Terrestrial Paleoecology & Global Change (Russian Academic Monographs, 1)

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Terrestrial Paleoecology and Global Change Y 80 60 40 20 10 C 80 60 40 20 10 B 80 60 40 20 10

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R USSIAN A CADEMIC M ONOGRAPHS , N o 1

Terrestrial Paleoecology and Global Change

Valentin Krassilov is the author of 20 books, among them the Palaeoecology of Terrestrial Plants, Cretaceous Period, Angiosperm Origins, Ecosystem and Egosystem Evolution, etc. The new book is based on his many year experience in the fields of palaeobotany, palaeoecology, structural geology and evolutionary biology.

V. A. KRASSILOV

This book critically evaluates the currently popular ideas of global change based on the plate tectonics, extraterrestrial impacts, greenhouse warming, etc. and offers alternative models. Krassilov presents ecosystem evolution as a sustainable oriented process with an increase in the biomass-to-dead mass ratio as a measure of progress. This general tendency is reversed the geobiospheric crises starting in the earth's interior and surfacing as the concerted geomagnetic, tectonomagmatic, geochemical and climatic events. These affect biota through turnovers of biotic communities and the adequate changes in population adaptive strategies, a major force under the species originations and extinctions, as well as the genomic evolution. The evolution of humans is envisaged as guiding this species to the role of the earth custodian. The book is important for evolutionists, ecologists, geologists, climatologists, geneticists, integrative biologists, botanists, zoologists, and the general educated person who is intrigued by the dynamic historical processes which shape biodiversity evolution. It could be used as a course book for undergraduate and graduate studies and is an excellent example of inspiring and creative interdisciplinary research of our planet.

Introduction

TERRESTRIAL PALAEOECOLOGY AND GLOBAL CHANGE Valentin A. Krassilov

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Valentin A. Krassilov. Terrestrial Palaeoecology and Global Change

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Introduction

Paleontological Institute, Russian Academy of Sciences Institute of Evolution, University of Haifa

Terrestrial Palaeoecology and Global Change by Valentin A. Krassilov

Sofia – Moscow 2003

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TERRESTRIAL PALAEOECOLOGY AND GLOBAL CHANGE First published 2003 ISBN 954-642-153-7

Russian Academic Monographs No. 1

© PENSOFT Publishers All rights reserved. No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form by any means, electronic, mechanical, photo copying, recording or othrwise, without the prior written permission of the copyright owner.

Pensoft Publishers, Acad. G. Bonchev Str., Bl.6, 1113 Sofia, Bulgaria Fax: +359-2-70-45-08, e-mail: [email protected], www.pensoft.net Printed in Bulgaria, January 2003

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CONTENTS

I. INTRODUCTION ...................................................................................................... xi I.1. The idea of global change ................................................................................... xi I.2. Global change: general ...................................................................................... xiii I.3. Driving forces: state of the art ......................................................................... xiv

II. TAPHONOMY .......................................................................................................... 1 II.1. Dead mass fossilization as a condition of life .................................................... 1 II.2. Self-fossilization and taphonomic enrichment .................................................... 8 II.3. Taphonomic bias ............................................................................................... 11 II.4. Taphonomy and evolution ................................................................................. 15 II.5. Tectonic setting ................................................................................................. 16 II.5.1. Intracratonal basins .................................................................................. 17 II.5.2. Concurrent fold belt basins ...................................................................... 18 II.5.3. Transcurrent basins .................................................................................. 19 II.6. Tectonic style .................................................................................................... 21 II.7. Taphonomic cycles ........................................................................................... 21 II.7.1. Coal-bearing cyclothems .......................................................................... 23 II.7.2. Variegate cyclothems ............................................................................... 26 II.7.3. Lacustrine cyclothems .............................................................................. 28 II.7.4. Carbonate cyclothems and black shales .................................................. 32

III. PALAEOECOLOGY ............................................................................................. 37 III.1. Life forms ........................................................................................................ 37 III.1.1. Multifunctionality ..................................................................................... 37 III.1.2. Life form inference: examples ............................................................... 42 III.1.3. Direct and indirect environmental correlates ......................................... 47 III.2. Plant communities ........................................................................................... 50 III.2.1. Community records ................................................................................. 51 III.2.2. Palaeosyntaxonomy ................................................................................ 59

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III.3. Basinal systems ............................................................................................... 61 III.3.1. Seral systems .......................................................................................... 63 III.3.2. Palaeocatena, general ............................................................................. 65 III.3.3. Plaeocatenas of regressive sequences ................................................... 67 III.3.4. Facial modifications ................................................................................. 72 III.3.5. Catenic systems over time and geography............................................. 73 III.3.6. Evolution over catenic systems .............................................................. 75

IV. PALAEOECOGEOGRAPHY ............................................................................... 79 IV.1. Biomes ............................................................................................................. 79 IV.2. Biome evolution ............................................................................................... 82 IV.2.1. Tropical rainforests .................................................................................. 82 IV.2.2. Temperate deciduous vegetation ............................................................ 84 IV.2.3. Conifer forests ......................................................................................... 87 IV.2.4. Xeromorphic vegetation .......................................................................... 90 IV.2.5. Grasslands ............................................................................................... 94 IV.2.6. Wetlands .................................................................................................. 95 IV.2.6. Aquatic tracheophytes .......................................................................... 103 IV.3. Biome chorology ............................................................................................ 105 IV.3.1. Permian biomes ..................................................................................... 107 IV.3.2. Mesozoic ecogeographic differentiation ............................................... 114 IV.3.3. Mid-Cretaceous biomes ........................................................................ 119 IV.3.4. Tertiary precursors of extant biomes .................................................... 120

V. TECTONIC FACTORS OF GLOBAL CHANGE .............................................. 123 V.1 Current situation in the geotectonics ............................................................... 123 V.2. Rotaion tectonics ............................................................................................. 125 V.3. Geoid ................................................................................................................ 127 V.4. Rotational stresses .......................................................................................... 128 V.5. Planetary fault system .................................................................................... 129 V.5.1. Mid-ocean rises and transcurrent share ...................................................... 133 V.5.2 Cretaceous O/D faulting and magmatic activity ..................................... 135 V.6. Mobile belts ..................................................................................................... 137

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V.6.1. Pacific arcs ............................................................................................. 138 V.6.2. Trenches and ophiolites .......................................................................... 139 V.6.3. Cretaceous circum-Pacific structures ................................................... 141 V.6.4. Cretaceous peri-Atlantic and peri-Arctic belts ........................................... 150 V.6.5. Cretaceous Tethys .................................................................................. 157 V.7. Tectonomagmatic phases ................................................................................ 162 V.8. Summary of tectonic factors .......................................................................... 163 V.9. Tectonic inferences from taphonomy and palaeoecology ............................. 165 V.9.1. Tectonics and vegetation of continental margins ................................... 166 V.9.2. Land connections and gateways ............................................................ 167 V.9.3.Palaeoecological constrains of continental rearrangements ................... 169 V.9.4. Permian landmasses and phytogeography ............................................. 170 V.9.5. Trans-Tethyan biotic links and collision of India to mainland Asia .......................................................................................... 173

VI. SEA-LEVEL FLUCTUATIONS ......................................................................... 177 VI.1. Sea-level models ........................................................................................... 177 VI.2. Hypsometric curve ........................................................................................ 180 VI.3. Tilt of the continents ...................................................................................... 184

VII. CLIMATE CHANGE .......................................................................................... 187 VII.1. Proxies ......................................................................................................... 187 VII.1.1. Lithologies ............................................................................................ 187 VII.1.2. Isotope ratios ....................................................................................... 193 VII.1.3. Fossil plants ......................................................................................... 194 VII.2. Causes and effects ...................................................................................... 207 VII.2.1. Orbital forcing ...................................................................................... 208 VII.2.2. Orogenic effects .................................................................................. 210 VII.2.3. Oceanic effects ................................................................................... 214 VII.2.4. Greenhouse effect ............................................................................... 218 VII.2.5. Stratospheric effects ........................................................................... 232 VII.2.6.Vegetation effects ................................................................................. 234 VII.3. Climate trends .............................................................................................. 237

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VII.4. Displacements of zonal boundaries as evidence of long-term climate change .............................................................................................................. 243 VII.4.1. Deciduous/evergreen zonal boundary through time ........................... 245 VII.4.2. Climatic interpretation ......................................................................... 246 VII.5. Early Tertiary climates and afforestation–deforestation cycles ................. 249 VII.6. Mid-latitude Tertiary cyclicity in eastern Asia ............................................ 254 VII.6.1. Succession of plant assemblages ........................................................ 255 VII.6.2. Four-phase climate change ................................................................. 262 VII.7. Summary of climate change ........................................................................ 265

VIII. ECOSYSTEM EVOLUTION ........................................................................... 267 VIII.1. Biotic communities as evolution units ........................................................ 269 VIII.1.1. Anticipation ........................................................................................ 269 VIII.1.2. Co-adaptation ..................................................................................... 271 VIII.1.3. Herbivory and plant evolution ............................................................ 273 VIII.1.4. Coenotic gene pool ............................................................................. 279 VIII.2. Biomass ...................................................................................................... 282 VIII.2.1. Biomass build-up and fluctuations ..................................................... 283 VIII.2.2. Biomass redistribution ........................................................................ 285 VIII.2.3. Assessment of Cretaceous terrestrial biomass ................................. 286 VIII.3. Diversity ..................................................................................................... 292 VIII.3.1. Diversity estimates ............................................................................. 292 VIII.3.2. Species richness and disparity ........................................................... 296 VIII.3.3. Arogenic communities ....................................................................... 298 VIII.3.4. Diversity and environmental stability ................................................. 304 VIII.3.5. Niche creating trend .......................................................................... 308 VIII.3.6. Diversity and niche overlap ............................................................... 309 VIII.3.7. Diversity and redundant density ........................................................ 311 VIII.3.8. Diversity evolution.............................................................................. 313 VIII.4. Sere as reiteration of ecosystem evolution ................................................ 316 VIII.5. Speciation cycles ........................................................................................ 317 VIII.6. Genetic assimilation .................................................................................... 320 VIII.7. Examples of palynomorphological evolution .............................................. 327

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IX. CRISES .................................................................................................................. 339 IX.1. Transboundary trends .................................................................................... 339 IX.2. Boundary events ............................................................................................ 341 IX.3. Causal model of the boundary events ........................................................... 349 IX.4. Mass mortality and mass extinction .............................................................. 359 IX.5. Replacement of ecological dominants .......................................................... 360 IX.6. Climax cut-off ................................................................................................ 363 IX.7. Post-crisis disparity ....................................................................................... 365 IX.8. Recovery ....................................................................................................... 368 IX.9. Example of plant community evolution at the KTB ..................................... 369 IX.9.1. Late Cretaceous assemblages .............................................................. 369 IX.9.2. Early Palaeocene assemblages. ........................................................... 371 IX.9.3. Vegetation change ................................................................................. 378 IX.9.3. Macropolymorphism and morphological innovation ............................. 381

X. CONCLUSIONS .................................................................................................... 387 INDEX ......................................................................................................................... 393 REFERENCES ............................................................................................................ 399

