. 45, C and D). In contrast to the fluidal matrix in LO units, most clasts in this olistostrome are reoriented in the direction of the prevailing flow. Moreover, most of the clasts are not deformed; only a few are plastically

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G. A. Pini low angle to the basal contact. They are gently folded in two directions, either parallel or normal to the longitudinal axis, following the distribution schematically illustrated in Figure 45A. Olistostromes: General characters

Figure 44. Panoramic views (A) and details (B) of a late Pliocene olistostrome (Castello di Fiagnano olistostrome, Plate 2A). A: The lensshaped olistostrome and its erosional basal contact on the upper Pliocene deposits. B: A closeup of the irregular basal contact.

deformed and strongly elongated. Elongation of clasts occurred in two directions, parallel and normal to the longitudinal axis of the body, as is evident in the two differently oriented polished cuts shown in Figure 45 (C and D). These clasts also display seriated extensional zones of displacement—mesoscopically ductile faults—and asymmeter boudinage, which developed mainly in the direction parallel to the longitudinal axis. The largest and most isodiameter of the competent, nondeformed clasts are accompanied by pressure shadows in the matrix that lack matrix foliation. These define typical σ-type structures. These structures formed by the interaction of clasts and matrix are observable mainly in the sections parallel to the longitudinal axis. Fluidal features and the orientations of clasts are parallel to continuous and striated, anastomosing planes, which define a centimeter spaced scaly fabric (Fig. 45B). A ghost scaly fabric is slightly evident on weathered rocks and is associated with the fluidal features (Fig. 45B), but it is not observable in polished cuts. The strong orientations of clasts and the scaly fabric are disposed at a

Type A and type B are the two textural end members of all the olistostromes. The passage between the two types is gradational. At the map scale, in the Visignano-Bordignano and Ca Nova Pagliarizzo olistostrome, the slabs are concentrated close to the Sillaro line and decrease in number and dimensions away from the Ligurian nappe. The provenance of material is the Ligurian nappe, and the direction of flow is from west-southwest to east-northeast. This direction of provenance is well supported by geologic data for the Pliocene olistostromes (Ca Nova Pagliarizzo and Castello di Fiagnano) as suggested by Ricci Lucchi et al., (1982), Ricci Lucchi (1986), and Pini (1987b). An origin of the Miocene olistostromes—such as the Visignano-Bordignano olistostrome—as lateral debris flows and slides from a lobe of Ligurian nappe forerunning on the Marnoso-arenacea basin was proposed by De Jager (1979), Pini (1987), and Landuzzi (1992). Thus, type B may be considered to characterize the proximal facies of olistostromes and type A to characterize a distal facies. Proximal facies make gradual transitions to distal facies. The brecciated matrix may be either isotropic or anisotropic, depending on the degree of microclast orientation. The anisotropic texture is concentrated in narrow belts throughout an olistostrome. It becomes progressively more abundant and important approaching the basal contact. This parallel orientation of microclasts may be accomplished by fluidal structures, resulting from plastic deformation of microclasts and from flame structures in the clays. Fluidal structures developed at the base of the olistostromes, close to the bottom contact. This flow structure is associated with simple-shear structures, such as extensional ductile faults and asymmeter boudinage, and with an incipient, continuous microscopic scaly fabric. The simple shear is associated with pure shear and thus defines a general shear, which produced the extension of clasts in two principal directions, parallel and normal to the flow. Genesis of the olistostromes All the characteristics described herein may be found in the olistostromes interbedded in the various successions of the Tuscan domains (see Abbate et al., 1970a, 1981a; Labaume, 1992) and in the olistostromes of the Marnoso-arenacea, which are not discussed here (see Abbate et al., 1970a; De Jager, 1979; Ricci Lucchi, 1975; Berti et al., 1994). The stratigraphic positions of almost all these olistostromes are shown in Figure 3. In general, these olistostromes imply a strong disruption and fragmentation of the source sequences, as well as the mixing of rocks of different ages, different degrees of deformation, and different conditions of consolidation. The chief differences are in the composition. Tuscan olistostromes are either type B, type A, or both; they mainly

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy

A

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B

SC 1 cm

C

D

1 cm 1 cm Figure 45. Distribution of the planar anisotropy of matrix in the late Pliocene Castel Nuovo di Fiagnano olistostrome, with respect to the longitudinal axis and the sense of transport (A). B: The strong planar anisotropy of the olistostrome matrix, due to clasts with a common planar orientation and fluidal features. A widely spaced scaly cleavage (SC) follows the planes of anisotropy. C and D: Polished cuts of sample in B made parallel and orthogonal to the longitudinal axis. The parallel section (C) shows high ductile elongation of microclasts (arrows), associated with asymmetric features due to seriated extensional ductile faults. Perturbations of the banded aspect of the matrix are related to the development of pressure shadows along the borders of isodiametric, competent clasts, similar to the syntectonic σ−type porphyroclasts (Passchier and Simpson, 1986) in ductile deformation. D: A view of the orthogonal section.

