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The tectono-stratigraphic development of the western oblique ramp of the south-central Pyrenean thrust system, northern Spain

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The tectono-stratigraphic development of the western oblique ramp of the south-central Pyrenean thrust system, northern Spain

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Content THE TECTONO-STRATIGRAPHIC DEVELOPMENT OF THE WESTERN OBLIQUE RAMP OF THE SOUTH-CENTRAL PYRENEAN THRUST SYSTEM, NORTHERN SPAIN. by Peter A. Bentham A Dissertation Presented to the FACULTY OF THE GRADUATE SCHOOL UNIVERSITY OF SOUTHERN CALIFORNIA In Partial Fulfillment of the Requirements for the Degree DOCTOR OF PHILOSOPHY (Geological Sciences) August 1992 Copyright 1992 Peter A. Bentham UMI Number: DP28600 All rights reserved INFORMATION TO ALL USERS The quality of this reproduction is dependent upon the quality of the copy submitted. In the unlikely event that the author did not send a complete manuscript and there are missing pages, these will be noted. Also, if material had to be removed, a note will indicate the deletion. UMI Dissertation Publishing UMI DP28600 Published by ProQuest LLC (2014). Copyright in the Dissertation held by the Author. Microform Edition © ProQuest LLC. All rights reserved. This work is protected against unauthorized copying under Title 17, United States Code ProQuest ProQuest LLC. 789 East Eisenhower Parkway P.O. Box 1346 Ann Arbor, Ml 48106- 1346 UNIVERSITY OF SOUTHERN CALIFORNIA THE GRADUATE SCHOOL UNIVERSITY PARK LOS ANGELES, CA LIFO R NIA 90007 This dissertation, written by Peter Arthur Bentham under the direction of h.i$....... Dissertation Committee, and approved by all its members, has been presented to and accepted by The Graduate School, in partial fulfillm ent of re quirements fo r the degree of DOCTOR OF PHILOSOPHY Dean of Graduate Studies Date DISSERTATION COMMITTEE Chairpersoi ACKNOW LEDGM ENTS I would like to acknowledge the following people who have all helped me during the completion of this dissertation research: my advisor Professor Douglas W. Burbank, for not only allowing me the chance to work and study in such a wonderful field area, but also for helping and guiding me during the definition and completion of this project. Without Doug’s boundless energy and enthusiasm, this work would not have been finished so promptly or so smoothly; Professor Steven P. Lund and Dr. Cai Puigdefabregas (Catalan Geologic Survey) for their time and assistance during the analysis of paleomagnetic data and the discussion of field relationships; Steven Vincent, Sr. Antonio Teixell, Sr. Antonio Barnolas, Sr. Manel Zamorano, Sr. Jaume Verges and Dr. Cai Puigdefabregas for many open discussions and exchanges concerning Pyrenean geology, and for allowing me to incorporate the salient parts their work into this research; the other ‘Burbank Boys’ Andrew Meigs, Phil Hogan, Peter Tailing and Julio Friedmann for their free-flowing discussions and critical reviews of my work; the people of Altoaragon, for taking me into their homes during the long hot summers of solitary study, most especially Rosa and her staff at “Casa Ames”, and Maria, Pepe, Nuri and Jose-Marie of “Casa Carrera”. Thank you all for making fieldwork such a pleasurable experience. This dissertation research received partial support from the following sources: The USC Dept, of Geological Sciences Graduate Student Research Fund; the USC Graduate School (Oakley Fellowship); A.A.P.G. Grants-in-Aid; and National Science Foundation Grant #EAR-8816181 to Prof. D.W. Burbank. Additionally, I would like to offer my appreciation to; Sally Henyey, Eileen O’Gorman, and Karen Young for their help during the long tedious hours spent in the USC Paleomagnetics Laboratory; the Dept, of Geological Sciences Office Staff, Rene, Virginia, Sue, Denise, Cindy and Desser, for smoothing my path through USC administration, and for making those more menial day-to-day tasks bearable and easy. Finally, I would like to offer my love and thanks to my friends and family: Kevin, Reese, Erik, Tuck and Sandy for supplying the many sporting diversions that helped me to maintain my sanity during the past 4 years; Jim, Sue, Malcolm, Semele, Kathy, Mary and Chris for just being around and making my life much more pleasant; my parents, Marjory and Frank, for their quiet strength and support, and for allowing me to pursue my own path for so many years, no matter how foreign it may have seemed; Phyllisa, for taking me into your heart and for giving me so many chances. I know I can never repay the debt I owe to you, but I hope that you will let me try. iii TABLE OF CONTENTS Chapter 1 INTRODUCTION: DEFINITION OF THIS RESEARCH AND REGIONAL PYRENEAN SUMMARY...........................................................1 Goals of Project................................................................................................ 2 Importance of this research............................................................................3 Regional Pyrenean Summary.........................................................................6 Thrust Belt.............................................................................................. 8 Foreland Basin...................................................................................... 14 Stage I (Late Santonian-Maastrichtian).................................14 Stage II (Late Maastrichtian-Paleocene).............................. 16 Stage III (Early and Middle Eocene)......................................17 Stage IV (Late Eocene-Oligocene)....................................... 19 Methodology......................................................................................................28 Chapter 2 THE CHRONOLOGY OF MIDDLE AND LATE EOCENE DEPOSITION AND DEFORMATION ACROSS THE WESTERN MARGIN OF THE SOUTH-CENTRAL UNIT, SOUTHERN PYRENEES, S P AIN .........................................................................................29 Abstract...............................................................................................................30 Introduction........................................................................................................ 31 Structural setting............................................................................................... 34 Stratigraphic Framework..................................................................................36 Sampling and Data Analysis...........................................................................43 Section Description, Measurement, and Sampling........................43 Laboratory A nalysis............................................................................. 45 Results................................................................................................................ 47 Pilot Study..............................................................................................47 Magnetostratigraphic R esults............................................................ 54 Ainsa Basin Magnetic Polarity Stratigraphies.................... 54 iv Mediano M PS............................................................... 54 Almazorre M P S ............................................................64 Eripol M P S ....................................................................68 Liguerre M P S ............................................................... 72 Tremp Basin Magnetic Polarity Stratigraphies .......... 77 Esera Valley Composite M P S ................................... 77 Lascuarre MPS.............................................................84 Sediment Accumulation Calculations............................................................88 Discussion......................................................................................................... 95 Lateral Correlation and Sequence Boundaries.............................. 95 Eocene Paleogeography.....................................................................98 Ypresian-Early Lutetian (-49-45 M a)................................... 98 Middle Lutetian (-45-43 M a)..................................................100 Late Lutetian (-43 M a)............................................................100 Bartonian-Priabonian (-42.5-36 M a)......................................103 Structural Development......................................................................103 Conclusion......................................................................................................... 106 Chapter 3 TEMPORAL AND SPATIAL CONTROLS ON ALLUVIAL ARCHITECTURE IN AN AXIAL DRAINAGE SYSTEM, LATE EOCENE ESCANILLA FORMATION, SOUTHERN PYRENEAN FORELAND BASIN, SPAIN............................................................................ 108 Abstract................................................................................................................ 109 Introduction........................................................................................................110 Eocene Structural and Stratigraphic Framework of the South- central Pyrenees...............................................................................................115 Structural Development......................................................................115 Stratigraphy........................................................................................... 117 Methods..............................................................................................................121 Stratigraphic Subdivision of the Escanilla formation.................................. 129 Ainsa Basin............................................................................................. 129 Lower M em ber......................................................................... 130 v Middle Member........................................................................... 135 Upper M em ber........................................................................... 135 Western Tremp B asin........................................................................... 137 Regional Correlation.......................................................................................... 141 Ainsa and Tremp Basins.......................................................................141 Correlation to the Jaca Basin............................................................... 142 Discussion...........................................................................................................148 Conclusions.........................................................................................................158 Chapter 4 A REVISED BRAIDED-STREAM DEPOSITIONAL MODEL: AN AGGRADING AND AVULSING LOW-SINUOSITY SYSTEM....................163 Abstract................................................................................................................ 164 Introduction........................................................................................................165 Regional framework........................................................................................... 167 Architectural Observations and Facies Descriptions...................................172 Facies Descriptions and Interpretations.............................................175 Gravel-Dominated Channel-Fill Facies:.................................175 Description......................................................................175 Interpretation.................................................................. 179 Sand-Dominated Channel-Fill Facies:...................................180 Description....................................................................180 Interpretation.................................................................. 182 Sheet-Splay Deposits................................................................ 184 Description....................................................................184 Interpretation.................................................................. 186 Pedogenically Modified Overbank Sediment........................ 187 Description.................................................................... 187 Interpretation.................................................................. 187 Additional Examples.......................................................................................... 188 Plio-Pleistocene of Southern New M exico.......................................188 Plio-Pleistocene of the Eastern Potwar Plateau, Pakistan 190 Ventura Red Beds, Methow Basin, W ashington.............................. 191 vi Proposed Depositional Model.......................................................................... 192 Discussion........................................................................................................... 195 Conclusions.........................................................................................................200 Chapter 5 DISSERTATION SUMMARY AND GENERAL CONCLUSIONS 201 Bibliography..................................................................................................209 Appendices.................................................................................................... 221 Appendix 1: Tables of Magnetostratigraphic Data A. Mediano Average Site Vector D ata..............................................222 C. Eripol Average Site Vector Data................................................... 225 D. Liguerre Average Site Vector Data...............................................227 E. Esera Valley Average Site Vector Data....................................... 229 Appendix 2: Detailed Lithologic Sections A. Legend for Detailed Lithologic Sections......................................233 B. Mediano Lithologic Colum n........................................................... 234 C. Almazorre Lithologic Column......................................................... 239 D. Eripol Lithologic Column................................................................. 241 E. Liguerre Lithologic Column.............................................................244 F. Esera Valley Composite Column - Santa Liestra....................... 247 F. Esera Valley Composite Column Continued - Meson de Pascual.....................................................................................................248 F. Esera Valley Composite Column Continued - Grustan 249 G. Lascuarre Schematic Column........................................................250 Appendix 3: Geohistory Data Chron 20 Geohistory Data and Graph................................................252 Escanilla Formation Geohistory Data and Graph.............................253 LIST OF FIGURES Chapter 1 Figure 1. Structural sketch map of the Pyrenees showing the location of the ECORS deep seismic reflection profile (taken from Munoz, 1991). 7 Figure 2. Strip geologic map and interpreted structural cross-section drawn along the ECORS deep seismic reflection profile (taken from Munoz, 1991)...................... 9 Figure 3. Crustal balanced and restored structural cross- sections along the ECORS deep seismic reflection profile (taken from Munoz, 1991). 13 Figure 4. Series of partially restored cross-sections showing the early Cretaceous extensional configuration of the Pyrenean system, and the subsequent thrust sheet geometries at a crustal scale during the progressive stages of foreland basin evolution (taken from Puigdefabregas et al., 1991)...... 15 Figure 5. Detailed partially restored cross-section of the southern Pyrenees during stage III of the foreland basin evolution. Section corresponds to the southern portion of the ECORS profile. The syntectonic sediments of this stage are shown shaded, and previous deposits are shown as stippled. UN and LN = Upper and Lower Nogueres units of the AZAS respectively. Taken from Puigdefabregas et al. (1991).....................................................................................................................18 Figure 6. Detailed partially restored cross-section of the southern Pyrenees during stage IV of the foreland basin evolution. Section corresponds to the southern portion of the ECORS profile. As in Fig. 5, syntectonic sediments of this stage are shown shaded, and previous deposits are shown as stippled (A = Aren Group. T = Tremp Group. AM = Ager and Montanyana Groups). Taken from Puigdefabregas et al. (1991)..........................................................................21 viii Figure 7. Apparent misfit between the Tremp and Ripoll basins during Late Ypresian time. Arrows indicate sediment dispersal directions. Shaded areas represent deep marine turbiditic basins. Numbers refer to the distances from a defined pin-line within the northern Ebro basin. Dashed lines delineate the margins of the SCU and Pedraforca thrust sheets. Taken from Nijman (1989)................................24 Figure 8. Tectono-sedimentary model for the Pyrenean orogenic basins given by Nijman (1989). From left to right the figures show the sequential evolution of the South- Central Pyrenees in schematic map form. Structural regime and significant structural events are indicated at the base of the figure, as well as the interpreted phases of basin evolution. Taken from Nijman (1989)...............................................................25 Figure 9. Summary of the paleomagnetic data showing rotational information from the central and eastern Pyrenees. Arrows represent the local mean declinations. Taken from Dinares et al. (1991).................................................................... 27 Chapter 2 Figure 1. The southern Pyrenean Foreland Basin. Inset map shows the approximate location of the study area within the Pyrenean system. The coverage of subsequent figures is also indicated. The important structural elements of the South-Central Unit (SCU) thrust system are also identified ......................................................................................................... 33 Figure 2. Simplified geologic map of the western area............................37 Figure 3. The pre-existing lithostratigraphic nomenclature of the western SCU. This figure incorporates data from Puigdefabregas (1975), De Federico (1981), Reynolds (1987), Mutti et al. (1988), and Cuevas Gozalo (1990)..............................39 ix Figure 4. Simplified geologic map of the western Tremp Basin. Villages and river valleys mentioned in the text are shown, as are the locations of the four magnetostratigraphic traverses taken within this portion of the study area (SLA = Santa Liestra. MDP = Meson de Pascual. GRU = Grustan. LAS = Lascuarre). The positions of more detailed sample location maps for the magnetostratigraphic sections are also indicated...................................................................................................... 42 Figure 5. Simplified map of the Ainsa Basin (based on Fig. 2) showing the locations of five detailed sample location maps (Figs. 7B(i), 7B(ii), 9B, 10B, and 11B) of the MPS within the syncline..............................................................................................44 Figure 6A. Plots of data for four samples showing representative Type I’ demagnetization behavior. GRU02A and MED43B represent non-marine siltstones, while MED01D and MED34E are marine mudstones. In the left- hand Zijderveld cartesian projection (As, 1960), open circles represent the vertical component of the magnetization direction, and the closed circles represent the horizontal component. NRM-H and NRM-V are the initial remanence components. The right-hand plot is a plot of intensity normalized to initial NRM, measured at each demagnetization step. The horizontal scale represents demagnetization level (either Oe or °C)........................................................49 Figure 6A continued...........................................................................................50 Figure 6B. Plot of data for two specimens showing representative Type IT demagnetization behavior (MED25A and GRU33B). Format is the same as Fig. 6A. Note the increase in relative magnetic intensity at higher temperatures. This is accompanied by highly variable magnetized directions as shown in the Zijderveld projection (As, 1960). The second two diagrams (MED25D) show a specimen subjected to mixed method demagnetization (see text for description). Note the lack on any increase in intensity and stable magnetized directions as demagnetization progressed.............................................................................. 52 Figure 6B continued...........................................................................................53 x Figure 6C. Plots of data for two samples showing representative Type III’ demagnetization behavior. Both represent non-marine lithologies. Again the format is the same as Figure 6A. In the Zijderveld projections (As, 1960), MED45A shows a southerly-directed declination but a positive inclination, while MED52B possesses a shallow west-dipping magnetized direction. Note the large percentage of initial remanence retained at 500 °C. in both normalized intensity plots....................................................................................55 Figure 7A. The Mediano Magnetic Polarity Stratigraphy (MPS). Schematic lithologic information, and lithostratigraphic nomenclature in indicated in the left diagram. The virtual geomagnetic pole (VGP) latitude is plotted against stratigraphic position in order to define behavior of the paleomagnetic field through time, and construct the MPS. Polar latitudinal error bars are calculated for each pole position. The legend for this, and all the following MPS diagrams is given in Fig. 9A.....................................57 Figure 7B. Detailed sample location maps for the Mediano MPS. The location of these detailed maps within the Ainsa Basin are shown in Fig. 5................................................................................. 58 Figure 7B continued.......................................................................................... 59 Figure 7C. Equal area stereographic projections of magnetic vectors from Class I and Class II sample sites and fold test for the Mediano MPS. a95 confidence cones on mean polarity vectors are indicated. Diamonds are used to present mean vector directions. Close circles represent lower hemisphere data points and open squares represent upper hemisphere data.....................................................................................61 Figure 8. Summary diagram incorporating biostratigraphic, lithostratigraphic and magnetic polarity correlations, linking the 6 MPS to the global Geomagnetic Polarity Timescale (Harland et al., 1990), and to each other. For the Esera Valley MPS, both the preferred and an alternative correlation are shown (see text for discussion). Single point, or poorly constrained reversals are indicated as zones of ‘possible normal polarity’................................................... 63 Figure 9A. The Almazorre MPS......................................................................65 xi Figure 9B. Detailed sample location map for the Almazorre MPS. The location of this detailed map within the Ainsa Basin is shown in Fig. 5 .....................................................................................66 Figure 9C. Equal area stereographic projections of magnetic vectors from Class I and Class II sample sites for the Almazorre MPS. See Fig. 7C for a legend............................................67 Figure 10A. The Eripol MPS. See Fig. 9A for legend...............................69 Figure 10B. Detailed sample location map for the Eripol MPS. The location of this detailed map within the Ainsa Basin is shown in Fig. 2.....................................................................................70 Figure 10C. Equal area stereographic projections of magnetic vectors from Class I and Class II sample sites for the Eripol MPS. See Fig. 7C for a legend.................................................... 71 Figure 11 A. The Liguerre MPS. See Fig. 9A for legend.............................74 Figure 11B. Detailed sample location map for the Liguerre MPS. The location of this detailed map within the Ainsa Basin is shown in Fig. 5.....................................................................................75 Figure 11C. Equal area stereographic projections of magnetic vectors from Class I and Class II sample sites for the Liguerre MPS. See Fig. 7C for a legend............................................... 76 Figure 12A. The Esera Valley Composite MPS. See Fig. 9A for legend....................................................................................................... 78 Figure 12B. Detailed sample location maps for the Esera Valley composite MPS. The locations of these detailed maps are shown in Fig. 4. (i) The Santa Liestra Section..........................79 Figure 12B. (ii) The Meson de Pascual Section......................................... 80 Figure 12B continued, (iii) The Grustan Section............................................81 Figure 12C. Equal area stereographic projections of magnetic vectors from Class I and Class II sample sites for the Esera Valley composite MPS. See Fig. 7C for a legend................................................................................................................... 83 Figure 13A. The Lascuarre MPS. See Fig. 9A for legend......................... 85 xii Figure 13B. Detailed sample location map for the Lascuarre MPS. The location of this detailed map is shown in Fig. 4..................................................................................................................86 Figure 13C. Equal area stereographic projections of magnetic vectors from Class I and Class II sample sites for the Lascuarre MPS. See Fig. 7C for a legend.......................................... ..87 Figure 14A. Geohistory diagram showing sequential basement depth and the component of tectonic subsidence for the Esera Valley (preferred correlation) and the Mediano MPS. Note the very different rates of tectonic subsidence during Chron 20 time. Thickness of section, and the amount of tectonic subsidence are given in kilometers................................90 Figure 14B. Geohistory diagram showing sequential basement depth and the component of tectonic subsidence for the Eripol and Liguerre MPS during Escanilla Formation time. Note, in general, the similarity of data from both sections, although in detail the Liguerre data MPS shows initially higher sediment accumulation and subsidence. Thickness of section, and the amount of tectonic subsidence are given in kilometers................................................................... 92 Figure 15. Syntectonic geometries and thickness variations along the western flank of the Mediano anticline in the region of Samitier. A. Geologic Sketch Map (see Figure 2. for map location). B. Panel diagram shows variations in thickness immediately beneath the lower reef-derived talus breccia. The map shows lateral correlation of these units into, and the height at which they appear in the Mediano MPS. Note the marked change in thickness between the talus breccia level and RS1 in the two figures............................................. 94 Figure 16. Four block-diagram showing sequential paleogeographic reconstructions of the western SCU during................................................................................................................... 99 Figure 16 continued. B. Middle Lutetian reconstruction.......................... 101 Figure 16 continued. C. Late Lutetian reconstruction......................102 Figure 16 continued. D. Bartonian-Priabonian reconstruction..................................................................................................... 104 xiii Chapter 3 Figure 1. The southern Pyrenean Foreland Basin. Inset shows the approximate location of the study area, along the western flank of the South-Central Unit thrust system, and the simplified configuration of the important structural elements discussed in the text...................................................... Figure 2. Simplified geologic base-map of the western area, the Ainsa Basin or Buil Syncline, situated along the western oblique ramp of the South-Central Pyrenean thrust system. The location of villages within the study area, and the magnetostratigraphic traverses are shown, as are the important structural features within the Ainsa Basin (ALZ = Almazorre. ERI = Eripol. MED = Mediano. LIG = Liguerre).... Figure 3. The adopted lithostratigraphy nomenclature employed during this study. The western area stratigraphic framework is essentially that of Puigdefabregas (1975), and the scheme for the eastern area is modified after Cuevas Gozalo (1990).................................................................................... Figure 4. Simplified geologic base-map of the eastern study area within the Tremp ‘piggy-back’ basin. The location of the Lascuarre magneto-stratigraphic traverse is shown, as is the lacustrine interval represented by the Escanilla Limestone. The two villages adjacent to the section location are also shown....................................................... Figure 5A. Stratigraphic subdivision and nature of the lower Campodarbe Group sediments exposed within the Ainsa Basin. The summary magnetic polarity stratigraphy (MPS) constructed within the syncline is shown. Black represent times of normal magnetic field orientation, and white represents reversed directions. The MPS is correlated with the magnetic polarity time-scale (MPTS) of Harland et al. (1990) in Fig. 7........................................................ Figure 5B. Simplified geologic map of the southern Ainsa Basin showing the spatial distribution of the three members of the Escanilla Formation. Lack of exposure at higher stratigraphic levels prevents division of the ‘middle’ and ‘upper’ members away from the core of the Buil syncline......... Figure 6. Graph showing the varying clast population data for the three members of the Escanilla Formation, and the overlying Oligocene conglomerates exposed within the Ainsa Basin. Four general groups have been delineated. The extreme right column represents the Oligocene strata sampled within the center of the Buil syncline, away from local input of Guara limestone off the Boltana anticline.............................125 Figure 7. Cross syncline litho- and magneto-stratigraphic correlation of the Escanilla Formation across the study area. Lithologic correlations confirmed in the field are shown as solid lines linking the magneto-stratigraphic traverses. Correlations based on the comparison of the MPS's with each other and with the MPTS, are shown as dashed lines. Average undecompacted sedimentation rates within each member of the Escanilla system are shown, and are calculated using the ages of chron boundaries taken from Harland et al. (1990).......................................................................................... 128 Figure 8A. Paleocurrent data from the upper Sobrarbe deltaic and basal Escanilla coastal plain fluvial sediments of the Ainsa Basin prior to the marine transgression at -42.7 Ma. Numbers within the paleocurrent roses show the number of measurements taken during a given interval. The location of the rose is used to indicate the region where the data was collected. This is true of Fig. 8B & Fig. 8C also........................................................................................................................133 Figure 8B. Paleocurrent data collected within the ‘lower’ Escanilla Formation between the short marine transgression and the amalgamated conglomeratic sheet deposited at ~41 Ma.............................................................................................134 Figure 8C. Paleocurrent data from the ‘middle’ and ‘upper’ divisions of the Escanilla Formation. Note the reversal and diversion of flow to the south and south-east, parallel to and away from the emerging Boltana anticline......................................................138 Figure 9. Stratigraphic subdivision and nature of the lower Campodarbe Group sediments exposed south of the Isabena Valley (see Figs. 1 & 4 for map location). Note the gross coarsening-up trend, and the regular occurrence of thick lacustrine limestone units throughout the vertical sequence............................................................................................................... 139 xv Figure 10. Regional W-E cross-sectional summary of the inferred stratigraphic relationships discussed within this study. The spatial relationships of the magneto stratigraphic sections are shown (LAS refers to the Lascuarre section shown on Fig. 4). The simplified structural configuration is also presented......................................................143 Figure 11. Simplified geologic map of the Sierras Exteriores between the Boltana anticline and Arguis. Detail is taken from the map of Puigdefabregas (1975). Important structural features are shown, and the interfingering relationships of the deltaic and continental facies can clearly be seen....................................................................................................145 Figure 12. Summary of the paleocurrent information collected within the Ainsa Basin showing the sequential evolution of the Campodarbe drainage system in response to regional tectonic development.......................................................................151 Figure 13. Four sequential paleogeographic reconstructions of the South-Central and Western Pyrenees during late Eocene to Oligocene times. Structures active during each interval are shown as solid, whereas inactive structural elements are shown as the lighter dotted pattern. Differences in fluvial style are shown, and can be related to their structural position with respect to active or inactive thrust sheets. Data from Puigdefabregas (1975), Reynolds (1987), Jolley (1988), McElroy (1990), and Hogan (1992) are incorporated into this summary.................................................................159 Chapter 4. Figure 1. The Southern Pyrenean Foreland Basin. Box shows the approximate location of the study area, along the western flank of the South-Central Unit thrust system, and the simplified configuration of the important structural elements discussed in the text......................................................................... 168 xvi Figure 2. Simplified geologic map of the western area, the Ainsa Basin or ‘Buil Syncline’, situated along the western oblique ramp of the South-Central Pyrenean thrust system. The location of villages within the study area and the magnetostratigraphic traverses are shown, as are any important structural leatures within the basin (ALZ = Almazorre. ERI = Eripol. MED = Mediano. LIG = Liguerre)........................170 Figure 3. The adopted stratigraphy applied during this study. The stratigraphic framework is essentially that of Puigdefabregas (1975), but has been modified in the light of recent biostratigraphic and magnetostratigraphic data (Bentham and Burbank, in review, and Cuevas-Gozalo, 1990)......................................................................................................................171 Figure 4. Sketch of a general view of the Escanilla Formation exposures immediately south of the village of Olson. Note the wide multi-storey channel body wholly enclosed within fine-grained overbank material, and the interfingering geometries between the conglomerates and overbank siltstones along the left margin of the channel system...................................................................................................................173 Figure 5a. Generalized vertical lithologic log through the Escanilla Formation fluvial sequence. Interpretations of the various lithofacies are shown at the left side of the stratigraphic column, while grain-size and sedimentary structures are shown to the right.....................................................................174 Figure 5b. Comparative general vertical lithofacies sequences showing the difference between the ‘sandy’ braided river facies models of Miall (1977, 1978) and a general sequence through the Escanilla deposits. The three columns are drawn at the same vertical scale. The Donjek and South Saskatchewan type examples are redrawn from Miall (1978), using the legend presented in Figure 5a...............................................................................................................176 Figure 6. Example of the gravel-dominated, channel-fill lithofacies association. Sand/Gravel lenses are present above a strong basal scour surface cutting into reddened overbank siltstones. Shallow foresets are present within the gravel lens at the lower left corner of the photograph. Channel body is 4 m-thick................................................................................... 178 xvii Figure 7a. Example of the sand-dominated, channel-fill lithofacies association. Sand lenses show complex trough and planar cross beds developed above a strong basal scour surface cutting into reddened overbank siltstones. Channel body is 3 m-thick....................................................................................181 Figure 7b. Distal view showing external morphology of a sand-dominated channel fill sequence. Note the strong ribbon geometry, and the lateral wing, traceable into fine overbank facies. This channel represents two stages of aggradation and filling. Channel sequence is 4 m-thick...........................183 Figure 8. Example of the sedimentary geometries present at the lateral margin of a sand-dominated channel fill sequence. Sheet-splay sandstones may be traced directly into lower channel body (lower left sheets traced to right). Splays commonly show a massive, tabular form with planar bases, while laterally equivalent channels erode down into underlying overbank materials............................................................................185 Figure 9. Generalized block diagram showing the proposed depositional model. Note the low-sinuosity channel belt, internally braided in character, entirely enclosed by vegetated flanks and floodplain environments. Architectural geometries are shown in the vertical views. The presence of the adjacent anticline is specific to the Ainsa Basin (the Boltana Anticline), and is not an integral part of the general depositional model. Rapid rates of subsidence are considered to be most important...........................................194 xviii LIST OF TABLES Table 1. Summary of the three general demagnetization behaviors shown by the samples during pilot study analysis. The types of sediments and specific examples characterizing the three types of behavior are given................................................................................. 