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I. INTRODUCTION Clues to the future are hidden in the past, but how deep in the past? Only a small fraction of palaeontological data is actually taken into consideration in respect to the ongoing environmental change. It is commonly assumed that global change, or most of it, is man-made while natural processes are too slow for our immediate concern. Only recent geological history, from the last glaciation on, enters environmental predictions by way of extrapolation. Yet a search for the clues might require a deeper digging in the past. We scarcely know how much of the recent environmental change is man-made. The geologically recorded global changes are greater than any present-day human activities can realistically inflict. Rather than gradually accumulating over millions of years, such prominent changes as at the Cretaceous/Tertiary or the Permian/Ttriassic boundaries might as well be rapid switches from one long-lasting standstill to another. The last 20 k.y. may be not a better guide to the future than 65 Ma or even 230 Ma. Moreover, when many variables are involved, extrapolation is the least efficient, if at all applicable, method of prediction. Rather we need an explanatory model based on a comparative analysis of the geologically recorded global change. The relevant fields of geology, ecology and evolutionary biology are presently governed by the paradigms of plate tectonics, greenhouse warming, and neo-Darwinian (“synthetic”) evolution theory, each of which has already surpassed the period of its utility to science. Their supporters claim no viable alternatives in view. In fact they thoroughly suppress, rather than seek, any possible alternatives. Prefaces often claim little consensus on a subject to be the incentive of a study. In the present study, it is more often too complete a consensus that gives the incentive. This book develops a concept of global change through ecosystem evolution, various aspects of which are discussed in my previous works (Krassilov, 1972a, 1976b, 1978b, 1992a, 1995a, 1997a, 1998a, etc.). Case studies in terrestrial palaeoecology have produced an enormous literature that is only sketchily reviewed for the purposes of the book. Acknowledgements. I thankfully acknowledge the generous help of my colleagues, my family and the editors of Pensoft Publishers.

I.1. The idea of global change Global change has recently become a common concern after a history of coming in and out of vogue. The Biblical Flood of about 4 k.y. BC should have been the last global change: “the waters shall never again become a flood to destroy all flesh” (Genesis, 9-15). Even an irrational notion of ongoing global change as solely manmade might have been tacitly inspired by the Biblical testimony of the post-diluvial

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covenant that warranted the survival of every living creature of flesh that is upon the earth. In the early geological thinking, marine fossils on land were brought by flood. In XVIII c. diluvianism was extended to explain the marine/non-marine alternations and their related biotic changes – the Cuvierian “revolutions” (“catastrophes” of the opponent’s writings). These ideas have formed the basic premises of biostratigraphy, the use of fossils for sequencing and correlation of sedimentary rocks. Later, the supposed diluvial deposits of northern Europe were reinterpreted as traces of a widespread terrestrial glaciation. Polar expeditions brought to light the older fossil floras with palms and fig trees (Heer, 1866-1882). Consequently, climatic change, abrupt as it might have been over the recent geological history, was recognized as the most significant global change. A competing school of natural philosophy adhering to the natura non facit saltum principle (epitomized in Leibnitz’s calculus) denied abruptness in favour of a gradual step by step progression that involved both the material and spiritual worlds, with fossils as an intermediate step between the living and non-living matter. The progressist ideas, shared by such influential figures as Voltaire, Turgot, Condorcet, Geoffroy Saint-Hilaire, Goethe, Erasmus Darwin and many of their contemporaries, formed a conceptual basis of the Enlightenment. The theory of biological evolution appeared as an outcome of the progressist philosophy. Biologists, such as Lamarck, only sought to provide a factual basis for it. Charles Darwin argued for a gradual cumulative advance of adaptive traits directed by natural selection. Such gradualistic theories were ever haunted by the host of Cuvier, drawing strength from the fossil record. Hence Lyell and Darwin attempted to discredit the fossil record, the notoriuos imperfection of which nourishes an illusion of abrupt worldwide changes. Although Darwin (1859) mentioned the possibility of evolution being enhanced by an occasional sea-level fluctuation, he more systematically tried to reduce the driving forces to competitive interactions alone. “Occasional” is a favourite word of the “Origin of Species” and the epigonal literature. Any geological change, whether of evolutionary importance or not, was certainly occasional. The recent revival of catastrophism is an outcome of globalization, in particular, of new technologies for planetary observations. Synchronous, globally interrelated changes are confirmed by anomalous sea surface temperatures of the equatorial Pacific bringing droughts to the Sahel and cold winters to Alaska. It is becoming increasingly clear that systemic, rather than reductionist, models are needed for predicting global change.

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I.2. Global change: general In complex systems, an external forcing is either damped by negative feedbacks or escalated by positive feedbacks. A switch from one mode to another makes evolution of such systems intermittent rather than gradual, hence the inconsistency of gradualistic theories. The system we deal with in global change research is the biosphere as defined by Vernadsky (1967 and earlier work), i.e. the shell of the earth encompassing all living matter and the products of metabolism, including those buried in sedimentary rocks. This is an open system perpetually reproducing organic life by recycling matter, with the main source of energy outside the planet. The biospheric functions are directed by the goal of biological production with minimal wastes. The biospheric structures are determined by their functions in the life-producing system. We are concerned with structural changes (such as the loss of biological diversity) but these are on the surface, conveying a deeper level of functional change. Neither structure nor function can be fully understood without unravelling the history. This is what palaeontology is about. The biosphere is superimposed upon the earth crust, hydrosphere and atmosphere. Heterogeneity of these substrates predetermines division of the biosphere into a great number of life-producing units, ecosystems, with their specific biotas. The ecosystems are linked by a perpetual flow of matter, through the continental runoff to the oceans and a backflow of oceanic production over trophic channels, providing for functional integration of the biosphere. An impact affecting any of the biospheric functions would inflict a reaction spreading over the mosaic of interrelated ecosystems. Evolution of functional systems is governed by the general system laws, which, for closed systems, are the classical thermodynamic laws. Closed systems evolve in the direction of a maximal entropy production. Conversely, an open system minimizes its entropy production (Bertalanfy, 1960; Prigogine, 1980; Kondepude & Prigogine, 1999) by drawing matter and energy from external sources and by exporting wastes. Each open system thus affects the neighbouring systems and is in turn affected by them. Any talk of biotic evolution going on independently of any environmental change is therefore meaningless. A history of the biosphere is the outcome of complex interactions with cosmic, geological and, recently, mental systems leaving their traces in the fossil record. The Darwinian metaphor of the geological record as a book written in an unknown language, with most pages torn away and the remaining showing a few lines each, is a frustrating reminder of the fact that the still popular theory of evolution was built on a notoriously unsafe theoretical, as well as factual, foundation. Not only is the record more complete presently than in the Darwinian time, but also room for speculations is narrowed by the modern methods of measuring evolution rates. The once popular notion

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that an abrupt change must be an artifact of the fossil record does no longer seem persuasive. Moreover, we need not any longer allow the prejudice of gradualistic philosophy to undermine our credence in the fossil record. To take a familiar example, the dinosaur-dominated terrestrial communities persisted, with minor overturns, over 160 m.y. They were replaced by the mammalian communities at about the Cretaceous/Tertiary boundary, about 65 Ma. Shortly before this date, the Mesozoic mammals, although taxonomically diverse, were small pioneering forms controlling a restricted range of trophic resources, mainly of the frugivorous and insectivorous niches (Lillegraven et al., 1979), scarcely overlapping any of the long established dinosaurian niches, hence no ground for a competitive exclusion. The replacement roughly coincided with a global regression of epeiric seas, a conspicuous climate change, the Deccan traps, and a metallic pollution (including an iridium spike). Isotopic anomalies reflect a major change in global biomass/dead mass production. An adequate structural change is the loss of biological diversity in both terrestrial and marine biotas. For the Cuverians such chronological coincidences meant that all plants and animals, or most of them, were suddenly flooded, frozen, poisoned or suffocated by the rising sea, adverse climates, volcanic eruptions or extraterrestrial impacts. Proponents of these views confuse mass extinction, a biospheric event, with mass mortality, a population event. Their simplistic explanations can only be sustained at the expense of a greater part of geological evidence. For the Darwinians, geological evidence is largely meaningless or irrelevant. Both fail to recognize ecosystem as the structural level at which global changes are most ostensibly manifested. Rather than being a cumulative result of occasional extinctions or occasional speciation events, evolution at the population, organismic and genomic levels is driven by ecosystem responses to global change. This evolutionary concept is developed in the following chapters.

I.3. Driving forces: state of the art Because of the complexity of global changes, a mere extrapolation of what is going on would scarcely tell us what to expect. To become predictable, global changes have to be explained. The explanations are drawn from planetology, geotectonics, climatology and evolutionary biology. However, the present state of the art is such that none of these disciplines is capable of a substantial contribution to the theory of biospheric evolution. In the late 1960’s, the advent of new global (plate) tectonics (Isaks et al., 1968) put an end to a long period of stagnation in earth sciences when all planetary models, including the briefly popular Wegenerian drift, seemed discredited. Regional geological research had been carried out with no reference to global situations. Younger generations might

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indeed find such strict regionalism depressing. Anyway, the brave new tectonics managed to rapidly overcome the initial scepticism that was soon replaced by a boundless enthusiastic support. The believers hastened to proclaim it the only scientific global model that explains everything worth explaining. Actually the model only ventured to explain what happened, with no reference to how and why. It therefore soon degraded to a body of ad hoc explanations linked by special terminology alone. In any regional reconstruction there are plates to be rifted, rafted, collided, subducted and obducted, with a ridge/trench system assumed for each particular episode. A global perspective has been lost once again in the multitude of case studies. As it is, plate tectonics is not a predictive theory. The Popperian principle of testing theories by attempts at refutation rather than confirmation (Popper, 1972) has been rarely if at all applied to the plate tectonics. Yet, as the recent experience has repeatedly asserted, even the thoroughly refuted theories can be kept afloat by a series of first-aid modifications, in which the original idea is either lost or made totally meaningless. As noted in (I.1), the original idea of global change has been prompted by the finds of marine fossils on land and even now the geological record is most convincing in the case of sea-level changes. Their plate tectonic explanation as the effect of an accelerated sea-floor spreading (elevated spreading axes) is only half-heartedly accepted even by the proponents for want of anything better. The consequences of sea-level changes for the climate and biota are commonly believed to be significant, yet the specific mechanisms of interaction are but dimly outlined. There is little doubt that climate affects both geological and biological processes and is in turn affected by them. Climatic events may then correlate with biotic turnovers, evolutionary innovations, growths and declines of civilizations, a possibility tackled by generations of palaeontologists, archaeologists, ethnologists, etc. However, a comprehensive theory of climate change is wanting. There are several theories, the orbital, geographical, CO2, etc., illuminating a few features rather than the whole picture. The choice of a leading force (CO2 at the moment) is arbitrary, mostly dictated by nonscientific considerations. Of all geological evidence, the records of climate change are the least direct. The alleged proxies are ambiguous: different combinations of climatic variables may leave similar isotopic, chemical or biotic signatures. A palaeoclimatic inference thus amounts to disambiguating the proxies for which no methodology exists so far. No wonder then that what we seem to firmly know about global climate change is not much. As for the biological aspects, it should be reminded that current evolutionary theory appeared as an organismic paradigm (Lamarck). It was then reformulated as a populational paradigm (Darwin and population genetics). Recent contributions came from molecular biology. The ecosystem level, though of primary importantce for the purposes of deciphering and predicting global change, yet remains the least explored.