consist of Subligurian rocks (Abbate et al., 1981a; Vai and Castellarin, 1993). The older olistostromes in the Marnoso-arenacea are type B bodies composed mainly of rocks from an external Tuscan domain and from the Epiligurian successions (Berti et al., 1994). Minor differences are textural and are mainly related to the presence of single and continuous beds, deformed by pinch-and-swell boudinage, as blocks inside the matrix of the type A facies (Abbate et al., 1981a; Labaume, 1992). Mixing of different lithotypes: Present-day and ancient models. The olistostromes of the Tuscan and Umbrian successions have been interpreted as mass gravitational deposits either in the foredeep at the front of the Ligurian nappe, or at a lateral ramp of the nappe (Merla, 1952; Abbate et al., 1970a, 1981a; Elter and Trevisan, 1973; Ricci Lucchi, 1975, 1986; De Jager, 1979; Labaume, 1992; Landuzzi, 1994; Berti et al., 1994). The Ligurian nappe constituted a morphological high, protruding in the Tuscan and Umbrian foredeeps basins (De Jager, 1979; Ricci Lucchi, 1986; Vai, 1987, 1988; Boccaletti et al., 1990a; Conti and Gelmini, 1994). These depositional relationships are indicated by the different sedimentary facies and bathymetry of coeval

deposits in the Epiligurian successions and in the foredeep turbidite sequences (e.g., see Ricci Lucchi, 1986). It is easy to imagine that slides and debris might have been detached from the front of the Ligurian nappe, intersected by several thrusts and faults (Castellarin et al., 1985), and from the deformational border of the Tuscan unit at the front of the Ligurian nappe (Berti et al., 1994). This explains why the olistostromes of the Tuscan successions are typically the precursors of the emplacement of the Ligurian nappe (Elter and Trevisan, 1973). The interpretation of the Epiligurian olistostromes is more problematic: the Epiligurian basins are satellite basins on the Ligurian nappe, the sedimentation in which was controlled by thrusts (Ricci Lucchi and Ori, 1985; Ricci Lucchi, 1986). Slope instability is characteristic of the submarine environment, particularly in an active tectonic setting. Even a very low angle of scarp can produce various kinds of mass-gravitational movements in non-consolidated or poorly consolidated sediments (Stow, 1990; Martinsen, 1994). In all the Epiligurian olistostromes, however, most of the materials were consolidated, and in several cases deformed, prior to being involved in sedimentary processes. All

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possible genetic mechanisms should involve the exhumation of large amounts of just-deformed and consolidated materials, which might have belonged to different structural levels. The same processes completely disrupt rocks from different units and formations and intimately mix them with non-consolidated and younger rocks. Possible processes may be related to slides, debris falls, debris avalanches, and debris flows related to thrust fronts, block faulting, and mud volcanoes. Slope failures at a thrust front. Slope failure may occur at the front of a thrust system, as reconstructed by seismic reflection profiles and by acoustic imagery along the border of the Cascadia accretionary margin (Davis and Hyndman, 1989). The slope instability is placed at the seaward side of the ridges related to the frontal anticlines of the thrust system. A single sedimentary body derived by these processes is a fan, or a lobe, with a maximum thickness of about 20–30 m, an area of ~60 km2, and an estimated volume of 1 km3 (Davis and Hyndman, 1989)(Fig. 46, A and B). A larger area, more than 100 km2, may be covered if several bodies interlace (see Fig. 46A), because of the persistence of subsequent events from the same zone of slope instability. These bodies have a proximal area with large blocks of sediments, the blocky zone in Figure 6 of Davis and Hyndman (1989) (Fig. 46B) that passes to smoother zones, which represent a distal, lessblocky, facies. The sediments coarsen and thicken upslope. Even relatively small slope failures may produce significant mixing of rocks of different age, deformation, and degree of compaction, considering the stacking of units and the exhumation of older rocks related to the thrust front. In addition, a previous tectonic fragmentation of the beds may be inferred and interpreted to have been related to the thrust activity. Debris and slide accumulations from submarine fault scarps. Other sedimentary bodies formed by the gravitational accumulation of slides and debris have been related to extensional tectonics during the Jurassic-Cretaceous sedimentological and structural evolution of the Southern Alps of Italy (see Castellarin, 1972, 1982a) and the Eastern Alps of Switzerland (Eberli, 1987). These accumulations, defined as megabreccia bodies, developed at the foot of submarine morphological scarps formed by synsedimentary normal faulting. These bodies are partly similar to the accumulations of debris and slides that have been extensively studied at the border of ancient and present-day carbonate platforms (see Tucker, 1990, and references therein). The main difference is that a megabreccia body derives from a single, catastrophic event of scarp collapse, triggered mostly by tectonic activity (Eberli, 1987). These deposits are organized in lobes that were sedimented in stratified pelagic limestones. The lobes are thick at the foot of the fault scarp and thin downslope, and the clasts coarsen upslope (Castellarin, 1972, 1982a). Longitudinal sections along the bodies indicated that different facies may be present in the same body. The bodies pass downslope from structureless and chaotic masses of blocks, to a disorganized, either clast-supported or mud-supported breccia, to stratified, normally graded and inversely to normally graded breccia, to blocks in a matrix characterized by fluidal structures and the orientation of