48 Table 2. Fisher Statistics on magnetic data from the 6 MPS sections. Both ‘in situ’ and structurally corrected data are given, as are the percentages of both Class I and Class II data sites generated during the bulk analytical study................................................... 60 ENCLOSURE Photogeologic map of the Escanilla Formation exposures within the southern Ainsa Basin. The amalgamated sheet sandstone and conglomerate unit present at the top of the lower Member of the Escanilla Fm. is located and traced across the basin. Geologic boundaries and important structures are also shown. (Map Pocket) xix ABSTRACT Regional reconstruction of the depositional and tectonic evolution of the western margin of the South-Central Unit (SCU) thrust system during middle-late Eocene time has been accomplished using detailed magnetostratigraphic correlation combined with established litho stratigraphic and sedimentologic information. The temporal framework supplied by the Magnetic Polarity Stratigraphy (MPS) was used to test and refine the existing lithostratigraphic scheme and to offer new syntheses of sequential Eocene basin development. Previous correlations across this region were hampered by marked lateral facies changes, poor biostratigraphic resolution, discontinuous exposure, and differential erosion of the Upper Eocene section beneath the overlying Oligocene conglomerates. The MPS time-control was also used to calculate rates of sediment accumulation and tectonic subsidence, thereby allowing the lateral comparison of these parameters across the western SCU margin. These data were then reconciled with local geologic map relationships in order to further document the Eocene structural development of the SCU western oblique ramp. The upper Eocene Escanilla fluvial system has been studied extensively within the southern Pyrenean basins, and detailed spatial reconstructions of fluvial environments and flow-patterns have been generated. These data show the strong influence of continuing structural development on fluvial deposition throughout the southern Pyrenean foreland basins. In particular, subsidence rate seems to have closely controlled the large-scale architecture of the Escanilla river system, with high rates of subsidence favoring the deposition of large volumes of fine grained immature pedogenic overbank material in association with coarse braided-stream channel sequences. Spatial variations in rates of tectonic subsidence, and therefore fluvial architecture, were largely a function of continued structural development, and the position of basin depocenters with respect to the SCU thrust system (hanging-wall, oblique margin, or footwall). xx CHAPTER 1 Introduction: Definition of this Research and Regional Pyrenean Summary 1 GOALS OF PROJECT This dissertation is the result of four years of study in the Spanish Aragonian Pyrenees. Nearly 7 months of fieldwork were completed from 1989-1991, usually during the early summer months of each year. The project was defined with the following main objectives: (i) To study the middle-late Eocene development of the western oblique ramp of the South-Central Unit (SCU) thrust system of the southern Pyrenees. This was to be done somewhat indirectly by describing the sedimentologic and geometric relationships between the middle-late Eocene depositional systems and adjacent structural elements within the Pyrenean foreland basin, most notably the N-S-trending Mediano and Boltana oblique ramp anticlines. (ii) To place this structural and sedimentologic development within an absolute temporal framework by constructing paleomagnetic reversal stratigraphies on the flanks of the anticlines, and to the east in the Tremp Basin. Correlation with the Harland et al. (1990) timescale would allow definition of temporally equivalent sedimentary environments across the oblique ramp system, and the sequential reconstruction of paleogeographic configurations during middle-late Eocene times. (iii) To complete a detailed sedimentologic analysis of the late Eocene Escanilla Formation fluvial sediments preserved across the oblique ramp system. This information will then be used to build a new depositional 2 model for braided stream fluvial systems flowing within rapidly subsiding tectonic environments, as opposed to many of the current models that have been derived from ephemeral braided streams draining regionally degrading terrains. (iv) To construct a number of simple models that may be used to reconstruct or predict the effects of on-going structural deformation within a foreland setting, paying particular interest responses in axial drainage systems as they encountered structures strongly oblique to the overall trend of the foreland. Such aspects of the later development of many modern and ancient foreland basin systems are not incorporated into quantitative basin-scale numerical models present in the geologic literature. IMPORTANCE OF THIS RESEARCH This study was partly designed to address regional stratigraphic correlations across an area showing complex, rapid lateral facies changes, and rather discontinuous exposures. The application of detailed field and photogeologic mapping, combined with the development of reliable magnetic polarity stratigraphies was intended to provide a framework in which to test the existing lithostratigraphic system (Garrido-Megias, 1968; Puigdefabregas, 1975; Nijman and Nio, 1975). The application of sequence stratigraphic concepts to the fill of the southern Pyrenean basins, in particular the western margin of the SCU (Mutti et al., 1985a; 3 1985b), while being a useful way to treat and group related packages of sedimentary facies, must still be placed and evaluated within in absolute temporal framework, if the complex structural and depositional history of the Pyrenean basins are to be reconstructed. These lateral facies changes seem to largely have been controlled by the position of important structural elements within the developing foreland basin, most notably the Mediano Anticline, which localized the regional coastline through much of early-middle Eocene time. The growth of this fold, along the western SCU oblique ramp, offers a chance to study the detailed interaction between a developing structure and the contemporaneous depositional systems within an active sedimentary basin. The description and timing of this development can yield additional information concerning the larger scale evolution of the SCU thrust system. Motion along this oblique ramp can then be reconciled with similar data from the frontal thrust system exposed to the southeast (Garrido-Megias and Rios, 1972; Mutti et al., 1985a; Nagtegaal et al., 1988; Burbank et al., 1992) in order to better understand and reconstruct thrust system development within the southern Pyrenees. Furthermore, the absolute ages and rates of structural and sedimentary development within the study area can be combined with the paleomagnetically constrained data from related studies along the Pyrenean system (Hogan et al., 1988; King Powers, 1989; Burbank et al., 1992; Hogan, 1992). Such a synthesis may then begin to compare the timing and rates of subsidence and sediment accumulation across spatially and structurally distinct sub-basins within the southern Pyrenean 4 foreland system. Given the strong diachrony in sedimentologic and structural regimes developed along the Pyrenean orogenic belt during on going collision, the temporally controlled reconstruction of regional relationships offers a possibly unique chance to study the large-scale variations in structural development along a complete foreland thrust belt and related basin. The Escanilla and related fluvial sediments are exposed across a number of these sub-basins within the southern Pyrenean system. Mainly flowing longitudinally along the subsiding foreland basin axis, the Escanilla rivers passed across a number of different late Eocene tectonic environments. In general, they have not been studied in great detail, and reconstruction of the whole drainage system, from proximal paleovalley sequences downstream to the distal coastal environments in the west has not been previously attempted. Once again, magnetic polarity stratigraphy offers a tool by which spatially distinct exposures of the Escanilla sequence may be linked and correlated. In this study, detailed fluvial lithofacies and architectural descriptions will be combined with the erected correlations in order to derive some general statements concerning the importance and effect of structural development on defining the nature of coarse fluvial systems within such dynamic tectonic environments. It is suggested that current braided stream facies and depositional models have been constructed in modern, regionally-degrading alluvial environments, most often in high-latitude cratonal settings dominated by glacial outwash streams (Moody-Stuart, 1966; Cant & Walker, 1976; Miall, 1977, 1978; Walker & Cant, 1984). Such river systems are not generally 5 representative of the fluvial systems that are preserved within the geologic record. Descriptions of the Escanilla system will be compared and contrasted with the existing vertical sequence facies models (Miall, 1977, 1978) derived from modern environments, and with examples of similar fluvial systems within the literature. In the light of these comparisons, I will build a new depositional model for coarse sandy braided river systems flowing within the actively subsiding tectonic regimes. Such systems have a far greater preservation potential than those presently flowing within regionally-degrading alluvial basins, and as such offer the most suitable devices for analyzing ancient coarse fluvial systems. They also represent our best chance of deriving predictive models of larger scale fluvial architectural relationships. REGIONAL PYRENEAN SUMMARY The Pyrenean orogenic belt trends more or less E-W and extends from the Cantabrian platform in the west, to the Provenge region of Southern France in the east (Fig. 1). It reaches a total length approaching 1500 km, and has an average width of some 200 km. Although the Alpine system developed coeval with the Pyrenees, major geologic structures are not continuous between the two orogenic belts. The Pyrenees were formed during a phase of late Cretaceous through Miocene convergence between the European plate and the Iberian microplate, and resulted from the 6 1 — J NORTHWARDS FACING STRUCTURES NORTHERN PYRENEES E 3 UPPER THRUST SHEETS C ) LOWER THRUST SHEETS SOUTHERN PYRENEES ] Figure 1. Structural sketch map of the Pyrenees showing the location of the ECORS deep seismic reflection profile (taken from Munoz, 1991). 7 closure of a small, short-lived Cretaceous transtensional basin (Puigdefabregas and Souquet, 1986; Roure et al., 1989). This main tectonic event led to the formation of a complex fold-and-thrust belt that developed north-vergent structures in the north part of the chain in southern France, and south-vergent structures in the southern Pyrenees of northern Spain. The development of a number of features related to this southerly-directed thrust system are the main focus of this dissertation. The following discussion of the important phases of both thrust belt and foreland basin evolution is based mainly in the region of the South-Central Unit (SCU) thrust system. The timing of similar deformational events was not synchronous along the length of the chain, rather it varied as a function of the lateral changes in crustal properties, as well as the general diachronous development of the orogenic belt (Puigdefabregas et al., 1991). Thrust Belt The Southern Pyrenees are made up of a series of basement- and cover-involved thrust sheets. Within the region of interest, in the central Pyrenees, these thrusts make up a complex system of a basement- involved antiformal stack flanked by imbricated foreland basin sediments (Fig. 2) (ECORS team, 1988). This Axial Zone Antiformal Stack (AZAS) involves only upper crustal rocks that were moved to the south along a 8 Figure 2. Strip geologic map and interpreted structural cross-section drawn along the ECORS deep seismic reflection profile (taken from Munoz, 1991). 9 north-dipping sole thrust as it developed (Munoz, 1991), and it was this wedge of crustal material that supplied a substantial part of the load that drove crustal flexure and foreland basin development. The northern part of the southern Pyrenean foreland basin flanking the AZAS in the central Pyrenees, is made up of cover thrust sheets, collectively referred to as the South-Central Unit (Seguret, 1972). Three main thrust sheets are preserved. From north to south, they are the Boixols, Montsec, and Sierras Marginales (Figs. 2 and 3). The thrust sheets have been transported southwards over autochthonous Paleogene and thinned Mesozoic sediments that sit directly upon basement of the Ebro basin (Camara and Klimowitz, 1985). The Boixols thrust sheet is the northernmost cover thrust, located to the north of the Tremp basin. The frontal structure of the thrust sheet is a complex feature buried and onlapped by syntectonic late Cretaceous sediments of the Aren sandstone formation (Garrido-Megias and Rios, 1972). The northern thrust contact between the Boixols and the AZAS is a passive-roof backthrust, related to stacking of basement thrust sheets below (Fig. 2) (Munoz, 1991). Motion on the Boixols thrust represents the structural inversion of an early Cretaceous extensional basin (Munoz, 1991). To the south, the Montsec thrust sheet shows a relatively simple synclinal structure forming the base of the Tremp Basin. The southern margin of the thrust sheet is defined by the arcuate trace of the Montsec thrust (Figs. 1 and 2). Based on syntectonic geometries along the flanks of the sheet, thrust emplacement initiated during the Ypresian stage (early 10 Eocene) (Williams and Fischer, 1984; Mutti et al., 1985b; Farrell et al., 1987). Out-of-sequence motion can also be shown to have occurred during a late Eocene-Oligocene phase of thrusting (Verges and Munoz, 1990). Estimates of minimum displacement along the Montsec thrust based on both surface cut-off, and well data suggest about 10 km of shortening has been accommodated since Ypresian fault initiation (Munoz, 1991). The southernmost Sierras Marginales thrust system is present between the Montsec thrust and the southern Pyrenean thrust front (Pocovi, 1978; Martinez and Pocovi, 1988). Made up of a several small imbricates (Fig. 2), the Sierras Marginales preserve syntectonic relationships between thrusts and late Eocene-early Oligocene continental sediments . Additionally, the later Oligocene section is also affected by thrusting. Stratigraphic and structural relationships suggest the onset of thrust motion during Eocene time, with later deformation occurring in latest Eocene-Oligocene times. Later deformation was partially the result of break-back sequence thrusting synchronous with continued thrust front development (Verges and Munoz, 1990). Estimates of total displacement in the Pyrenean system vary greatly. Based upon magnetic anomaly patterns in the Bay of Biscay, Boillot (1984) suggested 120-150 km of shortening had occurred in a NW-to-SE direction. Using balanced cross-section construction prior to the ECORS deep reflection seismic profile, N-S shortening estimates ranged from 115 km (Deramond et al., 1985) to 106 km (Williams and Fischer (1984), to as 11 low as 55-80 km (Seguret and Daignieres, 1986), and generally increased from W-to-E. Roure et al. (1989) presented the first orogen-scale balanced cross- section analysis using ECORS data, and suggests at least 120 km of shortening must have occurred, with 100 km or so accommodated within the lower crust. Munoz (1991) modified this initial ECORS profile interpretation and calculated a minimum total shortening estimate of 147 km (Fig. 3). This figure may be partitioned into 2 major components: 112 km of shortening was accommodated in the AZAS, while the remaining 35 km was accommodated by the imbrication of the basin fill (Munoz, 1991). This total value is in good agreement with the northward motion of the Iberian plate with respect to the European plate based on sea-floor anomaly data (Boillot, 1984; 1986). Munoz (1991) also noted a marked discrepancy between the length of upper crust versus lower crustal material preserved within the orogenic belt. Approximately 110 km of middle and lower Iberian crust is missing if the orogen is restored (Fig. 3), and the normal thickness crust presently beneath the foreland basin and external thrust system is reconciled with this restored section (Munoz, 1991). Early Cretaceous extension could account for this lack of middle and lower crustal material, but the degree of crustal thinning is difficult to determine with any confidence. One alternative explanation would be delamination of the upper crust from the remaining crustal pile, with the subsequent subduction of the underlying material (Daignieres et al., 1989; Munoz, 1991). This idea is supported by seismic reflection data from a 12 Jfci Z m . M Figure 3. Crustal balanced and restored structural cross-sections along the ECORS deep seismic reflection profile (taken from Munoz, 1991). 13 number of regions across northern Europe where extensional normal fault systems are not seen to penetrate to lower crustal depths; instead, they sole into lower-middle crustal detachments or discontinuities (Cheadle et al., 1987; Le Pichon and Barbier, 1987). It is suggested, therefore, that 110 km of lower crustal material remained attached to the lithospheric mantle and was subducted to the north, beneath the European plate (Daignieres et al., 1989; Munoz, 1991). The detached upper crustal rocks, along with their pre-Tertiary cover sequence were incorporated into the AZAS. Foreland Basin The South Pyrenean foreland basin is a triangular-shaped basin, located to the south of the main Pyrenean mountain system, and is bounded to the SE and SW by the Catalan and Iberian ranges respectively. The sedimentary fill of the basin can be split into a number of stages that were closely dependent upon the sequential development of the nearby orogenic wedge (Fig. 4) (Puigdefabregas et al., 1991). Stage I (Late Santonian-Maastrichtian) Extensional normal faults related to early Cretaceous rifting were inverted during this phase, forming the first active thrust faults in late Santonian time (Boixols thrust sheet) (Fig. 4). A deep, E-W-elongated turbidite trough formed ahead of the advancing thrust front, with thrust 14 APTIAN 114 Ma CENOMANIAN 92 Ma Stage CAMPANIAN 80 Ma 20 Stage II PALEOCENE 68 Ma 37 Stage LUTETIAN 47 Ma 1* 72 UPPER ' OLIGOCENE 30 Ma (present section) Stage IV 147 km Figure 4. Series of partially restored cross-sections showing the early Cretaceous extensional configuration of the Pyrenean system, and the subsequent thrust sheet geometries at a crustal scale during the progressive stages of foreland basin evolution (taken from Puigdefabregas et al., 1991). 15 geometries apparently closely controlling turbiditic deposition (Puigdefabregas et al., 1991). The narrow width of this initial foreland basin was probably due to the loading of warm, tectonically-thinned crust, thereby reducing the flexural wavelength of the sedimentary basin (Puigdefabregas et al., 1991; Karner et al., 1983). Subsidence may have been further enhanced due to the decay of any residual thermal anomaly inherited from the preceding early Cretaceous rifting event. The subsequent shallow marine fill was intermittently affected by the continuing thrust deformation (Garrido-Megias, 1973). These Maastrichtian sediments have been subdivided into a number of unconformity-bounded depositional sequences whose facies relationships, depocenters and angular geometries can be related to continued motion on the Boixols thrust (Nagtegaal et al., 1983; Simo and Puigdefabregas, 1985; Puigdefabregas and Souquet, 1986). Stage II (Late Maastrichtian-Paleocene) At this time, most of the south Pyrenean basin saw shallow marine depositional environments interfingering with continental fluvial and lacustrine sedimentation along the foreland basin axis. Conglomerates lie unconformably across the Boixols thrust and were deposited synchronous with piggy-back-style thrust propagation southwards into the south-central Pyrenees. Thickness changes and facies relationships within these continental sediments bear testimony to their syntectonic nature with respect to newly formed thrusts, most notably the Montsec thrust system (Fig. 4) (Puigdefabregas and Souquet, 1986). Further to the west in the 16 western Pyrenees, deeper water deposition was occurring on top of the earlier Cretaceous turbidites. This illustrates the along strike variability within the Pyrenean system, with Stage I type environments of the central and eastern Pyrenees persisting much later in the western Pyrenees. This diachrony is also present within the later phases of Pyrenean evolution. Stage III (Early and Middle Eocene) During this phase of development, the spatial arrangement of deep marine basins and contemporaneous continental and shallow marine environments was strongly controlled by large scale thrust sheet geometries. Piggy-back thrusting propagated further southwards into the subsiding foreland basin, and this is expressed by a thinning of the sedimentary fill within the associated thrust-sheet-top basins (Fig. 5) (Puigdefabregas et al., 1991). In turn, the longitudinal variation in thrust trajectory and geometry in the SCU of the southern Pyrenees, of particular interest for this study, was controlled by the position of Mesozoic extensional faults (Seguret, 1972). Lateral and oblique ramps in the main SCU thrust sheets (Boixols, Montsec and Sierras Marginales from north to south) initiated over these inherited structures, and they exerted strong influences on the facies patterns within the piggy-back depositional systems. The piggy-back Tremp basin trended parallel to the strike of these main structures, and showed enhanced subsidence to the north (Fig. 5). Passive roof backthrusting, and out-of-sequence thrusting within 17 Q . O ) C D 2 S’ • 2 (/> w Figure 5. Detailed partially restored cross-section of the southern Pyrenees during stage III of the foreland basin evolution. Section corresponds to the southern portion of the ECORS profile. The syntectonic sediments of this stage are shown shaded, and previous deposits are shown as stippled. UN and LN = Upper and Lower Nogueres units of the AZAS respectively. Taken from Puigdefabregas et al. (1991). 18 the AZAS at this time, as well as increasing the adjacent thrust load, served to generate greater relief, and increase the rate of erosion and sediment supply to the deforming foreland (Puigdefabregas et al., 1991). Coarser alluvial fan deposits dominate this northern margin of the basin, and can be shown to have prograded southwards, away from the emerging thrust system. Further to the south, these give way to finer- grained, axially-draining fluvio-deltaic systems that passed westwards into marine environments within the strongly subsiding Jaca basin (Mutti et al., 1985a). The longitudinal transition from continental to marine environments was controlled throughout this stage by the position of the western oblique boundary of the SCU thrust system, feeding clastic sediments from shallow platform of the Tremp basin to the west. The Jaca basin at this time showed deep marine turbidite sedimentation juxtaposed against widespread platform carbonate deposition along the basin’s southern margin (Puigdefabregas, 1975). This platform shows evidence for synsedimentary deformation by developing frontal thrust structures, as the Jaca basin began to detach as a piggy-back system (Almela and Rios, 1951; Puigdefabregas, 1975). Stage IV (Late Eocene-Oligocene) Largely characterized by the infilling of the deeper turbiditic Jaca Basin by the westward progradation of deltaic and continental sediments, this stage continues to show evidence for the control of sedimentation by the geometry of the SCU thrust system. Coarse alluvial fan deposition was occurring in the hanging-wall of the SCU while axial fluvial systems carried 19 sediments downstream across the oblique ramp zone to the west. These systems were mainly fed by erosion within the newly uplifted hinterland, but also were derived from local uplifts along the southern and eastern basin margins. The southern Pyrenean basin was closed to marine sedimentation at this time, and became internally draining. The basin center saw widespread evaporitic deposition at the end of the Eocene stage (Saez and Riba, 1986), which was then succeeded by a system of marginal alluvial fans feeding into a central lacustrine basin. This configuration persisted up to the end of Oligocene time The structural evolution of the orogenic belt at this time was largely controlled by the development of the AZAS, as well as out-of sequence motion on the SCU fault system (Fig. 6). As a consequence of this out-of sequence thickening in the internal and external parts of the thrust system, the overall taper of the Pyrenean orogenic wedge increased substantially in order to allow further southward translation (Davis et al., 1983; Puigdefabregas et al., 1991). This increase in orogenic relief is expressed sedimentologically by a strong upwards coarsening in contemporaneous late Eocene-Oligocene fluvial systems flowing along the axis of the foreland and related piggy-back basins (Puigdefabregas et al., 1991). The larger scale geometry and basin infilling of the south Pyrenean foreland basin seems to have been largely a function of surficial and subcrustal forces leading to tectonic subsidence. Subcrustal effects resulted both from the continuing thermal evolution of the extended Iberian 20 E C M Q. CL C E LU m cc CL CL O ) T3 > C D O ) C O C O Figure 6. Detailed partially restored cross-section of the southern Pyrenees during stage IV of the foreland basin evolution. Section corresponds to the southern portion of the ECORS profile. As in Fig. 5, syntectonic sediments of this stage are shown shaded, and previous deposits are shown as stippled (A = Aren Group. T = Tremp Group. AM = A gerand Montanyana Groups). Taken from Puigdefabregas et al. (1991). 21 crust, and from the flexure of that crust in response to the gravitational pull exerted by the subducting middle-lower crust and lithosphere (Molnar and Lyon Caen, 1988). Surficial loads were directly related to thrust emplacement, most importantly the AZAS, but also within the external parts of the thrust system. Although not contributing a major tectonic loads, these structures exert important local controls on facies distributions and drainage patterns at a sub-basin scale. The changing width of the foreland basin, and the different stages of deep marine or continental sedimentation can be related to the interaction of these surficial or sub-crustal loads (Puigdefabregas et al., 1991). The deep, narrow trough of Stage I was the result of loading thinned, warm crust, with subsidence being further enhanced by thermal cooling and contraction of this warmer material. The wider basin, with the more uniform facies distribution of Stage II represents a phase of crustal thickening, and the emplacement of surficial thrust loads. No major subcrustal load was important at this time. The return to elongate, deep troughs within Stage III and the rapid increase in tectonic subsidence at this time, are indicative of the onset of subduction of Iberian crust below the European Plate, and the development of the AZAS. This represents a rather significant surficial load. Continued growth of the antiformal stack in Stage IV inundated the basin with coarse clastic material. Although the surficial load increases through this stage, tectonic subsidence can be shown to decrease (Burbank et al., 1992). One must, therefore, invoke a subcrustal mechanism such as detachment of the subducting slab and the 22 loss of any slab-pull forces, or the thermal equilibration the subducted material, in order to decrease regional subsidence rates. The previous discussion of structural evolution and estimates of shortening within the central Pyrenees hinges on the assumption that regional thrust transport directions were essentially orthogonal to the overall strike of the orogenic belt. Balanced cross-sections have been offered by a number of authors that yield widely variable estimates of total shortening. Nijman (1989) synthesized a number of different datasets and offered a rather different interpretation of regional structural evolution. Instead of the linked thin-skinned thrust system favored by many authors (Deramond et al., 1985; Williams and Fischer, 1984; Seguret and Daignieres, 1986; Munoz, 1991), Nijman (1989) suggests the possible importance of gravity-driven tectonics and rotation during thrust sheet emplacement. Both mechanism largely invalidate the application of cross- section balancing methods. After palinspastic restorations of the Eocene depositional systems, Nijman (1989) saw a distinct misfit between the central and eastern Pyrenean Basins (Fig. 7). Most importantly, while the basins are juxtaposed and in near perfect alignment, both area show facies patterns that appear to suggest that both basins opened towards marine conditions to the west during early Eocene time (Rosell and Robles, 1975; Nijman, 1989). The proposed model suggests an en- echelon NW-SE arrangement of elongate synclinal basins, plunging towards to NW (Fig. 8). Subsequent anticlockwise rotational emplacement of underlying thrust sheets placed the early Eocene basins in structural 23 A LATE YPRESIAN MISFIT 50 km Pedraforca CSP Tremp Basin Ripoll Basin \ • piniine at Barbastro gypsum ridge Figure 7. Apparent misfit between the Tremp and Ripoll basins during Late Ypresian time. Arrows indicate sediment dispersal directions. Shaded areas represent deep marine turbiditic basins. Numbers refer to the distances from a defined pin-line within the northern Ebro basin. Dashed lines delineate the margins of the SCU and Pedraforca thrust sheets. Taken from Nijman (1989). 24 C /5 C O C O LLI Q C CL C O C O LLI Q C Q _ C O Q C C O LLI C O Q C Figure 8. Tectono-sedimentary model for the Pyrenean orogenic basins given by Nijman (1989). From left to right the figures show the sequential evolution of the South-Central Pyrenees in schematic map form. Structural regime and significant structural events are indicated at the base of the figure, as well as the interpreted phases of basin evolution. Taken from Nijman (1989). 25 alignment, leading to the apparent mis-match of sedimentary facies belts. The palinspastic restoration of the southern Pyrenean system shown by Nijman (1989) results in a north Iberian shelf showing a series of indentations and headlands of about 100 km wavelength (Fig. 8). The NW-SE orientation of these belts are related to compressive structures formed on a shelf previously subject to extensional normal faulting and sinistral motion between Iberia and southern France. The development of these structures mark the transition from a trans-tensive to a transpressive regime during latest Cretaceous time. Transpression was succeeded by more or less pure compression during early Tertiary time, and the subsequent rotational emplacement of the central and eastern Pyrenean thrust sheets (Fig. 8). However, this rather novel and provocative model for Pyrenean development is not generally supported by independent estimates of structural rotation derived from paleomagnetic study (Fig. 9) (Burbank and Puigdefabregas, 1988; King Powers, 1989; Dinares et al., 1991). Although structural rotations are present along the oblique margins of the South-Central Unit, these can be dismissed as only local effects related to translation of the thrust sheets over these oblique structures. In the southern Pyrenees, clockwise rotations are generally related to western (dextral) oblique margin, and counterclockwise rotations are seen along the eastern (sinistral) margin (Dinares et al., 1991). In the central portion of the SCU, away from ramp-related deformation, no consistent 26 Figure 9. Summary of the paleomagnetic data showing rotational information from the central and eastern Pyrenees. Arrows represent the local mean declinations. Taken from Dinares et al. (1991). 27 differential rotations are seen indicating that no significant counter clockwise rotation occurred during thrust sheet emplacement (Fig. 9) (Dinares et al., 1991). METHODOLOGY This research represents a multi-disciplinary study, employing the collection and analysis of a number of different types of geologic data. As such, a wide number of field and laboratory techniques were used, and will subsequently be described. Given the format of this dissertation, in that each of the 3 major chapters represents a presentation of a coherent piece of research with rather different focus, no single summary of the methods and analytical techniques employed by this study is offered. Instead, the relevant techniques are discussed within each individual chapter, in order to establish a framework in which the subsequent results may be evaluated. 28 CHAPTER 2 The Chronology of Middle and Late Eocene Deposition and Deformation across the Western Margin of the South-Central Unit, Southern Pyrenees, Spain 29 ABSTRACT Detailed correlation of Eocene sedimentary environments across the western oblique ramp of the SCU thrust system has traditionally been hampered by incomplete outcrop data, erosion at the base of the overlying Oligocene conglomerates, and rapid lateral variations in sedimentary lithofacies. With the application of well-constrained paleomagnetic reversal stratigraphies, this study: (i) places the middle to late Eocene tectonostratigraphic development of this region in an absolute chronologic framework; (ii) evaluates existing sequence stratigraphic analyses, where a reliable chronologic framework was not previously available; (iii) constrains the sequential backstripping of Eocene sedimentary systems, and the calculation of rates of sediment accumulation and tectonic subsidence; and (iv) revises earlier paleogeographic reconstructions of these complex, spatially variable Eocene sedimentary systems. Magnetostratigraphic dates from the Tremp and Ainsa basins, indicate the Mediano anticline developed during a rapid phase of growth and rotation from -48-42 Ma. An abrupt marine transgression between -45-43 Ma affected both marine and continental sedimentary systems. The transition from marine to non-marine sedimentation across the oblique ramp occurred at -42.9 Ma, with the onset of deposition of the Escanilla fluvial strata. Eocene fluvial sedimentation persisted until the onset of uplift and deformation of the Sierras Exteriores soon after -3 6 Ma. In the light of regional magnetostratigraphic correlations and the improved time-control this technique offers, the existing sequence 30 stratigraphic analysis of the region requires revision. Previous application of sequence stratigraphic concepts to this structurally dynamic environment has led to misleading interpretations concerning the importance of relative changes in sea-level in influencing Eocene deposition within the South Pyrenean foreland basin. Lateral variations in sediment accumulation have been strongly controlled by local tectonic development, instead of by regional changes in relative sea-level. INTRODUCTION Large-scale regional syntheses of depositional processes within actively deforming tectonic environments can significantly add to the description and understanding of both the structural and sedimentological evolution of the particular areas under investigation. However, such regional assimilation of stratigraphic information is strongly dependent upon the confidence with which geographically separated portions of a study area may be related directly to one another. In wholly marine basins, the presence of a detailed biostratigraphic zonation scheme can serve as a useful relative chronologic framework within which to correlate over large lateral distances. In internally deforming sedimentary basins, however, one often sees the complex juxtaposition of continental and marine depositional settings, which often diminishes the ability to use detailed biostratigraphic information as a correlation tool. Furthermore, in continental sedimentary basins, particularly those lacking palynological or tephrochronological data, reliable regional reconstruction of continental 31 depositional environments across adjacent structural provinces may be more or less impossible. In this study, we have used magnetic polarity stratigraphy (MPS) to erect an absolute chronologic framework within which we are able to regionally reconstruct the tectono-sedimentary history of two adjacent Tertiary sub-basins within the Southern Pyrenean foreland system of Northern Spain. Marked changes in middle Eocene sedimentary environments and associated marine and continental lithofacies were localized across the western oblique ramp of South-Central Unit (SCU) thrust system, as the thrust system developed (Fig. 1). Reliable stratigraphic correlation of coeval sedimentary packages from the western Tremp Basin, across the Mediano anticline, into the Ainsa Basin (Fig. 1) has been hampered both by these rapid lateral facies changes and by erosion of the middle-late Eocene stratigraphy at the base of the overlying early Oligocene conglomeratic succession. Several different studies (Mutti et al., 1985a; Crumeyrolle and Mutti, 1986; Mutti et al., 1988; McElroy, 1990) have attempted to apply sequence stratigraphic concepts (Vail et al., 1977) to the well exposed surface outcrop information in order to subdivide the Eocene stratigraphy into a number of unconformity-bounded depositional sequences within the Southern Pyrenean basins. Using aerial-photography and detailed surface mapping, Mutti et al. (1988) divided the middle-late Eocene sediments into two main sequences referred to as the Santa Liestra (late Ypresian-Lutetian), and Campodarbe (late Lutetian-Priabonian) Groups. 