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Hopefully, the period of stagnation will not be a long-lasting one. The disciplines that contribute to the conceptual foundation of the global change research may jointly overcome it by shearing a basic understanding of biosphere evolution as a multilevel top– bottom regulated process with the goal of efficient biological production, to which all the components are subordinated.

Chapter 2. Taphonomy

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II. TAPHONOMY Althouagh aware of biotic change over time, we still turn to the present-day biota for guidance in our inference from the fossil record. But our understanding of the presentday biota is also an inference from a limited set of sampling data. Likewise, fossil assemblages can be seen as samples of past biotas, the sampling “method” defined by taphonomy. While dealing with fossil material, the palaeontologist is inevitably confronted with the problems concerning transition from living to becoming part of the rock. The leading Enlightenment philosophers, such as Voltaire and Robinet, considered fossils as byproducts of nature’s creative force (William Smith’s biostratigraphy was founded on this assumption). Even now the process of fossilization is not fully understood. But with time our interest in fossilization has shifted from what fossils are to how they relate to organisms that once lived on earth. The relevant observations gradually accumulated in the course of palaeontological research until they emerged as a special geobiological discipline that can tell us something of both the former habitats and sedimentary environments (Efremov, 1950). The significance of taphonomy is not yet fully appreciated. Whether we trust the fossil record as an informative sample of past life or whether it is distorted out of credibility by a thin chance of preservation is a basic dilemma of evolutionary theory. Early catastrophic ideas of global change were based on a credulous attitude to the fossil record. In contrast, in the Darwinian tradition, any global change is considered as an artifact of imperfect recording. Needles to say, taphonomy is, or should be, a central issue of such controversies.

II.1. Dead mass fossilization as a condition of life Fossils were once perceived as bizarre creations of underground forces, and even modern palaeontology still tends to consider preservation of organic remains as accidental. Yet already in the 1920’s Vernadsky put forth the idea of a biogeochemical cycle, in which fossilization of organic carbon in one form or another is necessary for the maintenance of aerobic life. The fossil record thus conveys the functioning of the biosphere as a life-producing machine. All life forms of a biotic community contribute to dead mass production – if not body/trace fossils then a dispersed organic matter or geochemical, as well as isotopic, signatures (in the case of Archaea: Kuypers et al., 2001; Orphan et al., 2001). The fossil record is therefore a production of ecosystem rather than of individual species. Except the relatively rare remains of standing biomass (“fossil forests” of flooded tree stands), fossil plant material is a dead mass either of local origin (autochthonous/hypoautochthonous) or transported (allochthonous), preserved as coal, fossil soil (palaeosol), leaf

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Valentin A. Krassilov. Terrestrial Palaeoecology

Fig. 1. Dead mass export from terrestrial ecosystems to the atmosphere, continental traps and the ocean.

mats, spore-pollen assemblages or structureless particles. All ecosystems produce dead mass that is, in variable proportions, recycled, accumulated in a not readily accessible form and exported (Figs.1, 2). It is well-known that terrestrial ecosystems differ both in absolute and relative (to their standing biomass) amount of exported dead mass, which correlates with biological diversity and population strategies, in turn depending on environmental stability, climate (temperate ecosystems are, as a rule, more prolific dead mass exporters than tropical ecosystems) and developmental stage (pioneer/seral communities export more dead mass relative to their standing biomass than climax stages). The primaeval terrestrial ecosystems dominated by heterotrophic organisms might have depended not so much on their own primary production (of the cyanophytic–algal films on regolithic rock surfaces) as on the import of organic matter, such as air-borne spores of semiaquatic plants that appeared in the mid-Palaeozoic fossil record before the land plant macrofossils. The first traces of pedogenesis (known since the Ordovician or even earlier: Retallack et al., 1984; Retallack, 1986a) might be due to imported biomass and heterotrophic production. The role of primary production increased with land plant evolution, making terrestrial ecosystems the major exporters of organic matter,


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Fig.2. Organic-rich deposits: (a, b) mudflats of the North Sea coast; (c) lacustrine black shales, a non-marine oil source in the Daladze Formation, northeastern China, (d) huge coal seams in an open quarry, Victoria, Australia (a field excursion during the 3d Conference of the International Organization of Palaeobotany, Melbourne, 1990).

while in marine ecosystems the net biomass production typically exceeds the primary production. The relatively narrow coastal zone of estuarine–deltaic deposition is a major importer of organic matter accumulating most of the terrestrial runoff (Mantoura et al., 1991; Walsh, 1991; Wollast et al., 2001). Its own net production and further export to marine ecosystems varies in response to tectonic, eustatic and climatic factors controlling the total volume and composition of terrestrial input, in particular, the ratio of inorganic to organic nutrients. The estuarine ecosystems are either autotrophic (with the primary production exceeding the net metabolism) or heterotrophic (reverse), which defines their roles as, respectively, exporters or traps of organic matter (Oviatt et al., 1986; Smith & Hollibaugh, 1993; Kemp et al., 1997; Lin et al., 2001). Autotrophic functioning is enhanced by influx of inorganic nutrients exceeding that of organic nutrients, periodically causing phytoplankton blooms (Cosper et al., 1989) and eutrophication as a global phenomenon (Nixon, 1995), but the ratio of the nutrients is regulated by many factors, including the geological substrates (with a larger input of inorganic nutrients by rivers draining volcanic plateaux), vegetation, pedogenesis, disso-

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ciation of organic compounds on land or in the coastal wetlands, etc. The amount of buried organic carbon varies with continental runoff and the influx of nutrients that inflame marine productivity. A shifting balance of autotrophic/heterotrophic productivity is reflected in the relative contribution of herbivory versus detritivory to the food webs. The Palaeozoic arthropod communities were essentially detritivorous, a situation that gradually changed over the Mesozoic (Scott & Paterson, 1984; Llabandeira et al., 1988). However much is yet to be learned of the herbivore/detritivore relationships from the fossil record (VIII.1.3). A special case of imported dead mass utilization by aquatic organisms is a selection of construction material by caddis–flies mostly building their cases from mineral particles or ostracod shells but switching to plant debris with an increase of organic terrestrial runoff. A regular input of certain plant structures, e.g. of megaspores of aquatic lycopsids and ferns or water-borne seeds of wetland plants encourages the caddis-flies to acquire new construction techniques. Hence, when the caddis-fly cases are abundant, as in the Early Cretaceous lacustrine facies of Central Asia, their construction is a source of evidence of ecosystem functioning. Incidentally, some caddis-flies not only incorporated megaspores and small seeds in the tube walls but they also attached the larger ginkgophyte seeds (Karkenia) as floats (Fig. 3). On such evidence, the narrow-leaved karkenias were common waterside plants regularly shedding their seed crop to the water (Krassilov & Sukatcheva, 1979). The bulk of plant dead mass consists of shed organs, such as leaves, leafy shoots, pollen grains or propagules. Plant organs are abscissed for many reasons, including the maintenance of the optimal photosyntethic rates for tree crowns and canopies, developmental replacement, seasonality, as depots of metabolic wastes (accumulating in leaves), for reproduction and dispersal. How much of the dead mass is engaged in recycling, storage and export depends on a vegetation type. Storage of nutrients in dead mass is typical of tundras and deserts, whereas in broadleaved forests a larger part of the leaffall input is exported (Woodwell et al., 1978; Woodwell, 1984; Bolin, 1987). Pollen grains and disseminules are intended either for intracommunal reproduction or (in species that scarcely sprout under their own canopy and those that are preferentially pollinated by distant pollen sources) for export. Since an influx of dead mass is rarely fully reciprocated, the roles of exporters/importers are evident even in the case of interacting aquatic ecosystems. A rather typical example is provided by a mass accumulation of aquatic plants in the Late Cretaceous of Nemegt Formation, Mongolia (Krassilov & Makulbekov, 1995; Fig. 4). The fine-grained lacustrine shales contain Potamogeton pondweed shoots and an aquatic quilwart-like lycopod, Monilitheca, indicating clear oligotrophic water. The limnaceous seeds found with them represent a quite different type of aquatic community, eutrophic, with floating mats. Floating plants are rarely preserved in situ, unless they are buried by a sudden flux of alluvium. Yet their disseminules, caught by avulsing streems or high floods, are dispersed over the river basin. In the Nemegt example, the duckweed-like seeds are the


5

Fig.3. Caddis-fly cases made of small plant debris, with the larger seeds of Karkenia, a stream-side ginkgophyte, attached as buoys. The Early Cretaceous lacustrine facies of Mongolia (Krassilov & Sukatcheva, 1979).

commonest fossils in detrital levee sandstones indicating high floods. Their input in the oligotrophic facies suggests an abundant export of organic material from eutrophic ecosystems, the lavish aquatic dead mass producers. On a larger scale, Tympanicysta spike in the marine and non-marine Permian/Triassic transboundary sequences represents a dead mass produced by algal blooms and exported with eutrophic water spills from estuaries over carbonate platforms (Krassilov et al., 1999a, 2000). The significance of Tympanicysta spike as evidence of eutrophication is further discussed in IX.2. Highland ecosystems are exporters owing to their geomorphological position, but their impact on lowland ecosystems varies with relief, climate and vegetation. Lack of a definable upland material reflects either sparsely vegetated uplands, as in the Devonian, or an insignificant source area relative to the catchment area, as in the Early Carboniferous (the scleromorphic features of Carboniferous cordaites and peltasperms probably reflect helophytic, rather than xerophytic, adaptations; their facial occurrences may indicate a zonation of the mires rather than an upland source, see Gastaldo, 1989).

6


a

b

c

d Fig.4. The Late Cretaceous (Maastrichtian) aquatic assemblages of Mongolia (Krassilov & Makulbekov, 1996): (a) a seed compression of an extinct lemnaceous plant, supposedly from an eutrophic pond system with floating mats, abundantly represented as allochthonous fossils in a wide range of flood-plain and lacustrine facies, such as (b) clayey oxbow facies with an aquatic monocot, Potamogetonites, (c) the oligotrophic lacustrine facies with an extinct aquatic lycopsid, Monilitheca, represented by a giant megasporangium, and (d) laminated levee siltstones.