clasts (Castellarin, 1972, 1982a; Eberli, 1987). Locally present turbidite layers represent the downslope continuation of the megabreccia body (Castellarin, 1972, 1982a), similar to the debris flows from the carbonatic platform, which are not related to faults (Krause and Oldershaw, 1979). Megabreccia bodies are typically composed of rocks of different ages—the same rocks that were exposed on the fault scarp. The vertical extent of the stratigraphic interval involved in the megabreccia bodies depends on the height of the fault scarp (Castellarin, 1972). The distal part of the megabreccia bodies may also include hemipelagic mud, eroded from the substratum (Castellarin, 1982a). Mud volcanoes and mud diapirs. Mixing of rocks may also be produced in mud diapirs or mud ridges, as described in the literature (Barber and Brown, 1988; Orange and Underwood, 1995). Mud volcanoes are the external apparatus produced by extrusion of materials from a mud diapir that reached the sea floor (Fig. 46, C–E). Several examples have been studied with seismic reflection profiles and with sonar scans of the sea bottom on the Mediterranean Ridge of the Mediterranean Sea (Cita et al., 1981; Camerlenghi et al., 1992; various authors in Cita et al., eds., 1996b; Robertson et al., 1996), in the Black Sea (Ivanov et al., 1996), on the Nankai accretionary prism (Taira et al., 1992b), seaward of the Japan Trench (Ogawa and Kobayashi, 1993), in the Indonesian arc (Karig et al., 1987), on and seaward of the Barbados accretionary complex in the Atlantic Ocean (Stride et al., 1982; Langseth et al., 1988; Henry et al., 1990; Le Pichon et al., 1990). Some examples have been observed in subaerial outcrops along the south coast of Trinidad (Higgins and Saunders, 1967), in Timor and surrounding islands (Barber et al., 1986), and in the North Island of New Zealand (Stonely, 1962). The eruptive apparatus has a grossly conical shape and overlies, with a sharp contact, normal pelagic sediments (Heike et al., 1996; Galindo-Zaldivar et al., 1996)(Fig. 46D). The morphology is complicated by lobes and fans of debris deriving from creep and slumping of the eruptive apparatus and from slope landslides and debris flows (Henry et al., 1990; Heike et al., 1996; Ivanov et al., 1996; Robertson et al., 1996) (Fig. 46C). These debris fans and lobes can develop at different stages during mud-volcano activity, and they form sedimentary bodies interbedded in the normal sediments around the volcano (Henry et al., 1990)(Fig. 46E). The mud volcanoes are entirely composed of rocks lacking seismic reflectors. These materials, observed in cores, are mostly massive mud breccias. They are characterized by a homogeneous and isotropic distribution of angular clasts, supported by a matrix lacking graded bedding, common planar orientation of clasts, and laminations (Cita et al., 1994; Premoli Silva et al., 1996; Staffini et al., 1993; Robertson et al., 1996; van den Meer, 1996). The dimensions of the clasts range from tens of centimeters to some millimeters (Staffini et al., 1993; Cita et al., 1994, van den Meer, 1996). Clasts are mostly dark mudstones, and less commonly, limestones, sandstones, and siltstones (van den Meer, 1996; Akhamanov, 1996). The mud breccia also constitutes the lobes and fans, but here it has also an organized texture, as defined by Staffini et al., 1993, which features millimeter layering—lamina-

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy

Figure 46. Schematic map (A) of the distribution of the debris accumulations from slope failure at the front of a thrust system, and (B) schematic diagrams in plan and in section of a single debris splay from slope failure. Northern Cascadia subduction zone, Vancouver Island continental margin, after Davis and Hyndman (1989, redrawn). Schematic map (C) of the distribution of the morphologies and structures related to a mud volcano, the Bergamo dome, and (D) map and section of the Milford-Haven mud volcano, Mediterranean Ridge of the Eastern Mediterranean Sea. C is after Hieke et al. (1996, redrawn), and D is after Galindo Zaldivar et al. (1996, redrawn). Section of mud diapirs or volcanoes (E) at the front of the Barbados accretionary complex (after Henry et al. 1990, redrawn). Three groups of debrisflow bodies are associated with the volcano in the central part of the figure, even if the deeper group may be composed of sills. The front of accretion is on the right side of the section, to the west, indicated by the thrust that deformed the sediments.