32 C J C/5 c u c a CD CTJ II II II c / > a > O CD ca co m c u Figure 1. The southern Pyrenean Foreland Basin. Inset map shows the approximate location of the study area within the Pyrenean system. The coverage of subsequent figures is also indicated. The important structural elements of the South-Central Unit (SCU) thrust system are also identified 33 On the basis of recently created magnetic polarity stratigraphies, this study: (i) assesses the correlations that have been previously made; (ii) offers absolute ages for biostratigraphically dated strata within the middle- late Eocene lithostratigraphy of the western SCU; and (iii) evaluates the usefulness of the sequence-stratigraphic analyses. In the light of these new correlations, we discuss the middle to late Eocene depositional history across the western edge of the SCU. Absolute dates derived from the correlation of our local magnetic polarity stratigraphies (MPS) with the magnetic polarity time scale (MPTS) of Harland et al. (1990) have been used to calculate rates of sediment accumulation and tectonic subsidence spatially across the oblique ramp system. These data are used to evaluate the current models of development and timing of deformation across the western margin of the SCU thrust system. STRUCTURAL SETTING The Pyrenees are an E-W-trending mountain belt located at the NE part of the Iberian Peninsula, separating Spain from southwestern Europe (Fig. 1). The orogenic belt developed during a phase of late Cretaceous- Miocene convergence and limited northward underthrusting of the Iberian plate beneath Eurasia (Munoz, 1991). Collision began earlier in the east than in the west, and the onset of thrust deformation shows a strong diachrony as one traverses along the chain from east-to-west. In the southern-central Pyrenees, basins developed during Paleocene to early Eocene times in front of southerly-translating South-Central Unit 34 (SCU) thrust sheets, partly in response to thrust loading, and partly due to subduction-related flexure of the down-going Iberian Plate (Munoz, 1991). Collectively referred to as the South Pyrenean basin (Puigdefabregas, 1975), this region began to became increasingly partitioned or compartmentalized during the early Eocene epoch. This occurred as proximal parts of the foreland basin became incorporated into the developing South Pyrenean thrust wedge. A series of thrust-sheet top or piggyback basins, formed in the hanging-walls of SCU thrusts. The Tremp basin was formed in the hanging-wall of the Montsec thrust sheet, and the Ager Basin sits atop the younger Sierras Marginales thrust sheet (Fig. 1). Along the western margin of the SCU thrust system (Fig. 1), the Mediano and Boltana oblique ramp folds initiated during early Eocene time, at the onset of thrust motion and piggyback basin formation, and both have long protracted histories (Camara and Klimowitz, 1985; Farrell, Williams and Atkinson, 1987; Reynolds, 1987; Nijman, 1989; McElroy, 1990). Growth of Mediano anticline probably ended by Bartonian times, whereas Boltana anticline was possible active into the Priabonian Stage (Puigdefabregas, 1975). Sitting directly between these two folds is the Ainsa Basin (Figs. 1 and 2). The northern margin of this basin was first structurally defined during Ypresian times and is currently expressed as the E-W-trending Cotiella thrust system exposed along the foot of Pena Montanesa (Nijman, 1989). Deformation in the SCU continued into Oligocene times with out-of-sequence thrusting associated with an increase in or maintenance of thrust-wedge taper (Munoz, 1991). 35 This study attempts to reconstruct within a reliable chronologic framework the middle-late Eocene development of this western oblique ramp system, through the use of coexisting sedimentary systems as proxies of structural evolution. Thickness variations, changes in tectonic subsidence, and fades changes within time-equivalent strata were all largely controlled by the structural growth of the Mediano and Boltana oblique ramp anticlines. We will, therefore, evaluate the evolution and development of this portion of the SCU thrust system based on this largely stratigraphic information. STRATIGRAPHIC FRAMEWORK Due to the spatial changes in middle Eocene sedimentary environments across the study area, the correlation of shallow marine platform and continental sediments of the Tremp Basin with the deeper marine environments of the Ainsa Basin has traditionally been difficult. The stratigraphic framework prior to this study is summarized in Fig. 3. We adopt a modified version of the system given by Cuevas Gozalo (1990) for the eastern part of the study area located within the Tremp Basin including data from Mutti et al. (1988). For the western region, west of the Mediano anticline in the Ainsa Basin, we have slightly modified the terminology of Puigdefabregas (1975) in the light of more recent information (De Federico, 1981; Reynolds, 1987). In the Ainsa Basin (Fig. 2), the base of the middle Eocene section is present within the Hecho Group Turbidites (Mutti et al., 1972). These 36 Ainsa Arcusa • Mediano cd Pre-T MED B ; • Eripol Olson Escamlla m Barcabo k f t iifi Oligocene (Collegats) iirr Ml| Fig. 15 Conglomerates Pre-T r i r i i ] Pre-fold Hecho Grp. Eocene Sobrarbe Escanilla Magnetic Sequence Turbidites Lst. Fm. Fm. Section Figure 2. Simplified geologic map of the western area; the Ainsa Basin along the western oblique ramp of the SCU thrust system. The locations of the four magnetostratigraphic sections within the basin are indicated (ALZ = Almazorre. ERI = Eripol. MED = Mediano. LIG = Liguerre). 37 sediments have been variably assigned to the Campanue Fm. (Garrido- Megias, 1968), the San Vicente Fm. (Van Lunsen, 1970), the Montanana Group (Nijman and Nio, 1975), and the Santa Liestra Group (Mutti et al., 1988). Progressively onlapping the western flank of the Mediano anticline, these sediments were deposited during an important phase of fold growth. The siliciclastic Hecho Group turbiditic system is overlain by the shallow marine and deltaic facies of the Sobrarbe Fm. (De Federico, 1981) which thickens from south to north, forms a northward-prograding wedge of siliciclastic and carbonate sediments within the Ainsa Basin synclinal axis, and inter-fingers eastwards with the shallow marine reef and platform carbonate sediments (Fig. 3) of the Puy de Cinca Fm. (Garrido-Megias, 1968) (the Grustan Limestones of Cuevas Gozalo et al., 1985). These units shoal up and are overlain more or less conformably by the Escanilla Formation: an late Eocene fluvial system of coarse braided-stream character (Garrido-Megias, 1968). Mutti et al. (1988), however, suggested that the contact between the two is in fact an unconformity related to an abrupt relative sea-level fall at -39.5 Ma. Following their correlation of the Escanilla Formation with the Pamplona marls to the east in the Jaca Basin, Mutti et al. (1988) suggested that deposition of the Escanilla Fm. took place during Priabonian time and ended with uplift and erosion at about 38 Ma. The Escanilla sediments constitute a major (-1 km-thick) upwards- coarsening megasequence best exposed and developed within the Ainsa Basin. The complex internal sedimentologic and architectural relationships within the Escanilla Formation are discussed in detail in 38 Age Oligocene Bartonian- Priabonian Lutetian Ypresian WESTERN AREA Ainsa B asin R IO C IN C A Oligocene Conglomerates q_ , < 5 Q _ Q ■ § Escanilla < 3 - Fm C O O P uy d e Cinca Lst c /3 0 3 -2 l— CL r: o < 5 o sz o a ) X c o o > c o cr c 0 3 c o c o co Castisent ' 1 Sequence CD C O ’ ■ + -* c < o c X C D EASTERN AREA Tremp B asin ESERA IS A B E N A Oligocene Conglomerates Escanilla Fm Escanilla P uy de Cinca Lst Capeila & Castisent Fm. Figure 3. The pre-existing lithostratigraphic nomenclature of the western SCU This figure incorporates data from Puigdefabregas (1975), De Federico (1981), Reynolds (1987), Mutti et al. (1988), and Cuevas Gozalo (1990). 39 Bentham et al. (1991) and Bentham et al. (in prep). Early Oligocene alluvial conglomerates of the Campodarbe system (Puigdefabregas, 1975) unconformably overlie the Escanilla Formation across the study area (Reynolds, 1987). Along the flanks of the Mediano and Boltana anticlines, the Escanilla section has been completely eroded, and the Oligocene conglomerates can be seen directly overlying Lutetian sediments of the Sobrarbe and Guara Limestone Formations (Fig. 2). To the east, the equivalent middle Eocene strata within the Tremp Basin are dominantly continental or shallow marine in origin (Figs. 3 and 4). The Perrarua Fm. represents the lateral, marine equivalent to the Lower Campanue fan-delta conglomerates further to the north (Garrido- Megias and Rios, 1972; Crumeyrolle, 1987). Overlying and in part equivalent to both the Perrarua and upper Campanue formations, the Capella Fm. is interpreted to represent deposition on a tidally-influenced alluvial plain, down-system from the more proximal Campanue conglomerates (Fig. 3) (Cuevas Gozalo and De Boer, 1989; Cuevas Gozalo, 1990). This sedimentary system represents the Upper Montanana Group of Nijman and Nio (1975), later renamed the Santa Liestra Group by Mutti et al., (1988). Crumeyrolle (1987) discussed the internal relationships among the Campanue, Perrarua and Capella Formations along the Esera Valley, and was able to divide the section into 4 unconformity-bounded depositional sequences (SL1-SL4). Each begins with an abrupt seaward shift of shallow water or continental facies belts. The lower two sequences (SL1 and SL2) generally show near-shore sandstones and shelf mudstones (Perrarua Fm.) interfingering with 40 northerly-derived fan-delta conglomerates (Lower Campanue Fm.). The SL3 sequence documents a narrowing of the marine basin, related to uplift of the basins southern margin, while SL4 is wholly continental and corresponds to the Capella Fm. (Crumeyrolle, 1987) as defined by Garrido-Megias (1968). The upper boundary of the Capella Fm. (SL4) is marked by an abrupt marine transgression (the Biarritzian transgression of Puigdefabregas, 1975). This caused the local drowning of the Capella alluvial plain and the deposition of the Pano Fm. barrier-island complex (Figs. 3 and 4). Three separate barrier sequences are developed that show progressive eastward-stepping as the transgression continued (Cuevas Gozalo et al., 1985). The uppermost ‘Grustan’ barrier is present within the Esera Valley section. The base of the Pano Fm., also corresponds to the boundary between the Santa Liestra and succeeding Campodarbe Group depositional sequences defined by Mutti et al. (1988) (Fig. 3). Further to the east, south of the Isabena Valley (Fig. 4), this transgression is believed to be represented by a laterally extensive lacustrine limestone horizon at the base of the Escanilla formation (Cuevas Gozalo and De Boer, 1989). This “Escanilla Limestone” lies between the alluvial sediments of the Capella and Escanilla Formations (Fig. 3), reaches a thickness of up to 30 m, and may contain up to 16 separate lacustrine intervals (Nickel 1982). It has been suggested that a significant hiatus exists between the Escanilla Limestone and the overlying fluvial deposits (Cuevas Gozalo, 1990). However, no discrete omission, 41 r r r r / / / Capella Fm. Perrarua Fm. Cam panue Fm. Unconform ity Normal Faults fgrfrfl S 3 Oligocene Conglom erates Escanilla Fm. Pano & Puy de Cinca Fms. M agnetostratigraphy Figure 4. Simplified geologic map of the western Tremp Basin. Villages and river valleys mentioned in the text are shown, as are the locations of the four magnetostratigraphic traverses taken within this portion of the study area (SLA = Santa Liestra. MDP = Meson de Pascual. GRU = Grustan. LAS = Lascuarre). The positions of more detailed sample location maps for the magnetostratigraphic sections are also indicated. 42 erosion surfaces, or angular relationships are obvious. It should be noted that the base of the Escanilla Formation east of the Mediano anticline is placed at the base of the Escanilla Limestone purely for convenience of lithostratigraphic division. Correlation of these units across the study area is hampered by incision and erosion at the base of the overlying Oligocene conglomerates. The development of this basal paleovalley sequence along the Esera Valley caused the erosion and removal of the late Eocene strata, leaving the Oligocene conglomerates sitting directly upon middle Eocene sediments of the Capella and Puy de Cinca formations. SAMPLING AND DATA ANALYSIS Section Description, Measurement, and Sampling Eight magnetostratigraphic sections were described and measured across the study area. Four sections were located east of the Mediano anticline within the western Tremp Basin (Fig. 4). Three of these, Santa Liestra (SLA), Meson de Pascual (MDP), and Grustan (GRU) constitute lower, middle and upper parts respectively of the Esera Valley composite section. The other four sections were measured and described along strike from each other, within equivalent strata of the Escanilla Fm. exposed within the Ainsa Basin (Fig. 5). Stratigraphic thickness was measured using a Jacob’s staff and Abney level. Repeated measurements suggest that thicknesses are reproducible to within ±5%. 43 Figure 5. Simplified map of the Ainsa Basin (based on Fig. 2) showing the locations of five detailed sample location maps (Figs. 7B(i), 7B(ii), 9B, 10B, and 11B) of the MPS within the syncline. 44 Sediment samples for magnetic analyses within the sections were collected initially at approximately equal stratigraphic intervals. This sampling strategy was chosen to account for the following two factors: 1) The known frequency of magnetic reversal events during the time that the sections represent, i.e., the Middle to Late Eocene times. Obviously a section likely to contain many, frequent reversal events should be sampled more often than a comparable length section containing just one or two reversal events. 2) The overall estimated duration of the section in question. In practice, the actual sample spacing was also influenced by the availability of preferred or suitable lithologies within the vertical sequence and by the presence of any significant covered intervals where samples could not taken. Fine-grained siltstones and mudstones were targeted for collection. Where needed, at least one additional sampling pass was performed in order to more clearly constrain the magnetic reversal pattern preserved within the section, to fill areas lacking closely spaced data points, and to confirm the presence of shorter “single-site” reversals in magnetic polarity. At each site, 3-5 oriented specimens were collected according to the techniques of Hailwood (1989). Local bedding orientations were measured to allow tectonic correction of ‘in-situ’ magnetic data. Laboratory Analysis Measurements of the sample’s natural remanent magnetization (NRM) were made using an SCT cryogenic magnetometer. With a view to 45 determining the most suitable demagnetization strategy, a subset of 13 sites representing the major lithologies within the study area was subjected to two differing demagnetizaton procedures. The main sampled lithologies present in the section were gray siltstones, gray marls, cream mudstones/siltstones and reddened-mottled mudstones/siltstones. 2 specimens from each site were subjected to step-wise thermal demagnetization from 0-575 °C in 50° or 25° intervals. Based on the results of this procedure (see following discussion), 2 other specimens from the same sites were subjected to a more complex demagnetization process in which each sample was demagnetized via step-wise thermal demagnetization up to 200 °C (0°, 100°, 150°, 200°) and then was demagnetized by alternating field (AF) demagnetization from 0-500 Oe in 50 and 100 Oe intervals. The results of the pilot study were then used to defined intervals of both thermal and alternating field demagnetization during which stable characteristic magnetizations were measured. Bulk analysis of the remaining samples was then completed at two demagnetization levels within these treatment ranges, and the characteristic paleomagnetic field direction for each specimen was noted. The magnetization direction for a given sample location at a given temperature was calculated using the statistical methods of Fisher (1953). The statistical quality of each site was assessed and classified as “Class I” (statistically robust agreement of 3 or more individual specimens within a sample location, i.e., Fisher ‘kappa’ value > 10), “Class II” (apparent agreement individual specimens: kappa of > 5 and < 10 for 3 or more cubes, or kappa » 1 0 for only 2 46 cubes), or “Class III” (poor agreement of specimens at a site). Additional class II sites were defined by average site vectors that showed southerly declinations but shallow negative (upward) inclinations. Such directions are believed to represent reverse polarity data with a normal polarity overprint that was not fully removed during demagnetization. The average magnetic vectors from the Class I and Class II sites were used to calculate Virtual Geomagnetic Pole (VGP) paleolatitudes for each site. General analysis of the magnetized directions suggests the data show a pervasive post-depositional clockwise rotation of approximately 20°. This rotation was removed during the calculation of the site VGP paleolatitude. These data were then used to classify each site as either normal or reversed polarity. The errors in the average magnetic vectors were used to calculate an 0C 95 error envelope for the VGP paleolatitude, giving further assessment of the site data quality. Each VGP paleolatitude was then plotted against stratigraphic position in order to define a magnetic polarity stratigraphy for each section. RESULTS Pilot Study Results of the pilot study showed three general types of specimen behavior during sequential demagnetization (Table 1). Type I behavior (Fig. 6A) was described from samples of both marine and continental lithologies, but was more typical of sediments taken within the non-marine 47 Behavior Description Association Examples Type I (Fig. 6A) Type II (Fig. 6B) Type III (Fig. 6C) Decreasing intensities and stable directions during demagnetization between 200-400 °C or 200-400 Oe Type I behavior to ~350-400 °C, followed by increase in susceptibility, intensity and variable directions Atypical magnetic directions in conjunction with only partial demagnetization Both continental and marine facies Organic-rich shallow marine facies Continental Overbank facies MED01, MED43, GRU02 and MED34 MED25, and GRU33 MED45, and MED52 Table 1. Summary of the three general demagnetization behaviors shown by the samples during pilot study analysis. The types of sediments and specific examples characterizing the three types of behavior are given. 48 W, Up ■ o X Z D 0.01 L U S 00 280 °C NRM-V GRU02A NRM-H Characteristic Remanence Direction i • -0.02 -0.02 W, Up EMU x 10 -4 0.0 o 400 E, Down Temp °C S 0 0 500 Oe 500 Oe^g -0 .0 0 5 - 4- O . o® o® - 0 . 0 1 - -f- L L I NRM-H NRM -0 .0 1 5 N 0.0 0 .0 0 5 0.01 1 0 0 5 0 0 E, Down EMU x 10 4 Oersteds Oe Figure 6A. Plots of data for four samples showing representative Type I’ demagnetization behavior. GRU02A and MED43B represent non-marine siltstones, while MED01D and MED34E are marine mudstones. In the left-hand Zijderveld cartesian projection (As, 1960), open circles represent the vertical component of the magnetization direction, and the closed circles represent the horizontal component. NRM-H and NRM-V are the initial remanence components. The right-hand plot is a plot of intensity normalized to initial NRM, measured at each demagnetization step. The horizontal scale represents demagnetization level (either Oe or °C). 49 W, Up NRM-H 0.02 S o .o -l— - 0.02 NRM-V - 0.02 0.0 0.02 E, Down W, Up S 0 .0 H ------- NRM-H -0.05 NRM-V 0.0 0.05 E, Down Figure 6A continued. 1 "O I s E £ n ? 0 500 0 Oersteds Oe N 1 0 0 500 Temp °C 50 sequences of the study area. This behavior is typified by stable magnetized directions and progressively decreasing magnetic intensity in the range of 200-400 °C for the continental sediments, and 200-400 °C or 200-400 Oersteds (Oe) for the marine lithologies. In light of this data, bulk magnetic analysis of samples taken from the Escanilla and Capella Formations was performed at 280 °C and 320 °C. Type II behavior (Fig. 6B) was only seen in weakly magnetized specimens taken from shallow marine environments in the Ainsa Basin and the western Tremp Basin. These samples were often characterized by a change in magnetic behavior in the range of 300-400 °C, most notably by an increase in both magnetic intensity and susceptibility. This is believed to have been caused by the oxidation of a hydrated iron sulfide compound, such as greigite (King Powers, 1989; Burbank et al., 1992). Greigite is believed to grow after the dissolution of detrital magnetite, usually within 40 cm of the sediment-water interface (Karlin, 1990; Leslie et al., 1990a; 1990b). Provided that sedimentation rates are sufficiently high, this process can accurately record magnetic polarity near to the time of deposition, and may then be used to establish a MPS pattern (Karlin, 1990). However, the detailed structure of any field variation will tend to be smeared, or averaged while the sediment resides within this upper 40 cm. Data from other stratigraphic studies within the study area suggest that rates of Eocene marine sedimentation as constrained by magneto- or bio stratigraphic methods, are almost always greater than -1 0 cm/kyr (King Powers, 1989; De Boer et al., 1991; Cuevas Gozalo, 1990). Greigite in 51 W, Up 500 °C < s joo °c NRM-H 100 °c 500 °C < in cre a se in M a g n e tic S usce p tib ility - 0 .0 2 NRM-V 0.02 0.0 E, Down 5 0 0 Temp °C W, Up MED25D Initial h e a tin g to 2 0 0 °C . A F D e m a g n e tiz a tio n E, Down Oersteds Oe Figure 6B. Plot of data for two specimens showing representative Type IT demagnetization behavior (MED25A and GRU33B). Format is the same as Fig. 6A. Note the increase in relative magnetic intensity at higher temperatures. This is accompanied by highly variable magnetized directions as shown in the Zijderveld projection (As, 1960). The second two diagrams (MED25D) show a specimen subjected to mixed method demagnetization (see text for description). Note the lack on any increase in intensity and stable magnetized directions as demagnetization progressed. 52 W, Up 400 °C 400 °C - 0.01 Increase in Intensity - 0.01 0.0 0.01 E, Down 0.02 0 400 Temp °C Figure 6B continued. 53 such settings would probably grow less than ~5-10 kyr after magnetite deposition (Burbank et al., 1992). Therefore, it appears to be valid to use the polarity information preserved in the shallow marine environments of the southern Pyrenean basins provided we do not cause oxidation of this post-depositional greigite during demagnetization. Consequently, we performed bulk magnetic analysis of the shallow marine marls and siltstones by first heating and measuring the specimens at 200 °C, and then applying alternating field demagnetization at a field strength of 200 Oe. This mixed-method procedure is similar to that employed by a number of other authors for comparable marly marine lithologies (Galbrun et al., 1988; Odin et al., 1991). A small number of the non-marine lithologies exhibited Type III behavior (Fig. 6C). These samples show rather atypical magnetized directions (E-W-directed declinations, or southerly-declinations with positive inclinations), combined with the retention of greater than 50% of initial NRM at high demagnetization temperatures (> 500 °C). This is attributed to the presence of a significant amount of post-depositional, authigenic hematite. Magnetostratigraphic Results Ainsa Basin Magnetic Polarity Stratigraphies Mediano MPS The Mediano (MED) MPS (Figs. 5, 7 A and 7B) was constructed as a reference stratigraphy that traverses the major lithostratigraphic units 54 W, Up NRM-H 100 °c _ |_ 400 - 0.02 MED45A - 0.02 NRM-V -0.04 -0.02 0.0 0.02 0 500 E, Down Temp °C W, Up 0.02 100 °c NRM-H 400 °C S o ® 1 100 °c - 0.02 NRM-V - 0.02 0.0 0.02 N MED52B 1 0 0 500 E, Down Temp °C Figure 6C. Plots of data for two samples showing representative Type III demagnetization behavior. Both represent non-marine lithologies. Again the format is the same as Figure 6A. In the Zijderveld projections (As, 1960), MED45A shows a southerly-directed declination but a positive inclination, while MED52B possesses a shallow west-dipping magnetized direction. Note the large percentage of initial remanence retained at 500 °C. in both normalized intensity plots. 55 within this western area. Samples were collected at 92 sites through nearly 1700 m of marine and continental sediments, 46 % of which yielded Class I data (Table 2). The section begins within deep water marls and siliciclastic turbidites of the Hecho Group (Figs. 7A and 7B), and passes up through a upward shoaling megasequence of shallow marine, deltaic, and continental material. This section (Fig. 7A) shows 7 normal polarity magnetozones, three of which were defined on the basis of single point Class I or multiple Class II data. For the purposes of correlation, these short intervals of changing polarity were included within the better constrained zones of the same polarity either immediately above or below. A comparison of the bedding corrected data, with the ‘in-situ’ (un corrected) field magnetized directions shows that the data pass a fold test (Fig. 7C and Table 2) (McElhinny, 1964; McFadden and Jones, 1981). The bedding-corrected data consistently appear to define -2 0 ± 10° of clockwise rotation (Table 2). This is in agreement with existing data within the Tremp basin (King Powers, 1989; Dinares et al., 1991). Flowever, non-cylindrical folding and the plunging fold axis of the Mediano anticline may partly be the cause of this apparent rotation (Chan, 1988). Based upon the existing biostrati graphic information (De Federico, 1981; Schaub, 1981; Tosquella, 1988, Crusafont et al., 1966), the Mediano MPS (Fig. 7A) was correlated within the published Magnetic Polarity Time Scale (MPTS) of Harland et al. (1990) (Fig. 8) Although faunal information for the Ainsa Basin is limited, several biostratigraphic stages have been recognized. Macroforaminiferal assemblages within the 56 1500 1000 500 — Mediano Section Virtual Geomagnetic Pole Latitude -90 -45 0 45 90 ■m Middle Member Escanilla Fm. Lower Sheet Correlation Lower Member Escanilla Fm. (Bartonian) Sobrarbe Delta (Late Lutetian) n Hecho Group Turbidites (Late Ypresian) Platform-derived olistostromes and resedimented carbonates I N7 R6 N6 N5R5 R4 N4 R3 N3 I N2 R 1 N 1 R2 Figure 7A. The Mediano Magnetic Polarity Stratigraphy (MPS). Schematic lithologic information, and lithostratigraphic nomenclature in indicated in the left diagram. The virtual geomagnetic pole (VGP) latitude is plotted against stratigraphic position in order to define behavior of the paleomagnetic field through time, and construct the MPS. Polar latitudinal error bars are calculated for each pole position. The legend for this, and all the following MPS diagrams is given in Fig. 9A. 57 Figure 7B. Detailed sample location maps for the Mediano MPS. The location of these detailed maps within the Ainsa Basin are shown in Fig. 5. 58 Rio Susia 500 m f . a Javierre Figure 7B continued. ®£J c o « ™ ™ ^ 2 E J * 5 « £ I 8 u C l C D .E T 3 Q-S T 3 CO a > N o o » i - w i n ^ e ' e n o » i s ! c o f - ‘ D f ■ '* ao^fr ^ o co io i - i r > <o c n j o ^pcsi U cv i u> c n j ir> c\i ^ w in s s a s f 2 g s s s s O t - o CM O CM O CM O * - 0 5 r a 2 o C O ® S « > -S < I £ \ j ^ t o t t t t t t t t l o c o £ O O . C O o « o a s C O o a ti E a Z C O a 2 o ‘Nin o t - c o o > ® t n io w t o in i c i ■ v ® n i C M CO O C M 0 > - * - CD O CM CM £ 8 S2 8 5 8 8 zee zee z ae zee zee 3? v j S ' 3s 3? N» O ' 5 ? C O U ) o o 00 C M M- C M C M 3 5 V ? o ' v j S ' . f l S ' 3 ? 3 5 0 5 C O C O C O C O 'M - i r > <o s § . § .8. U J .* o ■ 6 3 .0 5 5 5 <j > U j c n . c £ ^ •2 o ® C O ■ 5 -c ® § E £ C V) o - — 0 5 r o c 5 * 0 5 ® c. -S t ® - 1 . = I ® c I E 0 5 1 1 i n c o q j C /5 C l « ® 00 m E 0 5 ° X - | M ^ s. o 2 ♦ = T 3 ( 0 ® &2 OS c o V ® E C Table 2 . Fisher Statistics on magnetic data from the 6 MPS sections. Both ‘in situ’ and structurally corrected data are given, as are the percentages of both Class I and Class II data sites generated during the bulk analytical study. 60 c o . . 4 - O Q 0) O 3 T O ? ■ M M I "D ^ o r * ° • C .2 r c J 3 h — » o cu 0) Q ^ - D o o t— ■ + - » .2 « ^ i i C D o s o 3 o ( 7) •) 3 o C O Figure 7C. Equal area stereographic projections of magnetic vectors from Class I and Class II sample sites and fold test for the Mediano MPS. ags confidence cones on mean polarity vectors are indicated. Diamonds are used to present mean vector directions. Close circles represent lower hemisphere data points and open squares represent upper hemisphere data. 61 Hecho Group turbidites exposed within the Ainsa Basin indicate a middle-late Cuisian age (local equivalent of the Ypresian Stage) (Schaub, 1981; Tosquella, 1988). The shallow water carbonate facies of the Puy de Cinca limestones are assigned to the middle Lutetian (Schaub, 1981), revising the late Lutetian date given by Biot (1962). The partly equivalent, and overlying Sobrarbe deltaic units are considered to be late Lutetian in age (De Federico, 1981). Furthermore, based on field relationships, rather than direct biostratigraphic data, Crusafont et al. (1966) proposed a Bartonian age for the succeeding continental sediments of the Escanilla Formation. Given this data, the normal magnetozone N1, at the base of the section is correlated with C21N of the MPTS (Fig. 7A and 8). Consequently, N2, and N3 are correlated with C20N, thereby confirming the middle-late Lutetian age for the Sobrarbe Fm. and P uyde Cinca Limestones. The top of C20N within the MPS lies either at 1040 m, or possibly at 1200 m, abqve the short, but poorly constrained normal magnetozone, N4. The former interpretation is employed within the correlation scheme (Fig. 8). N5 and N6 are correlated with C19N, and N7 is equated with the base of C18N. These data suggest that the Mediano section represents virtually un-interrupted deposition lasting approximately 7 myr, with the transition from deltaic to continental deposition occurring at -43 Ma (Fig. 8). Rather fortuitously, at the level of the base of N7, there is a laterally extensive layer of sheet sandstones and conglomerates that may be followed continuously across the Ainsa Basin. Using direct outcrop and 62 Eripol LI guerre Escanilla Formation * * • : ... . s \ i& ^ U p p e r Shoots interval Almazorre Mediano ’ w ir w w w tw ^ « « « W K X e ; * S S ¥ S :S :::S :: :: f S W S W S W I | < D 43 EscanHIa Ume st one LL-EB 43 v® ;^|Esera Valley I I ? Unsampled Interval © c d I MPTS Harland et al. (1990) ■ Normal Polarity I I Reverse Polarity Probable Normal Preferred Correlation — Alternative Correlation «— Lithologic Correlation Biostratigraphic |_yLate m l Middle ; i Late f r Early B Bartonian Information Ypresian Lutetian Lutetian Bartonian Figure 8. Summary diagram incorporating biostratigraphic, lithostratigraphic and magnetic polarity correlations, linking the 6 MPS to the global Geomagnetic Polarity Timescale (Harland et al., 1990), and to each other. For the Esera Valley MPS, both the preferred and an alternative correlation are shown (see text for discussion). Single point, or poorly constrained reversals are indicated as zones of ‘possible normal polarity’. 63 aerial photo relationships, this interval has been used to tie the other Ainsa Basin MPS directly to the Mediano sequence, thereby allowing correlation with the Harland et al. (1990) MPTS. A l m a z o r r e MPS The Almazorre MPS was constructed to test the lateral continuity of this sheet conglomerate correlation. The section crosses the lower 330 m of the Escanilla Fm along the eastern flank of the Boltana anticline, and begins immediately on the top of the Sobrarbe Fm, south of the village of Almazorre (Figs. 5, 9A and 9B). 18 sites were sampled, and the average vector data suggest some clockwise rotation (Fig. 9C) similar to that defined by the Mediano MPS data (Table 2). VGP data from the Class I and Class II sites define three normal and three reversed magnetozones (Fig. 9A), and it is possible to correlate the erected MPS with the MPTS of Harland et al. (1990) (Fig. 8). Given the presence of the lower sheet conglomerate horizon at -200 m above the section base, the lower reversed portion of the section is correlated with the reversed parts of C19 and C18, and the poorly defined normal magnetozone, N1, is interpreted to represent C19N. The upper, dominantly normal portion of the MPS (N2 and N3) is correlated with C18N (Fig. 8). This suggests that the lower sheet conglomerate is an isochronous interval within the Ainsa Basin, and a significant depositional event. 64 Stratigraphic Height (m) 300 - 200 1 0 0 1 1 1 1 Almazorre Section Virtual Geomagnetic Pole Latitude -90 -45 0 45 90 J Q LI LLI Lower Sheel Correlation < D L U I J N1 R1 Coarse Alluvial Conglomerates Sandstone/Conglomerate Channels Silt/Mudstone Overbank Sediments Marine Platform Carbonates Shallow Marine Deltaic Sediments Lacustrine Carbonates Shallow Marine Sands/Conglomerates Marine Marls and Turbidites • Class I Data O Class II Data -O— Polar Error Bars Normal Polarity Reverse Polarity Probable Normal Polarity "I I I I I • 5 . o ? - 5 . T . o C ? f ( t 3 O O - < / > O -ffl o Figure 9A. The Almazorre MPS. 65 oAlmazorre»V K f v ^ ' Amalgamated ’\ r . Sheet Amalgamated Sheet Figure 9B. Detailed sample location map for the Almazorre MPS. The location of this detailed map within the Ainsa Basin is shown in Fig. 5. 6 6 Almazorre Section North South Figure 9C. Equal area stereographic projections of magnetic vectors from Class I and Class II sample sites for the Almazorre MPS. See Fig. 7C for a legend. 67 E r ip o l MPS The 1000 m-long Eripol MPS is located entirely within the western limb of the synclinal Ainsa Basin (Fig. 5, 10A, and 10B). It documents the magnetic reversal pattern preserved within the Escanilla Fm. west of the Mediano anticline, and allows regional correlation with other Escanilla exposures across the study area. The measured section begins immediately above the Sobrarbe Fm. deltaic sediments, and ends at the unconformable base of the overlying Early Oligocene conglomerates (Fig. 10B). 63 samples were collected and 48 % of these generated Class I data (Table 2). The average site data plot antipodally (Fig. 10C), and possibly define a small amount clockwise rotation. The VGP data delineate 5 normal and 4 reversed magnetozones (Fig. 10A), and as with the Almazorre section, the Eripol MPS is tied directly to the longer Mediano MPS using the lower sheet correlation. The presence of the lower conglomerate sheet horizon at -150 m above the base of the sampled section, suggests that this lower portion of the MPS correlates with Chrons C18R and C19N (Fig. 8). The poorly defined reversal pattern at the base of the Eripol MPS showing two normal (N1 and N2) and one reversed magnetozone (R1) is wholly correlated with the C19N. The long normal magnetozone, N3, is correlated with the long phase of dominantly normal polarity represented by C18N and C17 (Fig. 8). None of the short reversed polarity zones of this interval were identified. This correlation suggests that the middle member of the Escanilla Fm. is largely Bartonian in age. 6 8 Stratigraphic Height (m) 1000 500 Eripol Section Virtual Geomagnetic Pole Oligocene Congloms. Upper Member Escanilla Fm. Upper Sheet Correlation Middle Member Escanilla Fm. Lower Sheet Correlation Lower Member Escanilla Fm. Figure 10A. The Eripol MPS. See Fig. 9A for legend. \ ) X • A w 7 T - \ 500 m Frontman % r s - n \ { °**0 O c e n e Unconform ity Figure 10B. Detailed sample location map for the Eripol MPS. The location of this detailed map within the Ainsa Basin is shown in Fig. 2. 70 Eripol Section North South Figure 10C. Equal area stereographic projections of magnetic vectors from Class I and Class II sample sites for the Eripol MPS. See Fig. 7C for a legend. 71 The upper part of the Eripol MPS is rather difficult to correlate confidently to the MPTS. The most reasonable correlation equates N4 to the topmost normal polarity zone of C16. R4 and N5 represent the Chrons 15R and 15N, respectively. This correlation implies that the upper member of the Escanilla Fm. is largely Priabonian in age. Unfortunately, due to the lack of suitable lithologies available for sampling at the very top of the Eripol MPS, we have no data to define accurately the age of the top of Escanilla Fm. or the base of the overlying Oligocene conglomerates. LIGUERRE MPS The Liguerre MPS was designed to date an adjacent part of the Escanilla Fm. in order to reconstruct variations in subsidence and sediment accumulation within the Ainsa basin. Beginning immediately on top of the Sobrarbe rrm., the section traverses over 1200 m of Escanilla Fm. strata preserved within the eastern limb of the syncline (Figs. 5, 11 A, and 11B). 58 sample locations yielded over 80 % Class I or Class II sites (Table 2). These data (Fig. 11C) show an antipodal relationship between northerly- and southerly-directed magnetized directions, combined with a small amount of clockwise rotation. Within the Liguerre MPS, the VGP data define 5 reversed and 5 normal magnetozones (Fig. 11 A). The short normal zone, N1, is poorly delineated, being defined by only one Class II data point, and subsequently we do not attempt equate this magnetozone with any stable polarity interval on the MPTS (Fig. 8). Correlation of the Liguerre MPS to the other MPS within the Ainsa Basin is constrained lithostratigraphically by three data points (Fig. 8). 