7

Upland runoff increases with a drop of sea level (depressing the basis of erosion) and with seasonality of precipitation causing high floods. A significant upland contribution to fossil plant assemblages is first recorded in association with the Late Palaeozoic Hercynian orogeny. Thick allochthonous coals are confined to the post-orogenic molassoid facies where peat accumulation was enhanced by rainwater runoff over vegetated slopes, bearing low sedimentary load and sustaining a permanently high water table in the foredeeps. Although some dead mass is exported from oceanic ecosystems to land (in the dung accumulations of sea birds, stranded algal mats, etc.), the terrestrial export is overwhelmingly larger. On balance, continents are exporters but these relationships might have been different before the advent of higher land plants or even for some time after that. Terrestrial vegetation is not only the major exporter of dead mass but it also controls transport (of different rates over vegetated or barren landscapes for both wind-borne and water-borne material) and deposition. Terrestrial organic material is scarcely incorporated in near-shore deposits unless it is trapped by littoral vegetation that reduces wave action. Since the Late Silurian on, littoral vegetation plays a key role in dead mass transfer from land to sea. Owing to its wide distribution and rapid recycling, terrestrial organic matter is a significant factor of oceanic productivity and sedimentation (De Leeuw et al., 1995; Opsahl & Bonner, 1997). Thus, lignine is found throughout the water column of both the Atlantic and Pacific oceans, its concentration correlates with riverine discharge, hence higher in the Atlantic. Palynological studies reveal a correlation of palynomorphs and structureless kerogens with oil/gas accumulation. In particular, the Botryococcus-type unicells and amorphous kerogens occur in oil source deposits, whereas terrestrial palynomorphs and “fibrous” kerogens associate with gas sources (Batten, 1982 and elsewhere). Differences in kerogen microstructures are conventionally ascribed to their marine or terrestrial origins. However, most kerogens are produced at river mouths that export many times more dead mass than is produced in the oceans. The fibrous “land plant” kerogens originate from organic debris of deltaic facies while the amorphous “algal” kerogens are estuarine production (algal blooms in hypertrophic estuarine ecosystems are the most abundant source of organic matter: Alimov, 1989). Thick deltaic deposits typically correlate with regressions, whereas estuarine contributions increase with ponding of river mouths by rising sea. In tectonically active marginal zones (foredeeps of coastal ranges, fore-arc depressions), terrestrial dead mass brought in with continental runoff is redeposited downslope by turbid flows inflicting short-term anoxic conditions and fluctuations of carbonate dissolution depths. The dead mass is intermittently deposited in organic-rich bands of turbiditic flysch sequences, sometimes containing structurally preserved plant fossils. Periodic influxes of suspended organic material enhance precipitation/diagenetic redistribution of silica over the banded silicites (cherts). The radiolarian cherts of ophiolite assemblages sometimes contain fossil wood (Abbate et al., 1980) and other ter-

8


restrial remains (also shallow-water benthic remains indicating a quasipelagic, rather than pelagic, origin). Fluctuations of carbonate dissolution depths and the concomitant divergence of calcite/aragonite lysoclines are reflected in a selective preservation of carbonate fossils (an alternation of ferruginous cephalopod limestones, amonitico rosso, and Aptichus facies with calcitic cephalopod mandibles in the absence of aragonitic shells: Bosellini & Winterer, 1975 and a Cretaceous example in II.7.4). Continental runoff stimulates oceanic dead mass deposition in the form of organicrich black shales marking widespread oceanic anoxic events (OAE). The well-known Mesozoic OAEs correlate with transgressions, global warming, positive δ13Ccarb anomalies, lacustrine black shale events (in the Mongolo-Okhotsk rift system: II.7.3) and thick coal deposition (the Aptian and Late Albian to Cenomanian coals of eastern Asia, Alaska, Fort St. John Group of western Canada, etc., reviewed in Krassilov, 1985, 1992b). The latter is evidence of a high rate terrestrial dead mass production that depletes 12C content of atmospheric carbon and an equilibrated carbonate reservoir. Prominent δ13C fluctuations, mostly negative, correlate with prominent orogenic events, high rate sea-level fluctuations and climate change (VII.1.2). This is what to be expected since both the ecological and taphonomic variables of dead mass production, export and deposition are affected by disruption of terrestrial ecosystems.

II.2. Self-fossilization and taphonomic enrichment Whatever the source of organic material, it is not just passively buried and fossilized but constitutes instead an active component of taphonomic environment. Thus, microbial mats are self-fossilizing in the sense that they lithify their own production. The rhizosphere – a community of underground plant organs, their fungal and animal symbionts – is self-fossilized by pedogenic mineralization. Submerged parts of wetland plants are protected from poisonous substances produced by anaerobic microorganisms by oxygen emitted from their aerenchyma to the anoxic environment. Such local oxydation induces precipitation of ferric oxydes as casts on roots and buried stems, hence ferruginous “fossil forests” (such as described in Nessov & Golovneva, 1995). Conversely, a reduction gradient over decaying plant material induces transport of silica, ferric compounds, carbonates and other fossilizing substances accumulating in the organic-rich zone. Coal-balls are familiar example of fine preservation owing to impregnation of cell walls by carbonates migrating from saline soil solutions to the rheotrophic peat. Fungal degradation is a common feature of coal-ball material (Stubblefield & Taylor, 1988), perhaps facilitating the mineralization by splitting cell walls. Bedding planes in organic-rich deposits are covered with electronically distinct microbial films that coat sedimentary grains (Noffke et al., 2001). The so-called plant impressions are, in fact, microbial films coloured with ferric precipitates mixed with


9

Fig. 5. Subcrustation, a preservation form commonly described as impression but actually resulting from deposition of a mineral film under the cuticle: a leaf “impression” of Ushia, a nothofagoid angiosperm from the Palaeocene of Cisuralia, on a coarse-grained polymictic sandstone.

residual organic matter. Such films are deposited on the surface of plant remains (incrustations) or subcutaneously (subcrustations: Krassilov & Makulbekov, 1996; Figs. 5, 6). Migrations of ions from organic-poor coarse-grained to organic-rich fine-grained components of sand/silt couplets (Awwiller, 1993) is as common as redistribution of lime over limestone/marl couplets (Einsele, 1982). Binding of ions by organic remains varies with the latter’s chemical composition. In a decaying leaf, the relief of mineral incrustation is defined by differentiation of photosyntetic and mechanical tissues over the pattern of costal/intercostal zones. A similar correlation of anatomical differences and mineralization is described for soft-bodied animals (Orr et al., 1998). Organic remains facilitate preservation of other organic remains by serving as foci for a concentric accretion of silicates, sulfides or carbonates. Terrestrial tetrapod remains from redbeds are embedded in nodular casts, the elementary composition of which differs from that of the surrounding rock (Rogers et al., 2001). In marine deposits, the fully permineralized megaspores, cones and even small flowers are often preserved in syderitic concretions with an ammonoid shell as the core (examples in Krassilov & Zakharov, 1975; Nishida, 1985). Even trace fossils, such as burrows, are preserved three-dimensionally owing to their trapped organic debris that enhance fossilization. In sediments enriched with buried organic material, fossilization is thus a self-accelerating process involving structures that otherwise escape preservation. The origins of the richest fossil plant/animal localities seem related to the phenomenon of taphonomic enrichment. Thus, fossils are not so much a gift of chance preservation as a normal product of interaction between dead mass and taphonomic environment. Freshly buried organic remains modify their taphonomic environment in such a way as to enhance their own fossilization. Structural preservation is owing to migration of ions along the chemical gradients imposed by the buried organic matter that is dehydrated, macerated, impregnated or molded by mineral substances. With the advance of palaeontological techniques, rocks apparently barren of fossils are becoming increasingly rare and, with time, may come to be considered as exceptional.

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c

a

d

b

e

Fig. 6. Subcrustation, a preservation form commonly described as impression but actually resulting from deposition of a mineral film under the cuticle: (a, b) a leaf “impression” of Ushia, a nothofagoid angiosperm from the Palaeocene of Cisuralia, on a coarse-grained polymictic sandstone; (c) SEM of the subcrusting film about 2 ìm thick, intruded with sandstone grains 20-30 ìm across; (d, e) an interior relief of the lower leaf cuticle imprinted on the subcrusting film showing costal cell rows of minor veins, with stomata in the areoles (Krassilov et al., 1996b).


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II.3. Taphonomic bias Taphonomy seeks to reveal the processes by which organic material is transferred from biotic communities to sedimentary rock. Transformation of dead mass into a buried and, eventually, fossil assemblage (the thanato-, tapho- and orictocoenoses of Efremov, 1950) is accompanied by a sequential loss of biological information compensated for by the gain of taphonomic information, which is then engaged in the reverse sequence of palaeoecological reconstruction. Organisms inhabiting sediment, such as aquatic infaunal or terrestrial soil invertebrates, as well as underground parts of vascular plants, are buried alive. Yet sediments are habitable at the time of non- or insignificant deposition while preservation results from a rapid burial that destroys the biotic community. Hence, even in the case of in situ burial, fossil assemblages lack a one-to-one correspondence to their source communities. It is assumed in this kind of analysis, that the major structural and functional aspects of source communities – their life form composition, dominance, productivity, seasonal changes, etc. – are reflected, in one way or another, in the amount and composition of their exported dead mass. Both plants and animals produce detachable structures in the course of growth and reproduction or as metabolic waste deposits. Fecal pellets, shed leaves, dispersed seeds and pollen grains that failed to germinate contribute significantly to the flow and deposition of organic matter greatly enhancing the preservation potentials of their producers. Over their life span, perennial plants may produce a larger dead mass than their standing biomass. For many plant species their contribution to dead mass is roughly proportionate to their significance as biomass producers. High-density populations are the major deadmass contributors. Larger plants shed more leaves, seeds, etc., than smaller plants. Consequently, the dominant trees are, as a rule, represented even in small samples while the taphonomic losses increase with rareness and the plant size: shrubs and herbs tend to be underrepresented relative to trees. However, some regional dominants, such as larch in the larch taiga of East Siberia, beech in the montane forests of Caucasus, maple in the broadleaved forests of Appalachians, etc., are notoriously underrepresented both in the macro- and microfossil assemblages. Even the dominance of Eucalyptus in Tasmania is blurred by the overrepresented waterside shrubs (Hill & Gibson, 1986). Representation is a function of many variables, such as basinal versus extrabasinal habitats, annual versus perennial habits, leaf mass production (crown size), phenology (evergreen/deciduous), leaf size and texture (sinking rates), pollination and dissemination modes (wind or insect pollination, anemochorous or zoochorous propagation), dispersal ranges (long or hort distance dispersers), seral status (pioneer/successional/climax forms), etc. In view of such multiple distortion factors, the species ratios in plant communities and their derived fossil plant assemblages inevitably differ. Yet at least some differences are predictable and analysable (III.3.2). For a basinal system of plant communities, the proximity to deposition site is of primary importance. Commonly, the adjacent (proximal) communities are better represented than the

12


remote (distal) communities (Fig. 7). Aquatic vegetation is an exception to this rule. Moreover, when both macro- and microfossils are recovered, purely autochthonous assemblages are quite rare. Depending on the facies, allochthonous material may come from a nearby source or represent a mixture from different sources, in which the proximal is not necessrily the dominant (Gastaldo, 1989). In addition to the differences in dead mass production rates, such factors as the geomorphology of river valleys, wind directions and the extent of the catchment area (Rhods & Davis, 1995) are crucial for a contribution from remote sources.

a

b

c Fig. 7. Preservation of intact inflorescences indicating proximity of Cretaceous angiosperms to the deposition sites: (a) Hyrcantha, a cymose inflorescence of perfect staminocarpellate flowers, (b) Caspiocarpus, a paniculate inflorescence of small crowded carpellate flowers from the Albian of Kazakhstan, (c) Sarysua, a cymose inflorescence of staminocarpellate flowers and (d) Taldysaya, a fasciculate inflorescence of small carpellate flowers from the Late Cretaceous of Kazakhstan (Krassilov et al., 1983).