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tion—due to the sorting by size of millimeter clasts, and a positive, fining-upward grading in clast size. Single sedimentary bodies from debris fans and lobes inside normal sediments consist of thin—3 to 10 cm—levels of organized breccia at the top and bottom of a thicker—20 to 80 cm—bed of massive breccia that support the largest clasts (Cohen sequence, as defined by van der Meer, 1996). Because of this distribution of clasts the entire body shows an inverse to normal grading. Fields of blocks have been observed in the lobes of mud volcanoes in the Mediterranean Ridge (Fig. 46C). This observation either suggests that large blocks from the source materials have been carried in and erupted with the mud breccias (Barber et al., 1986; Barber and Brown, 1988), or indicates that the slides affecting the eruptive apparatus produced blocks of mud breccia (Martinsen, 1994). The lithotypes and ages permit an inference about which part of the stratigraphic column contributed to the mud diapir: rocks of Miocene, Oligocene, Late and Early Cretaceous—Santonian, Cenomanian, Aptian, and Albian—ages have been recovered in the Holocene mud volcanoes of the Mediterranean Ridge (Premoli Silva et al., 1996). Depositional mechanism. Debris avalanches have been considered to be the mechanism responsible for the lobate bodies and some of the megabreccia bodies. Avalanching has been described as a sudden, rapid, and catastrophic process, related to a single event of slope failure, producing bodies that are ìmorphologically similar to subaerial landslides and slab avalanches of snowî as described by Davis and Hyndman (1989, p. 1470). The deposits are typically massive, non-graded, and structureless, rock masses, with scarce matrix, that pass downslope to massive, structureless and chaotic mud-supported breccia. The upslope coarsening of the deposits contrasts with the downslope coarsening that is distinctive of subaerial debris falls and avalanches (Martinsen, 1994). A possible distribution of the olistostrome facies related to this kind of depositional processes is shown in Figure 47A. Cohesive debris flow, as defined by Lowe (1979), is the mechanism involved in the distal part of the megabreccia bodies and in the debris lobes and fans derived from mud volcanoes. In submarine cohesive debris flows, the transport and emplacement are related to mostly laminar flow, concentrated at the base and at the top, while the rest of the body is translated as a rigid plug (Stow, 1990). Cohesive debris flows have a block in matrix texture. Clasts may vary from millimeters to tens of meters, and the matrix is typically fine grained. The larger grains are supported by mostly matrix cohesiveness (Lowe, 1979). Submarine debris flows are commonly rich in matrix, and the related matrix-supported deposits pass to prevalent muddy rocks in the case of mudrich debris flows. The texture is generally unordered, structureless, and massive, and the materials lack a significant orientation of clasts (Stow, 1990; Martinsen, 1994). However, a strong preferred orientation of elongated clasts parallel to the base of the bodies, an imbrication of clasts, and an inverse to inverse to normal grading have also been observed (Stow, 1990; Martinsen, 1994). The inverse and inverse to normal grading seems to develop downslope from the unordered and structure-

less texture (Krause and Oldershaw, 1979). The concentration of shear at the base of the body may develop a sheared fissilelensoid lamination (Stow, 1990). The coarsest debris is generally carried at the front of the flow, so that an upslope fining is observed (Martinsen, 1994). Figure 47B shows a possible distribution of the olistostrome facies in the case of a debris flow from debris and slides that accumulated at the front of a thrust system. The debris flows derived from mud volcanoes should be similar. Interpretation of the studied olistostromes. The distribution and relationships of the different facies in a single olistostrome are completely observable from upslope to downslope only in the Visignano-Bordignano and the Ca Nova Pagliarizzo olistostromes. In these olistostromes the type B proximal facies are upslope and make a gradual transition to the type A distal facies downslope. There is an upslope coarsening of clasts in a section parallel to the longitudinal axis of the lobe. The distribution of clasts in a section normal to the longitudinal axis changes from upslope to downslope. The type B facies containing abundant olistoliths shows an almost uniform and isotropic distribution of blocks, with a slightly positive grading of the coarsest fraction. The type B facies constituting slabs inside a matrix shows a random, isotropic distribution of blocks and a block in matrix texture. The type A facies displays an almost isotropic distribution of blocks, that, in the distal parts of the olistostromes, is accompanied by a slightly inverse grading of the coarsest fraction. The passage to the type A facies and to the inverse grading is associated with the development of fluidal features at the base of the olistostromes. Fluidal features may be interpreted as related to concentration of shear at the base, and may be an equivalent of the sheared fissile-lensoid lamination (Stow, 1990). This relationship suggests that the bodies in the Romagna-Umbria succession may be related to a single catastrophic episode of slope failure at the front of the Ligurian nappe. A diagram of the mechanisms responsible for these olistostromes is shown in Figure 47C: the proximal part is related to debris avalanches that pass downslope to a debris flow (Castellarin, 1972, 1982a). In this case the type B facies, and the part of the type A facies lacking fluidal structures, are related to debris avalanches. In contrast, the type A facies, and the type B facies characterized by the anisotropic orientation of clasts, may represent the deposits of cohesive debris flows. In the case of the EO units that have limited dimensions and are mappable bodies inside normal-bedded deposits, such as the EO1, EO2, EO3, EO4 , the dimensions of the sedimentary bodies are compatible with single slope failures at a thrust front (Fig. 47, A and B). The origin from slides at the front of a thrust system is consistent with the generally accepted structural-stratigraphic picture of the Epiligurian deposits: the more external Epiligurian successions were deposited in satellite basins, controlled and surrounded by thrust fronts (see Ori and Friend, 1984; Ricci Lucchi and Ori, 1985; Ricci Lucchi, 1986; Pini, 1987b; Castellarin and Pini, 1989). The thrusts might have exhumed Ligurian rocks, especially the older units. Thrusting might also have led to the mixing of rocks belonging to different structural units, and have also involved Epiligurian deposits (Fig.