72 The section begins on top of the Sobrarbe Fm.: therefore, the lower two reversed magnetozones (R1 and R2 of Fig. 11 A) represent C19R. The lower sheet conglomerate horizon ties into the Liguerre MPS at -480-500 m, and again is closely associated with a change from reversed to normal polarity which indicates that N2 and R3 are equivalent to Chron 19N and Chron 18R, respectively. As with the Eripol MPS, we correlate the long normal magnetozone, N3, with Chron 18N, Chron 17, and most of C16. A second laterally extensive sheet conglomerate interval is traceable from the Liguerre MPS (-650 m) to the Eripol MPS (-550 m) (Fig. 8 and 11 A). This lithostratigraphic tie is in good agreement with the magnetostratigraphic correlation within the normal portion of Chron 16 (Fig. 8). The top of the Liguerre MPS shows a slightly different magnetic reversal pattern than the Eripol sequence. The lack of suitable sampling localities within this interval may partly be the cause. We feel that the most reasonable correlation of the MPS to both the Eripol section and the MPTS equates the N4, R5, and N5 magnetozones to Chron 15N (Fig. 8). This implies that the long reversed magnetozone, R4, actually represents all of Chron 16 and Chron 15R, and that we have failed to identify any of the short normal polarity zones within Chron 16. 73 Stratigraphic Height (m) Liguerre Section Virtual Geomagnetic Pole Oligocene Congioms. Latitude -90 -45 45 90 1000 Upper Member Escanilla Fm. Upper Sheet Correlation Middle Member Escanilla Fm. 500 Lower Sheet Correlation Lower Member Escanilla Fm. o rm Figure 11 A. The Liguerre MPS. See Fig. 9A for legend. Figure 11B. Detailed sample location map for the Liguerre MPS. The location of this detailed map within the Ainsa Basin is shown in Fig. 5. Liguerre Section North South Figure 11C. Equal area stereographic projections of magnetic vectors from Class I and Class II sample sites for the Liguerre MPS. See Fig. 7C for a legend. 76 Tremp Basin Magnetic Polarity Stratigraphies E s e ra V a lle y C o m p o site MPS Across the Mediano anticline, within the western Tremp Basin, the Esera Valley composite MPS (Fig 12A) was developed in order to date stratigraphic units temporally equivalent to the lower Mediano sequence but in a very different depositional setting. Beginning at Santa Liestra (Figs. 4 and 12B), the section passes up through 900 m of shallow marine and coastal sediments of the Perrarua, Capella and Pano Formations (Fig. 12A). The section ends at the base of the Grustan limestone (Figs. 4 and 12B): a carbonate platform sequence considered to be laterally equivalent to the marine limestones exposed along the eastern flank of the Ainsa Basin (Puigdefabregas, 1975; Nijman & Nio, 1975) (Fig. 3). We collected magnetic samples at 71 locations within three short stratigraphic sections (SLA, MDP, and GRU). Over half of the sites generated Class I data points. These data define a bimodal, near antipodal distribution that again appears to show small amounts of clockwise rotation (Fig. 12C; Table 2). The Esera MPS shows 5 normal and 5 reversed magnetozones, a number of which are based on single point reversals of field polarity (Fig. 12A). The correlation of the Esera valley MPS with the MPTS is based both on faunal and previously determined magnetostratigraphic data. The presence of the Puy de Cinca limestones at the top of the Esera Valley MPS indicates a middle Lutetian upper limit for age of the stratigraphic 77 Effective Stratigraphic Height (m) Esera Valley Sections Virtual Geomagnetic Pole Latitude 500. Puy de Cinca "Grustan" Lst. (Mid-Late Lutetian) Pano Fm. (Middle Lutetian) Capella Fm. (Middle Lutetian) Section Correlation -90 1 Section Correlation ITTTI Perarrua Fm. (Ypresian) Class i Data G R U ■ M D P ♦ SLA Class II D ata O G R U □ M D P Figure 12A. The Esera Valley Composite MPS. See Fig. 9A for legend. 78 07099086 o 108m Unsampled MPS Section p r o j e c t io n t o M D P M p s To G r a u s Figure 12B. Detailed sample location maps for the Esera Valley com posite MPS. The locations of these detailed maps are shown in Fig. 4. (i) The Santa Liestra Section. To Perrarua/Capella S a n t a Boundary LIESTRA w Pe-Ca \ ° Boundary ■ rr\ Meson de Pascual 500 m so m m r / M Torre^e Esera 190 m ° 0 \ P r o j e c t io n t o GRU M p s Figure 12B. (ii) The Meson de Pascual Section. 80 o V ___ U _ Ll_ c z C L C Z .E a) O c CL H i CLZj Figure 12B continued, (iii) The Grustan Section. 81 sequence (Schaub, 1981). The macroforaminiferal assemblage within the Perrarua Fm. indicates a late Cuisan (Ypresian) age (Schaub, 1981). Using mammalian biostratigraphic data, Crusafont et al., (1966a) assigned a middle Lutetian age to the Capella Formation. The transgressive barrier island sequence of the Pano Formation is laterally equivalent to the upper Capella Formation, and therefore, would also be middle Lutetian in age. This then implies that the Esera Valley section would be laterally equivalent to the lower part of the Ainsa Basin stratigraphy, namely the Hecho Group turbidites and Sobrarbe Fm. The base of the Campanile Fm. (Fig. 4) has been dated magneto- stratigraphically (King Powers, 1989; Burbank, unpublished data), and is correlative with the lower part of Chron 21 (-5 0 Ma, Harland et al., 1990). The Esera Valley MPS begins approximately 135 m above the top of the previously sampled interval (Burbank, unpublished data). Therefore, the base of the MPS at Santa Liestra is correlated with the lower part of Chron 21N (Fig. 8). N5 is correlated with Chron 20N given the biostratigraphic data of Schaub (1981) which indicates that the Pano Fm. and Grustan Limestone (Cuevas Gozalo et al., 1985) at the top of the Esera Valley MPS are the same age as the Puy de Cinca Limestone and Sobrarbe Formation in the Ainsa Basin (Fig. 8). Given the pre-Bartonian age of these lithostratigraphic units, the top of the Esera MPS must predate Chron 19N. Hence, the topmost reversed magnetozone, R5, must correlate with Chron 19R, and N5 should represent Chron 20N. 82 Esera Valley Section North ( • □ □ South Figure 12C. Equal area stereographic projections of magnetic vectors from Class I and Class II sample sites for the Esera Valley composite MPS. See Fig. 7C for a legend. 83 Despite the limits on the top and bottom of the MPS, the correlation of N2, N3, and N4 remains problematic (Figs. 8 and 12A). These short normal zones are rather poorly defined by changing polarity data. With one exception, the polar error calculations associated with these data sites do not allow a definitive assignment of normal polarity. In the light of this uncertainty, the top of Chron C21N appears most likely to correlate with the top of N1 (Fig. 8). This correlation is employed in subsequent calculations of accumulation and subsidence rates. L a s c u a rre MPS As the most easterly of the six MPS, the Lascuarre section was constructed to permit reliable correlation of Escanilla Fm. strata preserved within the Tremp basin to the longer sequence found to the west of the Mediano anticline. Measured wholly within the Escanilla Formation preserved south of the Isabena Valley in the Tremp basin, the MPS traverses the lower 350 m of the fluvial sequence (Figs. 4, 13A and 13B). 25 sites were collected, and nearly 70 % of these yielded Class I data (Table 2). The resulting MPS reveals 5 normal and 5 reversed magnetozones. Five of these are defined on the basis of single point data reversals. No significant post-depositional rotation was suggested by these data (Fig. 13C; Table 2). Correlation of the Lascuarre MPS with the MPTS of Harland et al. (1990) is constrained by both biostratigraphic and indirect lithostratigraphic information. Crusafont and Pairo (1958) and Crusafont et al. (1966) 84 Stratigraphic Height (m) 4 0 0 - 3 0 0 ' 200. 100. Mmmmm w i p S Z S g g p s S f c B E B S 1 L U L L L U A U K f f | W | W ' ....... __ Ill III liLLU n n Lascuarre Section Virtual Geomagnetic Pole Latitude -90 -4 5 0 45 90 c o a ? .a E aj Q j C Q- iT " o L L - z p 6 'E "O to -O O ~ c/3 I > L U Lower Sheet Correlation 03 P E C LL CD C O ^ C O $ O o 0 0 _J LU Escanilla Lstj- (Late Lutetian -Early Bartonian) Figure 13A. The Lascuarre MPS. See Fig. 9A for legend. 85 Barranco de ia Riviera / / i / '/ } S p Castesillo / / o ° 7 o.° Scale 0 500 m Figure 13B. Detailed sample location map for the Lascuarre MPS. The location of this detailed map is shown in Fig. 4. 8 6 Lascuarre Section North South Figure 13C. Equal area stereographic projections of magnetic vectors from Class I and Class II sample sites for the Lascuarre MPS. See Fig. 7C for a legend. 87 suggested that the lignites within the basal limestones of the Escanilla Fm. are of Lutetian-early Bartonian age. A laterally continuous sheet conglomerate interval is present at -160-180 m. This level is interpreted to represent the eastern projection of the lower sheet conglomerate horizon within the Ainsa Basin. In both areas, the intervals show a similar internal architecture, they are associated with a change from reversed to normal field polarity, and clast population studies show the first appearance of a distinctive conglomeratic clast immediately below this level (Bentham et al., in prep). Consequently N2-N5 are correlated with Chron 18N, and the lower part of Chron 17, although the exact details of these correlations are unknown (Fig. 8). The short magnetozone, N1, close to the base of the Lascuarre MPS is tentatively correlated with Chron 19N (Figs. 8 and 13A): the same short normal zone preserved within the lower Escanilla Formation of the Ainsa Basin. SEDIMENT ACCUMULATION CALCULATIONS In order to estimate rates of sediment accumulation and tectonic subsidence, we used the time-constraints offered by our magnetostratigraphic correlations, combined with the geohistory analysis used by Burbank et al. (1992). This uses the standard algorithm for sediment decompaction (Sclater and Christie, 1980), but approaches the calculation of tectonic subsidence rather differently. A balance between tectonic load and local isostatic compensation is assumed at the beginning and end of each time-bounded increment of deposition. At the end of 8 8 each step, the difference between the isostatically predicted, and the modeled, decompacted position of the base of the stratigraphic column is attributed to tectonic loading. Incremental tectonic loads are then summed over time to give the curve of tectonically-induced subsidence. Given the lack of reliable paleobathymetric data, we assumed that the basins were filled to sea-level. Although this is essentially correct for the shallow marine units and distal alluvial deposits of the Puy de Cinca, Sobrarbe, and Capella formations, this is certainly not the case for the Hecho Group turbiditic strata. However, lacking reliable depth control, we chose to make such an assumption. Further uncertainties in the time control offered by magneto- stratigraphic dating and their subsequent effects in sediment-accumulation calculations are discussed in detail by Tailing and Burbank (in prep). Given the numerous sources of error including measurement error of stratigraphic thickness (probably at least +5%); non-unique correlation of our MPS with the MPTS of Harland et al. (1990); uncertainties associated with the exact placement of chron boundaries within our MPS; and inherent errors in the absolute ages of those boundaries as defined by Harland et al. (1990), we use the plot of sediment accumulation and tectonic subsidence qualitatively to assess spatial variation in a relative rather than absolute sense. Two sets of data are presented: firstly, geohistory calculations for the early-late Lutetian sediments preserved within the Mediano and Esera Valley MPS (Fig. 14A), and secondly data from the latest Lutetian-Priabonian fluvial sediments within the Eripol and 89 Cti C D E i- ^ 1 1 -t 1 1 1 1 i i i i i' C\l O C M ^ C D 00 O O O O O O O (LU>|) s s e u v i o j m CD o o § 5 'c -g Q O C O ^ % % ^ \ - CO c C D Q CD * = ■ ■ I co Q- “ CD CD 2 m Q I I I I CD O g q5 'c= ;g 1 1 1 2 0 0 W o -Q C O c d 1 3 LU I - CO < D C c C D If (fl H C O CD CD LU C O Q < D C LU Figure 14A. Geohistory diagram showing sequential basement depth and the component of tectonic subsidence for the Esera Valley (preferred correlation) and the Mediano MPS. Note the very different rates of tectonic subsidence during Chron 20 time. Thickness of section, and the amount of tectonic subsidence are given in kilometers. 90 Liguerre MPS (Fig. 14B). Error bars associated with each data point are considered to be at least ±10% of each thickness quoted. Although substantial, these uncertainties do not compromise the validity of this semi-quantitative analysis. A comparison of the geohistory curves for the Ainsa and Tremp basins (Fig 14A) shows a significant difference during Chron 20N (-45-43 Ma). Sediment accumulation west of the Mediano anticline was about 3-4 times as fast as in the Esera Valley, east of the structure, and tectonic subsidence rates were at least double those east of the fold. This thickening of section across the Mediano anticline occurs synchronously with the main phase of fold development. Detailed mapping along the western flank of the Mediano anticline, close to Samitier, shows a number of syntectonic geometries that may be linked directly to the Mediano MPS (Figs. 2 and 15A). Most importantly, a reef-derived talus slope breccia is both underlain and overlain by shelf marls of the Hecho Group turbidites. Within the underlying marls, beds of transported carbonate material (rich in nummulites, small corals, and oysters) are periodically developed. Individual beds, or small groups of these beds can be followed laterally down the paleoslope (RS1-RS3 of Fig. 15), and thicker groups can be traced directly into the Mediano MPS. The stratigraphic thickness of intervening marly sediment was measured at three adjacent localities immediately below the talus slope sheet (Fig. 15). Thicknesses between marker horizons increase rapidly away from the fold axis. RS1 and the talus breccia, for example, are separated by 91 CO C O C O CO CO 0 ) CO h - Q _ o - C M LU CO LU o CM O C M ^ C O 0 0 CD O O O O O O (LU>1) SS0U>jO!L|X CD o .9 5 C T 3 o -Q ^ (!) 3 CO O O c CD O £ co Q - Ct$ CD C Q Q C D O .9 S CD 2 . c o = 0-0 DC < D D LUh-CO O i i i CD I s — CO Q - CC c n c d LU C Q Q i Figure 14B. Geohistory diagram showing sequential basement depth and the component of tectonic subsidence for the Eripol and Liguerre MPS during Escanilla Formation time. Note, in general, the similarity of data from both sections, although in detail the Liguerre data MPS shows initially higher sediment accumulation and subsidence. Thickness of section, and the amount of tectonic subsidence are given in kilometers. 92 -7 0 m close to the fold. This increases to 90 m over just 300 m laterally, and to 175 m 2-3 km away within the Mediano MPS. The presence of exotic blocks and larger olistoliths of fossiliferous reef material throughout the lower 600 m Hecho turbidites of the Mediano MPS would suggest that the crest of the anticline allowed shallow marine platform carbonate deposition during early-middle Lutetian time (-48-44 Ma). These data indicate strong differential subsidence over short distances away from the flank of the growing fold. Subsidence calculations from the Eripol and Liguerre MPS, on opposite sides of the Ainsa Basin, are very similar (Fig. 14B). Both sections show initially rapid sediment accumulation rates that are essentially equivalent within the discussed uncertainties. After -4 2 Ma, the rates of sediment accumulation decrease by about 50%, and these slower rates are sustained into upper Escanilla Formation time (-3 7 M a ). These decreasing rates are a function of our preferred magnetic correlations. If the top of the Eripol and Liguerre MPS are older than our preferred correlations indicate the calculated rates of subsidence and sediment accumulation will subsequently increase. However, sedimentologic data in the Ainsa Basin, specifically a change in nature of fluvial overbank material from the base to the top of both sections would qualitatively agree with decelerating accumulation rates. The lower Escanilla Formation (43-41 Ma) within the Ainsa Basin is dominated by large volumes of immature, pedogenically modified overbank sediments (Bentham et al., 1991). The lack of mature calcic paleosol development is 93 LU \ o w w C O C O c c ffl < n < D izcccc \ c \ \ \ \ \ (l u) sseu>p!i|j_ m 1 = si < D < D l U l •§ 1 § I V A 0 . 0 . Q < 3 <D (D < o tn / / / / " C O < D co co m < d (D □ c c g._ c I § = 2 j§ «<a 8 *f“ M *D W Q O r e 2 r,re U -S 3 3 U j ££9-<?2a: f c O lttJ V J m Figure 15. Syntectonic geometries and thickness variations along the western flank of the Mediano anticline in the region of Samitier. A. Geologic Sketch Map (see Figure 2. for map location). B. Panel diagram shows variations in thickness immediately beneath the lower reef-derived talus breccia. The map shows lateral correlation of these units into, and the height at which they appear in the Mediano MPS. Note the marked change in thickness between the talus breccia level and RS1 in the two figures. 94 suggestive of rapid sediment accumulation (Retallack, 1986). Middle and upper Escanilla sediments (younger than 41 Ma) commonly show such calcic paleosol development. Additionally, across the Mediano anticline, within the western Tremp Basin, the sediments within the Almazorre MPS commonly show the development of strongly calcareous overbank material and nodular calcic paleosols. Although decompaction was not performed for this short section, regional correlation of the MPS (Fig. 8) suggests slower rates of sediment accumulation during lower and middle Escanilla time within the Tremp Basin, compared to the Ainsa Basin. DISCUSSION Lateral Correlation and Sequence Boundaries This study represents the first lateral correlation of time-bounded packages of sediment across the western oblique ramp of the SCU thrust system. Using the magnetostratigraphic correlations, we can evaluate the existing lithostratigraphy, and assess the usefulness of the previous applications of sequence-stratigraphic concepts within the study area. Mutti et al. (1988) split the middle-late Eocene sediments into two depositional sequences, the Santa Liestra Group and the Campodarbe Group. The two groups are separated by a transgressive surface present at the base of the Pano Formation (Mutti et al., 1988). These authors suggested the Pano Fm. is a transgressive deposit representing a substantial regional rise of relative sea-level (the ‘Biarritzian’ transgression of Puigdefabregas, 1975). In contrast, Reynolds (1987) postulated that 95 the Pano Fm. formed on the flank of the rising anticline in a ‘rim’ syncline: a salt withdrawal phenomenon related to the migration of the Triassic Keuper facies (the regional detachment) into the adjacent ‘diapiric’ fold. Reynolds, however, interpreted the Mediano anticline solely as a salt diapir, failing to recognize the structural control of the western ramp of the SCU as being the main reason for fold development. Salt migration has taken place along the Mediano anticline, but the timing of this motion is generally considered to be late, with respect to actual fold initiation (Teixell, pers. comm.). Nevertheless, because the main phase of fold development is synchronous with Pano deposition, the interpretation of the Pano transgression as a regional sequence boundary may well be incorrect. Instead, it may be a local phenomenon related only to continued development of the Mediano anticline. Furthermore, internal correlations made by Mutti et al. (1988) when attempting to force the sequence concept into the existing regional stratigraphic framework breakdown in the light of new paleomagnetic data. For example, Mutti et al. (1988) suggest that the Grustan Limestones (Puy de Cinca Fm.) represent the maximum flooding surface within the Campodarbe Group, and that this would correlate down system with the Sabinanigo sandstones of the NE Jaca Basin (see Fig. 1 for approximate location). However, after combining the magnetic data from this study with that of Hogan (1992) we see the Puy de Cinca Fm is ~43 Ma, and the Sabinanigo sandstones are dated at ~ 41 Ma. Additionally, the upper boundary of the Campodarbe Group, is defined in the Ainsa Basin to lie at the base of the Oligocene conglomeratic succession. Mutti et al. (1988) 96 date this surface as -3 8 Ma, whereas, the data from this study suggests it is probably no older than -36 Ma. Although representing a significant break within the Ainsa and Tremp basins, such a similarly well developed surface of the same age is not present throughout most of the Jaca Basin (McElroy, 1990; Hogan, 1992). However, within the Sierra Exteriores, an unconformity preserving a significant amount of paleorelief is observed (McElroy, 1990). As such, the break in sedimentation defined as a sequence boundary defined by Mutti et al., (1988) appears to be related to uplift and erosion during deformation of the southern Pyrenean frontal thrust systems (External Sierras and Sierras Marginales), rather than being the result of a regional eustatic base-level change. Although the correlations and subsequent interpretations made by Mutti et al. (1988) appear to be incorrect, the application of sequence stratigraphic concept as a tool to recognize equivalent stratal packages is useful. For example, they state the Campodarbe strata show marked thickening and facies changes across the axis of the Mediano anticline into the Ainsa Basin. The regional MPS correlation agrees with this change in sediment accumulation rates, and it can be related directly to fold development (Figs. 8 and 15). However, the interpretation that the stratal geometries across the SCU oblique ramp are the result of just relative sea-level changes seems to be misguided. This is further highlighted when the depositional sequences defined by McElroy (1990) from the Jaca Basin are reconciled with those of Mutti et al. (1988). Sequence boundaries for the Eocene sedimentary systems generally show little temporal coincidence along the length of the foreland basin (this 97 study, Hogan, 1992). The effects of regional and/or local tectonic development within this structurally dynamic environment can not be understated. Eocene Paleogeography Based upon our regional magnetostratigraphic correlations, revisions of the paleogeographic reconstructions provided by earlier studies, most notably Nijman and Nio (1975), show the sequential sedimentologic and structural evolution of the region surrounding the Mediano anticline during middle-late Eocene times. These depart in detail from the Nijman and Nio reconstructions in their exact correlation of middle-late Lutetian sedimentary environments. Ypresian-Eariy Lutetian (-49-45 Ma) Figure 16A represents Perrarua Fm. time (Upper Montanana Group of Nijman and Nio , 1975; Lower Santa Liestra Sequence of Mutti et al., 1988). The Mediano anticline was beginning to develop, and the region of the oblique ramp represented a very narrow shelf break separating shallow marine and continental environments of the Tremp Basin from deep marine turbiditic sedimentation to the west. Large volumes of sediment were supplied across the shelf break from the Campanue fan delta sequence to the submarine fan sequences of the Hecho Group turbidites within the Ainsa and Jaca basins. 98 A. West East Lower Campanue: Fan-Delta Coastline Santa Llestra Group Turbldltes Capella Fm, Perrarua Fm. S ub-M arine H ig h Paleoflow Direction Perrarua Fm. Castisent Fm. Incipient Thrust Motion Relative Subsidence Rate Reference Marker Thrust Transport, Direction i§ Isabena Ainsa Esera Basin Valley Valley Figure 16. Four block-diagram showing sequential paleogeographic reconstructions of the western SCU during: (A) Late Ypresian-early Lutetian; (B) Middle Lutetian; (C) Late Lutetian; and (D) Bartonian- Priabonian times. The reconstructions are based on those of Nijman and Nio (1975), and incorporate data from this study combined with information from Puigdefabregas (1975), Reynolds (1987), and Cuevas Gozalo (1990). The reference marker, shown in the schematic cross- sectional view, is intended to qualitatively indicate continuing fold development. Thicknesses are not shown to scale. 99 Middle Lutetian (-45-43 Ma) During Capella and Pano Fm. time (Fig. 16B), the facies across the study area record the main phase of uplift and rotation of the Mediano structure. Evidence within the Ainsa Basin suggests a carbonate platform developed upon the emerging structural culmination, and carbonate material was shed down slope into the surrounding shelf environments (Fig. 15). Interestingly, this would have been contemporaneous with Guara Limestone platform carbonate deposition to the west, on the Boltana anticline and within the southern Jaca Basin. West of the Mediano anticline, the Capella Fm. coastal alluvial deposits have migrated westwards, towards the Mediano anticline, in response to slowed subsidence and the filling of the western Tremp Basin. Late Lutetian (-43 Ma) The reconstruction during Puy de Cinca, or Sobrarbe Fm. deposition (Fig. 16C), shows the Sobrarbe Delta prograding northwards along the subsiding Ainsa Basin axis (Puigdefabregas, 1975; Bentham et al., in prep) synchronous with carbonate deposition on the Mediano fold axis, and continued growth of the Mediano structure. The platform carbonates upon the reef crest expanded towards the east, and this time of maximum transgression was probably coeval with the deposition of the Escanilla Limestone (‘Lake Lascuarre’) within the western Tremp Basin (Cuevas Gozalo, 1990). This correlation is consistent with the Esera Valley and Lascuarre MPS data. 100 T | C O C C D a o o 3 rH- Z D C C D Q . C D Q . Q . C D r“ c c d r» D C D O O D (I) r-4 « E o o' D Guara-Age Carbonate Platform Talus Slope B reccia O listostrom e HecnoGrp. , , Turbidites / Mediano Anticline Ainsa Basin O Upper Campa Conglomerate Pano Barrier Complex Capella Formation r . ... .. Alluvial Plain Perrarua Fm. Rapid Shortening On Oblique Ramp Thrust Transport D irectio n Esera Valley Isabena Valley 45-43 Ma Figure 1 6 continued. C. Late Lutetian reconstruction. "Lake Lascuarre Puy de Cinca Carbonate Platform Escanilla "™"— Limestone Slowed Subsidence ^ Base of Delta Continued Shortening on Oblique Ramp More Rapid Subsidence 43 Ma Isabena Esera Basin Valley Valley O K ) Bartonian-Priabonian (~42.5-36 Ma) The final reconstruction (Fig. 16D) shows the regional paleogeography during Lower Escanilla Formation time, after the main phase of growth of the Mediano structure. Coarse braided rivers fed directly across the fold axis, into the more rapidly subsiding Ainsa Basin, and short-lived lacustrine deposition occurred episodically within the Tremp Basin. This is believed to have been induced by period emergence of the Mediano anticline, during the final stage of fold growth, resulting in the ponding of the fluvial system within the western Tremp Basin (Bentham et al., in prep). From the Ainsa Basin, the Escanilla fluvial system fed westwards into marine environments of the Jaca basin (Bentham et al., in prep). Structural Development Using the along-strike magneto- and litho-stratigraphic correlations across the study area, we have temporally constrained a number of aspects of the regional tectonic development, most notably, the main phase of Mediano fold development occurred from about -48-42 Ma, . Holl and Anastasio (1990) derived a similar age for fold growth from the analysis of local unconformity geometries along the eastern flank of the Mediano structure. They dated rotation of adjacent coeval strata, but did not demonstrate the thickness variations that allow for a unique interpretation of syndepositional fold development. While their data is consistent with our analysis, like Reynolds (1987), they suggest folding was the result of differential loading and subsequent salt migration and 103 Figure 1 6 continued. D. Bartonian-Priabonian reconstruction. Sis PaleoValley Input u s £ /) f f l ' V f '/ Limited Lacustrine Deposition Escamlla System Floodplain Descreasing Motion on Oblique Ramp Ainsa Esera Isabena Basin Valley Valley O diapirism, caused by thrust emplacement and sediment progradation within the central Tremp Basin. We favor the interpretation of Camara and Klimowitz (1985), specifically, that the Mediano anticline is an oblique- ramp structure, related to the western, west-vergent margin of the South- Central Unit, and transferring displacement along a rather diffuse oblique ramp system between the Montsec and Monte Perdido thrust sheets. Although no discrete thrust fault can be mapped along the axis of the Mediano anticline, a through going structure is considered to be present at depth (Camara and Klimowitz, 1985). Furthermore, relatively minor offset dextral tear faults, have been mapped on the eastern limb of the anticline, that would be consistent with an oblique ramp interpretation for the Mediano structure (Fig. 15 of Nijman, 1989). Fold development was associated with strong facies changes and spatial variations in subsidence and sediment accumulation rates (Fig. 8 and 14A) across the study area. These changes in rates are not surprising when we consider the nature of the folding, and the structural positions of the MPS data. The Tremp Basin sediments sit in the hanging- wall of the Mediano fold, while the Ainsa Basin lies in the footwall of the structure. The rates of compacted sediment accumulation calculated for the western Tremp Basin (-37-45 cm/kyr depending upon exact correlation with the MPTS) are significantly different to those previously suggested for the same time period (-60 cm/kyr for Perrarua and lower Capella deposition; De Boer, Pragt, and Oost, 1991). During late Lutetian, and early Bartonian times, the Ainsa Basin continued differentially subside. On the western flank, adjacent to the 105 Boltana anticline, lower Escanilla sediments show a marked thinning, suggesting continued growth of the fold, rather than onlap of this already uplifted structure (Fig. 8). Thereafter, in Bartonian to Priabonian times, the axis of basin subsidence migrated westwards, towards the Boltana fold. This subsiding axis served to focus the later Escanilla fluvial system (Bentham et al., in prep), diverting late Eocene rivers southwards around the emergent fold. CONCLUSION Detailed lithologic and magnetostratigraphic correlation across the western oblique ramp of the SCU thrust system has, for the first time; (i) constrained the middle-late Eocene tectonostratigraphic development of this structurally dynamic region in an absolute chronologic framework; (ii) allowed the sequential backstripping of Eocene sedimentary systems, and the calculation of rates of sediment accumulation and tectonic subsidence; (iii) addressed the validity of applying sequence stratigraphic analysis in tectonically active environments, where a reliable chronologic framework does not exist; and (iv) aided revision of existing paleogeographic reconstructions for these complex, spatially-variable late Eocene sedimentary systems. Magnetostratigraphic dates from the Tremp and Ainsa basins, indicate Mediano anticline developed during a rapid phase of growth and rotation from ~48-42 Ma. This served to induce marked differential subsidence across the study area, and was largely responsible for the rapid, lateral 106 facies variations across the oblique ramp of the SCU at this time. An abrupt marine transgression between -45-43 Ma immediately east of the Mediano, in the uplifting hanging-wall of the SCU thrust system, was probably the result of local subsidence rather than a eustatic rise in sea- level. This subsidence may well, have been caused by lateral migration of Keuper facies evaporites into the growing fold core. A discussion of regional structural style supports the idea that the Mediano fold is an oblique ramp structure, related to southward translation of the SCU thrust system above its Triassic decollement horizon. The transition from marine to continental sedimentary environments across the oblique ramp occurred at -42.9 Ma, with the deposition of the Escanilla fluvial system on top of shallow marine and coastal sediments of the Puy de Cinca, Sobrarbe, and Capella formations. Eocene fluvial sedimentation persisted until the onset of uplift and deformation of the Sierras Exteriores soon after 36 Ma. This was marked by the formation of a complex sub-regional unconformity that placed early Oligocene conglomerates erosionally upon the middle-late Eocene stratigraphy. 107 CHAPTER 3 Temporal and Spatial Controls on Alluvial Architecture in an Axial Drainage System, Late Eocene Escanilla Formation, Southern Pyrenean Foreland Basin, Spain. 