d


13

Fossil samples of about 300-400 leaves usually represent all common trees in the sampling area (Krassilov, 1972a; Spicer, 1989a; Burnham et al., 1992). Regional vegetation is thought to have been better represented in the high-energy sedimentary facies while a low-energy deposition reflects local plant communities alone (Spicer & Wolfe, 1987). However, the opposite situation was found in the meandering river complex of Tsagajan Formation, Amur Basin (Krassilov, 1976a). Here the river channel and levee deposits represent almost exclusively the streamside trees or shrubs, such as Trochodendroides. Such fossil plant assemblages are monodominant or even single-species, with a dominant species obviously overrepresented. On the other hand, the oxbow clay lenses contain relatively diverse plant assemblages that are derived from several types of riparian and lower slope communites, including an aquatic community, Limno–Nupharetum, a grass-sedge marsh, Gramino-Carexetum, a swamp forest, Nysso–Taxodietum, a flood-plain broadleaved woodland, Platano-Trochdendroidetum, and a mixed canopy forest with Metasequoia, pines and Cyclocarya from the better drained sites upslope. The modern analogues show a quite similar differentiation of plant communities in respect to their flood tolerances (Gastaldo, 1987, 1989). In marine samples, an input from coastline vegetation prevails in shelf deposits while the mid-slope sedimentation integrates contributions from different sources, with pollen concentrations reflecting wind directions as well as the influx by rivers (Heusser, 1988). Composition of lacustrine palynological assemblages is also affected by the prevailing wind directions (Hill & Gibson, 1986). Since the dead mass production relative to biomass tends to decrease (inversely to structural complexity and efficiency of resource utitlization) over a succession of seral stages, the early stages tend to be overrepresented (Fig. 8). For the same reason, the

Fig. 8. Fallen leaves of maple forming a leaf-mat in pine forest with a sparse maple understorey, Bitza park, Moscow City, Russia.

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fossil record is biased in favour of generalist species that prevail at early seral stages and are more prolific dead mass producers than the relatively specialized species of the later stages. Thus, Czekanowskia is the most common plant fossil in the Jurassic deposits of Siberia (Heer, 1868-1982; Fig. 9). Its facial occurrences over the regressive sequences and the

b c

a

d Fig. 9. Dominant plants of a Late Mesozoic seral system, the Bureya Basin, Russian Far East (Krassilov, 1972c): (a) Phoenicopsis, a lowland climax dominant, (b) Czekanowskia, a pioneer stage dominant; (c) a single-species leaf mat of Czekanowskia with Leptostrobus cupules; (d) a two-species leaf mat of Czekanowskia and Pseudotorellia with a Leptostrobus seed-cone.


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series of coal-bearing cyclothems suggest a pioneer riparian vegetation that was successionally replaced by Phoenicopsis forests, regionally the most productive plant formation. Yet Czekanowskia is locally overabundant, forming coal seams up to 2 m thick. Overrepresentation of early seral stages not only distorts our vision of source vegetation, but also may influence our ideas of evolution rates, since these stages tend to be more conservative. Moreover, insofar as the climatically significant leaf indices are derived from the present-day climax vegetation while seral vegetation commonly differs from it in leaf characters (having either larger or smaller ratios of evergreen to deciduous species for example), overrepresentation of seral stages is a potential source of error in palaeoclimatological inferrences based on leaf morphology.

II.4. Taphonomy and evolution Insofar as sere is a brief reiteration of ecosystem evolution (Krassilov, 1992a, 1994b; VIII.4), taphonomic biases over seral sequences (above) mirror those occurring over geological time. Owing to their large dead mass production relative to their standing biomass and diversity, both early seral stages and early appearing ecosystems are relatively well-preserved. The Precambrian microbial mats seem better recorded than any later appearing assemblages of the kind. Early land plant preservation appears the most perfect of all terrestrial plant records both at individual and communal levels (recent discoveries in the field are mostly reinterpretations rather than finds of new forms while a crop of new forms comes each year from the Cretaceous, and most of the Tertiary floras are still inadequately known). Even the Carboniferous swamp communities are better studied than any of their succeeding wetlands. A search for phylogenetic roots of evolutionary lineages (and phylogenetic tips in the case of extinct lineages) is related to the same taphonomic problems. Both the first and last appearances are often considered as inadequately recorded because of their small population densities. It is claimed at the same time that new lineages typically arise as pioneer forms. The last survivors may also be the pioneer representatives of declining groups. If so, then both the first and last appearances should be fairly recorded. The notorious absence of ancestral forms is then a matter of interpretation rather than recording. Some proangiospermous forms, such as Baisia or Loricanthus, might have been actually overrepresented in the Early Cretaceous marsh assemblages (VIII.3.6). For a species, attenuation of the fossil record may reflect either actual decline or a change of habitat or seral niche and adaptive strategy that goes with it. Adaptive strategies differ in relation to environmental heterogeneities, or “grains” (Levins, 1968). A coarse-grain strategist feels environmental grains that pass unnoticed for a fine-grain strategist. The coarse-grain strategy typically associates with small populations even in a quite prosperous species. Attenuation of the fossil record may bring such species to apparent extinction before actual disappearance. The apparent extinctions tend to be

16


diachronous, mostly predating mass extinction at critical stratigraphic levels (VIII.3.4, IX.4). Conversely, those species that are recorded up to critical levels and disappear in mass extinction are fine-grain strategists that overcome mass mortality by high rate population growth until the ecosystem collapses (VIII.3.7, IX.4). For them, the records of extinction would coincide with actual disappearance. Further evolutionary problems arise in relation to preservation of population norm in the fossil record. At a naïve stage of palaeontological research it was thought that remains found together must be of a single individual. Composite skeletons exhibited in natural history museums are still commonly perceived as portraying perished individuals. However, a reconstruction on individual basis depends on intact finds that are exceedingly rare. Parts of a perennial plant shed during lifetime are usually scattered over a considerable area and are buried in different, even slightly diachronous, deposits. Conspecific remains found side by side more likely belong to different individuals of a source population. A composite plant or skeleton are thus products of population rather than individual. Since taphonomy deals not so much with individuals as with populations, statistical factors come into the picture. Statisitcal probability of entering the fossil record is in favour of common, rather than rare, population characters. A museum reconstruction based on dispersed parts can then be perceived as portraying a population norm rather than individual variation (though distortions due to the age-selective mortality may be significant in the larger mammals, in particular carnivores, including humans). These taphonomic considerations may have some bearing on interpretation of organic evolution, for they explain the notorious absence of transitional character states. Speciation is commonly perceived as gradual in case of large polymorphic systems diverging on the basis of different gene frequencies or it can be more or less abrupt when the founder effect is involved. Whatever the case, the fossil record tends to exaggerate abruptness owing to the normalizing taphonomic effect.

II.5. Tectonic setting Taphonomic processes responsible for import and deposition of organic matter are in various ways affected by tectonic settings of sedimentary basins (Fig. 10). The rates and modes of organic deposition, either in a concentrated form, as coal or bituminous shales, or dispersed are controlled by tectonic processes. The larger concentrations, about 90% of total organic matter, are confined to estuarine and deltaic facies of marginal taphrogenic basins (Mantura et al., 1991; Cosper et al., 1989; Walsh, 1991; Nixon, 1995; Wollast et al., 2001). Tectonic developments and volcanism of marginal fold-belts bear on both the total terrestrial input and the ratio of inorganic to organic nutrients that determines the mode of autotrophic or heterotrophic ecosystem functioning and the amount of organic matter exported to marine environments (II.1).


17

Fig. 10. Intracratonic and marginal sedimentary basins controlled by concurrent and transcurrent fault systems.

Both offshore and lateral, along-shore, spread of nepheloid layer (with suspended particulate matter) depends on slope relief controlling suspension currents and the coupling of pelagic and benthic processes (Wollast et al., 2001). The temperate zone estuaries deposit during a spring freshet up to 0.5 m thick organic-rich sediment at the turbidity maximum zone, the location of which depends both on freshwater outflow and flocculation induced by saline intrusions (Woodruff et al., 2001). Suspended material is conducted along the fault-bound canyons, with major deposition foci at their heads. Hence continental margin geomorphology, defined by tectonic developments, is of primary importance to the biospheric carbon cycle. Active tectonic zones produce a greater diversity of taphonomic environments owing to their heterogeneous landscapes, earthquake-generated debris flows and volcanism. Taphonomic effects of volcanic activity are multiple, such as vegetation burning, leaf fall, in situ burial by ash fall, transportation (occasionally of tree trunks in upright position) by pyroclastic flows, stream ponding, etc. (Spicer, 1989a; Scott, 1990; Archangelsky et al., 1995). Differentiation of positive volcanic features – coastal ranges, island arcs – divided by the fault-bound foredeeps and transverse troughs results in an intricate mosaic of taphonomic environments. In contrast, an epeiric subsidence generates a pattern of few extensive taphonomic zones recurrent over sedimentary sequences.

II.5.1. Intracratonal basins Intracratonal basins are taphrogenic depressions of various shapes and dimensions, sometimes traversing entire continents. They are persistent over times as negative, intermittently subsiding features. Their origins are due either to ongoing rifting or to crustal sagging over the sutures of an older rifting phase. They are rift basins or epirift basins, respectively.

18


A familiar example of the former is the East African rift system with deep stratified lakes, typical of which are high biotic production rates that are compensated for by the equally high rates of organic deposition rendering the water body oligotrophic. The analogous Early Cretaceous lacustrine basins of Transbaikalia, Mongolia and northern China are confined to a transcontinental rift system of the Mongolo-Okhotsk zone extending from Tibet to the Pacific coast (V.5.2). Here the latest Jurassic/Early Cretaceous taphrogenic phase was accompanied by extensive basaltic volcanism. The volcanic plateaux were cut by rift valleys, with the drainage runoff deposited in stratified lakes. Their sedimentary fill of black shales and clastic/carbonate cyclothems contains abundant remains of lacustrine biota, as well as terrestrial debris. These localities are famous owing to proangiosperms, feathered dinosaurs, early birds, etc. (Krassilov. 1982; Krassilov & Bugdaeva, 1999, 2000; Sun et al., 1998; Guo et al., 2000). Black shales provide an industrial source of oil in northern China. Epirift basins are shallow saucer-shaped depressions developed over the buried rift zones. Typical examples are the Late Mesozoic North Sea – northern Europe, West Siberian and North American interior basins (V.5.2, VI). Their underlying rift systems were laid down in the Triassic or earlier (Ziegler, 1980; Surkov et al., 1981; Ocola & Meyer, 1973). Over the Jurassic and Cretaceous the gradually sagging epirift depressions were repeatedly inundated by shallow epeiric seas, accumulating bituminous shales (e.g., the Bazhenov Formation of West Siberia, a major oil source: Ushatinsky, 1981) at the estuarine phase and coal at the following deltaic phase. Terrestrial plant material was abundantly deposited at this latter phase. The richest mid-Cretaceous floras of West Siberia, Kazakhstan and North America came from epirift basins.