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy 48A). The EO1, EO2, EO3 and EO4 deposits are either type B or type A, but the two facies are not observed together in the same body, so it is not possible to attempt a general interpretation of the genetic mechanism. The type B olistostromes are structureless and seem more likely related to debris avalanches; on the contrary, the type A olistostromes are characterized by very evident fluidal features at the base and may be more easily related to debris flows.

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The various outcrops of the EO1-2 unit consist either entirely of type A facies, or of both type A and B facies (Figs. 11 and 40A). In the latter case, the type B facies characterizes the bottom part of the olistostrome and passes upward and horizontally to the type A facies. The matrix that includes the slabs in the type B facies is characterized by the absence of fluidal features, even directly on the basal contact, where a pronounced parallel preferred orientation of the microclasts is present (Fig.

Figure 47. Interpretive diagrams of the origin of olistostromes from thrust fronts. A: Origin from a single, instantaneous and catastrophic event of slope failure, causing a debris avalanche. B: Origin by a debris flow from a zone of accumulation of numerous small slides and debris falls or flows. C: Origin from a single, instantaneous and catastrophic event of slope failure, causing a debris avalanche and, downslope, a debris flow.

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Figure 48. Schematic diagrams of origin of the Epiligurian olistostromes, in piggy-back, thrust-bordered basins. Origin is from thrust-related slope failure (A), or mixed , from mud diapir volcanoes and slope failure (B).

42B). Where the type A facies constitutes the base of the EO12 olistostrome, a well-defined fluidal texture can commonly be observed in the matrix (Figs. 43, A and B). Because of its thickness and wide areal extent, the EO1-2 may derive from a single large slope failure; alternatively, it may be the result of several failures from the same source, or nearer source points (Fig. 46A), at a thrust front. In the first case, the mechanism of emplacement may be related to debris avalanches, which produced: (1) the parts of the olistrostrome with type B facies at the bottom and type A facies at the top and (2) the parts with type A facies lacking either grading or fluidal features at the base, which pass downslope (?) to debris flows that deposited the portion of the olistrostrome with type A facies characterized by fluidal features and grading of blocks. In the second case, the type B facies may correspond to a single, proximal body deposited by debris avalanches and subsequently covered by type A bodies. The latter are related either to debris avalanches or to debris flows, depending on their internal textures. In the case of a single, large depositional event, the fact that the base of the resulting body rests on different Epiligurian and Ligurian formations (Plate 2A and Fig 15) may be explained by the strong erosion of the substratum.

However, an origin from a mud volcano, with associated sills, debris fans, and lobes, may also explain the great thickness of the body and the irregular relationships with the underlying Ligurian and Epiligurian formations. The EO1-2 unit, however, contains not only Ligurian rocks of every age but also Epiligurian rocks that are very close in age to the possible age of olistostrome emplacement. Moreover, these Epiligurian components appear to have been nonconsolidated when they were included in the olistostromes. These relationships are completely different from what is expected for mud volcanoes: in that case, the clasts, which are derived from rocks that have been deeply buried, are completely consolidated and markedly older than the age of the emplacement of the mud volcano (Premoli Silva et al., 1996). For this reason, I exclude an origin from mud diapirs alone. A possible interaction of slope failures and mud volcanoes can be hypothesized, considering that mud diapirs and ridges can coexist at thrust fronts, as in the present-day examples in the Mediterranean Sea (Langseth et al., 1988; Henry et al., 1990; Le Pichon et al., 1990; Cita et al., 1996a). An idealized reconstruction of the two possible origins for the Epiligurian olistostromes is suggested by Figure 48 (A and B). The figures also indicate the different kinds of clasts that are to be expected from the different models.

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy The erosive base that is observable in several examples of the EO and Romagna-Umbria olistostromes is difficult to interpret. In the literature, an erosive potential is only associated with debris avalanches, and has been excluded for the mud-cohesive debris flows with laminar flow at the base (Bosellini et al., 1989). The presence of blocks of the substratum inside the type A olistostromes with fluidal structures—e.g., Pliocene bodies in the Castello di Fiagnano olistostrome—may have resulted from the erosive activity of an inferred, upslope debris avalanche (Castellarin, 1982a). However, all the olistostromes or parts of olistostromes that are suspected to have been deposited by cohesive debris flows with laminar flow fill channels and have erosional morphologies at the base (Figs. 44 and 45A). The convex and irregular top surface that characterizes almost all the studied EO and Romagna-Umbria olistostromes has been observed and described for several other olistostromes in the Tuscan unit and Marnoso-arenacea in the Northern Apennines. This peculiar morphology is often accompanied by a cover of hemipelagic marlstone and thin turbidites, called the marl drape (Ricci Lucchi, 1986), which is present in place of the normal turbidites. This facies gradually passes upward to the classic facies of basin-plain turbidites, which are present beneath and all around the olistostrome body. These features suggest that the olistostromes in the successions of the Tuscan and Umbria domains constituted morphological highs in the basins (Abbate et al., 1981a; Ricci Lucchi, 1975; De Jager, 1979). Such sea-floor topography can explain the onlap of the normal deposits on the olistostromes, and the discontinuous Epiligurian sedimentation on the thicker bodies (e.g., EO1-2). The origin of the LO units seems more likely to be related to debris avalanches, considering the high proportion and wide diffusion of slabs therein. The thinner parts might have originated as cohesive debris flows, in view of their fluidal texture. The problem is that the fluidal features developed in zones of type A facies that are enclosed between horizons of type B facies, or that constitute the matrix surrounding slabs in the type B facies. The interpretation of the Ligurian olistostromes still remains obscure, because it is not possible to see the bottom and top stratigraphic contacts. Fluidal deformation could also have been produced by tectonic deformation of the LO units when they were not completely consolidated. A partial genesis from mud volcanoes cannot be excluded. SUMMARY OF DEPOSITIONAL AND DEFORMATIONAL HISTORY AND SETTING OF TECTONOSOMES AND OLISTOSTROMES Being able to distinguish the olistostromes from the tectonosomes has completely changed the cartography in this part of the Northern Apennines (cf. Plates 1 and 2). Two consequences have resulted: (1) the tectonosomes and the Ligurian olistostromes, LO, are viewed as coherent, if strongly deformed, stratigraphic sequences, instead of a chaotic mass, and (2) the Epiligurian successions are considered to be more widespread and continuous