108 ABSTRACT In young or currently active foreland basins of the world, along- orogen variations in structural deformation and/or depositional environments are common elements of the later ‘molasse’ phase of basin development. Such longitudinal changes are not presently incorporated within most of the quantitative and predictive synthetic models of such sedimentary systems. The late Eocene Escanilla Formation of the South- Central Pyrenean foreland basin represents an ancient drainage system in which such variability can be studied in detail using high-resolution magnetostratigraphy. Downstream changes in the nature of the alluvial system were strongly influenced by the on-going Eocene structural partitioning of the basin as it began to become incorporated into the southward-advancing South Pyrenean thrust system. Lower subsidence rates within these allochthonous ‘piggy-back’ sub-basins served to increase channel-body interconnectedness of sheet-like alluvial conglomerates, to inhibit the preservation of significant volumes of fine grained overbank material, and to promote the extensive development of pedogenic calcrete horizons. During the phase of coastal progradation along the subsiding basin axis, a number of N-S-trending anticlines impeded the westward regression of the alluvial system, producing the strong diachrony in the age of a Lutetian-Priabonian-aged deltaic system long the orogen. Fold growth across the western oblique ramp of the South-Central Unit (SCU) thrust system dramatically influenced middle to late Eocene drainage 109 patterns and lithofacies’ distributions. Within the portions of the drainage system upstream of these active folds, the alluvial deposits were periodically ponded allowing the deposition of micritic lacustrine limestones. Rapid fluctuations in sea level seem to have exerted significant control on the drainage system well upstream into the alluvial drainage basin. Base-level rises caused short reversals in the longer- term westward regression of marine environments across the foreland basin, while base-level falls produced regional sheet conglomerate deposition IN TR O D U C TIO N The middle to late Eocene sediments of the South Pyrenean Foreland were deposited in an internally deforming system of thrust-sheet-top or ‘piggy-back’ basins. The depositional systems active at this time showed great lateral variability and diachrony along the length of the orogen. Along-strike variations in structural deformation or depositional environments, while being important aspects of the later history of foreland basins, are not addressed or incorporated within most quantitative synthetic models concerning the stratigraphy of such sedimentary systems (Heller et al., 1988; Flemings & Jordan, 1989; 1990). If one is to identify correctly and interpret the sedimentary fill of such basins in the geologic record, one should be aware of such possible lateral variability, and of the factors that may produce it. The late 110 Eocene Escanilla Formation of the South-Central Pyrenean foreland basin offers a chance to study such variability in great detail, in a region where fortuitous preservation of the important geometric relationships has allowed the description of the thrust deformation through the effects it exerted on the coeval depositional systems within the basin (Garrido- Megias, 1973; Puigdefabregas, 1975; Mutti, Seguret & Sgavetti, 1988; Burbank et al., 1992). This paper addresses this phase of Pyrenean foreland history, when the previous flexural basin deposits began to be incorporated into the southward-advancing south Pyrenean thrust wedge as plate convergence continued. The study area is situated across one of the major N-S-oriented structural trends within the generally E-W-striking south Pyrenean thrust system (Fig. 1). Excellent exposures of the Escanilla Formation fluvial deposits have allowed the description of downstream and temporal variations in sedimentary facies and alluvial architecture as this major Eocene drainage system flowed longitudinally along the basin from E to W, progressively encountering different structural regimes as it did so. The South-Central Unit (SCU) (Fig. 1) of the southern Pyrenees represents a linked system of southward-directed cover-involved thrust sheets and associated piggy-back basins. Thrust motions show a complex pattern of forward- and hindward-imbrication, as well as significant phases of out-of-sequence fault reactivation (Puigdefabregas et al., 1991). Three main thrust sheets are present, the Boixols, the Montsec, and the Sierras Marginales, and they link along their eastern 111 < D n j C 3 * - ® |E < j z O £ o ® m v ) c S Figure 1. The southern Pyrenean Foreland Basin. Inset shows the approximate location of the study area, along the western flank of the South-Central Unit thrust system, and the sim plified configuration of the important structural elements discussed in the text. 112 boundary into the Segre Fault Zone oblique ramp (Fig. 1) (Verges & Munoz, 1990). Along their western boundary, a more poorly defined, diffuse oblique ramp system is seen, with the development of a number of transport-oblique anticlines (Boltana and Mediano). This study is concerned mainly with the later development of this western oblique ramp and with the controls that its growth exerted on the middle-late Eocene syntectonic depositional systems. The western oblique ramp system separates three small, distinct sedimentary basins (Nijman, 1989). In the east, in the hanging wall of the South-Central Unit thrust system, one sees the E-W -elongate Tremp Basin (Fig. 1). Directly upon the oblique ramp, in between the Mediano and Boltana anticlines, is the smaller N-S-elongate Ainsa Basin (Figs. 1 & 2). Further to the west of the oblique ramp, sitting in the footwall of the SCU, is the Jaca basin. The remnants of the Escanilla drainage system are exposed across all three basins, preserve the migration of alluvial facies belts, and were deposited during important phases in basin development and delineation. First, we will describe the internal stratigraphy and large-scale architecture of the Escanilla Formation from two areas exposed across the western oblique ramp of the South-Central Pyrenean thrust system, in the Ainsa and Tremp Basins (Fig. 1). A three-fold subdivision of the Escanilla Formation based upon these observations is constructed for 113 Ainsa • km S o Q_ 1 0 3 28 So Arcusa • / \ — • Mediano ^ I -I6 0 \ tv [ 1 P re -T \ So M E D • Almazorre Pre-T O lson: LIG E R i Barcabo ALZ Oligocene (Collegats) Conglomerates P re - 3 H I 0 Ainsa Grp. Eocene So Magnetic Sobrarbe Escanilla Pre-fold Sequence Turbidites Lst. Fm. Fm. Section Figure 2. Simplified geologic base-map of the western area, the Ainsa Basin or Buil Syncline, situated along the western oblique ramp of the South-Central Pyrenean thrust system. The location of villages within the study area, and the magnetostratigraphic traverses are shown, as are the important structural features within the Ainsa Basin (ALZ = Almazorre. ERI = Eripol. MED = Mediano. LIG = Liguerre). i 114 the Ainsa Basin, and this is carried to other parts of the ancient alluvial drainage network, in order to reconstruct the nature of this Eocene river system. Our Escanilla Formation reference stratigraphies will then be contrasted with work of other authors in the downstream equivalents of the alluvial system preserved within the Jaca Basin (Puigdefabregas, 1975; Jolley, 1988; McElroy, 1990; Hogan, 1992). Such regional correlations have been made possible by the use of magnetic polarity stratigraphy: a tool that has proven useful for constraining the tectono- stratigraphic development of this region within an absolute, rather than a relative temporal framework (Hogan, Burbank, & Puigdefabregas, 1988; King Powers, 1989; Burbank et al., 1992). These descriptions and comparisons will then be used to develop general comments concerning the spatial changes in alluvial architecture and alluvial drainage pattern, and how these appear to have been closely controlled by lateral variations in the structural development, and by the internal partitioning of this major foreland basin system. EOCENE STRUCTURAL AND STRATIGRAPHIC FRAMEWORK OF THE SOUTH-CENTRAL PYRENEES Structural Developm ent The Pyrenees are an E-W-trending mountain system developed at the NE corner of the Iberian Peninsula, separating Spain from the rest of Europe. Formed during a phase of late Cretaceous-Miocene 115 convergence and limited northward underthrusting of the Iberian plate beneath Eurasia (Daignieres et al., 1989; Munoz, 1991), the Pyrenean collision began earlier in the east than in the west, and the onset of thrust deformation shows a strong diachrony as one traverses along the chain from east-to-west. Three distinct structural provinces are differentiated within the southern Pyrenean system (Fig. 1). The eastern Pyrenees, situated to the east of the Segre Fault Zone, are made up of a series of cover and basement involved thrusts and their associated piggy-back basin that is usually referred to as the Ripoll basin. This area is not considered to have been part of the same dispersal system as the Escanilla drainage basin during late Eocene time, and therefore, is not discussed further in any detail. In the southern Pyrenees, basins developed during Paleocene to early Eocene times, ahead of the southerly translating thrust sheets, partly in response to thrust wedge loading, and partly due to subduction- related flexure of the down-going Iberian Plate (Munoz, 1991). Collectively referred to as the South Pyrenean basin (Puigdefabregas, 1975), this region began to partition or compartmentalize during the early Eocene epoch. This occurred in response to the incorporation of proximal parts of the foreland into the developing South Pyrenean thrust wedge, as thrust-sheet-top or piggy-back basins were formed in the hanging-wall of the advancing thrust system. The Mediano and Boltana oblique ramp folds initiated during early Eocene time, at the onset of thrust motion and piggy-back basin formation, and both have long protracted histories (Camara & Klimowitz, 116 1985; Farrell, Williams & Atkinson, 1987; Reynolds, 1987; Nijman, 1989; McElroy, 1990). Growth of the Mediano anticline probably ended by Bartonian times, whereas the Boltana anticline was active into the Priabonian Stage. Sitting structurally between these two folds is the Ainsa Basin (Fig. 2), the northern margin of which was first defined during Ypresian-Lutetian times and it is currently expressed as the E-W-trending Cotiella thrust system exposed along the foot of the Pena Montanesa mountain (Nijman, 1989). Across the Boltana anticline to the west, the Jaca Basin showed a two-phase history. Prior to the growth of the Sierras Exteriores, and the definition of the southern structural margin of the Jaca Basin, a flexural foreland basin existed that saw largely deep-marine turbidite deposition during early to middle Eocene times. During, and subsequent to thrust detachment in late Eocene and early Oligocene time and the formation of the piggy-back Jaca Basin (Almela and Rios, 1951; McElroy, 1990; Hogan, 1992), regional sedimentary systems were being strongly modified by ongoing structural deformation (Puigdefabregas, 1975). This deformation continued through into Miocene time when discrete thrust motion largely ceased within this region (Crusafont, Riba & Villena, 1966; Soler & Puigdefabregas, 1970), and the current basin configuration was fixed. S tratigraphy The stratigraphic framework of the southern Pyrenean basins has become particularly complex and confused, showing little or no 117 standardization across research groups. Here we adopt a modified version of the system used by Cuevas Gozalo (1990) for the eastern part of the study area located within the Tremp Basin (Figs. 3 and 4). For the western region, west of the Mediano anticline in the Ainsa and Jaca Basins, we slightly modify the terminology of Puigdefabregas (1975) and Nijman and Nio (1975) (Fig. 3). In the east, the Lutetian tidally-influenced alluvial deposits of the Capelia Formation are overlain by the Escanilla Limestone and succeeding Escanilla fluvial deposits representing the lower Campodarbe Group (Puigdefabregas, 1975) (Fig. 4). It is acknowledged, however, that the base of the Escanilla Formation is placed at the base of the Escanilla Limestone purely for convenience. This stratigraphic level may not necessarily represent the same chronologic level as the base of the Escanilla Formation defined within the Ainsa Basin. Across the Mediano anticline, correlation is difficult, partly due to lateral facies changes and partly as a function of incomplete and discontinuous exposure. On the fold axis, the Escanilla Formation thins and lies on top of deltaic deposits of the Sobrarbe Formation, and unconformably upon shallow marine platform carbonates of the Puy de Cinca Limestones, bearing testimony to the unit's syntectonic depositional history. Further to the west, in the Ainsa Basin between the two surface anticlines, the Sobrarbe deltaic facies thicken and build a northward-prograding wedge of siliciclastic and carbonate sediments within the synclinal axis (Figs. 2 & 3). In this area, the upward transition into the Escanilla fluvial sequence, although conformable and gradual, is 118 Age Oligocene Bartonian- Priabonian Lutetian Ypresian WESTERN AREA A in s a B a s in R IO C IN C A Oligocene Conglomerates CL 0 C D - Q 1 _ C O x > o CL E C O O Escanilla Fm P u y d e C in c a L s t w CD ' • g lo 3 I — C L 3 O O o sz o CD X < 0 £ = C D & g « s c o cr c ® CO CO C O C a s tis e n t ' 1 S e q u e n c e C D C O " ■ + — • c < o c eg ID C D EASTERN AREA T re m p B a s in E S E R A IS A B E N A Oligocene Conglomerates cL * _ O C D X) C O • o o C L E C O O P u y d e E s c a n illa C in c a L s t i ' . ' / . ' i I.-’ ? T ^ S 5 & S i Escanilla Fm CD O c z CD 3 cr CD C O s V ) CD CO Capella Fm. u S’ Campanue Perarrua Fm. Castisent Fm. Figure 3 . The adopted lithostratigraphy nomenclature employed during this study. The western area stratigraphic fram ework is essentially that of Puigdefabregas (1975), and the scheme for the eastern area is modified after Cuevas Gozalo (1990). 119 Sheet Conglomerate Horizon * Escanilla Limestone yiiigpi ^ ^ ' I * ? W i , , , . ' ''V,, Collegats Group 0 ^ km 1 “ v . ^ Magnetic Section Escanilla Formation „ __ Rivers _ II r- uapGila Formation village Perrarua Formation 18 Dlp/Strlke Normal Fault Figure 4 . Simplified geologic base-map of the eastern study area within the Tremp 'piggy-back’ basin. The location of the Lascuarre m agneto stratigraphic traverse is shown, as is the lacustrine interval represented by the Escanilla Limestone. The two villages adjacent to the section location are also shown. 120 interrupted by a rapid and short-lived marine transgression. The greatest thickness (> 1 km) of Escanilla Formation preserved in the study area is developed within this syncline (Fig. 5A), and it may be subdivided, on the basis of alluvial style and architecture, into 3 distinctive units, informally referred to as the ‘lower’, ‘middle’ and ‘upper’ members (Fig. 5B). The nature of these will be described and discussed in detail in the following sections. East of the Mediano anticline, in the Tremp Basin (Figs. 1 and 4), correlation of the lower parts of the Escanilla sequence was attempted, but because of erosion and removal of Escanilla section at the base of the overlying Oligocene conglomerates, a comparison of the complete stratigraphy cannot be made. These Oligocene-aged conglomerates unconformably overlie the Escanilla Formation across the whole study area, and in places, the complete Escanilla section has been removed and the conglomerates sit directly upon the Lutetian tidal, or carbonate platform deposits of the Capella Formation, Puy de Cinca and Guara Limestones respectively. M ETH O D S In order to constrain the structural and stratigraphic development of this part of the Pyrenean foreland, five magnetostratigraphic sections were constructed across the SCU western oblique ramp system, in regions of particularly good and continuous exposure (for the positions of Oligocene Conglomerates 1000 C/3 w <33 C . O I- o Ic Q . CO g > CO G O 500 b P - O V - o ° • o* . c » • w Large-scale upwards-coarsening and thickening of sheet channel bodies. Increased presence of calcareous overbank and calcrete development UPPER EXTENSIVE SHEETS Interval dominated by overbank silts and low interconnectedness of sheet sandstones - AMALGAMATED SHEET SANDS AND CONGLOMERATES First appearance of distinctive red-matrix quartz-pebble conglomerate clasts Interval dominated by overbank silts and low interconnectedness of sheet and ribbon sandstones Fine delta-top sequence, high sinuosity channels and gypsiferous overbank, Deltaic and shallow marine nummultic marls and siliciciastic sandstones 100-. MPS Hecho Group Figure 5A. Stratigraphic subdivision and nature of the lower Campodarbe Group sediments exposed within the Ainsa Basin. The sum mary magnetic polarity stratigraphy (MPS) constructed within the syncline is shown. Black represent tim es of normal magnetic field orientation, and white represents reversed directions. The MPS is correlated with the magnetic polarity tim e-scale (MPTS) of Hariand et al. (1990) in Fig. 7. 1 2 2 Escanilla Stratigraphic Subdivision Lower Lower Middle f Upper . Figure 5B. Simplified geologic map of the southern Ainsa Basin showing the spatial distribution of the three members of the Escanilla Formation. Lack of exposure at higher stratigraphic levels prevents division of the ‘m iddle’ and ‘upper’ members away from the core of the Buil syncline. 123 these magnetostratigraphic traverses refer to Figs. 1, 2 and 4). Lithologic and sedimentologic data were collected at a sub-m eter scale. Paleocurrent directions were collected from both uni-, and bi-directional flow indicators, including planar cross-sets, trough axes, gutter structures and large clast imbrications. When found in conjunction with uni directional indicators, bi-directional data were reinterpreted to be more explicitly recording paleoflow. Although some of this data was taken from two-dimensional, vertical exposures and basal scour surfaces, the common occurrence of exhumed upper boundaries allowed the collection of the majority of this data from excellent three-dimensional flow indicators. Clast populations were systematically recorded at a number of locations throughout the main exposures of the Escanilla system (Fig. 6). The first appearances of any particularly distinctive clast lithologies were also noted, and recorded on the stratigraphic logs. Mapping of sedimentary facies and geometric or structural relationships was completed, in order to describe the spatial correlation of sedimentary units, and the relationships of these units to the important contemporaneous structures. Lateral facies analysis was performed in the field, using both land and aerial photography. Channel and overbank lithofacies were distinguished, and the three-dimensional spatial arrangement and relative proportions of these two main facies groups were then used to characterize the changing alluvial architecture of the Escanilla system. 124 oono yaddn anaaiw c o £ C D yaddn H — • _C o aiaaiw 1 c a_ C O da/won D) " 5 H — • yaddn C O anaaiiAi d3MO"l o o o L O eouepunqv eBEiueojed Figure 6 . Graph showing the varying clast population data for the three members of the Escanilla Formation, and the overlying Oligocene conglom erates exposed within the Ainsa Basin. Four general groups have been delineated. The extreme right column represents the Oligocene strata sampled within the center of the Buil syncline, away from local input of Guara limestone off the Boltana anticline. 125 Paleomagnetic samples were collected at regular intervals, in order to best constrain the paleomagnetic reversal pattern preserved in the sequence. Oriented block samples were collected for the magnetic analyses, and the sampled surface and local bedding orientations were recorded, in order to remove any post-depositional tilting and yield a preserved paleomagnetic field direction. At least four specimens were cut and analyzed for a given sample location. Prior to wholesale magnetic analysis, a pilot study employing both thermal and alternating-field (AF) demagnetization techniques was completed. These data suggested that the continental deposits typically yield their characteristic natural remanent magnetic directions between 250-400 °C, and the marine lithologies between 200-400 Oersteds (after the low temperature removal of any viscous overprints). Bulk sample measurement was completed at three demagnetization levels within these ranges, and a characteristic paleomagnetic field direction for each specimen was noted. The magnetization direction for a given sample location at a given temperature was calculated using the statistical methods of Fisher (1953). The statistical quality of each site was assessed and classified as ‘Class I’ (statistically robust agreement of 3 or more individual specimens within a sample location), ‘Class IT (apparent, but not statistically robust agreement of 2 or more individual specimens), or ‘Class III’ (poor agreement of specimens at a site) using the same scheme of Burbank et al. (1992). Additional class II sites were defined by average site vectors that showed southerly declinations but negative (upward) inclinations. The average magnetic vectors from the Class I 1 26 and Class II sites were then used to calculate virtual geomagnetic pole (VGP) paleolatitudes. These were then used to classify each site as either normal or reversed polarity. The errors in the average magnetic vectors were used to calculate an ag5 error envelope for the VGP paleolatitude, giving further assessment of the site data quality. Magnetic chrons were ideally defined by two or more adjacent data points of similar polarity, but one point reversals were noted. The erected reversal stratigraphies were then correlated with the Magnetic Polarity Time- Scale of Harland et al. (1990), using independent biostratigraphic and magnetostratigraphic age control within the underlying shallow marine units across the study area (Cuevas Gozalo, 1990; King Powers, 1989). Applying the time constraints given by the careful construction of reliable magnetic chronologies, it was possible to calculate sediment accumulation rates across the study area (Fig. 7). The data presented in this paper represent crude undecompacted sediment accumulation rates averaged over the interval of a well-defined magnetic polarity chron. Because the application of decompaction and subsidence modeling within such an alluvial system is not well constrained for basins lacking extensive subsurface porosity information, such an analysis was not completed here. This should not, however, affect our conclusions, as undertaking such a study would only significantly change the absolute magnitudes of our accumulation data, whereas, the relative orders would remain constant. As all we require in this study are spatial comparisons of rates, we feel confident that our conclusions are not simply a function of our sim pler first-order approach. 127 I — CO < LU ■ > 1 -36 -37 -38 -39 -40 -41 3 ? 3 1 1 1 NVINOSVIUd N V I N O l U V a | N V l l H i m m m m m m f t ; ; > i C“ / / O 0 9 S& '^S s s *V j f t S S l .2 S J » M W m L U mem sjbox u«||!^ u i ejBOS suiLL Figure 7. Cross syncline litho- and magneto-stratigraphic correlation of the Escanilla Formation across the study area. Lithologic correlations confirm ed in the field are shown as solid lines linking the magneto stratigraphic traverses. Correlations based on the comparison of the MPS’s with each other and with the MPTS, are shown as dashed lines. Average undecompacted sedimentation rates within each member of the Escanilla system are shown, and are calculated using the ages of chron boundaries taken from Harland et al. (1990). 1 2 8 STRATIGRAPHIC SUBDIVISION OF THE ESCANILLA FORMATION Ainsa Basin The upper Eocene fluvial sediments of the Escanilla Formation were first described and named by Garrido-Megias (1968). As mentioned previously, the late Eocene Escanilla Formation lies conformably upon deltaic deposits of the Sobrarbe Formation (Figs. 2 & 5A). This system prograded northwards along the synclinal basin axis during latest Lutetian to earliest Bartonian times, and the package of sediment thins markedly to both the east and west, as the flanking structures are approached. Rotational slumps and slides mapped by De Federico (1981) indicate mass movement away from both fold axes into the syncline suggesting that the folds were actively growing at this time. On the basis of the coarseness and internal geometries of the channel-fill, the Escanilla formation has been interpreted as the record of low-sinuosity stream systems generally flowing axially along the partitioning late Eocene foreland basin (Garrido-Megias, 1973; Jolley, 1988; Cuevas & Puigdefabregas, 1991). They were fed by paleovalley sequences flanking the southern margin of the uplifting core of the Pyrenean Axial Zone (Puigdefabregas et al., 1985; Reynolds, 1987; Vincent and Elliott, in prep). Although the sediments do not fit classic braided stream facies models, this interpretation is essentially correct. However, the Escanilla deposits are considered anomalous because of the high proportion of fine-grained overbank material preserved in 129 association with coarse conglomeratic channel-bodies (Bentham, Tailing and Burbank, 1991; in prep). On the basis of systematic changes in alluvial architecture the 1 km-thick Escanilla Formation in the Ainsa basin, we divided the Escanilla Formation into three distinct units, hereby referred to informally as the ‘lower’, ‘middle’, and ‘upper’ members (Figs. 5A, 5B, and 7). Lower Member Based on lithostratigraphic and magnetostratigraphic correlations, this lower unit varies in thickness laterally across the Ainsa Basin (Fig. 7), and reaches a maximum thickness of about 400 m in the synclinal axis. The unit then thins westwards to about 200 m thickness in the region of Almazorre, along the eastern limb of the Boltana anticline. The units lower boundary is defined at the first appearance of reddened continental overbank mudstones and siltstones, and the subsequent lack of marine fauna, whereas the upper boundary is drawn at a thick, laterally extensive horizon of amalgamated sheet sandstones and conglomerates. Based on our magnetostratigraphic correlations with the Harland et al. (1990) timescale, this interval ranges in age from -42.7- 41.4 Ma (late Lutetian-early Bartonian), and is typified by high sinuosity channels developed on a low gradient coastal alluvial plain. Channel sand-bodies often show evidence for channel-scale lateral accretion. Associated fine-grained overbank deposits are commonly gypsiferous. Approximately 25-30 m above the traditionally mapped base of the Escanilla Formation we noted a very planar-based sheet sandstone that 130 was not obviously channelized. Exposed along the synclinal axis, and immediately underlain by marine or lagoonal gray marls with bivalves, this sandstone shows bi-modal, small-scale ripple and trough-cross stratification and is capped by a horizon rich in shallow marine nummulitic foraminifera. Above this sandstone, fluvial sediments return and gypsiferous overbank deposits are common within the lower 150 m of this 300 m- thick package. This lower unit typically shows a much high proportion of overbank fines, enclosing much coarser channelized sand and conglomerate bodies. Filled by coarse-pebbly sandstones and subordinate volumes of gravel, these channels are most often sheet sands and conglomerates with rare sandstone ribbons. On the whole, they usually exhibit low degrees of vertical channel interconnectedness or amalgamation. Channel bases are generally strongly erosive, and the succeeding aggradational channel fills occasionally show sand or gravel lateral accretion surfaces in their lower parts. Subsequent channel deposits are most typically coarser trough and planar cross-bedded coarse sandstones and conglomerates that preserve a general fining- upward trend. Enclosing overbank are strongly burrow or root mottled, and tops of channel sequences often preserve extensive vertical root structures suggestive of tall grass and reed stands. Paleocurrents within this lower unit vary considerably, but systematically through time. Initial channels show flow directed off the top of the Sobrarbe Delta towards the NW-NNW (Fig. 8A), thereby mimicking the progradational advance of the delta along the axis of the 131 actively subsiding syncline (Nijman & Nio, 1975; Puigdefabregas, 1975; Barnolas, pers. comm.). These directions are quickly succeeded by paleocurrents that define a broad sweeping arc, flowing westwards from the Mediano anticline then progressively rotating SW and SSW as one traverses the Ainsa Basin (Fig. 8B). The lowest channels along the western flank of the syncline show flow dominantly directed towards the SSW approaching parallelism with the axis of the Boltana anticline. The top of this unit is defined immediately below a very prominent continuous horizon up to 30-40 m-thick, made of both laterally and vertically stacked coarse sand and conglomeratic sheet channel-bodies. This unit can be followed laterally from east-to-west across the entire Ainsa Basin (Cuevas & Puigdefabregas, 1991) (Fig. 5B & 7). It is also very conspicuous on aerial photographs where it may be traced continuously from west of the village of Escanilla to the region of Barcabo, along the eastern flank of the Boltana anticline. Fine-grained material is much less voluminous in this interval, and the degree of channel body interconnectedness is much higher than in the units both above and below. Approximately 20-30 meters below this laterally continuous interval, clast composition studies show the first appearance (in significant volumes) of a rather distinctive clast lithology (Fig. 5A). These red matrix, quartz pebble conglomerate clasts are seen throughout the rest of the Escanilla section, but are not seen in proportions above a few tenths of one percent below this point (Fig. 6). 132 A. Uppermost Sobrarbe Fm. & basal Escanilla Fm. (-43 Ma) Tidal Reworking on Delta-too V'VVm Figure 8A. Paleocurrent data from tne upper Sobrarbe deltaic and basal Escanilla coastal plam fluvial sediments of the Ainsa B asn prior to the marme transgression at -4 2 .7 Ma. Numbers within the paleocurrent roses show the number of measurements taken during a given interval. The location of the rose s used to indicate the region where the data was collected. This is true of Fig. 8 B & Fig, 8 0 also 133 B, Lower Escanilla Fm. (-42.5-41 Ma) Santa Maria del Buil K * * ' 4 k I i w ® Figure 8 6 . Paleocurrent data collected within the lo w e r’ Escanilla Formation between the short marine transgression and the am algamated conglom erate sheet deposited at ~41 Ma, 134 Middle Member Above the laterally continuous level of sheet sand and conglomerate channels, one sees a return to fluvial sediments sim ilar to those immediately below. Coarse conglomeratic sheets and rare ribbon sand- bodies are seen wholly enclosed within contemporaneous overbank material (Fig. 5A). Based on our erected magnetostratigraphic correlations, this unit ranges in age from 41.4--39.0 Ma (Bartonian) and appears to thicken slightly to the west reaching a thickness of -300 m (Fig. 7). The member's upper limit is difficult to place and is defined on the basis of upwards-coarsening and thickening trends. It is placed at the next laterally continuous level of conglomerates that may be traced across the core of the syncline. Due to the presence of post-folding erosion and removal of section prior to the deposition of the Oligocene Collegats group, the higher parts of the Escanilla Formation are removed along the flanks of the Buil syncline, and wider regional correlation at this stratigraphic level is not possible. This member shows a general coarsening-up sequence, and rare thin calcrete horizons have been observed. Overbank material remains very strongly mottled, and much finer-grained than any observed channel fill material. Thin (<11 cm- thick) micritic limestones are occasionally preserved within this overbank interval, but these are quite rare and are not generally laterally inextensive (normally less than few 100's of meters) 135 Upper Member Only extensively preserved within the center of the Buil syncline, this unit is characterized by thick (>5 m), coarse, wide (100's m) conglomeratic sheets and comparable intervening thicknesses of calcareous, finer-grained vertical accretion deposits (Fig. 5A). Conglomerates show poorly defined structures such as planar and trough cross-bedding, and often show laterally extensive scour surfaces. Within the reddened overbank siltstones, calcretes are more commonly developed, and thin micritic limestone layers also show an increase in abundance. The true thickness of this unit is unknown due to erosional truncation below the basal Collegats Group unconformity (Fig. 7), and so its absolute age range is also unknown. This unit coarsens upwards, with maximum clast diameters increasing from about 20 to 40 cm. Uppermost clast compositions preserved within the Escanilla Formation are very similar to the overlying Collegats Group populations, suggesting common source areas and clast provenance existed across the unconformity (Fig. 6). Within the core of the syncline, the unconformity is marked by a slight angular discordance in southerly dip between the upper Escanilla sheets and the basal Collegats Group. However, along the western flank of the syncline, an extremely strong angular unconformity exists with the lower Escanilla, Sobrarbe, and Guara Limestone Formations (Fig. 2). This motion may be related to both continued growth of the Boltana anticline, and to uplift and emergence of the Sierras Exteriores in Priabonian and early Oligocene times. 136 Similar to flow within the underlying unit, the paleocurrents from the middle and upper members define a sweeping arc from east-to-west crossing the Ainsa Basin, furthermore, the change from west-directed top south-directed flow is even more dramatic. Rather dispersed flow indicators along the western flank of the Ainsa Basin generally show a progressive diversion into parallelism with the Boltana anticline (Fig. 8C), however, immediately south of Almazorre, sandstone channels suggest flow E and SE, away from the rising, or already emergent fold. Western Tremp Basin Immediately south of the Isabena Valley, in the region of Lascuarre (Fig. 4), the Escanilla Formation reaches a thickness of only -500-550 m (Fig. 9). The base is defined by the laterally extensive Escanilla Limestone. This unit can reach a thickness of up to 30 m, and may contain a maximum of 16 separate lacustrine intervals (Nickel, 1982). In this region, it has been suggested that a significant hiatus exists between the Escanilla Limestone and the overlying fluvial deposits (Cuevas Gozalo, 1990). However, no angular relationships are seen and no discrete omission or erosion surfaces are present (Cuevas & Puigdefabregas, 1991). Coals (lignites) exist in association with this basal Escanilla sequence (Garrido-Megias, 1968; Cuevas Gozalo, 1990). Overlying this, the terrestrial alluvial deposits representing the rest of Escanilla Formation show a similar upwards-coarsening and -thickening, 137 C. Middle-Upper Escanilla Fm. (-42.5-41 Ma) Figure 8C. Paleocurrent data from the ‘m iddle’ and ‘upper’ divisions of the Escanilla Formation. Note the reversal and diversion of flow to the south and south-east, parallel to and away from the emerging Boltana anticline. 138 Oligocene Conglomerates ' • “ i ifo - O ’ C o " J 5 E o LJ- 1 5 E 03 o C O LU W Q > .Q E Q ) Q ) a . a . « > T 3 T 3 a > x> E a > 500 _ * P o o rly e xp o se d up w a rd s c o a rs e n in g s e q ue n ce of co n glo m e ra te s a n d s sts, w ith in c re a s e d p re se n ce of n o du lar calcrete h orizons. T h in la cu strin e lim esto ne h o rizon s are ra re . 400. - MPS 300. 200. 100 Capella Fm. Interval of th in , laterally e xte n sive sheet ssts a n d thicker (~1.5 m ) w hite, m icrtic lim e s to n e s . S ed im e n t is finer-grained th a n u nits a b o ve or b e lo w . C ha n n e lise d a n d sheet-like ssts. In te rb e d d e d w ith c o m m o n , laterally e xte n sive m icritic lim e sto n e s. S he e ts a re coarser a n d thicker th a n th o s e in ove rlyin g unit In terva l of laterally a n d vertically a m a lg a m a te d cqarse^sts a n d co n glo m e ra tic s h e e ts . - 1 st ap pe a ra n ce of re d q uartz p e b b le c o n glo m e ra te w ithin co a rse c h a n n e l b o d ie s . C oa rser inte rva l, w he re sa n d s a re typically w hite a n d arkosic, w ith granitic c la sts a s rare clasts w ithin th e th in co n g lo m e ra te s . L im e s to n e s a re ra re . Escanilla Limestone Figure 9. Stratigraphic subdivision and nature of the lower Campodarbe Group sedim ents exposed south of the Isabena Valley (see Figs. 1 & 4 for map location). Note the gross coarsening-up trend, and the regular occurrence of thick lacustrine limestone units throughout the vertical sequence. 139 regressive mega-sequence as seen further to the west. Here, however, the presence of thin limestone sheets (generally <1.5 m) throughout the section is extremely striking. Such units are very rare in the Buil syncline, and when seen, they are significantly thinner (<20 cm) than in these eastern exposures. In the east, the limestones occur with greater frequency in association with both the fine overbank and coarse conglomeratic channel-fill. Approximately 170 m above the base of the Escanilla Formation is a very prominent, laterally amalgamated sheet of wide conglomeratic channel bodies. This unit may be traced easily across the Escanilla exposure front, south of Lascuarre, until it dives westwards below the overlying Oligocene unconformity south of the village of Laguarres. As within the Ainsa Basin, 20 meters or so below this laterally extensive interval, one sees the first appearance of reworked, red matrix quartz pebble conglomeratic clasts within the channel fill deposits (Fig. 9). Conglomeratic channels in this area are comparable to, although somewhat coarser and thicker than, those seen above and below the amalgamated interval within the Ainsa Basin. The section also shows that the channel-fill coarsens upwards more rapidly, and wholly sand-filled channels are extremely rare. Ribbon sandstones are less evident than further to the west, and wider sheets dominate the hillside exposures. Overbank material is generally much more calcareous, and nodular caliche paleosol horizons are much more common here than across the Mediano anticline to the west.. 140 REGIONAL CORRELATION Ainsa and Tremp Basins Based on clast population studies (the associated first-appearance of the Triassic quartz-pebble conglomerate), fluvial facies comparisons, and magnetostratigraphic correlation (Fig. 7), the amalgamated unit exposed within the Tremp Basin is interpreted as the same laterally consistent horizon that may be traced across the Ainsa Basin, and as such represents a regionally significant depositional surface. The lower 170 meters of the Escanilla section in this region would, therefore, be the eastern expression of the ‘lower’ member described between the Mediano and Boltana anticlines. The remaining Escanilla strata would then represent the ‘middle’ and ‘upper’ members, but no attempt is made to differentiate the two, because of the uncertainty concerning the amount of section that has been removed in this area by erosion. A general summary of the re-interpreted stratigraphic and geometric relationships preserved across the study area are shown in Fig. 10. Correlation of the underlying Escanilla Limestone with other units, even in the light of the magnetostratigraphic data from Lascuarre, remains inconclusive. There are two possible interpretations. Nijman & Nio (1975) suggested that the lacustrine limestone sequence, with it’s associated coals, would correlate to the west with a marine transgression during Puy de Cinca limestone sedimentation (Middle Lutetian, ~ 45-44 Ma). This transgression would have induced a rise in the regional water table and formed limestone and coals within the alluvial regions of the 141 Tremp Basin. Ponded facies within the lower parts of the Sis and Grup paleovalley feeder systems to the north and northeast of the Lascuarre region are believed to have been deposited at this time also (Vincent, pers. comm.). However, in the light of the paleomagnetic data available at this time, this interpretation requires non-deposition or erosion immediately after Escanilla limestone formation, prior to the onset Escanilla fluvial deposition. Given the lack of evidence for any unconformity between the Escanilla limestone and succeeding fluvial deposits, a second, favored interpretation is possible. This would suggest that Escanilla limestone deposition was largely correlable with Sobrarbe delta progradation within the Ainsa Basin. No major break is then required at the top of the lacustrine sequence. This might suggest that lacustrine deposition occurred as ponded alluvial deposition upstream of the emerging Mediano anticline. Correlation to the Jaca Basin Combining the mapping and facies descriptions of earlier studies (Puigdefabregas, 1975; Jolley, 1988; McElroy, 1990) with the chronologic control of Hogan (1992), it is possible to correlate and compare units down the alluvial system, across the Boltana anticline, into the eastern Jaca Basin (Fig. 11). Puigdefabregas (1975) described a rapid deepening in marine environments, shown by the drowning of the Guara Limestone platform carbonates, and subsequent deposition of the Arguis Marls (Fig. 11). 142 L L 1 §1 S o Q C L U £= D C uj O z Z j C D O 2 ? <2 end z ^ C D < Z <z S < -J i- g -J L. 0 5 « < o c C O y , n J Q ) a. c o Figure 10. Regional W-E cross-sectional sum mary of the inferred stratigraphic relationships discussed within this study. The spatial relationships of the magneto-stratigraphic sections are shown (LAS refers to the Lascuarre section shown on Fig. 4). The simplified structural configuration is also presented. 