II.5.2. Concurrent fold belt basins Fold belts of continental margins typically comprise coastal ranges, back arc basins, island arcs, their foredeeps and the bordering trenches forming a system of concurrent sedimentary/volcanic/metamorphic zones. On the western Pacific margin such appeared in the Permian at least and were fully developed since the mid-Cretaceous and over the Tertiary (V.6.3). Here the terrestrial volcanoclastic deposits of Sikhote Alin, Okhotskian coast and Chukotka were replaced to the east by the paralic to shallow marine coalbearing foredeep deposits over the Tartar Strait and western Sakahlin. The melange belt of accreted island arc/trench deposits extended over eastern Sakhalin, Lesser Kuril Islands and Hidaka zone of Hokkaido. Over the coastal volcanic belt, plant localities occur in intermontane depressions, the larger of which are confined to transcurrent fault zones (below). The foredeep of the coastal ranges accumulated most of the terrestrial runoff, with coal and fossil plant localities in deltaic to proximal marine turbidite facies and with oil/gas sources in the distal turbidite zone. In the melange zone, leaf fossils and well-preserved palynomorphs


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occasionally occur in the turbiditic forearc deposits emplaced upon the pillow basalts, volcanic breccias and agglomerates, the latter containing silicified fossil wood (Krassilov et al., 1988). Similar zonal structures are recognizable in the eastern Pacific and Tethyan fold belts (V.6), with organic-rich deposits accumulating in the fordeeps of coastal ranges and with silicified wood occurring in the melanges (as in the Great Valley foredeep – Franciscan melanges of California, V.6.3). Fossil wood is found also in cherts of typical ophiolite assemblages (Abbate et al., 1980) that are traditionally interpreted as pelagic deposits. Though terrestial plant material is occasionally transported to pelagic deposition sites, its structural preservation requires a far more rapid sedimentation than is typical of pelagic environments. As in the melange zone of the Lesser Kuril Islands, terrestrial plant remains might have been deposited in silicitic turbidites of a forearc basin.

II.5.3. Transcurrent basins Major coastal/offshore deposition loci are confined to transcurrent fault basins cutting across continental margins and incised in the mainland. Many such basins are productive of oil/gas and coal. They are filled with estuarine or deltaic sediments redeposited as offshore turbidites. Basins of this type arise by pull-apart at intersection of transcurrent fault zones with concurrent shear zones of continental margins, marginal fold belts or mid-ocean ridges, or otherwise at stepovers between displaced segments of the latter. Their morphology, subsidence rates and persistence over times depend on the kinematics of intersecting strike-slip fault systems (Barnes et al., 2001 and references therein). The Bay of Biscay – Aquitaine Basin and the opposite Gulf of St. Lawrence – Orphan Basin of North Atlantic margins are typical examples, with median grabens extending as continental fault zones. The nearly evenly spaced coastal basins of the African and South American margins (Figs. 11, 12) are formed by a series of transcurrent faults that were active in the Late Jurassic and Cretaceous (V.6.4). Their sedimentary fill consists of volcanomictic redbeds and black shales overlain by the shallow-water Maastrichtian carbonates (Zambrano & Urien, 1970; Malumian & Baez, 1976; Da Rosa et al., 2000). Such basins persist as negative features over several geological periods and some of them are still expressed geomorphologically as bays or inlets. At the western Pacific margin, a series of transcurrent fault basins, from Anadyr Bay in the north to Posyet Bay in the south (V.6.3), cuts across the coastal range/trough system and projects inland as intermontane depressions. These are filled with Neocomian flyshoids unconformably overlain by coal-bearing paralic deposits replaced inland by volcanomictic redbeds, ignimbrites and lava flows (Krassilov, 1989b).

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Fig. 11. Marginal basins of South America: (Am) Amazon, (Ca) Campos, (Co) Colorado, (Es) Espirito Santo, (Ma) Magallanes, (Mg) Marajo, (Pa) Parnaiba, (Pe) Pelotas, (Pn) Paraña, (Pt) Potiguar, (Re) Recife, (Rt) Reconcavo-Tucano, (Sa) Salado, (Sn) Santos, (Sf) São Francisco, (Sl) São Luis, (Sg) Sergipe, (Sj) San Jorge, (Va) Valdés.

Fig. 12. Marginal basins of Africa: (Be) Benue, (Cb) Cabinda, (Ca) Caoko, (Cu) Cuanza, (Ga) Gabon, (Ld) Lüderitz, (Mo) Mosamedesh, (Se) Senegal, (SL) Sierra Leone. Also shown: (Cm) Cameroon Line, (Ah) Ahhagar and (Ti) Tibesti plateaux. (1 ) Marine and lagoonal deposits, (2 ) Volcanic structures.

A general similarity of transcurrent basins extends also to the margins of the Indian Ocean. The eastern coast of India is cut by the deeply incised Mahanadi, GodavariKrishna, Palar and Cauvery basins (V.5.2). Their central grabens are filled with the Permian to Jurassic coal-bearing deposits disconformably overlain by the marine to paralic Cretaceous, transgressive over the cratonic borders (Sastri et al., 1981). Black shales accumulated in the deeper parts of the basins, fringed by the rifal limestones. On Australian margin, the Fitzroy graben and its parallel median troughs of the Bonaparte, Carnarvon and Canning basins (Mollan et al., 1970; Forman & Wales, 1981) seem to belong to the same system of transcurrent faults as the Indian basins. The Otway, Gippsland and parallel basins of the southern margin are pulled apart at intersection with the Bass Strait shear zone. They contain Cretaceous fossil plant localities of “Wealden” aspect.


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II.6. Tectonic style Insofar as taphonomic environments are defined by tectonism, deposition of organic material reflects tectonic style of the epoch. Thus, at the Jurassic/Cretaceous boundary, most of terrestrial runoff went to the downfaulted marginal basins of the Atlantic, Pacific, Indian and Arctic oceans, with depocentres in the pull-apart trancurrent fault zones. Their thick estuarine to deltaic Wealden-type deposits accumulated organic material preserved as structural fossils or kerogens (Batten, 1982). Rich Early Cretaceous plant localities including the British Wealden flora, the Potomac flora of North American Atlantic coast, the Nican flora of Partisansk Basin, Russian Far East, the Coonwarra flora of Gippsland Basin, southeastern Australia, etc.(Fontaine, 1889; Krassilov 1967a: Douglas 1969; Watson & Sincock. 1992; Watson 2001) occur in the marginal transcurrent basins. Thick organic-rich deposits also accumulated in the concomitantly developing transcontinental rift basins of central-eastern Asia, with lacustrine black shales and associated fossil plant localities of Transbaikalia, Mongolia and northeastern China (II.7.2). Subsidence of continental crust in the Late Cretaceous was accompanied by a nearly four-fold decrease in sedimentation rates. Coal deposition was insignificant in comparison with the Lower Cretaceous reserves while marine accumulations of hydrocarbons were considerably reduced, mainly at the expense of gas source rocks with kerogens of terrestrial origin. Both the marginal transcurrent basins and transcontinental rift zones lost their significance as major traps of terrestrial dead mass. With transgression of epeiric seas over cratonic areas, terrestrial runoff was channelled to the slowly sagging epirift basins. The Late Cretaceous coal and the bulk of fossil plant material came from the marginal clastic facies of epirift basins in northern Siberia, Kazakhstan, northerncentral Europe and North America (reviewed in Vakhrameev, 1991). The mid-Cretaceous is generally considered as the time of high-rate evolution in many terrestrial groups of plants and animals changing the general aspect of land biota. Angiosperms, rare before the Albian, spread over all landmasses forming new types of plant communuties as well as enhancing innovations in the co-evolving groups of arthropods and tetrapods. Yet the abruptness of the mid-Cretaceous biotic turnover is partly owing to the change in tectonic style and taphonomy, with the major part of terrestrial runoff redirected from marginal to intracratonal basins.

II.7. Taphonomic cycles Since fossilization depends on taphonomic variables, such as deposition rates, mineral composition, pH, etc., and since these variables undergo either directional or fluctuant changes in the course of sedimentation, fossils of various dimensions, weight and durability are unevenly distributed over sedimentary sequences. In situ burial of dead mass

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accumulations, such as oyster banks or peat, depend on episodic massive fluxes of sediment, as in pyroclastic flows or tempestites. More commonly, organic remains are transported and deposited as sedimentary particles. Air-borne remains fall as pollen (insect) rain and are incorporated in the soil or peat or further transported by water currents to their destination in the flood-plain, lacustrine or marine deposits. The floating remains settle down with the fine fraction of suspended clastics in the fining-upward sequences or during a pause of clastic deposition as organic films on the bedding planes or as accumulations of allochthonous peat. Megaspores and seeds of a size of coarse sand particles are transported and buried as such. Fossil wood and bone fragments commonly occur in beach conglomerates or debris flow gravels. Those captured by high tides are repeatedly redeposited by cannibalistic waves devouring their own sediment until they become anchored in the tidelites. Such taphonomic assemblages recur over sedimentary sequences as a regular feature of sedimentary cycles. Sedimentation is a biospheric phenomenon reflecting periodicity of global processes governed by rotation and revolution of the planet, with multiple side effects and feedbacks generating a hierarchial cyclicity. Cyclic structures result from both the intrinsic dynamics of sedimentary systems (autocycles) and from external forcing (exocycles). Thus, a fining-upward sequence of fluvial deposits, with a coal bed or fossil soil on top, might have been owing to the autocyclic avulsion, i.e. abandonment of a stream channel in favour of a new course (Allen, 1965, 1970; Stouthamer, 2001). Avulsion of a leeved channel starts with breeching the levees by unchannelled flows that leave poorly-sorted sandy crevasse-splay deposits at the base of the cycle, followed by the newly accumulated well-sorted levee deposits. The bayous and oxbow lakes of the abandoned stream belt are filled with silty flood-plain alluvium and are overgrown by wetland vegetation, the rhizosphere of which reworks the waterlogged silt depositing pedogenic clay with ferruginous rhizocretions. The vegetation is gradually decoupled from ground water supply to develop as a raised bog, or is dried up and calcified and eventually cut by the channel. Fossil soil is a pacer of the process developing at the end of each accumulation phase and truncated by the next erosion phase (Retallack, 1985b). Remarkably, the buid-up of soil profiles is truncated at about the same developmental grade over a long succession of fluvial cyclothems, indicating a nearly equal duration of recurrent pedogenic events. This means that episodic sedimentation is regularly cyclic. The autocycles are overlapped by or included in exocycles of various ranks, of which the decimetre to few metre-scale mesocycles are commonly attributed to the orbital – Milankovitch – cycles. Statistics of the mesocyclites in a variety of facies realms, from continental redbeds to pelagic carbonates, invariably reveals one or another of the Milankovitch periodicities (about 23, 45, 100 k.y. periods of precession, obliquity and eccentricity and the combined periods of ca. 59 and 400 k.y.). The orbital cycles affect sedimentation owing to their climatic (fluctuations of mid-latitude precipitation with precession, of mean annual temperature with eccentricity and of the equatorial–polar temper-


23

ature gradient with obliquity) and tectonic effects. Since those effects are significant for virtually any type of depositional environment, the Milankovitch cycles are ubiquitous. Incidentally, the fluvial autocycles are enhanced by the precession-driven precipitation exocycles periodically increasing the drainage. Both are included in the 10 m-scale erosional exocycles inflicted by a tectonic downfaulting and/or by sea-level fluctuations, with the basis of erosion periodically depressed as sea-level falls (Best, 1997). Unconformities and sedimentological changes – of detrital composition, palaeocurrent directions, diagenesis, etc. – tend to be more prominent over the exocycle than autocycle boundaries (McCarthy et al., 1999; Miall & Arush, 2001). Alternation of basinal facies over the shoaling/deepening sequences is partly autocyclic, caused by decompensation of subsidence by deposition. The system is self-regulated by the negative feedback of decompensation on deposition rates. It is at the same time affected by a variety of eustatic, climatic and biotic factors. Although a leading cyclic process is evident in the case of seasonal varvites, tidal couplets, flysch triplets of seismic fault zones, etc., for many types of cyclic sequences there are arguable alternatives. Thus, the millimetre-scale bands of black shales might reflect either seasonal fluctuations of biotic production or terrestrial runoff or both, superimposed and partly masked by the 4-year El Niño cycles of surface water temperature. The classical carbonate/coal/shale cyclothems are described as glacioeustatic but they might have been also generated by the tectonically driven sea-level fluctuations or climate change (below). Yet the variety of periodic processes might reflect a common pace-setting cyclicity of a more general significance. Cyclic sedimentation is a major source of taphonomic inference. Though fossil assemblages never recur in a strictly identical form, their types tend to recur with sedimentary cycles, such as fluvial autocycles, attesting to persistence of their source communities. Such observations are the best evidence of biotic communities being sustainable structural units. On the other hand, a major change in taphonomic assemblages over a succession of sedimentary cycles indicates an intervention that inflicts a restructuring of source communities. Thus the concerted sedimentary and biotic changes occur over the climate-driven exocycles Cyclicity makes taphonomic observations repeatable, the inference varifiable.