61

than they formerly were, because they encompass large masses of olistostromes previously grouped together with the tectonosomes and LO units as argille scagliose. These consequences have opened two fronts of interest: the study of the tectono-stratigraphic, strongly disrupted Ligurian units (tectonosomes and LO), and the study of the sedimentological and regional meaning of the olistostromes within their newly interpreted, Epiligurian setting. These studies try to reconstruct the depositional and deformational history of the Ligurian units and Epiligurian successions in this part of the Apennines. The sedimentary and structural evolution of the Bologna Apennines is discussed in the two following sections: the first treats the preEpiligurian history, when the Ligurian tectono-stratigraphic units were deposited and deformed during the Eoalpine-Mesoalpine tectonic phases; the second considers the evolution during and after the deposition of the Epiligurian successions. Pre-Late Eocene (Eoalpine-Mesoalpine tectonics) The tectonosomes have been subdivided into two different groups of units, on the basis of their different ages, dispositions, and deformational styles: the Lower Cretaceous tectonosomes, (LC) derive from the deformation of lithostratigraphic units extending in age from Neocomian to Albian; and the Upper Cretaceous–middle Eocene tectonosomes (UCE) derive from the deformation of upper Campanian–middle Eocene lithostratigraphic units. These groups of units are separated by a long gap in sedimentation that extended for about 15 m.y., from the Cenomanian to the early Campanian. The LC tectonosomes are laterally continuous, narrow belts, whereas the UCE tectonosomes are everywhere disposed as smaller, lozenge shaped bodies inside the LC belts. The UCE and LC units also differ in internal structures and fabric. Disruption of strata in LC tectonosomes was mainly caused by tectonic phenomena, such as boudinage and macroscale to mesoscale duplication and stacking. These processes affected consolidated sedimentary materials and resulted in the complete transposition of the bedded sequences in a strain regime where simple shear predominated. I call the overall process brittle, discontinuity-dominated deformation. Mesoscopically ductile deformation, of nonconsolidated materials, is also preserved in the LC units, but it is difficult to observe because of the overprint of brittle deformation. The mesoscopically ductile deformation can be related either to very shallow tectonic activity such as thrusting, or to gravity-induced deformation, such as gravitational spreading or laterally confined sliding. Mesoscopically ductile deformation is better preserved in the UCE units. This deformation might have developed independently from the deformation of LC, mostly in unit UCE1 (upper Campanian–Maastrichtian?). This kind of deformation is probably related to synsedimentary slumps and slides. However, the most abundant structures are geometerally and kinetically concordant with the brittle deformation in LC and may be related to a phase of thrusting of LC and UCE units.