143 Hogan (1992) was able to date this event as occurring immediately prior to normal magnetic polarity chron 19 (-42.6 Ma after Harland et al., 1990), and correlated it with a major sea-level rise right at the Lutetian- Bartonian boundary shown by the Haq et al. (1987) sea-level curve. We suggest that the marine transgression observed at the base of the Escanilla Formation within the Ainsa Basin may represent the coastal expression of this eustatic event (dated at 42.7±0.1 Ma in the Ainsa Basin). To the west of the Boltana anticline, the continental deposits of the lower Campodarbe Group sit conformably upon the shallow marine and deltaic facies of the Belsue-Atares group, however, immediately along the western flank of the fold, the basal Campodarbe sediments lie in strong angular unconformity against early Eocene sediments independently documenting the earlier phase of fold growth during the early Lutetian (De Federico, 1981) (Fig. 11). Close to Nocito, west of the Boltana anticline, one can clearly see an interfingering of deltaic and continental deposits reflecting small, short-term variations in relative sea- level in a way analogous to the significant marine transgression identified within the Ainsa Basin. The facies boundary separating the Belsue and lower Campodarbe clearly migrates about 5 km eastwards, reversing the longer-term westerly-directed regressive trend. Slightly further to the east, in the region of the village of Bara (-1 5 km west of the Boltana anticline), the facies transition appears to have been localized for a significant time, possibly in response to the growth of the N-S- 144 Ill o c o o C D O — c = > J £ ? .o < u o C D C / 3 C l O o ® i-P Q -ra — T 3 . Q . r CL ■ a c- " C 3 = > O — 0 3 O *— 5 t - 3 c 3 O O ) a > c 5 Figure 11. Simplified geologic map of the Sierras Exteriores between the Boltana anticline and Arguis. Detail is taken from the map of Puigdefabregas (1975). Important structural features are shown, and the interfingering relationships of the deltaic and continental facies can clearly be seen. H M 145 ^9999999 trending anticline Alcanadre anticline (Fig. 11). in a sim ilar way, but over a longer timescale, the middle Eocene sediments of the Montanyana deltaic system were localized along the western edge of the Tremp Basin by the development of the Mediano anticline. Fold growth may serve to stabilize facies belts for long periods of time (De Boer, Pragt, & Oost, 1991), preventing the lateral migration of depositional environments. As a result, the transition from deltaic to continental sedimentation is strongly diachronous. This diachrony may be quantified if the MPS data from this study are combined with those of Hogan (1992). His sections constrain the age of the top Belsue deltaic strata in the region of Arguis (Fig. 11) to be 5 myr younger than within the Buil syncline. The delta, therefore, can be shown to change age from Lutetian in the region flanking the Mediano anticline, to wholly Priabonian in the central Jaca Basin, across a horizontal distance of less than 100 km. The Eocene shoreline moved at an average progradation rate of about 10 m/kyr, and although this is strongly dependent on factors including sea-level variation and sediment supply, the most important control seems to have been the continued growth of a number of N-S-trending folds (Almela and Rios, 1951; Puigdefabregas, 1975; McElroy, 1990), blocking the westward regression of marine environments. In the region of Bara, at the northern end of the Alcanadre anticline (Fig. 11), one can clearly see the effects of structural development. The coastline was localized for a significant period of time, and the marine-continental facies boundary did not migrate laterally. The spatial distribution of the syn-depositional folds probably resulted in a rather episodic, step-wise migration of the 146 shoreline, as the shoreline would abut against a growing structure until sediment supply outstripping local subsidence (Puigdefabregas, 1975). The coastline would then prograde rapidly westwards until it met the next structural barrier. Because of this diachrony, correlation of time-equivalent packages of sediment downstream from the Escanilla Formation within the Ainsa Basin into the deltaic system, and the overlying lower Campodarbe Group is not simple. In the region of the Sierras Exteriores and the Jaca Basin, the Campodarbe Group is commonly subdivided using a tripartite system into lower, middle, and upper members (Puigdefabregas, 1975; Jolley, 1988). It must be stressed that these divisions in no way relate to those within the Escanilla Formation. Higher parts of the Escanilla are only time equivalent with the lower Campodarbe member within the Jaca Basin. The lower and middle Escanilla deposits would correlate with times of marine Arguis Marl deposition and Belsue-Atares delta progradation. Hogan’s work suggests that average undecompacted sediment accumulation rates in the region of Arguis were of the order of 35 cm/1000 yrs within the basal 500 m of the Campodarbe. This would then correlate with the upper Escanilla Formation within the Buil syncline. The alluvial architecture in this region was described in detail by Jolley (1988). Jolley’s Arguis section is typified by her “Type 1” channel-fill sequences. These mainly ribbon channels show low degrees of interconnectedness, and are preserved in association with high proportions of overbank fines that show immature paleosol development. Above this, Jolley (1988) described a shift to more laterally extensive 147 thick conglomeratic sheets that are very similar to the uppermost sheets further to the east. Although developed well after the lower Escanilla Formation, the lower Campodarbe Group shows an extremely similar alluvial architecture. Both of them show transitions up-section to wide, coarse sand and gravel sheets, due probably to increased sediment supply, and/or slowing regional subsidence. The strong coarsening-upwards trend observed within the Escanilla Formation has largely been related to intra-basinal tectonics active within the central and eastern Pyrenees during late Eocene to Oligocene times (Puigdefabregas et al., 1991). This major reactivation of earlier thrust surfaces was expressed as rapid growth and basement involvement in the Axial Zone antiformal stack, synchronous with break-back thrusting within the thin-skinned thrust system of the South-Central Pyrenees (Burbank et al., 1992), that served to increase the taper of the southward- directed thrust wedge. This resulted in the development of significant structural relief, exposing internal parts of the mountain belt, and increasing the erosion and supply of clastic sediments to the deforming foreland basin (Puigdefabregas et al., 1991). However, the presence of a major sub-regional unconformity at the top of the Escanilla Formation, suggests that the development of the alluvial fill of the basin at this time was influenced by spatial variations in structural development, changes in basin configuration, or by a significant base-level fall. 148 DISCUSSION The development of MPS across the study area, used in association with new and existing lithostratigraphic data, has also allowed the spatial and temporal correlation of structural development and variations in alluvial architecture during this complex period of Pyrenean foreland basin history. The three-fold lithostratigraphic division of the Escanilla Fm. lends itself well to the erected MPS. Escanilla fluvial deposition prograded across the study area approximately 43-42.7 Ma. Uplift and erosion of the upper Escanilla deposits are believed to have occurred sometime soon after ~36.5 Ma. This upper limit, due to omission of section by erosion, is obviously difficult to constrain accurately. The strong reversal in drainage pattern, from the NW to SSW during basal Escanilla time (-42 Ma) in early Bartonian time is interpreted to result from out-of-sequence reactivation of the Pena Montanesa thrust system at the northern margin of the Ainsa Basin. This phase of motion was described by Farrell et al. (1987) and resulted in footwall tilting and folding forming a slight SW-dipping regional slope on top of which Escanilla rivers were diverted southwards around the southern end of the Boltana anticline (Fig. 12). However, this interpretation is not unique, and the evidence cited by Farrell et al. (1987), namely the folding and tilting of the Ainsa channels, may have occurred by tilting of the western limb of the Mediano anticline. Paleocurrents throughout most of the Escanilla Formation in the Ainsa Basin define a sweeping arc from east-to-west crossing the Ainsa Basin. 149 This shows a 90° rotation from westerly to more southerly paleo-flow directions. This is especially pronounced in the middle and upper members when the southerly-directed axis of the fluvial system seems to migrate eastwards into the intervening syncline (Fig. 12). Furthermore, along the western flank of the Ainsa Basin, south of Almazorre, sandstone channels suggest flow E and SE away from the Boltana fold axis, possibly the reflecting a phase of late Eocene growth of the Boltana anticline. This is supported by the westward thinning of the lower member of the Escanilla Formation defined within the Ainsa Basin exposures. In the synclinal axis, undecompacted sedimentation rates are generally in the region of -30+5 cm/1000 yrs, and subsequently decrease westwards to -1 7 cm/1000 yrs in the region of the Almazorre (ALZ) MPS. Increased proportions of sandstone are present in the more rapidly subsiding region, reflecting a concentration of channel sand- bodies into the synclinal axis (Cuevas & Puigdefabregas, 1991). Nonetheless, rates of subsidence where rapid enough that channel-body vertical interconnectedness remained low (Bridge & Leeder, 1978; Alexander & Leeder, 1987) Schuster & Steidtmann (1987) report similar occurrences of low- sinuosity stream channels in association with very fine enclosing overbank material from the Cretaceous of the Green River Basin. Although the geometry of their stream deposits appears more typical of an anastomosed river system and not braided like the Escanilla rivers, Schuster & Steidtmann suggest that the channels were fixed laterally, 150 Cotiella Thrust s^System Ainsa Basin W. Tremp Basin E. Jaca Basin a Sobrarbe and Basal Escanilla Fm. Lower Escanilla Time Middle-Upper Escanilla Time Figure 12. Summary of the paleocurrent information collected within the Ainsa Basin showing the sequential evolution of the Campodarbe drainage system in response to regional tectonic development. 151 unable to migrate across the floodplain because of the rapid local rates of subsidence and sediment accumulation within their subsiding foreland basin. They cite rates of sediment accumulation (undecompacted, -170 mm/1000 yrs) comparable to those calculated here within the Escanilla Formation. Furthermore, based on comparison of alluvial architecture from two sequences formed in differentially subsiding regimes in Idaho and Wyoming, Kraus & Middleton (1987) suggest that rapidly subsiding floodplains prevent the lateral migration of stream channels and subsequent reworking of finer alluvium along the fluvial network. This is due to the dominance of vertical accretion of the channel belt in order to maintain a graded profile as the region actively subsides. Such fluvial systems tend to avulse more frequently, prior to extensive lateral migration or amalgamation of units within the meander belt. Additionally, the deposition of fine, cohesive sediments, primarily along the flanks of the active channel belt can help to confine the system laterally, as the erodibility of the bank material would decrease. The presence of vegetation on these flanks would also significantly inhibit lateral channel migration by bank erosion (Smith, 1976). Such observations may serve to explain the variations in Escanilla Formation alluvial architecture, as lateral and temporal variations in subsidence and sedimentation rates can be discerned by use of the magnetostratigraphic correlations. (See Fig. 7 for undecompacted sedimentation rates calculated for the Escanilla formation within the Ainsa Basin). In summary, the development of an anomalous low- sinuosity fluvial system, dominated by large volumes of 152 uncharacteristically fine-grained overbank sediments (>40 % silt or clay), seems to have been facilitated by the presence of rapid lateral changes in subsidence rate, which in turn influenced the nature and rate of sediment accumulation. The subsiding Buil syncline helped to focus the fluvial system, and this focusing effect was not only expressed by the peculiar alluvial architecture, but also by the changes in paleoflow at this time and the progressive southward rotation of the fluvial system into parallelism with the major structural elements within the study area. The laterally amalgamated sheet at the top of the lower member of the Escanilla Formation, although not easily traceable downstream into the marine system to the west of the Boltana anticline, does seem to be a regionally significant interval. Occurring at about ~41.4-41.0 Ma, it was interpreted by Cuevas & Puigdefabregas (1991) as a response to regional base-level lowering, resulting in erosion and reworking within the proximal part of the Tremp Basin. They suggest that the fines- dominated intervals enclosing this lower part of the Escanilla section were deposited during phases of gradual base-level rise, allowing net vertical aggradation of the alluvial plain. Kraus & Middleton (1987) describe a similar interval of strongly amalgamated sheet channel bodies within the Willwood Formation of the North Bighorn Basin, Wyoming. Instead of invoking a regional base-level fall, they use pedological evidence to suggest that this was a time of lowered subsidence and decreased sediment accumulation. Such an interval, within a sequence dominated generally by vertical accretion deposits could be used to invoke a period of foreland basin stability and slowing of basement 153 subsidence. However, the amalgamated unit within the Escanilla Formation does not appear in association with increased soil maturity within the enclosing vertical accretion deposits. Gardner & Cross (1991) suggest that fluvial geometries, especially within the distal reaches of a drainage basin, may be solely modulated by variation in base-level and accommodation space availability. By detailed correlation within the Cretaceous Ferron Sandstone of Utah, they link variations in fluvial architecture directly to changes in base-level and subsequent delta-lobe progradation or retrogression. In the light of this study, we suggest that the coarse interval deposited across the oblique ramp system, and within the piggy-back Tremp Basin, was the result of a lowering in regional base-level, rather than being due of a period of reduced basin subsidence as suggested by Kraus & Middleton (1987). It is, however, acknowledged that the unit was deposited during a period of decreasing subsidence rates between lower and middle Escanilla times. Unfortunately, the lack of resolution in our magnetic chronologies at this time prevents the calculation of sediment accumulation rates for the amalgamated interval itself. However, due to consistent immaturity of pedogenically modified overbank material, we would suggest that the rates of subsidence remained high during the deposition of the coarse amalgamated sheet, and then subsequently decreased within middle Escanilla time. This eventual decrease in rate could well be related to the same base-level fall that generated the regional conglom eratic sheet. 154 Based upon the preferred MPS, middle and upper Escanilla times seem to have been a period of slowed subsidence and sediment accumulation. Calculated rates fall to about 150±30 mm/1000 yrs in the middle Escanilla, and although they are a little more loosely constrained, seem to fall to -70-80 mm/1000 yrs in upper Escanilla time. Sedimentologic evidence, such as paleosol development, preferential channel amalgamation, and increased proportions of coarse gravels all may reflect this slowing of subsidence. These changing rates, when combined with the paleocurrent information collected within the Ainsa Basin give a remarkably coherent description of the basin history during middle to late Eocene times, just prior to the major influx of coarse gravels into the region of the western oblique ramp. The input of these gravels was, in part, due to renewed tectonism within the upper reaches of the Escanilla Formation drainage basin to the east and north in the Pyrenean Axial Zone and in the deforming foreland basin (Burbank et al., 1992). This late Eocene-Oligocene deformation represents a major phase of out-of-sequence thrusting and shortening within the central Pyrenees, resulting in significant renewed relief in the uplifted source regions. Vast amounts of conglomerate were shed down the alluvial systems to the South, and were then diverted into the westerly-flowing axial system of the Campodarbe Group. Correlation of units across the Mediano anticline has highlighted striking differences between the Escanilla exposures within the piggy back basin and those across and outside the western South-Central Unit oblique ramp system, east of the Mediano fold, the Escanilla Limestone 155 represents a protracted lacustrine phase that we believe to be equivalent to shallow marine and deltaic Sobrarbe deposition within the oblique ramp system (Fig. 10). Anticline development caused a periodic ponding of the alluvial system within the piggy-basin, shutting down clastic supply to the delta front. There is no direct evidence to suggest a significant hiatus either above or below the Escanilla limestone, and we interpret it to represent a substantial condensation of section rather than a true omission of time. The succeeding fluvial deposits of the remaining Escanilla Formation continue to show evidence for the episodic ponding of the fluvial system. The general paucity of these limestones, in both frequency and extent across the Mediano anticline to the west, would suggest they were developed predominantly upstream of this episodically active structural trend. As such, the Escanilla Formation alluvial architecture within the piggy-back basin was being controlled not by variations in relative sea-level or strand-line migration (Posamentier, 1991), but by the Mediano anticline periodically acting as a local base- level. The presence of strongly calcareous overbank deposits, a thinner time-equivalent section, and increased volumes and interconnectedness of coarse conglomeratic sheets would all suggest that the region on top of the translating thrust sheet, east of the Mediano anticline, was subsiding less rapidly than either the region of the oblique ramp, or the autochthonous foreland of the Jaca Basin. As such, we can explain the differences in alluvial architecture within time-equivalent regions of the 156 Escanilla system as being closely related to basin structural setting. Local subsidence rates seem to have been the most important control of architectural development at times of stable base-level. During Bartonian time, while wide conglomeratic sheet channel-bodies were being preserved within the Tremp “piggy-back” basin, “fines-dom inated” sections were being preserved across the more-rapidly subsiding oblique ramp system. Similarly 5 myr later in upper Escanilla time (Priabonian), when the oblique ramp experienced reduced rates of subsidence, the essentially autochthonous flexural foreland of the Jaca Basin was subsiding more rapidly, and predictably the spatial differences in the alluvial deposits across this region reflect this westerly-increase in tectonic subsidence. These spatial relationships are summarized in a series of paleogeographic reconstructions of the study area that sequentially describe the regional evolution through late Eocene into Oligocene times (Fig. 13). The late Lutetian reconstruction (Fig. 13A) shows the phase of deltaic progradation along the subsiding Ainsa Basin axis, and the coeval deposition of the Escanilla Limestone, west of the growing Mediano anticline. This may be correlated with deep marine deposition within the Jaca Basin to the west. By early Bartonian times (Fig. 13B), the coastline had migrated across the Boltana anticline. At this time, the fluvial system changed in nature across the axis of the Mediano anticline, and this can be related to the respective structural positions of the two regions. In the Tremp Basin, slower subsidence was expressed by reduced preservation of fine-grained material, and the development of 157 strongly calcareous overbank material. At the same time, the fluvial system within the Ainsa Basin is reflecting increased subsidence rates across the SCU oblique ramp, and a fines-dominated fluvial section was developed. In Priabonian times, when the Boltana anticline was strongly emergent, and the Ainsa Basin was incorporated into the hanging-wall sequence of the SCU, a similar distribution of fluvial style was seen (Fig. 13C). A fines-dominated section was developed downstream within the Jaca Basin, while the Ainsa Basin experienced slowing rates of subsidence and a coarser, more calcareous section was deposited. Thin limestones in the Ainsa Basin at this time may have preserved short-lived ponding of the Escanilla drainage system upstream of the Boltana fold. The final reconstruction (Fig. 13D) shows the post-Escanilla configuration, after further motion of the Boltana anticline, wholesale translation of the Jaca Basin, and growth of the Sierra Exteriores. No major west-directed fluvial system existed at this time. Instead, a series of I ate rally-adjacent large alluvial fans were developed. These appeared to enter the Oligo-Miocene Ebro foreland basin at significant structural re entrants (Hirst, 1985; Hirst & Nichols, 1986) along the Sierras Exteriores and Sierras Marginales thrust fronts. C O N C LU S IO N S Down-stream changes in alluvial architecture within the late Eocene continental depositional system of the southern Pyrenean foreland basin seem to have been closely controlled by spatial variations in thrust 158 C O C O S o O ) *S « o c si I I <Q 2s 5 ■ o o < < n C O c o Q. Figure 13. Four sequential paleogeographic reconstructions of the South-Central and W estern Pyrenees during late Eocene to Oligocene times. Structures active during each interval are shown as solid, whereas inactive structural elements are shown as the lighter dotted pattern. Differences in fluvial style are shown, and can be related to their structural position with respect to active or inactive thrust sheets. Data from Puigdefabregas (1975), Reynolds (1987), Jolley (1988), McElroy I (1990), and Hogan (1992) are incorporated into this summary. 159 deformation along the orogenic belt. Variations in local tectonic subsidence rates, related to differences in structural position within the deforming foreland basin affected fluvial channel morphologies, channel- stacking geometries, the exclusion or preservation of fine-grained material, and the development and extent of pedogenic carbonates. Lower subsidence rates within the allochthonous piggy-back Tremp basin and the autochthonous foreland basin served to (i) increase channel-body interconnectedness of the wide sheet conglomerates, (ii) prevent the preservation of significant volumes of fine vertical accretion material, and (iii) allow the widespread development of pedogenic calcretes and calcareous fines. Across the actively deforming western oblique ramp, within the intervening synclinal Ainsa Basin, sheet and ribbon channel-bodies were excellently preserved, laterally confined by and entirely enclosed within increased proportions of overbank fines. Calcretes were much less pervasively developed. The middle-late Eocene partitioning of the linear foreland into a number of structurally distinct sedimentary basins, bounded by oblique ramp tip-line folds such as the Mediano and Boltana anticlines, also exerted close control on the dispersal patterns and facies development. Upstream of the oblique ramp, within the piggy-back basin, the alluvial deposits were periodically ponded or dammed, allowing the deposition of micritic lacustrine limestones. During the phase of deltaic progradation along the subsiding basin axis, these developing folds served to localize the middle-late Eocene shoreline, as local fold emergence balanced regional flexurally-driven subsidence, preventing 160 the oceanward progradation of the delta-front. A number of these N-S- trending anticlines impeded the westward regression of the alluvial system, producing the strong diachrony in the age of the Belsue deltaic system across the oblique ramp into the Jaca Basin. No systematic relationship between fold position and paleoflow is obvious. Whereas the Boltana anticline strongly modified the paleocurrent pattern throughout Escanilla time, the Mediano fold may not have affected the streams in the same way. Instead, fold development seems to have sporadically ponded the fluvial system within the western Tremp Basin, inducing lacustrine deposition. However, the Boltana anticline was much more strongly emergent at a time when the motion of the Mediano anticline was largely decreasing, and it appears to have controlled local surface gradients, thereby diverting streams southwards into the subsiding Buil syncline. In such a tectonically active foreland setting, however, rapid fluctuations in eustatic sea level seem to have exerted significant influence on large-scale architectural geometries well upstream into the alluvial drainage basin. Eustatic base-level fall induced reworking of portions of the alluvial basin, which delivered increased volumes of coarse alluvial deposits downstream into coastal plain settings. This developed regionally extensive horizons of laterally- and vertically- stacked channelized sands and conglomerates. This study has shown that, using carefully constructed magnetic chronologies and the application of detailed regional correlations, an ancient depositional system can be reliably reconstructed as it crossed a 161 number of different structural regimes. Such high spatial variability of tectonic environments and sedimentary facies is not presently incorporated within the existing two dimensional models addressing foreland basin development (Heller et al., 1988; Flemings & Jordan, 1989; 1990). As soon as the proximal regions of the flexural basin begin to be incorporated into the advancing thrust wedge, current predictive models break down. Application of such models should, therefore, be undertaken with care. One must be able to identify any along-orogen variability of structural development, timing of thrust motion and possible effects of eustatic or local base-level variation before the sedimentary fill of the foreland basin as a whole may be employed as an index of thrust- wedge development. It is very important to view the stratigraphic record within proximal deforming foreland basins in the context of the advancing thrust-wedge as the two become intimately linked as deformation and shortening continue. 162 CHAPTER 4 A Revised Braided-Stream Depositional Model: An Aggrading and Avulsing Low-Sinuosity System 163 ABSTRACT Models of braided stream deposition have largely been developed from studies of regionally degrading alluvial environments. Glacial outwash streams, in particular, have supplied important and widely cited descriptions of intra-channel processes. Such systems have low long term preservation potential and are unlikely to be present in the geologic record in large quantities. Therefore, the study of these modern degradational systems may not provide holistic analogs of the larger- scale alluvial architecture developed in braided river environments in the ancient. The Escanilla Formation of the Spanish Pyrenees provides a well-exposed example of an Eocene fluvial system flowing axially within the Pyrenean foreland basin. Sedimentologic study shows coarse channelized deposits of braided character wholly enclosed within large amounts of fine-grained overbank mudstones and siltstones (> 40% by volume), with both being deposited coevally across the Escanilla floodplain. A new depositional model is proposed that combines facets of existing models derived from other fluvio-morphologic systems. The model consists of a laterally-confined channel belt, internally preserving a braided-stream character, capable of rapid vertical aggradation on short geologic time-scales (~ few kyrs). Avulsion processes are used to explain finer sediment deposition in interfluve settings, as well as the large-scale architecture geometries within the lower Escanilla Formation. 164 This new model illustrates that discrete channel-belt avulsion and the preservation of thick sequences of overbank material are not exclusively characteristics of higher-sinuosity fluvial systems. INTRODUCTION The development of braiding in fluvial systems can be shown to be the result of the complex interaction of a large number of independent variables (Leopold and Wolman, 1957; Miall, 1977). Most importantly these include the amount and variability of stream discharge; the width, depth, and velocity of flow; the volume and grain-size distribution of the sediment load; and the slope and roughness of the stream bed. Extra channel factors including bank material, the presence of bank vegetation, and large variability in discharge, may also help to control channel planform and any resulting depositional geometries. Braiding is the consequence of a stream's episodic inability to move certain clast sizes of its load. Sorting occurs, and deposition of the coarser bed-load results in the initiation of within-channel bar-forms (Leopold and Wolman, 1957). Most studies of lithofacies developed within modern braided rivers have focused upon the channel deposits whereas the associated overbank or vertical accretion deposits have been largely neglected. Although in many widely cited modern examples (e.g. Miall, 1978; Cant and Walker, 1978), such deposits are volumetrically unimportant, a growing number of recent studies of ancient 165 fluvial successions have described significant volum es of fine-grained material preserved in association with coarse channel systems of braided character (Raynolds, 1980; Mohrig, 1986; Desloges and Church, 1987; Jolley, 1989; Reynolds, 1989; Mack and Seager, 1990; Mack and James, in press). The presence of large volumes of fine-grained material within braided stream deposits would be atypical according to most interpretations that rely upon modern systems as analogs (Moody-Stuart, 1966; Cant and Walker, 1976; Miall, 1977). W alker and Cant (1984) went so far as to say that the lack of fine-grained vertical accretion deposits may be a useful criterion when attempting to identify such ancient low-sinuosity systems. Most current braided stream depositional models are derived from modern fluvial systems that, on long timescales (>1000 yrs), are actively degrading (glacial outwash streams in mountainous areas, for example). Although capable of aggradation for shorter periods in response to small changes in discharge or local base- level (Schumm, 1991), the deposits of such systems are unlikely to be preserved on geologic timescales. W hether the deposits and alluvial geometries preserved in the short-term are truly representative of coarse fluvial systems that make up the geologic record will be addressed in the following discussion. Fluvial sedimentologic study of the Late Eocene Escanilla Formation of the central Spanish Pyrenees has highlighted weaknesses inherent within established fluvial low-sinuosity depositional models (Miall, 1978). Excellent exposures present along the western oblique ramp of the 166 Pyrenean South-Central Unit thrust system have allowed the detailed description of the character and architectural geometries of coarse low- sinuosity channels and associated overbank material and have facilitated the construction of vertical and lateral variations of lithofacies (Bentham et al., 1991). The Escanilla Formation represents a particularly well exposed example that appears to contradict braided stream models that are characterized by limited fine-grained sediment deposition (e.g., W alker and Cant, 1984; Rust and Jones, 1987). As demonstrated by Allen (1983) and Miall (1985), high degrees of lateral variability of lithofacies in fluvial environments preclude the wholesale application of vertical facies sequence models. Vertical and lateral facies sequence description within the Escanilla Fm., however, may be used to propose a different depositional model than those currently used to characterize braided stream deposition (Bentham et al., 1991). REGIONAL FRAMEWORK The study area is situated between the Mediano and Boltana anticlines. These represent two major N-S-trending structural boundaries within the E-W-striking Pyrenean foreland basin system of northern Spain (Puigdefabregas, 1975; Camara and Klimowitz, 1986) 167 © 0 0 Q ] 0 ) < 0 < c Figure 1. The Southern Pyrenean Foreland Basin. Box shows the approximate location of the study area, along the western flank of the South-Central Unit thrust system, and the sim plified configuration of the important structural elements discussed in the text. 1 6 8 (Fig. 1). Excellent exposures of the Escanilla Formation fluvial deposits within the small, 20 km-wide, Ainsa Basin (Fig. 2), have allowed the detailed description of sedimentary facies and alluvial architecture within this Eocene drainage system (Bentham et al., in prep). The Escanilla system in the Ainsa Basin is underlain by the shallow marine and deltaic units of the Belsue Fm. (Fig. 3), and unconformably overlain by coarse alluvial conglomerates of Oligocene age (Garrido-Megias, 1973). Constituting the lower part of the Campodarbe Group of Puigdefabregas (1975), the Escanilla sediments represent continental deposition within the south-central Pyrenees. Biostratigraphic and magnetostratigraphic data (Cuevas Gozalo, 1990; Bentham and Burbank, in prep) suggest that the Escanilla Formation ranges from latest Lutetian to early Priabonian in age, and was deposited during a structurally dynamic phase of Pyrenean foreland basin development. At this time, the pre-existing flexural foreland began to detach as a series of laterally distinct piggy-back basins, and proximal parts of the foreland basin were incorporated into the southerly-directed South Pyrenean thrust system. Alluvial architecture within the Escanilla Formation can be seen to vary spatially and temporally, and it is believed that this was largely controlled by lateral variations in tectonic subsidence rates (Puigdefabregas, 1975; Jolley, 1988; Bentham et al., in prep). These variable subsidence rates can, in turn, be directly related to the continuing late Eocene structural development of the south-central Pyrenees. 169 Ainsa • km so o So Arcus a • • Mediano ^ b6 0 \ s . I I Pre-T \ So 18 MED • Almazorre; Pre-T Olson EFU < ■ * 1 1 ~ f? LIG: Ts Oligocene (Collegats) Conglomerates Pre-T ” « ] m c So I □ Pre-fold Ainsa Grp. Eocene Sequence Turbidites Lst. Magnetic Section Escanilla F m . Sobrarbe Fm. Figure 2. Simplified geologic map of the western area, the Ainsa Basin or ‘Buil Syncline’, situated along the western oblique ramp of the South- Central Pyrenean thrust system. The location of villages within the study area and the magnetostratigraphic traverses are shown, as are any important structural features within the basin (ALZ = Almazorre. ERI = Eripol. MED = Mediano. LIG = Liguerre). 170 Oligocene • c C 05 . S c C o O _Q tH 05 05 'C CD Q- c : a * 4— » C D ■ + — ' 3 Ypresian AINSA BASIN Puigdefabregas (1975) RIO CINCA Coilegats Group Escanilla Fm GO CD ■ » — « "g '_ Q w 3 f- Q. 3 O < 5 0 3 CO c < 0 < 0 i CD 3 c r C a ) c a C O C O ■*; C D ® c . C O 0 ) w =- O § o c o Figure 3. The adopted stratigraphy applied during this study. The stratigraphic framework is essentially that of Puigdefabregas (1975), has been modified in the light of recent biostratigraphic and magnetostratigraphic data (Bentham and Burbank, in review, and Cuevas-Gozalo, 1990). ARCHITECTURAL OBSERVATIONS AND FACIES D E S C R IP TIO N S Four vertical stratigraphic sections (Fig. 2) were measured and described in lower Escanilla strata exposed within the synclinal Ainsa Basin. These vertical sequences were combined with two-dimensional facies analysis based on aerial photography, photography of outcrops, and physical tracing of extensive exposures in order to describe the overall lateral variability and geometries within the lower Escanilla alluvial system. In most exposures of the lower Escanilla Fm., the outcrops are dominated by laterally extensive sandstone bodies interleaved with sub equal amounts of fine overbank material (Fig. 4). Vertical sections typically traverse multiple channel systems and show nearly half of the section to be finer in grain size than a fine-to-medium sandstone (Fig. 5a). Channel sand bodies are seldom wider than about 500 m and are typically 5-20 m thick, yielding width:depth ratios in the range of 25- 100:1. As such, they would be generally referred to as “sheet” sandstones (Friend, 1983). These sheets exhibit a complex internal architecture of vertically and laterally stacked lenses of coarse-to-pebbly sandstone and gravel, and in this sense they are very similar to the multi storey, multi-lateral lenticular sheets of Marzo, Nijman, and Puigdefabregas (1988). Although individual channel fills can be placed 172 CO O C / D C / D C / D O o C / D o > ■o 0 5 O ) Figure 4. Sketch of a general view of the Escanilla Formation exposures immediately south of the village of Olson. Note the wide multi-storey channel body wholly enclosed within fine-grained overbank material, and the interfingering geometries between the conglomerates and overbank siltstones along the left margin of the channel system. 173 Y YY Y Y Y Y Y Y Y Y “ Sedimentary Structures Y Y Y Root Mottling & Bioturbation Current Ripples Sheet Splay m iU jT O - TOW vvS. Y Y Y Y YY/ 10 — Minor Channel Splay Horizontal Parallel Lamination Planar Cross-Stratification Trough Cross-Stratification Gravel Scours & Lenses (Massive & Cross-Bedded) Gravel Lateral Accretion Surfaces Scour Surface Major Gravel-Dominated Channel-Fill Sequence with Basal Lateral Accretion Clay Plug ■ o © C O c e o C/5 © © 'o o w C L l_ © © c (X © © E Lithologies _L CO C © = O <5 05^ Q C < 0 5 - * = ■- S O 5 P L a V i Reddened Overbank Siltstones and Mudstones. Fine-Grained Sandstones and Coarse Siltstones Medium to Coarse-grained Sandstones Gravels and Pebbly Sandstones Figure 5a. Generalized vertical lithologic log through the Escanilla Formation fluvial sequence. Interpretations of the various lithofacies are shown at the left side of the stratigraphic column, while grain-size and sedimentary structures are shown to the right. 