II.7.1. Coal-bearing cyclothems Peat accumulation is a major carbon sink for terrestrial ecosystems controlled by a variety of tectonic, climatic and vegetational factors. Two major types of peat and their derived coals are traditionally designated as “paralic”, of marginal basins, and “limnic”, of intrerior basins. The tectonic settings are molassoid foredeeps of orogenic zones or saucershaped intracratonic depressions over deep-rooted fault zones respectively. A change in coal type with cratonization of orogenic zone is illustrated in Gould & Shibaoke (1980).

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The present-day examples of paralic peat formation are mostly tropical, such as on the Caribbean coast of Panama (Phillips & Bastin, 1990) or in the Orinoco delta, with several types of herbaceous to arboreal peat-producing wetlands, notably the fern marshes and fern-covered mires (Wagner & Pfefferkorn, 1997), possible ecological equivalents of Mesozoic fern-marsh communities. Mangroves occupy 75% of low energy tropical to subtropical coastlines extending to about 30ºN in Japan, Florida and Luisiana. They rank among the most productive tropical ecosystems and prolific dead-mass exporters (Westlake et al., 1998). The mangrove peat consists of roots mostly, with a variable contribution of leaf litter and the turf macroalgae. A model of “phase communities” by Anderson (1983) has been developed in respect to the tropical peat swamps of Malaysia and is presently adopted for other areas of mangrove vegetation. The succession starts with stabilization of beach deposits by mangroves and, with a sea retreat, proceeds over several seral stages to the peatforming subclimax. Mangal growth initially depends on a variety of factors, such as coastal geomorphology, substrate, tidal influence, rainfall, storm frequency, sedimentation rates, etc. (Tomlinson, 1994). Any of this or their combination can determine both seaward and landward limits of mangroves, as well as their zonation and succession. A typical zonation finding their analogy in the fossil record include an Avicennia – Sonneratia growth several hundred meters wide followed by Rhizophora zone, palm swamp, mixed bog forest to sawgrass swamp and a variety of heathy woodlands on sandy terrains (Bird, 1986; Phillips & Bustin, 1990). Initially the zonation is unstable, but with establishment of dense pneumatophore network the mangroves increasingly control sedimentation rates and acquire a long-term stability reflected in a fairly thick (several meters) peat accumulations. The old-growth mangroves are sedimentary traps of slow decomposition rates that store more organic carbon (about 60% of total input) than other woody communities (Alongi et al., 2001). Redox oscillations owing to high tides, seasonal delution of soil pore-water by monsoonal rains, root respiration and bioturbation are among the variables that influence the microbial decomposition rates in the mangrove peatlands. Fluctuations of water table, ground waters, pH and the nutrient levels are reflected in the peat composition and accumulation rates. A retreat of mangroves leaves peat covered with marine sediment. Tropical mangrove peat alternating with shallow platform carbonates is a plausible extant analogue of the Carboniferous paralic coal sequences with prevailingly autochthonous peat accumulations in the flow-fed rheotrophic mires that evolve into the rainwater-fed ombrotrophic bogs with a depauperate zonal vegetation (Davies, 1921; Scott, 1989). The subtropical to warm-temperate deltaic vegetation with Taxodium–Nyssa swamp forests are relatively meagre autochthonous peat producers. Allochthonous peat periodically accumulates in the sediment-starved lows bypassed by the main distributary channels. Thick Tertiary molassoid coals of downfaulted alluvial basins


25

of Central Europe and western North America are mostly derived from allochthonous peat accumulations in high-energy deposition environments. Their accumulation is owing to a dense vegetation cover that reduced the clastic runoff at humid phases of climate cycles. In the temperate zone, ombrotrophic peat bogs prevail in the low-energy environments of slowly sagging intracratonic basins, such as the present-day alluvial plains of northern Europe and West Siberia. Giant Siberian coal reserves owe their origins primarily to this type of peat deposition (see an early example from Kuznetsk Basin in Zalessky & Tchirkova, 1934). Similar combinations of climatic and tectonic variables control startigraphic successions of coal types. Coals derived from rheotrophic peat are characterized by a far-going microbial decomposition of plant dead mass, with abundant macerals that are occasionally amassed as cutinite, resinite or exinite. Rheotrophic fusenites are primarily related to oxidation by occasional influxes of acidic surface water. In ombrotrophic coals, decomposition of plant material is variable, the little altered palaeosol and leaf beds associating with fusenite from bog fires. While the rheotrophic coals penetrate deep in the summer-dry zone, the ombrotrophic coals are mostly restricted to summer-wet climates. Both types of coal are involved in sedimentary cyclicity forming either the sandstone–shale–coal triplets of fluvial sequences or the shale–coal–carbonate triplets of littoral sequences. Both tectonic and eustatic factors have been inferred as explanations of such cyclic sequences. In respect to Palaeozoic cyclothems, a tectonic model (Weller, 1930) was replaced by eustatic model (Wanless & Shepard, 1936) and the derived eustatic–climatic models (e.g., Cecil, 1990) related to Gondwana glaciations. A weak point of the latter is that vast glaciations were episodic while the coaly cyclothem patterns persisted over the larger part of the late Palaeozoic to Cenozoic sequences. Of the multiple factors engaged in the orgin of coal cyclothems, the following seem of general significance: (1) Shift of peat-producing wetland zones related to tectonic (earthquake generated) subsidence of coastal mires (with shallow carbonates prograding over peat, as described for the extensive tropical mangrove peat on the Caribbean coast of Panama: Phillips & Bustin, 1990) or eustatic sea-level fluctuations, with derived effects of ground water table/salinity, (2) Fluctuations of clastic runoff periodically overwhelming peat accumulation, related to sea level (the basis of erosion), climate (the precession-driven precipitation cycles) or a filtering effect of marginal vegetation that changes with vegetational successions (e.g. Typha – Acrostichum succession: Collinson, 1983), hurricane intensity and other disruptive effects, (3) Changes in dead mass degradation rates related to vegetational succession of peat-forming/non-forming communities (such as the sawgrass/spike rush communities in the Everglades: Collinson, 1983) and detritivory levels in turn related to alternation of

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eutrophic (rapid dead mass disintegration at low pH values) and oligotrophic (rapid dead mass accumulation at high pH values) regimes. The Milankovitch-type cyclicity might have been involved in each of these groups of factors.

II.7.2. Variegate cyclothems In the alternations of light/dark, green/brown, bluish/purple deposits, the colour is given mainly by the iron and manganese compounds, silica and dispersed organic matter. It can be primary, related to the chemistry of sediment and the pore water, or secondary, related to ground water chemistry and fluctuations of water table. The brown/green alternation is due to fluctuation of the Fe (III) – Fe (II) redox boundary (Lyle, 1983). Both Fe and Mn redox boundaries shift with influx of organic matter. Mobilization, accretion and scavenging of metals across redox boundary has been studied in detail in relation to Mn nodules (Stumeyer & Marchig, 2001). The cyclicity of nodule accretion and metallic concentrations is related to the surface water productivity. In particular, the nodule growth rates reflect short-term – of El Niño time scale – vertical oscillations of redox zonation (König et al., 1999). The redox potential is also affected by a rapid deposition of turbidite or volcanic ash. Oxygen diffusion beneath tephra is affected by manganese concentrations and, on evidence from the Mount Pinatubo fallout, is completely prevented by a tephra layer more than 3 cm thick (Haeckel et al., 2001). In mixed water bodies, organic deposition is low while iron is precipitated as ferric compounds (brown to purple bands). In contrast, the ogranic-rich sediments of stratified water bodies are grey to black, with different shades of green owing to accumulation of ferrous compounds. Fe++ concentrations are high in the hypolimnion at low pH and redox potential, but rapidly decrease with water mixing and oxygenation. A flow of phosphrus over the sediment/water interface is closely linked to Fe++ concentrations. Ferruginous deposits adsorb phosphates that are released with anoxy and circulate in the hypolimnion or, with episodic mixing, fertilize the epilimnion. Thus, colour banding is related to fluctuations of aquatic ecosystems, for which the terrestrial runoff, the depth of thermocline, and the biotic production to dead mass deposition ratios are the critical variables. Tectonic settings and volcanism control the input of iron and other metallic ions in sedimentary environments. Both bedrocks and volcanic fallouts affect pH of surface waters and soil, while climatic effects prevail in postsedimentary ferrous/ferric replacements and the origins of massive redbeds (usually containing residual greenish inclusions). A vast realm of Permian–Triassic variegate (red) deposits extends from the Urals over the Russian Plains to central Europe. In the European Rotliegenden, their repeated sequences commonly include eolian sands, playa lake mudstones and calcareous fossil soil, with a Milankovitch-scale cyclicity (Bailey, 2001). The transboundary Permian–