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A possible interpretation of the stratigraphic and structural characteristics of these units incorporates the effects of the Eoalpine tectonic phases on the paleogeographic domain of the Ligurian paleo-ocean, which is where the tectonosomes originated. The Eoalpine phases (Eo-Alpine events of Trümpy, 1973) in the Northern Apennines consist of the Austrian phase of Cenomanian-Turonian age and the Gosau phases of Santonian–early Campanian age (Vai and Castellarin, 1993), so the age of activation of the phases exactly corresponds to the stratigraphic gap between the LC and UCE tectonosomes. Following this interpretation, the different deformational styles that distinguish the two group of units may be related to Eoalpine tectonics. Severe deformation of the Lower Cretaceous lithostratigraphic units during these tectonic phases produced embryonic LC tectonosomes. This deformation first affected nonconsolidated materials at a shallow depth, near the sea floor. The tectonic phases might have produced the almost complete consolidation of the rocks, which I propose was caused by the rapid burial due to thrusting and stacking of the sequences, and to the long duration of the deformational interval, from Cenomanian to late Campanian, about 13 m.y. Compaction and water expulsion might have been helped by the development of shear zones and thrust surfaces, as suggested by drilling in present-day accretionary prisms (Moore et al., 1986). After the last Eoalpine phase, the late Campanian–middle Eocene strata were deposited on the just-deformed and almost completely compacted LC tectonosomes. These strata draped over an earlier system of tectonic wedges. The deposition continued until the early part of the middle Eocene, the age of the younger UCE tectonosomes. Both the embryonic LC tectonosomes and the late Campanian–Eocene deposits were deformed together after the earliest part of the middle Eocene and before the latest middle Eocene; the latest middle to late Eocene base of the Epiligurian Monte Piano Formation unconformably overlies the tectonosome structures. This tectonic phase produced the definitive setting and internal structures of the LC and UCE tectonosomes. It also led to the development of tectonic wedges, which are evident between the valleys of Quaderna and Idice Rivers (Plate 2A) in the structures referred to as the Idice-Quaderna wedge system in Figure 16B. A middle Eocene tectonic phase is the classic Ligurian phase of several authors (e.g., Merla, 1952; Abbate et al., 1970b; Reutter et al., 1978; Zanzucchi, 1978, 1980). Reutter et al. (1978), Vai and Castellarin (1993), Castellarin (1994), and Mutti et al. (1995) suggested that the Eocene phase may correspond to the so called Mesoalpine tectonic phase in the Alps (Meso-Alpine event of Trümpy, 1973). Some bodies with an olistostrome texture participated in the Eoalpine-Mesoalpine wedge system. These bodies are older than the latest middle Eocene, having been deformed together with the tectonosomes in that time. These olistostromes, here defined as Ligurian olistostromes, or LO, include slabs and blocks of just-consolidated and disrupted LC tectonosomes and of slightly disrupted, nonconsolidated Late Cretaceous, probably Cenomanian-Turonian,

lithostratigraphic units of a basal complex. However, they do not include, as far as is known, younger Ligurian rocks (CampanianMaastrichtian and Paleogene flysches), so a reasonable age for these olistostromes is Santonian–early Campanian. These bodies, which were involved in the first system of tectonic wedges, accumulated after the deformation of the Cenomanian-Turonian tectonic phase and were deposited during or between the Gosau tectonic phases. Moreover, the presence of olistostromes reinforces the interpretation that the deformations related to the Eoalpine phases occurred at shallow depths. In conclusion, if my reconstruction is correct, the tectonosomes and the LO units of the Bologna Apennines can be considered as the product of the pervasive and polyphase deformation of an entire single Ligurian stratigraphic succession. This succession was deposited and deformed in a paleogeographic domain distinct from those of the other supergroups shown in Figure 5. The lithostratigraphic components of the LC units are distal equivalents of the Lower Cretaceous parts of the basal complexes—the Palombini Shales or Varicolored Shales. The LC units are interpreted as a distal facies because the turbidite beds are typically less abundant and thin. The stratigraphic gap between the LC and the UCE units, and the presence of the Ligurian olistostromes testify to the strong impact of the Eoalpine phases on sedimentation. A thin section of turbidites and overlying hemipelagic sediments was deposited on the deformed Lower Cretaceous units from the late Campanian to the earliest middle Eocene. These deposits are the chronostratigraphic equivalent of the Ligurian flysch. The upper Campanian–Maastrichtian deposits consist of thin mixed calcareous-siliciclastic turbidites that may represent a distal equivalent of the Helminthoid flysch. The Paleocene(?) to lower-middle Eocene deposits represent a chronostratigraphic equivalent of the Paleocene-Eocene flysch, even though the lithofacies are very different. Reconstructing the original position of the paleogeographic domain in which the tectonosomes originated is a problem of regional geology, a topic that is not a target of this book. According to the hypothesis introduced here, the tectonosomes may derive from a part of the Ligurian ocean that was deformed earlier and more strongly than all the other domains. This domain is inferred to have been a zone of basins containing scarce overbank and hemipelagic deposits, from Cenomanian to middle Eocene time. This kind of deposition may have taken place in a bypass environment, such as a slope, which was probably related to the emergence of a deformational front in the Ligurian basin. Late Eocene–recent (Neoalpine tectonics) During younger deformational events, the EoalpineMesoalpine chain, together with the other Ligurian units in the Ligurian nappe, was translated onto the units of the Adria margin. During the translation of the Ligurian nappe, the Epiligurian deposits accumulated in piggy-back basins controlled by thrusts. The age of the thrusts has been reconstructed from their effects on the Epiligurian deposits, which are mainly expressed as angu-