174 within one or two of Miall’s (1977; 1978) vertical fining-upwards facies sequences resulting from deposition in a “sandy” braided stream system (South Saskatchewan and Donjek facies models) (Fig. 5b), a more complete description of the whole sequence, including intervening overbank sediments, highlights a significant difference. Simply, the Escanilla system preserves far too much fine-grained sediment compared to existing braided stream facies models. Facies Descriptions and Interpretations Four important lithofacies are recognized in the deposits of the lower Escanilla Formation within the Ainsa Basin: (i) gravel-dominated channel-fill sequences; (ii) sand-dominated channel-fill sequences; (iii) sheet sandstone splay deposits; and (iv) pedogenically modified overbank sediment. The minor lithofacies include white micritic limestones, and mature calcic paleosols or pedogenic calcretes. These are not commonly present within the lower Escanilla Fm. and are not discussed in detail below, except to note the implications of their absence or paucity. Gravel-Dominated Channel-Fill Facies: Description This gravel-dominated facies (Lithofacies i) is preserved within both wide (100’s of m) sheet and narrow (10’s of m) ribbon channel bodies 175 X g L L l X g x CL < X g i- < X f — CO Donjek Type 15 — South Saskatchewan Type This Study — _ Y Y Y r Y Y r r r r Y Y Y ] Y r Y > r r r — r r r - y y y | Y Y Y >- i >• »- I .... / Figure 5b. Comparative general vertical lithofacies sequences showing the difference between the ‘sandy’ braided river facies models of Miall (1977, 1978) and a general sequence through the Escanilla deposits. The three columns are drawn at the same vertical scale. The Donjek and South Saskatchewan type examples are redrawn from Miall (1978), using the legend presented in Figure 5a. 176 that erosively overlie reddened, root-mottled siltstones and mudstones (Lithofacies iv). Clasts range in diameter from 2-15 cm, and are typically sub- to well-rounded. They most commonly occur as lenses of clast- supported conglomerates (Fig. 6), or as low- (< 20°) and high-angle (> 30°) cross-sets. Low-angle sets are commonly 0.5-1.0 m high and consist of alternating conglomeratic and coarse pebbly sandstone layers. They are preserved in the lower portions of the channel-fill sequences, immediately above the basal erosion surface, and the foresets usually dip at high angles to basal paleocurrent indicators, such as gutter or groove casts. The more steeply dipping foresets are defined by coarse pebbly sandstones and subordinate gravels, and are preserved within irregular, erosively-based lenses. These are surrounded by similarly sized lenses of coarse planar and trough cross-bedded sandstones which vary in dimensions, but are typically about 0.5 m deep and 4-5 m wide (Fig. 6). Paleocurrent directions derived from these more common foresets have a greater variability, but normally reflect the channel trends, and are in agreement with other paleoflow indicators. The massive, clast-supported conglomerates also occur in similar lensoid geometries, with erosive upper and lower bounding surfaces (Fig. 6). Coarse sandstone lenses become more abundant towards the top of the channel-fill sequence, and a fining-upward trend is often present within them (Figs. 5a & 6). Occasionally these channels preserve fine silt or clay ’plugs' within low, laterally restricted scours, or within minor channels at the top of the channel-fill sequence. In general, 177 Figure 6. Example of the gravel-dominated, channel-fill lithofacies association. Sand/Gravel lenses are present above a strong basal scour surface cutting nto reddened overbank siltstones. Shallow foresets are present within the gravel lens at the iower left corner of the photograph. Channel body is 4 m-thick. 178 however, even though such fine-grained material wholly encloses the channelized bodies, it is rarely preserved within the channel-fill sequences themselves. Interpretation This facies and the overall facies association are the result of major trunk braided-stream deposition. The lower-angle cross-bedded conglomerates most often preserved at the base of the channel-fill sequences are interpreted as coarse gravel transverse or bank-attached bar deposits, formed during the earliest stages of channel history. The dip of the cross-beds, at high angles to the apparent paleoflow direction, suggests they represent coarsely defined lateral accretion surfaces developed during a minor short-lived phase of higher channel sinuosity. The massive conglomeratic lenses represent channel-lag deposits formed by shifting channel bars moving above a basal channel scour surface (Miall, 1977). The lenses of crudely cross-bedded conglomerates and pebbly sandstones oriented parallel to paleoflow direction are interpreted to reflect migrating linguoid bar deposition (Walker and Cant, 1984; Bristow, 1987). These are in turn erosively overlain by the sandier lenses that are preserved within minor channels and scours. The major channels are interpreted to have been cut during the channel flood-stage and then filled during the succeeding phase of waning flow. The overall fining-up trend represents an phase of channel- system aggradation. Some channels show repeated aggradational cycles, rather than a single fill sequence. All channels are wholly 179 enclosed within finer-grained lithofacies. The fine clay plugs are interpreted to represent abandonment of the active channels within the braided stream channel belt and probably correspond with the final stages of an avulsion event, as the stream discharge is wholly diverted into its new course. Suspension deposition and in-filling of existing topography within the old channel belt occurs prior to the establishment of overbank deposition in this abandoned region. Sand-Dominated Channel-Fill Facies: Description This facies (Lithofacies ii) is most commonly preserved within smaller, narrower channel bodies (width-depth ratios range from 15-25:1), and is distinguished by the lack of coarse, conglomeratic material above any basal lag deposits (Fig. 7a). Medium- to coarse-grained sandstones are most common, and are preserved in a number of different geometries. Most frequently the sands are preserved as lenses of trough cross bedded sandstone, typically thinner (<5 m, rather than 5-20 m) and narrower (10’s m, as opposed to 100’s m) than those described in the previous section. This lithofacies association also includes laterally continuous layers of both trough and planar cross-stratified sandstones (Fig. 7a). Troughs vary in size, but tend to decrease upwards from about 2-3 m wide near the base to less than 1 m at the top of the channel-fill deposit. Thicknesses for both the troughs and planar sets are almost always less than 0.5 m. Paleocurrent directions derived from the dip of 1 8 0 Planar Cross-sets Pebbly Sandstone and Conglomerate Trough Cross-sets Figure 7a. Example of the sand-dominated, channel-fill lithofacies association. Sand lenses show complex trough and pianar cross beds developed above a strong basal scour surface cutting into reddened overbank siltstones. Channel body s 3 m-thick. 181 foresets of planar cross-sets, and the plunge directions of trough axes are sub-parallel to the trend of the channels and the paleoflow direction indicated by basal scour features. Upper parts of these sand-dominated channel sequences are typically made up of low-relief, wide trough cross-bedded sands lying above a laterally extensive low-angle scour surface. The medium-coarse sandstones show a high degree of biogenic disturbance. They are pervasively color-mottled, as a result of extensive root and/or burrow invasion, and exhumed upper surfaces often show many cylindrical silt- or clay-filled root-casts weathered out in negative relief. The sand-dominated fill sequences also show a similar aggradational fining-up pattern to their gravel dominated equivalents. Interpretation These narrower sand-bodies are interpreted to represent short-lived splay channels developed during the later stages of a major trunk-stream avulsion event (Stage 1 1 1 splays of Smith et al., 1989) (Figs. 5 and 7b). Developed as part of an evolving avulsion channel system and made up of stable but short-lived channels, these fill sequences aggraded vertically rather than migrated laterally. The laterally extensive, low-relief scour surfaces present near the tops of many channel-fill sequences are interpreted to be “Type 3” surfaces of Miall (1985), and represent low- stage reworking of larger bed-forms or pre-existing channel belt topography. The planar and trough cross-stratified sands within these scours are interpreted to reflect the in-filling of basal channel scours by the migration of transitory dune bedforms (Singh and Bhardwaj, 1991). 1 8 2 Figure 7b. Distal view showing external morphology of a sand- dominated channel fill sequence. Note the strong ribbon geometry, and the lateral wing, traceable into fine overbank facies. This channel represents two stages of aggradation and filling. Channel sequence is 4 m-thick. 183 The planar cross-sets represent the preservation of transverse bar forms or straight crested dunes that were migrating parallel to the channel axis (Cant and Walker, 1976; Singh and Bhardwaj, 1991). Upper surfaces of these channel-fill sequences are often pervasively bioturbated by vertical, tube-like root structures. Purple, yellow and green color mottling is commonly associated with these features, and they are interpreted to represent casts of the roots of ‘reed-like’ grasses growing on the tops of abandoned channels. Sheet-Splay Deposits Description Developed as 10 cm- to meter-scale sheets of homogeneous medium-fine sandstone and siltstone, this lithofacies (Lithofacies iii) shows gradational to planar, horizontal to slightly irregular erosive basal contacts above finer-grained lithofacies (Fig. 8). Most commonly these sheet-splay deposits are seen as laterally extensive sheets that thin and pinch out laterally into the enclosing fines away from adjacent channels (Fig. 8, lower left sheet). Any internal sedimentary structures have generally been destroyed by extensive biogenic activity, although small asymmetric current ripples have been observed within some thin sandy sheets. Some fine sand sheets may also be traced laterally towards the coarse sand- and gravel-dominated channel sequences, and can be shown to be contiguous with sandstone lithofacies (Lithofacies ii) of the channel-fill. The extent of erosion along the basal contact of these 184 1 E > .■ *£ j Figure 3. Example of the sedimentary' geometries present at the lateral margin of a sand-dorrm ated channel fill sequence. Sheet-spiay sandstones may be traced directly into lower channel body (lower left sheets traced to right). Splays commonly show a massive, tabular form with planar bases, while laterally equivalent channels erode down into underlying overbank materials. 185 sheets seems, in part, to be related to their proximity to the adjacent channel, and they are also seen to fine markedly from sandstone to fine siltstone over a few 10's of meters from the major channel. Interpretation These massive sandstone and siltstones are interpreted as unconfined tabular sheets developed during overbank flooding (Mohrig, 1986), or during the early stages of major channel-belt avulsion (Stage I and II splays of Smith et al., 1989). The laterally continuous thin sand and silt lenses resemble the crevasse-splay sandstones of some authors (e.g., Allen, 1965; Kraus, 1987). The lack of sedimentary structures is the result of extensive post-depositional reworking of the depositional surface by flora and infauna (Mohrig, 1986), although rare preservation of current ripples suggests the coarser splays were deposited as shallow, partially-confined flows across limited parts of the Escanilla floodplain. Thicker units (0.2 to 1.0 m) represent multiple overbank depositional events that have subsequently been homogenized after deposition. The sheets contiguous with channel bodies represent break-out crevasse- splay events that fine into and interfinger with overbank sediments. These relationships suggest rapid variations in flow velocity away from the active channel margin onto the surrounding floodplain. Flow was probably baffled or inhibited by the surrounding vegetation cover upon the floodplain and channel margins. The presence of vegetation is supported by the common rootlet development and color mottling at the upper surface of abandoned channel-fill sequences. The siltier units are 1 8 6 interpreted to be evidence of splay abandonment and waning flow across the floodplain, or of deposition in more distal portion of the splay systems, well away from the causative channels (Smith et al., 1989). Pedogenically Modified Overbank Sediment Description Extremely important by volume within the fluvial system (> 40%), this facies association (Lithofacies iv) can be thought to be vertically and laterally gradational with the finer-grained components of the overbank sheet splay events. The dominant lithofacies are reddened, pervasively bioturbated and mottled fine siltstones and silty mud stones. Calcified and sediment-filled root casts have been observed, and these are often combined with the pervasive mottling. Mottling is expressed as purple, green or orange discoloration intermixed within dominantly red and brown silt and mud stones. Bedding is very rarely preserved within these sediments, although a faint color banding is often recognizable across hillslope outcrops. Mature calcic paleosol horizons, although present in this lithofacies association, are generally quite rare. Thin micritic limestones (< 10 cm) are preserved irregularly within the reddened overbank material, but they are volumetrically insignificant and of limited lateral extent. Interpretation Deposition of this facies association is interpreted to occur as largely unconfined sediment-charged flows during the waning stage of a 187 widespread overbank flooding event. Rates of deposition are likely to be extremely rapid during such events. Extensive wetland areas can also be established during avulsion events, and these may greatly contribute to trapping and deposition of large volumes of fine-grained material (Smith et al., 1989). Smith et al. (1989) also stress that fine-grained deposition in interfluve regions can occur during all stages of flow across the floodplain during channel avulsion. The presence of heavily vegetated channel flanks, as suggested by the presence of root-casts and color mottling, would serve to stifle flow away from the main channel belt, inducing the fining trends away from the channels while also favoring fine-grained deposition across wide parts of the floodplain. The lack of mature calcic paleosols within this lower part of the Escanilla formation is taken to indicate relatively limited post-depositional pedogenesis of the fine portion of the section. The presence of such immature paleosols is usually used to infer high rates of subsidence and sediment accumulation within the drainage basin (Kraus and Middleton, 1987). Floodplain regions were generally not subject to long periods (> 10,000 yrs) of subaerial exposure and non-deposition (Retallack, 1986). ADDITIONAL EXAMPLES Plio-Pleistocene of Southern New Mexico A depositional system similar to the Escanilla fluvial system is present in the Plio-Pleistocene strata of the southern Rio Grande rift. The fluvial 188 lithofacies and larger-scale architecture of the Palomas and Camp Rice Formations of southern New Mexico have been described and interpreted by Mack and Seager (1990) and Mack and James (in press). These authors developed a model in which alluvial architecture, the nature of the channel deposits, the ratio of channel to floodplain facies, and calcareous paleosol development were all strongly dependent upon local tectonic environments. Mack and James (in press) suggested that these characteristics of the ancestral Rio Grande fluvial system were chiefly dependent on whether the rivers were flowing in a symmetrically or asymmetrically subsiding extensional graben. The most analogous strata to the Escanilla Fm. studied by Mack and James (in press) are the Plio-Pleistocene strata preserved within the symmetrically-subsiding Hatch-Rincon basin of south-west New Mexico. Approximately 40% of the stratigraphic succession preserved in this basin is made up of overbank sediment finer in grain-size than medium- grained sandstone. In this sequence, channel sand bodies are 100- 1000’s meters wide, and comprise single and multi-storey sheets eroding into the underlying floodplain deposits. They generally range in thickness from 2-18 m. Channel-fill is most commonly medium sandstone, although coarser pebble lenses and laminae are also preserved. Basal scours can also be overlain by up to 50 cm of conglomeratic channel-lag. Internally, although generally finer grained than the Escanilla channels, these fill sequences are essentially very similar. Gravel-filled troughs and scours are succeeded upwards by 189 stacked lenses of trough cross-bedded, planar cross-bedded and planar laminated sandstones. Upper parts of the channels comprise smaller- scale trough and planar cross-bedded sandstones that may display variable paleocurrent directions. Mack and James (in press) also interpret these as the result of low flow stage modification of larger pre existing bed-forms. Intervening floodplain deposits enclosing the channelized sand bodies are dominantly variably bedded mudstones and coarse siltstones to fine sandstones, and thus they appear to be very similar to the Escanilla overbank sediments. Calcic paleosols are generally common, more so than within the Escanilla deposits, and these may be traced laterally for up to a kilometer (Mack and James, in press), suggesting sediment supply across the floodplain was laterally variable. Mack and James (in press) envisaged a largely unconfined channel belt, able to migrate and avulse its way across the Plio-Pleistocene floodplain over time. They invoke traditional overbank processes, previously not discussed in detail within a low-sinuosity fluvial system, to deposit the significant volumes of fine-grained material in conjunction with the coarser channelized material. Plio-Pleistocene of the Eastern Potwar Plateau, Pakistan The Upper Siwalik Group sediments of the Eastern Potwar region of Pakistan were described in detail by Raynolds (1980), and were interpreted as the record of low-sinuosity river systems flowing south from 190 the emerging Himalayan thrust front. As in the Escanilla Formation, coarse conglomerates and sheet sandstones are present, wholly enclosed within the finer floodplain deposits of the Plio-Pleistocene drainage system. Internally, the channelized sheets show abundant local erosion surfaces and the low paleocurrent variance that is considered indicative of braided river deposition. Raynolds (1980) noted the atypical association of coarse channel-fill sequences with much finer- grained adjacent overbank material and posed the question “...How is it possible to interdigitate these extremely high energy facies with the low energy overbank sandy mudstones?..” (Raynolds, 1980, p.218). He envisaged an alluvial fan surface in a humid climate where vegetation would help to stabilize and preserve any overbank sediment next to the active braided stream channels. This idea of bank stability and abundant overbank sediment preservation is not currently a significant component of low-sinuosity fluvial depositional models. Ventura Red Beds, Methow Basin, Washington Described in detail by Mohrig (1986), this Cretaceous fluvial sequence shows an very similar association as that preserved within the Escanilla Formation. Coarse-grained, low sinuosity channels alternate with thick overbank accumulations, and these fine-grained overbank sediments comprise up to -7 0 % of the total vertical lithostratigraphic sequence. The author described erosively-based channel deposits that fine upwards into overbank sequences which are characterized by interbedded sandstones and massive siltstones. Additionally, thin sheets 191 (< 20 cm) of medium to pebbly sandstones are also preserved and these are interpreted to represent single flood events laid down in a network of shallow overbank channels. Mohrig (1986) noted that this overall facies association is unlike most models of coarse-grained, braided river systems, and stresses the importance of the overbank sedimentation. Mohrig (1986) suggested that the overbank deposits were preserved because the Ventura channel belt only occupied a relatively small portion of the total alluvial surface, and because unconfined overbank flooding events frequently inundated large areas of the Ventura floodplain with sediment charged flows. Thick overbank sequences were able to accumulate prior to migration of the active channel belt into any particular part of the floodplain. PROPOSED DEPOSITIONAL MODEL The lower Escanilla Formation represents a coarse braided stream system dominated by the preservation of large amounts (> 40% by volume) of fine-grained siliciclastic sediment (Figs 4 and 5a). This finer material can be shown to have been sourced from the coarser channel- fill sequences, implying a strongly mixed-load depositional system (Schumm, 1977). Although in basic description, the lower Escanilla Formation may superficially appear similar to an 'anastomosing' fluvial system, we feel that the characteristic coarseness and sheet geometry of the channel-bodies, in combination with their complex internal architecture are all consistent with a braided stream interpretation. 192 Anastomosed systems seldom preserve such coarse bed-load material, and typically show channel-bodies with narrow, ribbon geometries (Smith 1976; 1983; Smith and Smith, 1980; Schuster and Steidtmann, 1987). The depositional model of the Escanilla fluvial system is best represented by a wide channel-belt, internally showing a braided morphology of minor channels, wholly enclosed within finer-grained overbank sediments (Fig. 9). Within the surrounding interfluve regions, aggradation is accomplished episodically, in response to unconfined overbank sheet flows (Mohrig, 1986) and channelized splay deposition (Smith et al., 1989). This aggradational style implies significant lateral temporal heterogeneity of sediments across wide floodplains. The position of the active channel belt will vary through time, as the channel belt aggrades and then avulses (Fig. 9). Rates of vertical accretion and avulsion, based on the lack of mature paleosols, the low vertical and lateral interconnectedness of the trunk stream channel belts, and the large number of major channels preserved within the section, were high throughout lower Escanilla time. This is supported by the calculation of effective (undecompacted) sediment accumulation rates across the study area using magnetostratigraphy (Bentham et al., in prep). These rates are typically -25-35 cm/1000 yrs and are comparable with rates derived from modern fluvial settings showing similar gross architectural relationships (Leeder, 1978). 193 0 5 “ 8 « < s ^= < 4 - f c © < D If * 05 C * * © C ?*> © Q C C Figure 9. Generalized block diagram showing the proposed depositional model. Note the low-sinuosity channel belt, internally braided in character, entirely enclosed by vegetated flanks and floodplain environments. Architectural geometries are shown in the vertical views. The presence of the adjacent anticline is specific to the Ainsa Basin (the Boltana Anticline), and is not an integral part of the general depositional model. Rapid rates of subsidence are considered to be most important. 194 Stratigraphically analogous with the present-day Brahmaputra system, but on a very different scale, the complex internal arrangement of sand and gravel lenses within the Escanilla channel sequences resulted from the rapid switching and avulsion of smaller channels within the major trunk channel belt and by the lateral or downstream migration of low relief bar-forms (Bristow, 1987). Small basal lateral accretion surfaces may be present, suggesting high-sinuosity morphologies immediately after new channel initiation. These are overlain by coarser- grained lenses and scours that represent rather more energetic, less- sinuous minor channel activity. These, in turn, are followed by fine-scale planar, and trough cross-bedded sandstones deposited during low-stage flow. Channel-fill sequences terminate upward either with wholesale abandonment and deposition of fine silt and mud plugs within small- scale channels, or they show low-stage reworking of larger bed-forms producing the small complex stratification above a laterally extensive irregular scour surface (Miall, 1985). D IS C U S S IO N Intra-channel processes active during Escanilla Formation deposition appear to have been comparable with those described by many other authors from modern ‘sandy’ braided stream systems (Moody-Stuart, 1966; Miall, 1978; Cant and Walker, 1978; Desloges and Church, 1987). However, the application of such models to the Escanilla system 195 begins to break down when one considers the relationship between the channel-fill sequences and their enclosing overbank material. The interfingering of channel deposits and the fine overbank is clearly seen along the channel margins where sheet splays can be traced laterally into the interfluve regions (Figs. 7b and 8). Fine-grained deposition within floodplain settings was, therefore, occurring at the same time that much coarser-grained material was being moved within the major trunk stream. The deposition of finer, more cohesive sediments, primarily along the flanks of active channels helped to confine the system laterally, as the erodibility of the bank material was decreased. Additionally, the presence of vegetation on these flanks would further prevent lateral channel migration by bank erosion (Smith, 1976). Interestingly, two 'essential' conditions are required if braiding is to develop within a natural channel reach, namely, sediment transport and 'low threshold' to bank erosion (Leopold, Wolman, and Miller, 1964). If cohesive, non- erodible banks serve to confine the channel width, the capacity of a channel reach to transport bed-load material will be enhanced, and deposition and subsequent intra-channel bar development will be inhibited. The Escanilla Formation and the other examples cited previously, therefore, represent peculiar associations that have not been widely identified by fluvial geomorphologists studying modern systems. The study of Smith et al. (1989) offers an possible explanation of the apparent partitioning and subsequent preservation within the Escanilla river system of fine-grained material out of the braided channels into the 196 interfluve settings. The short-lived splay systems of the South Saskatchewan River rapidly deposit large volumes of fine-grained material across the floodplain (Smith et al., 1989). The longer-lived splay channels showed different hydro-dynamic conditions for much of their history, and only the most stable, long-lived splay channels would evolve to form the new trunk stream after a discrete avulsion event. The Escanilla splay lithofacies associations ((ii), (iii) and (iv)), represent a large volume of the total fluvial sequence preserved, and are interpreted to have been rapidly deposited during episodic channelized avulsion (Smith et al., 1989) or overbank flooding (Mohrig, 1986). Both mechanisms seems to have been important, and one need not preclude the other. The lack of mature paleosol development suggests that significant breaks in deposition did not exist for very long periods (probably less than 10,000 yrs, Retallack, 1986). In overall geometry, the Lower Escanilla fluvial system seems to have produced deposits similar to those described by Desloges and Church (1987) in the Bella Coola River of British Columbia. Their "wandering gravel-bed" river produces similar channel zones within finer associated overbank deposits. In both cases, the floodplain appears to have accreted vertically, at the same time that the channels were actively transporting much coarser-grained material. However, the overbank deposits of the Escanilla system are significantly finer-grained than those described by Desloges and Church. While gravel-sized sediment was 197 being moved within the Escanilla channels, fine silt and muds were being deposited across the proximal floodplain. Low degrees of channel belt interconnectedness (Fig. 9) would suggest relatively high rates of subsidence and sediment accumulation (after Bridge and Leeder, 1979). The observed rates of -25-35 cm/1000 yrs are of the same order as average un-compacted vertical accretion rates derived from the studies of floodplain deposits in a number of modern river systems (Leeder, 1978). Schuster and Steidtmann (1987) also reported low-sinuosity channels enclosed within finer overbank material, and they suggest the channels were fixed laterally, unable to migrate across the floodplain because of the rapid local rates of subsidence and sediment accumulation within their subsiding foreland basin. On the basis of comparison of the alluvial architecture from two related sequences formed in differentially subsiding regimes, Kraus and Middleton (1987) concluded that more rapidly subsiding floodplains tend to prevent the lateral migration of active channels, inhibiting the reworking of finer alluvium further down the fluvial network. This is because the channel belt and adjacent area must aggrade vertically in order to maintain grade as the region rapidly subsides. It is worth pointing out that rates of subsidence and vertical aggradation are still much slower that the potential rates at which a braided river system may laterally expand. This still, therefore, requires that the Escanilla rivers were laterally confined and unable to significantly modify their widths during bankfull discharge. 198 This significant departure from “normal” or classic braided stream depositional models may in part be a function of our currently available, and possibly biased, modern analogs. With the notable exception of recent studies from the Brahmaputra River (Bristow, 1987), braided stream models have been based largely on the study of modern glacial out-wash streams or low-sinuosity rivers flowing in regionally degrading alluvial environments, where erosion and transport are the dominant processes acting over geologic time. Due to the availability of exposures, and accessibility (many are drawn from examples on the North American continent), it appears that the braided stream data set has become unintentionally biased by examples from these degradational settings. Overbank material may be deposited within such a system for short time-scales, but slow aggradation rates and rapid channel migration promote the reworking and transportation of this material further downstream. It will seldom be preserved in association with coarser channel deposits (Cant and Walker, 1976). Relatively rapid sedimentation and system aggradation within lower Escanilla time, combined with the presence of cohesive bank material, allowed the preservation of these overbank lithofacies, immediately adjacent to coarse, braided stream channel bodies. This preferential preservation is highlighted when a vertical facies sequence from the Escanilla deposits is compared to existing published facies models derived from streams flowing in such regionally degrading environments (Fig. 5b). 199 C O N C L U S IO N S Fluvial sedimentologic study of the Escanilla Formation highlights limitations of existing braided stream depositional models. In rapidly subsiding sedimentary basins, streams of braided character can produce deposits that have many characteristics of higher-sinuosity river deposition, and only when the channel-fill sequences and their associated overbank materials are related to each other temporally, does this become evident. A modified model for braided stream deposition is proposed, where a channel-belt of internally braided character aggrades vertically rather than expanding laterally overtim e. Episodic avulsion of this channel system causes floodplain aggradation and the periodic switching in trunk stream position. Modern braided streams flowing in presently degrading regions, therefore, while helping us to monitor and understand intra-channel processes, such as bar formation and subsequent bar and channel evolution, are not useful as holistic analogs for the aggradational fluvial systems that are preserved within the geologic records of active tectonic regions. We suggest that future studies of such systems be focused upon modern or recent fluvial systems deposited or flowing within young sedimentary basins. They may supply our best devices for analyzing ancient coarse fluvial systems, and our best chance of deriving general predictive models of larger scale alluvial architecture. 2 0 0 CHAPTER 5 Dissertation Summary and General Conclusions 201 While discussions and summaries of the important aspects of this research are presented within chapters 2, 3, and 4, one single synthesis of all this information is considered to be of use. In the following chapter, I offer a re-iteration of these preceding conclusions, as well as providing a somewhat broader discussion of the notable parts of this study. Detailed lithologic and magnetostratigraphic correlation across the western oblique ramp of the SCU thrust system has, for the first time allowed the following: (i). Constraint of the middle to late Eocene tectonostratigraphic development of this structurally dynamic region in an absolute chronologic framework; (ii). Sequential backstripping of Eocene sedimentary systems, and the calculation of rates of sediment accumulation and tectonic subsidence; (iii). Evaluation of the application of sequence stratigraphic analysis in this technically active environment, where a detailed reliable temporal framework did not previously exist; and (iv). Revision of pre-existing paleogeographic reconstructions for the western SCU oblique ramp during middle-late Eocene time. Furthermore, magnetostratigraphic dating from the Tremp and Ainsa basins allows us to make a number of new, significant conclusions concerning the absolute temporal development of the study area: (i). The Mediano anticline developed during a rapid phase of growth and rotation from -48-43 Ma. This served to induce marked differential 2 0 2 subsidence across the study area, and was largely responsible for the rapid, lateral facies variations observed across the western margin of the SCU at this time. Continental and shallow marine sedimentation was generally confined within the Tremp Basin, and this changed laterally into shelf, and deep marine deposition to the west; (ii). An abrupt marine transgression occurred at about 43 Ma (late Lutetian, rather than Bartonian) immediately east of the Mediano, in the uplifting hanging-wall of the SCU thrust system. This was probably the result of local subsidence rather than a eustatic rise in sea-level. This occurred synchronous with deltaic sedimentation within the Ainsa Basin; (iii). The transition from marine to continental sedimentary environments across the oblique ramp occurred at -42.9 Ma, with the deposition of the Escanilla fluvial system on top of shallow marine and coastal sediments of the Puy de Cinca, Sobrarbe, and Capella formations. Eocene fluvial sedimentation persisted until the onset of uplift and deformation of the Sierras Exteriores soon after 36 Ma; (iv). Marked lateral changes in sediment accumulation and subsidence can be documented within the Escanilla deposits of the Ainsa Basin after the application of litho- and magneto-stratigraphic correlation. In particular we can discern the shifting of the basin depocenter in response to continued development of the Boltana anticline. After detailed sedimentologic analysis, combined with regional magnetostratigraphic correlation across adjacent sedimentary basins, I have been able to begin to reconstruct a major southern Pyrenean fluvial 203 drainage system. Documenting the along-system changes in facies occurrence and architecture, leads us to conclude that the upper Eocene Escanilla depositional system was closely controlled by spatial variations in thrust deformation along the orogenic belt. Lateral changes in local tectonic subsidence rates, related to differences in structural position within the deforming foreland basin strongly influenced: the internal channel morphologies and channel-stacking patterns, the relative proportion of coarse channel deposits to fine-grained overbank material, as well the development and extent of pedogenic and lacustrine carbonates. When a particular depositional setting, such as the Ainsa basin, was experiencing lowered subsidence rates within an allochthonous piggy back basin, this served to: (i). Increase channel-body interconnectedness of the wide sheet braided-stream conglomerates; (ii). Inhibit the preservation of large volumes of fine-grained overbank siltstones and mudstones; and (iii). Allow the widespread development of pedogenic calcretes and calcareous overbank fines between coarse channel sequences. Across the actively deforming western SCU oblique ramp in Bartonian time, or within the autochthonous Jaca Basin during the Priabonian stage, sheet and ribbon channel-bodies were more extensively preserved, laterally confined by and entirely enclosed within increased proportions of overbank fines. Calcretes were much less pervasively developed within the more rapidly subsiding regions. 204 The middle-late Eocene partitioning of the linear foreland into a number of structurally distinct sedimentary basins, bounded by oblique ramp tip- line folds such as the Mediano and Boltana anticlines, also exerted close control on the dispersal patterns and facies development. Upstream of the oblique ramp, within the piggy-back basin, the alluvial deposits were periodically ponded or dammed, allowing the deposition of micritic lacustrine limestones. During the upper Eocene regressive phase along the deforming foreland basin axis, developing folds served to localize the upper Eocene shoreline, as fold emergence was able to balance regional flexurally-driven subsidence. In a similar way, a series of N-S-trending anticlines present along the southern flank of the Jaca Basin impeded the westward regression of alluvial facies, producing the strong diachrony in the age of the Belsue-Atares deltaic system across the oblique ramp to the west. In such a tectonically active foreland setting, however, fluctuations in eustatic sea level seem to have also exerted significant influence on large- scale architectural geometries well upstream into the alluvial drainage basin. Eustatic base-level fall induced reworking of portions of the alluvial basin, which delivered increased volumes of coarse alluvial deposits downstream into coastal plain settings. This led to the formation of regionally extensive horizons of amalgamated channelized sands and conglomerates preserved across the Ainsa and Tremp basins. Such horizons have proved useful for laterally correlating regions of Escanilla exposures within the adjacent basins. 