Tectonosomes and olistostromes, argille scagliose, Northern Apennines, Italy lar unconformities (see Sestini, 1970; Ricci Lucchi and Ori, 1985; Ricci Lucchi, 1986; Pini, 1987b; Bettelli et al., 1989b; Castellarin and Pini, 1989; Mutti et al., 1995). Another possible effect of the tectonic activity was to induce the emplacement of olistostromes at the front of thrust belts. Olistostromes may be used to date the age of tectonic pulses. 1. The EO1 unit may be connected to early Oligocene tectonic activity, but this inference must be confirmed with new evidence. 2. The EO2 unit indicates late Oligocene–early Miocene tectonic activity, which is also indicated by unconformities at the base and inside the Antognola Formation and at the base of the Bismantova Formation. 3. A possible late Serravallian–Tortonian phase is marked by olistostromes in the Bismantova Formation (EO3) and in the preevaporitic marls (EO4), as well as by olistostromes in the Marnoso-arenacea (Visignano-Bordignano). 4. The Pliocene emplacement of the Ligurian nappe at the Po Plain is marked by olistostrome bodies along the Sillaro line (Ca Nova Pagliarizzo, Castello di Fiagnano olistostromes). The other effect of upper Oligocene–Neogene tectonics was to rotate and partly reactivate the Eoalpine-Mesoalpine structural arrangement of the tectonosomes and LO olistostromes: the ancient structures and fabric were reoriented along a new system of tectonic wedges or along the thrust fronts. This rearrangement resulted in the large-scale folds shown in Figure 25A and induced the Chattian-Burdigalian wedges shown in Figure 16B. The thrusts affecting the Ligurian nappe correspond in age with the tectonic phases reconstructed in the substrate of the Ligurian nappe (the Tuscan and Romagna-Umbria units), as well in the foreland and foredeep (Marche Adriatic and Po Plain fold and thrust belts) (see Castellarin and Pini, 1989). Moreover, the geometry of these thrusts is very similar to the geometry of thrusts in the substrate (AGIP data in Cremonini and Ricci Lucchi, 1982; Pieri and Groppi, 1981; Pieri, 1983), of those in the Romagna-Umbria fold and thrust belt, and of the buried thrusts in the Po Plain (Castellarin et al., 1985; Castellarin and Pini, 1989). The thrusts are not limited to shallow depths but instead seem to cross the entire Ligurian nappe, forming a continuation of the structures recognized and extrapolated in the substrate. These reasons suggest that the emplacement of the Ligurian nappe at the boundary between the Apennines and Po Plain is caused by thrust tectonics, rather than by gravitational tectonics (see Castellarin et al., 1985; Castellarin and Pini, 1989). The younger olistostromes are concentrated close to the lateral termination of the Ligurian nappe, toward the Marnoso-arenacea and the Santerno successions (Romagna-Umbria succession and fold and thrust belt). Along this lineament (Sillaro line), the structural axes (thrusts and foliation of the tectonosomes) are rearranged in a northeast-southwest direction and depict a structural arc. The sense of tectonic transport is consistently the same (northeast), so that this arcuate structural system can be considered as the expression of a dextral transpression (Sillaro arc, Castellarin and Pini, 1989).

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ACKNOWLEDGMENTS The first version of this book was reviewed by Darrel S. Cowan and Gian Battista Vai. Peter Ballance and other two anonymous reviewers gave valuable contributions, and together with Abhijit Basu, Science Editor, GSA Books, stimulated a substantial revision of the paper. This final version derives from the last critical review of Gian Gaspare Zuffa and Darrel Cowan. I am greatly indebted to Darrel S. Cowan of the Department of Geological Sciences, University of Washington, Seattle; without his constant help, and assistance with English, this paper would never have been published. Alberto Castellarin made the first suggestions on the tectonic nature of the tectonosomes and proposed that I do a Ph.D. thesis on this topic. His contributions to understanding the structural evolution of the tectonosomes, the Neogene tectonics (Sillaro arc), and to dating the competent blocks have also been extremely helpful. Anna Maria Borsetti and Alessandra Negri contributed by dating several samples of tectonosomes and Epiligurian rocks. Giorgio Crestana and Ernesto Rabbi greatly contributed to the study of the argille scagliose, from 1983 to 1987. Giorgio Crestana also contributed to the geological survey of the Bologna Apennines (as shown in Plate 2A). This paper has been greatly improved by discussions with and the suggestions of Peter F. Ballance, Abhijit Basu, John D. Bradshaw, Marco Campos-Venuti, Luigi Cantelli, Alberto Castellarin, Darrel S. Cowan, Cipriano De Libero, Ken-ichiro Hisada, Hans Peter Laubscher, Corrado Lucente, Juiro Ogawa, Franco Ricci Lucchi, K. Bernhard Spörli, Fabio Tateo, Gian Gaspare Zuffa, and Gian Battista Vai. All the schemes, maps, and line drawings in this book were drawn on Microstation 95, Educational License. Luigi Cantelli greatly contributed to the development of the computer hardware and software system. Angelo Guerriero constantly solved hardware problems. Gabriele Tonelli proposed many shortcuts. Paolo Ferrieri made the photographs at the scanning electron microscope (Figs. 33, A–D, and 42) and several other photographs of hand samples (Figs. 13B, 29A, 42, A and B, 43A, and 45, B–D), and Remo Gamberini made the thin sections and helped in making the polished cuts of the olistostrome matrix. This work derives from my Ph.D. thesis (Pini, 1987b); it also presents evidence and interpretations that have resulted from subsequent work and from new unpublished research. Field and laboratory work and the publication of this book have been supported by the Ministero dell’Universitá e della Ricerca Scientifica e Tecnologica (MURST) 60% and 40% grants (to A. Castellarin and G. B. Vai) Comitato Nazionale delle Ricerche (CNR) “Sismotettonica” and CROP 1-1A grants (to A. Castellarin) and CNR Gruppo Alpi grants (to G. Simboli). The computer graphic software and hardware were funded by grants from CNR Gruppo Alpi (to G. Simboli) and MURST 60% and 40% (to A. Castellarin).

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