205 Interestingly, the fines-dominated regions of Escanilla fluvial sediments are considered to be atypical, in the light of the published fluvial facies models addressing ‘sandy’ braided stream river deposition. As such we see the preservation of large volumes of immature, pedogenically-modified overbank material in direct association with coarse channel-bodies of an internally braided character. A modified model for such braided stream deposition is proposed, where a braided stream channel-belt aggrades vertically rather than expanding laterally over time. Episodic avulsion of this channel system causes floodplain aggradation and the periodic switching in major trunk stream position. The mismatch of the Escanilla deposits with existing facies models possibly highlights a fundamental weakness in the way these reference sequences have been established. Largely based upon the analysis of modern braided streams flowing in presently degrading regions, such studies offer an excellent chance to monitor and understand intra-channel processes, such as bar formation and subsequent bar and channel evolution. However, I believe they are not useful as holistic analogs of the aggradational fluvial systems that are preserved within the geologic records of active tectonic regions. Such modern systems, while preserving sedimentary structures and facies sequences on short (~100’s- 1000’s yr) time-scales, have little or no preservation potential on realistic geologic time-scales (~myr-100’s myr). The long-term aggradation of a fluvial system within a subsiding sedimentary basin offers a fundamentally different control on the river system than short-term system aggradation in response to local variations in run-off, sediment supply, regional or local 206 base-level variation within a degrading alluvial basin. To use the structures and sequences preserved in the latter to interpret and understand the former seems to be rather misguided. Therefore, I suggest that future studies of such potential analogs be focused on modern or recent fluvial systems deposited or flowing within young sedimentary basins. They may supply our best devices for analyzing ancient coarse fluvial systems, and our best chance of deriving general predictive models of larger scale alluvial architecture. This study has shown that, using carefully constructed magnetic chronologies and the application of detailed regional correlations, an ancient depositional system can be reliably reconstructed as it crossed a number of different structural regimes. Such high spatial variability of tectonic environments and sedimentary facies is not presently incorporated within the existing two dimensional models addressing foreland basin development (Heller e ta l. 1988; Flemings & Jordan 1989, 1990). As soon as the hinterlandward regions of the preceding flexural basin begin to be incorporated into the advancing thrust wedge, as a series of piggy-back basins, current predictive models break down as this region becomes part of the tectonic load. No further consideration of deposition within these regions is included. Application of such models within most foreland basin settings should, therefore, be undertaken with care. One must be able to identify any along-orogen variability of structural development, timing of thrust motion, and related formation of piggy-back sedimentary basins. Only then can the sedimentary fill of the foreland region as a whole may be employed as an index of thrust-wedge 207 development. 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(1984) Sandy Fluvial Systems, in Walker, R.G., ed., Facies Models 2nd Edition, Geoscience Canada Reprint Series, 1, 71-90. Williams, G.D. (1985) Thrust tectonics in the south central Pyrenees. Journal of Structural Geology, 8, 11-17. Williams, G.D., and Fischer, M.W. (1984) A balanced section across the Pyrenean orogenic belt. Tectonics, 3, 773-780. 220 Appendix 1: APPENDICES Tables of Magnetostratigraphic Data 221 Mediano Magnetostratigraphic Section AZ1 ,DP1 Azimuth and Dip of Sampled Surface AZ2.DP2 Azimuth and Dip of Local Bedding Eff. Ht Effective Stratigraphie Height in Composite Section Field Information 1 1 Average Vector & Statistical Data v q p Lat Specimen Ht tm) £21 DPI £22 DP2 Irealm fiol Level Inc Dec K ifin a A £5 Class 1 Class II Error MED01A,B,C,D 2 254 25 254 24 MED81 A,B,C,D 3 252 27 254 24 Ther 320 °C 58.0 17.0 16 32.1 84 17 MED82A,B,C,D 16 253 24 263 27 MED02A,B,C,D 47 224 11 263 27 AF 200 Oe 71.4 -49.3 30 23.0 46 28 MED83A,B,C,D 82 210 7 240 27 AF 200 Oe 41.1 22.1 36 20.8 69 14 MED04A,B,C,D 87 222 16 240 27 MED03A,B,C,D 101 230 24 253 25 AF 200 Oe 63.7 -74.7 88 132 27 22 MED05A,B,C,D 111 250 28 250 30 MED84A,B,C,D 116 192 19 250 30 AF 200 Oe 46.6 9 2 9 94.9 71 62 MED85A,B,C,D 151 262 24 240 39 MED06A,B,C,D 157 224 23 240 30 AF 200 Oe -10.5 -145.2 19 29.1 -49 34 MED07A,B,C,D 204 240 23 237 26 MEDOSA,B,C,D 227 172 17 245 27 AF 200 Oe -10.8 -141.5 37 42.5 -48 50 MED09A,B,C,D 272 226 22 248 25 MED86A,B 280 245 22 248 25 AF 200 Oe -38.1 -130.6 62 15.8 -56 15 MED86C.D 280 265 25 248 25 MED10A,B,C,D 300 250 23 280 24 MED11A,B,C,D 340 252 19 263 21 MED12A,B,C,D 371 218 10 256 27 MED87A,B 391 352 2 252 24 MED87C.D 391 145 36 252 24 MED13A,B 409 105 40 252 24 AF 200 Oe -57.7 -104.0 1485 6.5 -48 8 MED13C.D 409 95 45 252 24 MED88A,B,C 420 264 21 252 24 AF 200 Oe 52.9 0.6 1 1 39.2 71 26 MED88D.E 420 235 27 252 24 ME014A,B,C,D 425 356 35 234 32 AF 200 Oe 12.1 -112.6 139 10.6 •24 19 MED89A.B 439 5 14 234 32 MED89C.D 439 272 15 234 32 MED15A,B,C 465 295 12 215 35 MED15D.E 465 238 19 215 35 MED90A,B,C 469 250 10 225 35 MEO90D 469 240 23 225 35 MED16A,B,C,D 500 237 17 237 24 MED91 A,B,C,D 506 243 24 237 24 AF 200 Oe 45.7 -13.6 7 49.2 58 45 MED17A,B,C 531 300 8 245 24 AF 200 Oe 62.7 -11.0 18 29.1 68 21 MED17D.E 531 300 4 245 24 MED92A,B,C,D 537 240 20 245 24 AF 200 Oe 56.9 32.3 15 32.6 78 18 MED93A,B,C,D 552 242 21 242 21 AF 200 Oe 44.6 5.6 23 26.1 68 18 MED18A,B,C,D 571 226 21 238 18 MED19A,B,C,D 598 200 66 255 18 AF 200 Oe 64.7 -19.7 610 5.0 62 4 MED20A,B,C,D 623 243 15 255 19 AF 200 Oe 59.8 -32.8 27 50.5 52 54 MED21A,B,C,D 646 20 20 224 19 AF 200 Oe 45.5 48.1 204 8.7 61 7 MED22A,B,C 666 285 41 230 23 MED22D.E.F 666 215 18 230 23 MED23A,B,C,D 695 224 17 235 20 AF 200 Oe 60.3 6.1 2863 2.3 79 1 MED24A,B,C,D 717 226 14 250 19 AF 200 Oe 71.9 53.6 172 9.4 44 12 ME025A.B 742 335 20 235 20 MED25C.D 742 340 13 235 20 MED26A.B 772 217 20 250 14 AF 200 Oe 503 43.3 62 15.7 68 11 MED26C.D 772 245 17 250 14 MED27A,B 810 242 10 240 15 AF 200 Oe 39.9 35.6 70 14.8 65 12 MED27C,D,E 810 232 11 240 15 MED28A 839 254 5 253 15 AF 200 Oe 31.8 24.1 38 20.1 62 17 MED28B,C,D 839 294 23 253 15 MED29A,B,C,D 868 274 20 244 23 AF 200 Oe 564 28.8 57 16.5 80 9 MED30A,B,C,D 903 112 12 235 18 AF 200 Oe 55.8 11.6 382 6.3 80 3 MED31 A,B,C,D 929 310 23 225 24 AF 200 Oe 583 69.0 10 40.2 54 41 MED32A,B 968 150 41 245 23 AF 200 Oe 66.9 2.9 52 17.3 77 10 ME032C.D Q £ fl 9 0 0 133 45 245 23 MED60A,B,C,D 973 122 21 243 18 AF 200 Oe 58.2 40.2 31 22.6 74 14 MED61 A,B,C,D 983 83 15 232 15 AF 200 Oe 60.8 45.7 38 20.4 71 13 MED33A,B,C,D 997 98 56 232 15 AF 200 Oe 17.2 66.8 228 16.6 36 25 A. Mediano Average Site Vector Data 222 ME062A,B,C,D 1009 229 12 232 15 AF MED34A.B.C 1024 220 12 253 12 MED34D.E 1024 119 16 253 12 MED63A.B.C 1034 83 19 253 12 AF MED63D 1034 164 19 253 12 MED35A,B,C,D 1078 39 14 230 22 MED64A,B,C,D 1081 212 13 235 20 AF MED36A,B,C,D 1099 180 8 241 14 AF MED65A,B,C,D 1104 223 17 241 14 MED37A.B 1134 249 8 242 14 Ther MED37C 1134 165 5 242 14 MED37D 1134 10 8 242 14 MED66A.B 1161 104 40 236 7 MED66C 1161 158 36 236 7 MED660 1161 98 13 236 7 MED38A,B,C,D 1165 327 41 236 7 Ther MED67A.B 1178 227 8 236 7 Ther MED67C.D.E 1178 220 11 236 7 MED39A,B,C,D 1194 227 12 227 12 Ther MED41A.B 1225 252 12 231 11 Ther MED41C.D 1225 223 9 231 1 1 ME040A,B,C,D 1234 82 13 243 11 Ther MED42A,B,C,D 1241 280 24 231 11 Ther MED68A,B 1245 93 27 246 10 ME068C.D 1245 76 48 246 10 MED69A,B,C,D 1256 69 34 212 9 Ther MED43A.B 1275 315 1 212 9 ME043C.D 1275 110 19 212 9 MEO70A.B 1290 203 8 199 12 Ther ME070C,0,E 1290 184 6 199 12 ME044A.B 1293 194 6 199 12 Ther MED44C.D.E 1293 180 6 199 12 MED45A,B,C,D 1298 45 33 224 10 MED71 A,B,C,D 1299 59 16 224 10 ME046A.B 1321 22 65 211 9 Ther MED46C 1323 69 12 211 9 Ther MED46D 1323 12 3 211 9 ME072A,B,C,D 1326 65 7 211 9 MED47A,B,C,D 1357 224 10 224 10 Ther MED48A,B,C,D 1375 119 15 234 9 Ther MED49A.B 1410 127 11 230 7 Ther MED49C.D 1410 125 10 230 7 MED50A,B,C,D 1468 195 8 195 8 MED 73 A 1474 280 23 195 8 ME073B,C,D 1474 300 12 195 8 ME074A,B,C,D 1481 198 7 186 7 Ther MED51 A,B,C,D 1483 42 11 186 6 MED75A,B,C,D 1515 124 26 186 6 Ther MED52A,B,C,D 1516 200 1 186 6 MED76A,B,C,D 1549 216 38 175 6 MED53 A,B,C, D 1551 254 30 175 6 MED77A,B,C,D 1561 75 32 175 6 Ther MED54A,B,C,D 1582 260 11 143 8 Ther MED54E.F 1582 288 13 143 8 MED78 A,B,C, D 1605 74 51 130 5 Ther MED55A.B 1618 251 6 130 5 Ther MED55C.D 1618 120 10 130 5 MED79A,B,C,D 1643 306 18 135 7 MED56A.B 1648 115 6 135 7 MED56C.D 1648 80 7 135 7 MED80A,B,C,D 1659 274 44 140 6 Ther MED94A.B 1680 250 42 140 6 Ther MED94C.D.E 1680 50 18 140 6 MED95A,B,C,D,E 1690 78 6 140 6 Ther 200 Oe 14.3 30.0 18 29.8 52 32 200 Oe 46.9 32.1 10 42.0 71 28 200 Oe -13.7 -130.7 129 10.9 -44 14 200 Oe 59.7 63.8 713 9.4 58 9 320 °C -44.2 -173.1 238 8.0 -68 6 320 °C -25.5 -152.0 214 17.1 -58 16 280 °C -17.1 -164.9 1405 3.3 -54 3 280 °C -18.2 -151.7 1 721 3.0 -54 3 280 “C 7.8 3.0 176 18.9 47 23 280 “ C -7.1 -146.1 10 40.4 -47 48 280 °C 44.0 -13.1 5 63.2 57 59 320 “C 63.1 23.9 27 24.2 87 12 280 “C 22.3 54.1 11 40.0 46 49 320 °C 48.9 49.6 365 6.5 62 5 280 “C 51.8 13.2 90 26.7 77 15 280 °C -57.2 -148.2 50533 1.1 -79 1 320 “C -36.6 -142.0 116 11.5 -61 10 320 “C -6.4 141.0 603 10.2 -24 18 280 “C -64.3 -151.6 58 16.3 -84 8 280 °C -8.3 -143.0 15 33.4 -47 40 280 “C -3.1 159.2 81 13.8 -34 21 320 “C 31.1 -142.2 14 34.2 -26 58 280 “C 47.8 29.5 26 24.8 72 16 280 °C 42.7 29.1 69 35.1 69 15 280 °C 45.2 17.8 46 18.4 72 12 280 °C 3 38 12.1 14 72.7 63 59 320 “C 53.7 -1.0 24 26.0 71 17 320 °C 56.0 59.1 7 51.5 60 46 Mediano Data Continued Total Number of Sites 92 % of Class I Data Sites 46 % of Class II Data Sites 17 223 Almazorre Magnetostratigraphic Section AZ1 ,DP1 Azimuth and Dip of Sampled Surface (°) AZ2.DP2 Azimuth and Dip of Local Bedding (°) Field Information ~] | Average Vector & Statistical Data v g p Lat. Specimen * Ht (mi AZ1 EE1 AZ2 DP2 Treatment Level los Dec Kappa A-95 Class 1 Class II Error ALZ01A,B 25 348 21 114 9 Ther 320 “ C -25.4 -140.0 125 11.1 •54 11 ALZ01C.D.E 25 77 78 114 9 ALZ02A,B,C,D,E 33 249 45 114 9 ALZ03A.B 50 279 1 120 11 Ther 280 “C 5.2 -60.2 282 7,4 9 14 ALZ03C,D,E 50 283 26 120 11 ALZ04A.B 65 30 13 120 11 Ther 320 °C -42.4 -175.9 5 59.0 -66 44 ALZ04C.D 65 78 15 120 11 ALZ04E 65 60 27 120 11 ALZ05A,B,C,D,E 79 319 44 120 11 Ther 280 °C -16.4 -166.5 110 24.1 -53 25 ALZ06A,B,C,D,E 1 1 1 238 24 120 10 ALZ07A.B 138 220 28 140 10 Ther 320 °C 51.0 -10.9 13 74.3 63 60 ALZ07C.D 138 195 22 140 10 ALZ08A,B,C 162.5 146 45 140 10 ALZ08D.E 162.5 239 36 140 10 ALZ09A,B,C,D,E 183 45 6 140 10 Ther 320 *C -7.2 -138.0 18 29.8 -45 38 ALZ10A,B,C,D,E 187 290 54 140 10 Ther 280 °C -13.5 -105.1 27 24.1 -29 39 ALZ11 A.B 199 305 37 140 10 ALZ11C,D,E 199 210 20 140 10 ALZ12A.B 217 217 14 140 10 Ther 320 'C 62.1 17.91 10 40.9 88 21 ALZ12C.D 217 185 10 140 10 ALZ13A.B.C 222 194 19 140 10 Ther 320 *C 65.5 25.4 58 16.3 85 8 ALZ13D.E 222 181 21 140 10 ALZ14A,B,C 226 213 33 125 9 Ther 280 “C 28.3 178.3 16 31.7 -27 54 ALZ14D,E 226 219 36 125 9 ALZ15A,B 270 151 17 125 9 Ther 280 *C 28.3 16.3 5 66.1 60 58 ALZ15C,D,E 270 192 28 125 9 ALZ16A,B,C,D 282 32 55 125 9 Ther 280 “C 44.2 25.5 9 43.7 71 29 ALZ17A,B,C,D 290 268 16 130 8 ALZ18A.B.C 300 88 31 130 8 Ther 280 “C 42.4 18.5 525 5.4 70 4 ALZ18D.E 300 66 30 130 8 Total Number of Sites 18 % of Class I Data 39 % of Class II Data 33 B. Almazorre Average Site Vector Data 224 Eripol Magnetostratigraphic Section AZ1 ,DP1 Azimuth and Dip of Sampled Surface (°) AZ2.DP2 Azimuth and Dip of Local Bedding (°) Field Information 11Average Vector & Statistical Data v q p Lat. Specimen * Htfm) M l DP1 M 2 DP2 Treatment Level Inc I2SS Kappa A-95 Class 1 Class II Error ERI40A.B 3 337 19 115 10 EFU40C,D 3 20 12 115 10 ERI01 A,B,C,0 12 120 24 113 24 Ther 280 °C 59.4 -164.0 58 33.4 -5 66 ERI02A,B,C,D 28 135 20 113 24 Ther 280 "C 48.4 -164.0 76 29.0 -16 55 ERJ41A.B 48 18 6 120 24 Ther 280 "C 66.3 52.1 529 10.9 68 8 ERI41C.D 48 218 23 120 24 ERJ03A,B,C 53 205 15 120 24 Ther 280 "C 23.1 63.7 24 53.3 40 73 ERJ03D 53 256 26 120 24 EFU42A,B,C,0 69 117 6 120 24 Ther 280 °C 16.5 172.9 11 40.0 -36 60 ERJ04A,B,C 77 328 34 120 24 Ther 280 °C 62.1 -29.7 148 20.6 55 21 ERI04D.E 77 130 2 120 24 ERI43A.B 99.5 179 38 120 24 EFU43C 99.5 291 6 120 24 ERI43D 99.5 240 27 120 24 ERI05A,B 102 355 17 120 24 ERI05C.D 102 80 14 120 24 ERI44A,B 118 10 21 120 24 ERI44C,D,E 118 340 8 120 24 ERI06A 123 111 25 160 16 ER106B.C 123 191 35 160 16 ERI06D.E 123 146 24 160 16 ER145A.B 138 60 6 160 16 ERI45C,D 138 334 20 160 16 ER107A, B,C,D 149 323 31 160 16 ERI46A,B,C,D 165 53 43 140 12 Ther 280 °C -10.0 167.7 17 31.1 -41 42 ERIOBA.B 167 35 30 160 12 Ther 280 °C 5.2 -155.7 26 24.6 -43 32 ERIOBC.D 167 16 58 160 12 ERI09A,B,C,D 188 217 17 150 9 Ther 280 °C 49.4 -76.4 11 38.6 17 72 ERI47A.B 215 224 6 155 10 Ther 280 °C 51.7 12.3 49 36.5 76 21 ERI47C.D 215 156 7 155 10 ERI10A,B,C 216 154 24 155 10 ER110D 216 64 6 155 10 ERI11A,B 240 61 35 155 10 320 “C 54.6 -3.2 22 26.9 70 18 ER111C,D 240 69 20 155 10 ERI12A,B,C,D 263 1 37 155 10 320 “C 59.7 -19.2 28 24.8 61 20 ERI13A.B 284 179 9 140 9 320 °C 39.8 75.3 26 24.8 40 34 ERI13C,D 284 181 10 140 9 ERI14A.B 306 214 62 140 9 Ther 280 °C 67.6 60.2 20 58.7 63 48 ERI14C,D 306 174 9 140 9 ERI15A,B,C 327 10 43 130 10 Ther 280 °C 53.0 64.2 10 89.0 54 90 ERI15D,E 327 50 4 130 10 ERI48A,B,C,D 337 164 6 130 10 Ther 280 °C 40.3 39.1 48 17.9 63 14 ERI16A,B,C,D 348 173 2 130 10 Ther 280 “C 53.3 59.0 16 32.1 58 29 ER149A,B,C 357 54 44 130 10 Ther 280 °C 53.4 78.3 14 34.8 45 44 ERI490.E 357 86 64 130 10 ERI17A,B,C 365 100 3 130 10 Ther 280 ®C 47.8 36.3 7 116.3 70 79 ERI17D,E 365 158 18 130 10 ERI18A,B,C,D 396 333 27 150 11 Ther 280 °C 49.7 4.5 30 22.9 71 15 ERI19A,B,C,D 415 316 29 150 11 Ther 280 °C 49.3 29.8 10 41.9 74 26 EHI20A.B 443 250 20 150 11 Ther 320 “C 60.5 26.1 36 20.7 84 11 ERI20C,D 443 255 18 150 11 ERI21A,B,C,D 467 131 10 150 11 Ther 320 °C 36.9 26.5 184 9.1 65 7 ERI50A,B,C,D 484 52 50 165 13 Ther 2 80 'C 49.9 15.6 15 32.7 76 19 ERI22A,B,C,D 486 160 10 165 13 ERI23A,B,C 504 280 7 165 13 Ther 280 °C 55.2 40.4 57 16.4 72 11 ERI23D.E 504 110 13 165 13 ERI51 A,B,C 509 102 40 165 13 Ther 280 “C 62.8 -15.5 36 21.0 65 16 ERI51 D,E 509 165 16 165 13 ERI24A,B,C,D 528 176 13 145 10 Ther 320 “C 47.3 28.3 563 5.2 72 3 ERI25A,B,C 564 162 6 145 10 Ther 320 °C 43.8 75.3 459 5.8 42 8 ERI25D.E 564 154 9 145 10 C. Eripol Average Site Vector Data 225 ERI26A,B,C,D 593 171 20 155 13 ERI52A,B,C,D 595 258 48 145 10 Ther 280 °C 38.4 19.8 97 12.6 67 9 ERI53A,B,C,D 608 138 19 145 10 Thor 280 “C 43.0 29.7 8 47.3 69 33 ERI27A.B 609 326 8 155 13 ERI27C.D 609 24 25 155 13 ERI28A,B,C,D 621 304 25 155 13 ERI54A.B 624 61 54 155 13 Ther 280 “C 4.2 -137.3 7 48.6 -39 38 ER154C,0,E 624 99 13 155 13 ERI55A.B 637 113 22 155 13 ER155C.D 637 225 5 155 13 ERI29A,B,C 644 163 17 155 13 Ther 320 “C -17.9 -158.0 14 74.3 -54 75 ERI29D.E 644 116 7 155 13 ERIS6A,B,C,D 645 230 13 155 13 Ther 280 °C -40.4 -173.8 36 20.7 -66 16 ERI57A,B,C,D,E,F 649 10 52 155 13 Ther 280 “C -21.3 -170.7 6 56.1 -55 56 ERI30A,B,C,D 677 24 5 155 13 ER158A,B,C 677 320 39 155 13 Ther 280 “C 26.1 36.3 28 23.6 56 23 ERI58D 677 315 46 155 13 ERI59A,B,C,D 688 300 51 155 13 Ther 280 "C 46.8 74.8 76 14 J 44 18 ERI31A.B 703 252 10 155 13 ERI31C.D 703 354 30 155 13 ERI60A.B 703 22 60 155 13 Ther 280 "C 15.2 -180.0 18 29.6 -35 45 ERI60C.D 703 6 74 155 13 ERI32A.B 732 270 24 153 13 Ther 280 "C -40.9 -164.1 21 57.7 -68 41 ER132C.D 732 290 22 153 13 ERI33A,B 748 344 15 160 8 Ther 280 "C -48.0 136.3 129 2 22 -38 32 ER133C.D 748 340 62 160 8 ERI61 A,B 748 333 41 160 8 Ther 280 "C -44.4 -172.1 959 4.0 -69 3 ERI61C.D 748 340 50 160 8 ERI34A,B,C 794 354 21 160 8 Ther 320 "C 30.9 23.1 16 32.1 62 27 ERI34D.E 794 5 48 160 8 ERI62A.B.C 794 359 35 160 8 ERI62D.E 794 2 31 160 8 ERI35A.B 806 292 24 160 8 Ther 280 "C 50.5 -10.3 11 39.5 63 32 ERI35C.D 806 295 26 160 8 ERI63A,B,C,D 806 321 69 160 8 Ther 280 °C 53.2 44.0 39 19.9 69 14 ERI64A.B.C 820 268 50 160 8 Ther 280 °C 45.5 72.6 56 16.6 45 21 ERI64D.E 820 263 63 160 8 ERI36A.B 821 353 62 140 7 Ther 280 “C 63.0 19.4 50 17.7 89 9 ERI36C 821 45 73 140 7 ERI36D 821 31 19 140 7 ER137A.B.C 845 246 41 140 7 Ther 280 °C 51.2 12.7 155 9.9 76 6 ERI37D 845 175 6 140 7 ERI38A.B.C 877 242 15 140 7 ERI38D.E 877 253 14 140 7 Total Number of SamDles 63 % of Class I Data Sites 48 % of Class II Data Sites 25 Eripol Data Continued 226 Liguerre Magnetostratigraphic Section AZ1 ,DP1 Azimuth and Dip of Sampled Surface (°) AZ2.DP2 Azimuth and Dip of Local Bedding (°) Field Information ~ l jAverage Vector & Statistical Data v g p Lat. | SDecimen # Ht (ml AZ1 DPI AZ2 DP2 Treatment Level Inc £ ££ Kappa A-95 Class 1 Class II Error UQ40A.B 0 205 13 208 12 UG40C.D 0 198 15 208 12 LIG01 A,B,C,D 10 260 20 208 12 Ther 280 “C -58.0 -142.8 1078 3.8 -76 2 LI041 A,B,C 10 160 48 208 12 Ther 280 °C -42.0 -155.6 108 11.9 -69 8 UG41D.E 10 228 29 208 12 LIG02A,B,C,D 33 220 57 208 12 UG42A.B.C.D 48 190 9 225 13 Ther 280 °C -31A -159.4 829 4.3 -62 4 UG03A,B,C,D 57 148 38 225 13 Ther 280 °C 10.8 171.0 19 61.5 •34 95 UG43A.B.C 75.5 79 24 230 10 Ther 280 °C 53.1 -28.5 5 59.3 51 64 UG 430 75.5 266 1 230 10 UG04A.B.C 68 180 6 230 10 Ther 280 °C 10.3 -107.4 25 52.0 -22 94 UG04D.E 88 161 19 230 10 UG44A,B,C,D 99 310 16 230 14 Ther 280 "C -31.7 -140.8 63 15.6 -58 14 UQ05A,B,C,D 119 216 12 230 14 Ther 280 "C -11.9 176.7 9 99.7 -46 121 UG45A,B,C,D 135 330 8 228 14 Ther 280 "C 58.0 -41.2 22 27.2 45 34 UG06A.B 152 45 12 228 14 Ther 280 "C 25.8 -93.0 32 45.8 -6 91 UG06C.D 152 225 36 228 14 UG46A.B 174 325 26 228 14 Ther 280 "C 67.0 -12.2 96 25.8 68 18 UG46C.D 174 325 21 228 14 UG46E.F 174 245 8 228 14 UG07A.B 185 64 4 236 10 Ther 280 °C 35.4 66.9 25 25.3 44 32 LIG07C 185 46 5 236 10 UG07D.E 185 291 13 236 10 UG06A,B.C,D 209 357 4 224 8 Ther 280 "C -14.5 -136.0 169 9.5 -47 11 UG09A,B,C,D 253 344 32 224 8 Ther 320 °C 5.1 175.1 38 41.7 •40 57 UG10A,B,C,D 288 205 13 234 1 1 Ther 280 "C -19.8 -143.9 114 11.6 -53 12 UG11 A,B,C,D 303 203 16 234 1 1 Ther 280 °C -45.8 -126.8 6 130.4 -58 119 UG12A,B,C,D 324 76 15 234 1 1 Ther 320 °C 10.4 -106.3 160 19.9 -21 37 UG13A,B,C,D 354 81 56 230 8 UQ14A,B,C,D 384 151 27 230 8 Ther 280 °C -33.8 -146.9 776 9.0 -62 8 UG47A,B,C,D 385 16 10 230 8 Ther 280 "C 12.5 -164.0 29 23.3 -51 25 UG48A,B,C,D 400.5 275 54 230 8 Ther 280 “C 56.2 40.3 8 47.6 73 30 UG15A 404 61 32 230 8 UG15B,C,D 404 344 25 230 8 UG49A 422 167 11 230 10 Ther 280 "C 65.6 18.8 28 49.4 87 25 UG49B.C 422 220 13 230 10 UG49D 422 227 3 230 10 UG16A.B 423 344 26 230 10 Ther 320 “C 36.2 -56.6 23 54.6 23 96 UG16C.D 423 73 22 230 10 LIG50A,B,C,D 449 315 12 230 10 Ther 280 °C 41.3 11.0 221 8.3 68 6 UG17A,B,C,D 460 258 39 230 10 Ther 280 "C 34.6 -48.3 70 30.4 29 50 UG51 A,B,C,D 469 210 18 230 10 UG18A,B,C,D 491 235 14 210 8 Ther 280 °C 63.0 4 2 22 27.2 79 15 UG19A.B 518 348 36 240 8 Ther 320 "C 79.7 50.0 159 19.9 61 17 UG19C.D 518 358 35 240 8 UG20A.B 530 241 14 240 8 Ther 280 “C 48.9 1.6 23 54.9 69 38 UG20C.0 530 298 50 240 8 UG21A.B 564 355 23 240 8 Ther 320 °C 7.4 15.1 19 29.1 49 34 UG21C.D 564 262 10 240 8 UG22A,B,C,D 578 168 9 220 11 Ther 320 “C 53.9 26.1 43 19.0 79 11 UG23A.B 604 199 6 220 1 1 Ther 320 °C 0.3 3.5 12 82.1 43 107 UG23C.D 604 155 5 220 11 UG23E.F 604 125 7 220 11 UG24A,B>C,D 619 11 47 220 11 Ther 320 °C 54.2 39.4 27 24.1 72 15 UG25A,B,C 640 335 24 190 6 Ther 280 °C 42.6 35.7 20 59.0 66 44 UG2SD.E 640 70 29 190 6 D. Liguerre Average Site Vector Data 227 UG52A,B,C,D UQ53A.B.C UQ53D.E UG26A.B UQ26C,D,E UG54A UG54B,C,D UG27A,B,C UG27D.E UG28A,B,C UG28D.E UG29A.B UG29C.D UG55A,B,C UG55D.E UG56A,B,C UG56D.E UG30A,B,C,D UG31A UG31B.C LIG31 D,E LIG59A,B,C,D UG59E.F UG57A,B,C UG57D.E UG32A.B UG32C.D UG33A LIG33B UG33C.D UG34A UG34B UG34C.D UG58A,B UGS8C.D UG35A,B,C,D UG60A.B.C UG60D.E UG36A.B UG36C.0 UG61A.B UG61 C,D 640 660 660 684 684 699 699 712 712 736 736 764 764 767 767 781 781 791 805 805 805 824 824 835 835 836 836 852 852 852 867 867 867 867 867 918 938 938 980 980 983 983 25 340 355 346 6 67 58 142 83 91 121 195 208 300 311 243 101 155 153 191 235 6 19 118 174 331 33 115 195 61 126 240 114 208 260 135 168 160 330 180 202 195 19 52 24 23 36 45 33 10 22 17 3 15 24 9 18 14 51 3 21 7 1 38 37 18 8 16 44 47 26 12 26 16 30 28 31 15 21 9 42 13 10 45 190 190 190 190 190 200 200 190 190 170 170 140 140 140 140 140 140 140 155 155 155 170 170 155 155 155 155 155 155 155 170 170 170 155 155 170 170 170 165 165 165 165 6 6 6 6 6 10 10 6 6 7 7 8 8 8 8 8 8 8 11 11 11 15 15 11 11 11 11 11 11 11 15 15 15 11 11 15 16 16 16 16 16 16 Ther 280 °C 59.5 20.6 36 20.9 86 22.0 16.8 50 17.6 57 Ther 320 °C 15.9 13.6 307 14.3 53 Ther 320 °C -53.8 158.0 16 66.8 -56 Ther 280 °C -20.6 -138.6 52 17.3 -51 Ther 280 °C -25.6 -112.9 17 31.1 -39 Ther 280 °C -48.0 132.8 4737 1.8 -36 Ther 280 ® C -44.0 -169.8 410 6.1 -70 Ther 320 °C -24.6 -160.0 33 21.8 -58 Ther 280 °C 62.2 -17.0 7 125.2 64 Ther 280 “C 41.8 51.9 12 36.5 57 Ther Ther 320 °C 31.2 280 °C 63.1 52.4 28.7 213 11 17.2 38.2 84 51 Ther 320 “C 48.1 3t.O 11 86.3 72 Ther 280 °C 65.0 -22.1 79 28.9 61 Total Number of Sites 58 % of Class I Data % of Class II Data 41 40 1 1 17 19 43 20 35 19 20 55 24 Liguerre Data Continued 228 Esera Valley Composite Section AZ1 ,DP1 Azimuth and Dip of Sampled Surface AZ2.DP2 Azimuth and Dip of Local Bedding Eff. Ht Effective Stratigraphic Height in Composite Section Grustan Section Field Information 1 (Average Vector & Statistical Data v g p Lat. Specimen Ht fml E ff. ht ( m ) A Z 1 DP1 A Z 2 DP2 Treat't Level Inc Dec Kappa A-95 Class 1 Class 1 1 Error GRU61 A,B,C -5 445 40 1 0 184 9 QRU61D -5 445 340 1 0 184 9 QRU01A,B 1 451 115 1 1 104 9 AF 200 Oe -29.1 -166.9 38 2 0 .2 -60 18 GRU01 C,D 1 451 2 1 0 32 184 9 GRU60A.B 6 456 60 25 184 9 AF 200 Oe -35.1 -129.0 59918 1 . 0 -54 1 GRU60C,D,E 6 456 19 27 184 9 GRG02A,B 10 460 13 82 184 9 Ther 280 °C -63.5 157.2 108 13.4 -74 8 GRU02C.O 10 460 92 47 184 9 AF 200 Oe -42.0 -145.9 320 14.0 -67 1 0 GRU03A,B,C,D 13 463 248 4 184 9 AF 200 Oe 38.5 -139.2 24 25.5 -20 46 GRU04A,B,C,D 28 478 216 11 190 1 0 AF 400 Oe -3.5 -142.1 126 22.5 -41 24 GRU05A.B 39 489 39 34 176 7 AF 200 Oe 60.6 1 58.9 27 24.4 4 49 GRU05C.D.E 39 489 105 26 176 7 GRU07A,B,C,0 44 494 272 69 176 7 AF 400 Oe 22.6 86.6 51 17.4 25 30 GRU06A,B,C,D 50 500 350 2 2 176 7 AF 200 Oe 72.6 120.1 37 20.5 32 32 GRU08A,B,C,D 57 507 06 38 176 7 AF 400 Oe 4.3 -132.5 15 32.5 -37 47 GRU08E,F,G,H 57 507 73 2 0 176 7 GRU09A 68 518 354 2 188 8 GRU09B 68 518 58 2 0 188 8 GRU09C 68 518 58 69 188 8 GRU090 68 518 38 5 188 8 GRU10A.B 74 524 320 35 188 8 GRU10C.E 74 524 58 64 188 8 GRU10D 74 524 68 9 188 8 GRU11 A,B,C,D 85 535 44 40 190 9 Ther 280 °C -33.8 -169.1 7 53.3 -62 44 GRU12A,B,C,D 99 549 22 38 190 9 AF 200 Oe -16.7 -156.7 7 123.1 -54 126 GRU40A.B 105 555 240 33 190 9 AF 200 Oe -54.1 -104.8 69 30.6 -47 37 GRU40C.D 105 555 262 2 0 190 9 GRU13A 109 559 248 28 190 9 GRU13B 109 559 340 32 190 9 GRU13C 109 559 94 12 190 9 GRU13D 109 559 225 53 190 9 GRU14A,B,C,0 120 570 198 18 190 9 GRU41A.B 124 574 159 2 0 220 5 AF 200 Oe 67.2 -44.6 199 17.8 47 21 GRU41C.0 124 574 120 14 2 2 0 5 GRU15A,B,C,D 128 578 1 0 59 2 2 0 5 GRU16A,B 138 588 132 11 2 2 0 5 GRU16C.0 138 588 84 15 2 2 0 5 GRU42A,B,C,D 144.5 595 185 6 2 2 0 5 AF 200 Oe -35.2 -156.1 20 28.2 -65 22 GRU17A,B,C,D 149 599 32 68 2 2 0 5 AF 400 Oe -4.4 -153.0 59 33.0 -47 39 GRU43A 158 608 190 21 230 6 AF 200 Oe 25.9 68.7 27 50.3 38 71 GRU43B.C 158 608 137 19 230 6 GRU43D 158 608 179 2 0 230 6 GRU18A,B,C,D 160 610 53 34 235 8 GRU19A,B,C,D 170 620 75 26 235 8 AF 400 Oe 14.7 -174.0 233 16.4 -36 24 GRU44A.B 174.5 625 27 5 235 8 AF 200 Oe -73.8 142.3 359 13.2 -53 14 GRU44C.D 174.5 625 265 5 235 8 GRU20A.B 178 628 43 60 2 0 0 6 GRU20C 178 628 42 63 2 0 0 6 GRU20D 178 628 34 40 2 0 0 6 GRU21 A,B,C,D 195 645 55 81 2 0 0 6 AF 200 Oe -9.6 -160.0 826 4.3 -50 5 GRU22A,B,C,D 200 650 25 24 204 4 AF 400 Oe -4.6 -154.8 7 52.8 -47 63 GRU23A,B,C,D 209 659 210 5 204 4 AF 200 Oe -5.3 -147.3 82 13.7 -47 17 GRU24A,B,C,D 218 668 356 78 205 4 AF 200 Oe -4.8 -168.5 31 22.5 -45 28 GRU25A,B,C,D 226 676 351 44 205 4 AF 400 Oe 6.8 -159.4 15 32.5 -42 43 GRU26A,B,C,D 236 686 311 10 206 6 AF 400 Oe -11.7 -154.5 45 18.6 -51 20 GRU27A,B,C,D 249 699 50 57 210 8 Ther 280 "C -29.1 -126.2 9 46.8 -36 69 GRU28A,B 259 709 84 56 210 8 GRU28C.D 259 709 90 53 210 B GRU29A.B 271 721 145 10 210 8 AF 200 Oe -12.2 -165.4 21 27.8 -51 30 GRU29C.D 271 721 122 20 210 8 E. Esera Valley Average Site Vector Data 229 QRU31 A,B,C,D 291 741 89 68 210 6 QRU45A,B,C,D 293 743 141 2 210 6 AF 200 Oe 6.0 -10.4 13 77.1 41 105 GRU30A.B.C.D 295 745 245 5 210 6 AF 400 Oe 60.2 15.6 37 20.5 85 10 GRU32A,B,C,D 296 746 224 12 232 10 GRU46A,B,C 300 750 231 11 210 6 AF 200 Oe 19.6 -127.5 109 24.1 -28 40 GRU46D 300 750 295 4 210 6 GRII33A.B 307 757 252 21 233 10 Ther 280 "C 23.9 -5.8 63 17.6 57 16 GRU33C.D 307 757 294 8 233 10 AF 200 Oe 44.8 -28.7 11 86.9 47 104 GRU47A,B,C,D 312 762 235 8 235 8 GRU34A,B,C,D 322 772 244 45 233 10 AF 200 Oe 11.9 97.8 29 23.2 13 45 GRU48A,B,C,D 326.5 777 64 24 245 9 AF 200 Oe -38.4 140.2 6 53.5 -36 79 GRU35A,B,C,D 343 793 241 13 270 11 GRU50A,B,C,D 343 793 208 25 270 11 AF 200 Oe -59.6 -144.7 36 20.8 -78 12 GRU36A.B 353 803 270 23 270 12 GRU36C.D 353 803 280 23 270 12 AF 400 Oe 41.0 -6.9 4 67.8 60 60 GRU49A,B,C,D 360 810 29 58 270 12 AF 200 Oe -7.2 -161.0 17 31.3 -49 36 Meson de Pascual Section Held Information | |Average Vector & Statistical Data v g p Lat. Specimen Ht (m) EfLHlim l AZ1 DP1 AZ2 DP2 Ireaimeni Level Inc Dec Kappa A-95 Class 1 Class II Error MDP01 A,B,C,D 4.5 305 178 7 203 15 Ther 280 “ C 25.3 13.6 6 58.1 58 53 MDP02A,B,C,D 10 310 175 6 203 15 MDP03A,B,C,D 31 331 161 15 203 15 Ther 280 “ C 24.5 -33.3 20 28.2 35 43 MDP04A.B.C.D 50 350 67 35 192 12 Ther 280 “C 44.3 59.7 22 27.1 53 28 MDP05A,B,C,D 91 391 26 7 213 12 Ther 280 “C 46.1 -32.7 493 5.6 45 7 MDP06A.B 99.5 400 81 35 213 12 Ther 280 “ C 63.1 7.0 22 27.0 81 15 MDP06C 99.5 400 8 47 213 12 MDP06D.E 99.5 400 170 17 213 12 MDP07A.B 113.5 414 165 5 250 7 MDP07C.D 113.5 414 55 35 250 7 MDP09A.B 113.5 414 294 14 250 7 Ther 280 °C 22.3 -24.3 17 65.7 40 91 MDP09C.D 113.5 414 240 25 250 7 MDP08A,B,C,D 126.5 427 98 13 250 7 Ther 280 °C 59.7 -53.8 355 6.6 37 9 MDP10A,B.C,D 139 439 275 22 220 6 Ther 280 °C -36.6 -152.0 7 51.5 -65 40 MDP11A 149 449 109 19 220 6 MDP11B.C 149 449 175 7 220 6 MDP11 D,E 149 449 22 5 220 6 MDP12A,B 156 456 326 58 200 8 MDP12C 156 456 335 57 200 8 MDP12D.E 156 456 327 54 200 8 MDP13A,B,C,D 165 465 206 10 206 10 Ther 320 °C -19.7 174.9 22 27.2 -49 31 MDP14A,B 170 470 22 70 208 10 Ther 320 °C -56.7 110.1 19 29.5 •25 51 MDP14C.D 170 470 22 74 208 10 Santa Liestra Section Field Information | (Average Vector & Statistical Data v g p L at. Specimen HI (mi Eff. Ht fmi AZ1 DPI AZ2 DP2 Treatment Level Inc Dec Kappa A-95 Class 1 Class II Error SLA01A 1 1 183 20 320 9 AF 200 Oe 51.5 14.4 11 38.9 77 23 SLA01 B,C 1 1 202 24 320 9 SLA 01 D,E 1 1 244 5 320 9 SLA02A 15 15 142 4 315 10 AF 200 Oe 35.1 1.3 14 33.7 60 29 SLA02B.C 15 15 200 11 315 10 SLA02D 15 15 168 19 315 10 SLA03A 30 30 240 11 315 10 AF 200 Oe 61.6 12.5 21 27.8 84 14 SLA03B 30 30 226 14 315 10 SLA03C.D 30 30 306 32 315 10 SLA04A.B 42.5 43 150 45 315 10 AF 200 Oe 37.7 1.6 245 16.0 62 13 SLA04C.D 42.5 43 141 40 315 10 SLA05A,B,C,D 53 53 57 25 315 10 AF 200 Oe 28.6 12.9 50 36.1 60 32 SLA06A, B 81 81 166 15 315 10 AF 200 Oe 36.8 -47.4 467 11.6 30 19 SLA06C.D 81 81 248 4 315 10 SLA07A,B,C,D 88 88 148 16 315 10 AF 200 Oe 44.1 7.5 44 18.8 69 13 SLA08A,B,C,D 115 115 332 15 315 10 AF 200 Oe 52.3 29.1 377 6.4 76 4 Total Number of Sites 71 Esera Valley Data Continued % of Class I Data Sites 51 % of Class II Data Sites 27 230 Lascuarre Magnetostratigraphic Section AZ1 ,DP1 Azimuth and Dip of Sampled Surface (°) AZ2.DP2 Azimuth and Dip of Local Bedding (°) Field Information ~| [Average Vector & Statistical Data vg p Lat SDecimen « Ht C m ) A il DP1 AZ2 DP2 TrvaunHi Level Inc Dec Kanoa Class I Class II Error LASOIA.a.C 1 162 32 175 12 Ther 320 “C -28.1 -176.9 8 47.4 -60 4 1 LAS01D.E t 189 13 175 12 LAS CCA 6 117 14 175 12 Ther 320 “C 26.0 44.0 12 37.6 42 50 LAS 02 B 6 248 3 175 12 LAS02C,D,E S 213 5 175 12 LAS03A,B,C 25 216 10 175 12 Ther 280 “C 29.2 -167.1 24 25.7 -29 43 LAS COD,E 25 224 41 175 12 LAS04A,B,C 32 40 62 175 12 Ther 280 “ C -58.9 -136.8 44 38.6 -58 36 LAS040.E 32 43 48 175 12 LAS06A 40 357 39 175 12 Ther 320 °C 23.4 13.0 36 43.1 56 42 LAS05B 40 27 69 175 12 LASOSC 40 18 70 175 12 LASOSD 40 66 10 175 12 LAS06E 40 17 83 175 12 LAS06A.B 48 20 20 175 12 LASOSC, D,E 48 98 14 175 12 LAS 07 A 67 7 24 175 12 Ther 320 “C 50.2 28.3 17 31.3 64 25 LAS07B,C,D,E 67 356 20 175 12 LAS0SA,B,C 74 306 68 175 12 Ther 280 °C -70.7 -179.7 45 18.5 -80 10 LAS0SD,E,F 74 222 10 175 12 LAS 09 A, B 90 115 16 196 15 Ther 280 “ C -28.0 -126.1 9 43.3 -36 64 LAS09C,D,E 90 54 37 196 15 LAS10A.B 105 1 1 42 196 15 Ther 280 “ C -34.0 -154.1 6 56.9 -56 55 LAS10C,D,E 105 356 42 196 15 LAS11 A,B,C,D,E 120 190 17 196 15 Ther 280 “C -42.2 -177.1 13 35.9 •70 24 LAS12A.B 125 19 72 196 15 Ther 320 “ C -19.1 142.6 5 58.8 -43 77 LAS12C,D,E 125 12 64 196 15 LAS13A 160 125 a 196 15 Ther 280 “C 19.0 143.0 15 33.5 -26 57 LAS13B.C 160 220 6 196 15 LAS13D 160 15 37 196 15 LAS13E 160 1 41 196 15 LAS14A.B 170 357 71 196 15 Ther 280 °C 69.0 33.7 55 16.8 67 12 LAS14C 170 350 70 196 15 LAS14D.E 170 261 46 196 15 LAS15A,B,C,D,E 185 184 28 196 15 LAS16A,B,C,D 200 50 12 196 15 Ther 280 “C 56.5 -0.9 53 17.2 82 9 LAS16E 200 126 6 196 15 LAS17A,B,C,D,E 215 181 21 196 15 Ther 320 “C 62.5 -2.9 4344 1.9 88 1 LAS1BA.B 230 94 16 196 15 Ther 280 “C -54.1 165.5 16 32.2 -75 19 LAS1SC.D.E 230 60 11 196 15 LAS19A.B 240 306 47 196 15 Ther 320 “ C 66.5 -28.0 7 48.7 70 33 LAS19C 240 266 6 196 15 LAS19D.E 240 186 74 196 15 LAS 20A, B 260 281 67 196 15 Ther 280 “ C -58.4 153.6 16 31.5 -70 21 LAS20C 260 265 78 196 15 LAS20D.E 260 255 80 196 15 LAS21A.B.C 273 303 42 196 15 LAS21D.E 273 52 61 196 15 LAS22A,B,C,D,E 295 150 21 196 15 Ther 320 “ C 41.8 -34.4 51 17.5 55 17 LAS23A.B 305 311 81 196 15 Ther 320 “ C -50.3 -161.5 7 52.4 -70 35 LAS23C.D 305 332 77 196 15 LAS23E 305 310 64 196 15 LAS24A.B 320 352 43 196 15 Ther 280 “C 31.8 -15.5 85 13.5 60 12 LAS24C 320 125 8 196 15 LAS24D.E 320 127 3 196 15 LAS2SA,B,C,D,E 330 326 76 196 15 Ther 280 “ C 60.8 -7.0 86 13.4 84 7 T otal N um ber of S ites 2 5 % of C lass I D ata 6 8 % of C lass II D ata 2 0 Lascuarre Average Site Vector Data Appendix 2: Detailed Lithologic Sections 232 Sedimentary Structures Lithofogies o o o Y Y Y A A A A Nodular Calcrete Horizons Root-Casts and Bioturbation Gypsiferous Overbank Interval Nummulitic Foraminifera Bivalve Shells or Debris Gastropod Shells Basal Flute and Groove Structures Gutter Casts Internal Scour Surface Mudcracked Surface Ripple Cross- Stratification Planar Cross- Stratification Sand Trough Cross-Stratification Gravel Filled Troughs & Scours Gravelly Planar Cross-Stratification X 11 \ 1 1 3 0 Reddened/Mottled Overbank Mudstones & Siltstones Fine-Grained Sandstone Sheets & Coarse Siltstones Medium to Coarse-Grained Channel Sandstones Channelized Pebbly Sandstones & Conglomerates Fine-Grained Micritic Limestone Horizons ] Gray Marls and Marine Siltstones Reworked Bioclastic Limestones Siliciclastic Sandstone Turbidites Shallow Marine Sandstones Uni-Directional Paleocurrent Indicator (num ber of m easurem ents) Bi-Directional Paleocurrent Indicator (num ber of m easurem ents) Paleomagnetic Sample Location Grain-Size & Scale A. Legend for Detailed Lithologic Sections 233 200 8 4 o 5 0 82q < £ & « » < Y Y Y SLUMP Y Y Y Y Y Y XXX X XXX / <S2>«*>© 400 300 10 o 200 ® < » © Olistostrome * Y Y Y Corals Olistostrome Y Y Y Y Y Y < © « » © B. Mediano Lithologic Column 234 600 180 910 500 13q 400 < 2> < * > © Y Y Y Y Y Y Olistostrome Mud Rip-Up Clasts 800 Y Y Y Y Y Y Y Y Y 250 Y Y Y Road Bridge 700 220 Y Y Y Y Y Y Y Y Y 600 Mediano Lithologic Column Continued. 235 H echo G roup Turbidites**— j— ► Sobrarbe Fm. 1000 33q 61 o 31q 29 o 800 1200 - 39 q ® mm & ® © ® ® •© ® ® * o ® 4 ® <»© Y Y Y Y Y Y 67 q 38q 660 37 o E L L _ IS ~ 650 § 1100- O 360 LU 6 4 0 _ 4 • 35 O - — E LJL 0 _ Q c o 1 — . -Q O ^ 63 q 340 6 2 0 1000 £ E E 3 £ B f t . j g 2 2 2 2 2 l Z Z t 02 £nmm 37 ®»® * ®<»© ® 4 D © ® < » @ m Mediano Lithologic Column Continued. 236 1400 4 8 0 _ 47Q J 7 2 0 46 0 l l 8 1300 44 0 70 0 43 0 69 0 420 J 40Q 41 O K S S S S & Lateral Projection /QZ 2Z t 2S 2 Z Z 2T / X X Z 2 8 B & B & SMS. \ 3 fmr mn A A A A Y Y Y 1600 540 770 53 0 76 0 52Q 75 0 1500 51 n 74 0 73 0 50Q 1200 49Q 1400 Mediano Lithologic Column Continued. 237 Additional Samples Taken at: 1685 m (# 9 4 ) 1694 m (#9 5 ) 1680 4 Mediano Lithologic Column Continued. 238 100 5 0 2 11 200 o 10o 90 4 O 5 0 3 O '*2 * ^ 8 O 1 5 0 7 O n«s«R 10 20 1 O 60 100 rg g g sft V fg C. Almazorre Lithologic Column. 239 140 13q 120 200 rrm Vertical Facies Sequence through a typical Gravel- Dominated Channel-Fill (Lithofacies i). Almazorre Lithologic Column Continued. 240 D. Eripol Lithologic Column 241 600 A 5 2 0 2 6 0 800 2 5 0 m 24 o 21 o 2 0 Q 1 9 0 \ 14 '* * 2ags 7 4 400 - P t t t t i Eripol Lithologic Column Continued. 6 2 0 3 4 0 33' 32q 60 00 s 7 T 59q 58i 30 27 53 600 T O g s S - o o o o o o grnga. ^ 10 o o o 242 7511 Oligocene Conglomerates 1020 -I _ Angular "Unconformity 1000 960-4 900-4 3 8 0 m s . 37 o o o o 3 360 6 4 0 8 800 T i m Eripol Lithologic Column Continued. o o o 243 74918 200 70 4 6 Q 6 O 45 q 5 0 100 44 o 4 o 430 30 420 20 4 , ’8 40q X Y Y Y Y Y Y A A A A Y Y Y E = \ * -fe zr^b Y Y Y Y Y Y Y Y Y Y Y Y Y Y Y 13q 1 2 o m 10o 9 o \ Y Y Y 1 Y Y Y AT 1 3 ~*±SX£*&- 13 A A A A 22 E. Liguerre Lithologic Column. 244 600 - I 220 210 20 o 1 9 0 500 180 1 7 0 51 O 50q = " 490 4 T \ 1 6 q = 400 2 §§§§§_ Y Y Y \ 2 355^- X S F S S F S 17 — u— ^ Y Y Y 3 10 28 o 270 5 " 26 o 53 o rffn 24 o 23 o 600 12 "^g£> zzri > iS *£ g > s,^ S S ^ 1 y y y g 16 \ 4 Liguerre Lithologic Column Continued. 245 iooo A 61 o 36 q o o o 60 Q 590 35 o 900 H 58 Q 3 4 0 3 3 0 = 5 7 Q 3 2 0 310 800 3 T V 1 NO EXPOSURE Oligocene Conglomerates TTTT 1160 i Angular Unconformity 1000 Liguerre Lithologic Column Continued. 246 84 200 ® < ■ » © F. Esera Valley Composite Column 300 200 Santa Liestra. < & > @ <^ 9 © © ® © < * » © < S J > < 0 © (§ ) a g p @ © > < * ► © © •© 247 Perarrua Formations | >Capella Formation 05Q 190 140 130 03Q 0 2 Q “ NO EXPOSURE 3 < © < ■ » © ® < a ,® © < a > © < Z S > © 120 10Q 08 0 !8 i NO EXPOSURE 100 Esera Valley Composite Column Continued - Meson de Pascual. 248 18q 4 3 0 16 q 13Q 1 1 O 04q 0 1 1 mixm f sm m 5 JrBMUm ^ 22 Y Y Y Y Y Y Y Y Y Y Y Y Puy de Cinca Lst. . 1 — 4oa-| C o ro E l . o LL O C (0 a. 8 c o ro E o LL i2 0 Q. (0 O 3 3 q 4 7 0 8 29q Massive Reef < 3 3 > « ■ » © m u jrm 7 7 T lg L L U l * r ' *&m > ^ o © finnan ssss^. >436^. 03 Esera Valley Composite Column Continued - Grustan. 249 Oligocene Conglomerates f- 500. 400. _01 o CaDella Escanilla Limestone Lascuarre Schematic Column. Appendix 3: Geohistory Data 251 ESERA VALLEY Int. Tod Int. Base Thickness Base. Deoth Tect. Sub Total Rate Tect Rate (Ma) (Ma) (km) (km) (km) (cm/kyr) (cm/kyr) 43.13 44.57 0.910 0.910 0.389 2 1 44.57 47.01 0.865 0.880 0.379 9 3 47.01 48.51 0.565 0.661 0.300 44 20 48.51 48.51 0.000 0.000 0.000 MEDIANO Int. Top Int. Base Thickness Base. Deoth Tect. Sub Total Rate Tect Rate (Ma) (Ma) (km) (km) (km) (cm/kyr) (cm/kyr) 41.31 42.14 1.455 1.455 0.581 18 6 42.14 42.57 1.215 1.304 0.533 17 6 42.57 42.90 1.105 1.230 0.509 23 7 42.90 43.13 0.995 1.154 0.485 12 3 43.13 44.57 0.950 1.126 0.478 45 18 44.57 47.01 0.310 0.480 0.224 20 9 47.01 47.01 0.000 0.000 0.000 Time (Ma) 49.00 48.00 47.00 46.00 45.00 44.00 43.00 42.00 41.00 0 .0 0 0 ■T.T.O 0.400 0.600 1.000 - 1.400 - 1.600 -------■— ESERA ------ O-------ESERA ^ MED Base. --- * — MED Tect. Base. Tect. Sub. Depth Sub. Depth Chron 20 Geohistory Data and Graph. 252 ERIPOL InLTop Int. Base Thickness JBasg, D.eptti TecLSuk Total Rate Tect Rate (Ma) (Ma) (km) (km) (km) (cm/kyr) (cm/kyr) 36.54 36.93 0.930 0.930 0.385 18 7 36.93 37.16 0.845 0.861 0.359 15 6 37.16 37.31 0.805 0.827 0.346 23 9 37.31 41.31 0.765 0.793 0.333 10 4 41.31 42.14 0.335 0.401 0.179 7 3 42.14 42.57 0.275 0.342 0.156 23 10 42.57 42.90 0.190 0.245 0.112 74 34 42.90 42.90 0.000 0.000 0.000 LIGUERRE lnL_Iop Int. Base Thickness Base. Depth Tect. Sub Total Rate Tect Rate (Ma) (Ma) (km) (km) (km) (cm/kyr) (cm/kyr) 36.54 37.31 0.865 0.865 0.360 15 6 37.31 41.31 0.725 0.747 0.314 8 3 41.31 42.14 0.395 0.446 0.195 23 10 42.14 42.57 0.200 0.252 0.116 2 1 10 42.57 42.90 0.125 0.163 0.075 49 23 42.90 42.90 0.000 0.000 0.000 Time (Ma) 44.00 43.00 42.00 41.00 40.00 39.00 38.00 37.00 30.00 -0 .2 0 0 i . i .. . i ■ i i .... ■ i i ,. i, i l * l i - I i I Escanilla Formation Geohistory Data and Graph. 253

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Creator Bentham, Peter A. (author)

Core Title The tectono-stratigraphic development of the western oblique ramp of the south-central Pyrenean thrust system, northern Spain

Contributor Digitized by ProQuest (provenance)

Degree Doctor of Philosophy

Degree Program Geological Sciences

Publisher University of Southern California (original), University of Southern California. Libraries (digital)

Tag OAI-PMH Harvest,Plate Tectonics

Language English

Permanent Link (DOI) https://doi.org/10.25549/usctheses-c29-354771

Unique identifier UC11219075

Identifier DP28600.pdf (filename),usctheses-c29-354771 (legacy record id)

Legacy Identifier DP28600.pdf

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Document Type Dissertation

Rights Bentham, Peter A.

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Source University of Southern California (contributing entity), University of Southern California Dissertations and Theses (collection)

Access Conditions The author retains rights to his/her dissertation, thesis or other graduate work according to U.S. copyright law. Electronic access is being provided by the USC Libraries in agreement with the au...

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