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Triassic sequence at Nedubrovo, Vologda Region of European Russia, begins with grey gravels and coarse sands deposited over the eroded surface of variegate Permian marls. These basal deposits contain reworked rhizocretions as well as scattered bone fragments of terrestrial tetrapods. They are overlain, sequentially, by purple clay, black silty clay, greyish green siltstones, pink siltstones and brown clay, cut by conglomerates of the next erosion cycle (Krassilov et al., 1999a, 2000). Plant debris is found at a few levels in the green horizon. The black silty clay is mm-lamellate, smectitic, with pyrite nodules, charcoal, wood fragments and abundant leaf cuticles (also occasional insect cuticles) scattered over the lamellae or amassed as coaly lenticles. The palynological assemblage of predominantly terrestrial origin contains few planktic unicells, as well as Tympanicysta, the abundant filaments of cylindrical to barrel-shaped cells previously assigned to fungi but reinterpreted (Krassilov et al., 1999b) as green algae related to Spirogyra (Zygnematophyceae). This latter forms dense mats in fresh-water ponds, occurring also in brackish waters and mineral springs. In Tympanicysta, a brackish environment might have fostered formation of thick-walled akinetes that were preferentially preseved in estuarine deposits. While the redeposited rhizocretions of the basal gravels suggest a caliche over the adjacent drylands, the abundant plant debris of the black clay horizon and above is evidence of a dense plant cover (xeromorphic conifers and peltasperms) exporting an abundant dead mass. The purple clay – black clay succession reflects a switch from lagoonal (with inflowing surface waters – outflowing bottom waters) to estuarine (reverse) circulation (green silt horizon) and back (pink horizon). Anoxy of the black clay facies might have been related to the following factors (or their combination): (1) ponding of estuaries by the rising sea; (2) influx of terrestrial organic material enhancing eutrophication; (3) ashfall or redeposited ash from a nearby volcanic area (the Timano-Petchorian volcanic province simultaneous with Siberian traps) as a source of smectite, Fe++ enrichment enchancing a release of phosphorus from sediment fertilizing surface waters. The smectitic clay with Tympanicysta is a characteristic feature of the Permian/ Triassic transboundary sequences (Jin et al., 2000) indicating that, at the moment of frontal river ponding and eutrophication at an early stage of the end-Permian transgression, the organic-rich waters were splayed offshore generating an estuarine-type circulation in the advancing epeiric seas (more on this in IX.2-3). Variegate cyclothems commonly contain primary or redeposited fossil soil (palaeosol) horizons. A considerable depth of calcification in the redbed palaeosols abounding in stacked carbonate rhizocretions indicates a dry to seasonally dry/wet climate. However, the excessive abundance of carbonate nodules, as in the Lower Eocene of Bighorn Basin, Wyoming (VII.5), indicates reworked, rather than authentical, palaeosols. In such cases, the associated plant assemblages are poor guides to redbed climates. The in situ root beds of variegate cyclothems contain either horizontal or vertical roots suggesting repetitious developments from waterlogged to well-drained soil and the corresponding vegetational succesion (of Eospermatopteris wetlands to Archaeopter-

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is dryland forests in the Devonian: Driese et al., 1997). Yet wetland trees, such as swamp cypress, also produce taproots, and the Nysso-Taxodietum assemblage is a constant associate of redbed cyclothems. In ombrotrophic environments, penetrating roots are functional before and in the process of decoupling from ground water feeding. Alternatively, in rheotrophic environments, the ferruginous sheaths on vertical roots indicate penetration beneath anoxic layer to avoid a microbial damage of root tips. Alternation of these trophic types might have been responsible for certain types of variegate cyclicity. Pedogenesis itself shows a regular periodicity of soil development/truncation cycles related to the vegetation cover (sparse at the time of truncation), precipitation, upland runoff and fluctuations of the ground water table (Retallack, 1985b). In the Palaeogene of Bighorn Basin, Wyoming such cycles might have been related to avulsion with a periodicity of about 2.000 years (Wing et al., 1991), that is, of the time-scale of a seral vegetation/soil development. As in the case of coal-bearing cyclothems (above), current interpretations of Palaeozoic variegate cyclothems with evaporites and pedogenic horizons are strongly influenced by the idea of Gondwanic glaciations. In particular, glacioeustatic fluctuations are inferred as a cause of Pennsylvanian evaporites (regression, dry equatorial climate) alternating with silicoclastics, dolomites and black shales (transgression, humid climate). Such alternations might have been related to precession-driven climatic fluctuations even in the absence of considerable continental ice sheets. The opposite correlation of climatic phases with glacioeustatic events has been postulated by Miller et al. (1996) for the Permian (Wolfcampian) carbonate/variegate mudstone cyclothems, with carbonate deposition under an arid climate at the sea-level rise – high stand phase (deglaciation) switching to silicoclastic deposition with pedogenic horizons under a monsoonal dry/wet climate at the sea-level fall – low stand phase (glaciation). However, no evidence of glaciation is known for the Wolfcampian and chronostratigraphic correlates. The lithofacies alternation can be fully explained by periodic fluctuations of river runoff plus the filtering effect of coastal vegetation. Since Potonieisporites is a common palynomorph of evaporitic sequences while Lycospora occurs in silicoclastic deposits (Rueger, 1996), a replacement of pteridosperm–coniferoid mangroves by freshwater lepidophytic marshes with an increase of river runoff may seem a parsimonious interpretation of their succession. Similarly, a reduction of coastal vegetation (mangroves) has been invoked as a factor of continental runoff and a consequential increase in palynomorph diversity over the Milankovitch-scale sedimentological–palynological cyclicity in the Palaeogene (Rull, 2000).

II.7.3. Lacustrine cyclothems Biochemical cyclicity is well studied for lacustrine ecosystems (Henderson-Sellers & Markland, 1987). Depending on the type of water body and climate, the cycles are


29

annual or of longer periodicities, reflecting lake level fluctuations, biotic production/deposition rates and the stratified/mixed structure of the water column. The low to mid-latitude lacustrine sediments commonly show cycles of several ranks. They reflect, in descending order, (1) The sea-level bound erosional cycles marked by influx of coarse-grained material, (2) The Milankovich-scale precipitation cycles affecting both the lake level and the drainage runoff, (3) Their associated pH and biotic cycles, (4) The El- Niño scale climatic periodicities expressed primarily as the monsoon intensity cycles, (5) Ponding by uplifted tectonic blocks or lava flows (as in the rift lakes: Johnson et al., 1996), (6) Seasonal cycles of drainage runoff and lascustrine biotic production expressed as mm-scale light/dark bands, (7) Their superimposed trophic autocycles due to a periodically disbalanced production and deposition of organic matter. In the rift lakes of East Africa, fan deltas are formed over the bordering fault zones at low stands while gravity currents dominate sedimentation at the high stand phases. Alkaline water chemistry and evaporite deposition associate with low stands (Scholz et al., 1990). Volcanic aerosol emissions and ash deposits affect water chemistry through acid rains and chemical weathering of magmatic bedrock, a source of aluminium and fluorine. Their effect on fish at the top of trophic cascades may alter trophic conditions of the lake. In the rift lake Turkana, Kenya, the 4-year microcycles reflect hydrologic fluctuations related to El- Niño (Hoffman & Johnson, 1988). Remarkably, this scale cyclicity was found also in lacustrine deposits of Cretaceous rift system in eastern Asia (Fig. 13). The Milankovitch scale erosion cycles correlate with charcoal from vegetation burning, an increase of pH and the prevalence of early successional species. Over a seral succession, pH tends to decrease owing to the development of a deeper soil profile and an increasing uptake of cations by vegetation (Rhods & Davis, 1995). The lacustrine black-shale deposition in the Cretaceous of central-eastern Asia is an example of a relatively stable rift lake system of high biotic productivity, yet remaining oligotrophic owing to a rapid subsidence of the dead mass to the anoxic zone of a stratified water body. In Transbaikalia, the lacustrine deposits rest on the weathered granitic bedrock, with coarse clastics (volcanoclastics) at base, followed by the irregular siltstone – shale – marl cyclothems intervened by the coarse sandstone – gravel beds marking the larger erosional cycles (Fig. 14). The sequence is about 300 m thick spanning about 500 k.y. of the latest Hauterivian to early Barremian (Scoblo et al., 2001). Thick black shales of mm-thick lamellae occur in the middle part of the sequence.

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a

b

c

d

Fig. 13. Varvite sequences in the Late Albian – Early Cenomanian shales of the Daladze Formation, northeastern China: (a, b) a quasiperiodic alternation of thick/thin varvite bundles of ca. 4 years (El Niño scale) and ca. 11-22 years (solar activity scale), (c) a sequence of the El Niño-scale cycles interrupted by two nonlaminated intervals reflecting episodes of water mixing, (d) an angiosperm leaf on a light lamina covered by a dark lamina with small plant debris.

Thick black shale deposition reflects a stable thermocline in the deeper part of the lake. Thin lamination might have been owing to cyanophyte mats covered with lime that was rapidly dissolved under low pH of the hypolimnion. Sand grains on organic lamellae indicate erosion pulses that might correspond to precipitation maxima of the same scale as in the clastic/carbonate cyclothems of the shoals. The coarse-grained members represent tectonically driven erosion events accompanied by volcanic activities. The siltstone/marl couplets about 1.5-2 m thick are assigned to the Milankovitch-scale precipitation cycles, with the marl beds marking a drier phase. The marls show microscopic dark/light bands reflecting seasonality of organic deposition. Occasional marl/siltstone bedding planes, traced over several hundred metres, are densely covered with impressions of aquatic insects, mostly larvae of Hemeroscopus dragonflies, Ephemeropsis mayflies and Coptoclava beetles. The imagos are lacking, except for a few dragonflies. Though such accumulations might be due to moulting, the association of different species suggests mass mortality. Fish remains (Lycoptera) are rather common. Plant debris and terrestrial insects are scattered over the bedding plane indicating an advance of wetlands.


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b

a

d

c

Fig. 14. Early Cretaceous (mid-Neocomian) lacustrine black-shale facies with proangiosperm remains and an aquatic Lycoptera– Ephemeropsis assemblage: (a) an outcrop at Basia, Vitim River, Transbaikalia, (b) lamellate black shales, (c) marly wetland palaeosol with roots and Hemeroscopus dragon-fly imago and larvae, (d) pappate cupules of Baisia, a semiaquatic proangiosperm, and the aquatic larvae of a may-fly Ephemeropsis and a dragon-fly Hemeroscopus on a marly palaeosol slab.

Of several factors potentially involved in the episodes of mass mortality, the following might have been of a critical importance: 1. As in the present-day European lacustrine ecosystems devastated by technogenic acid rains, pollution by volcanic eruptions in the surrounding highlands might have inflicted sharp pH fluctuations, their direct impact accentuated by an input of aluminium (affecting respiration in aquatic organisms) from the granitic bedrock.

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2. With a rise of lake level at the transition from dry to humid phase of precipitation cycle, the oligotrophic limnobiota might suffer from eutrophication effect of a massive input of organic nutrients from wetlands. 3. A similar eutrophication effect might have been rendered by turnover of organicrich bottom deposits under a depressed thermocline. In rift basins, a concerted effect of such factors would follow from a coincidence of high rate tectonic activity and climate change.

II.7.4. Carbonate cyclothems and black shales The basinal cyclic sequences of dolomites, limestones, chalk (with or without silicite nodules), chert and black shales are ascribed to a variety of factors, such as sea-level or climatic fluctuations, tectonic subsidence rates, biological productivity versus dead mass deposition rates, shifts of thermocline or picnocline (density interface), carbonate dissolution depths, hydrothermal activity, etc. (Worsley & Martini, 1970; Heezen et al., 1973; Bosellini & Winterer, 1975; Hsü, 1976: Demaison & Moore, 1980; Arthur et al., 1988; Wignall, 1994: Prokoph et al., 2001; Wilson & Norris, 2001). Secondary processes of decalcification, chertification, ferromagnetization or diagenetic redistribution of lime (or silica) might have been largely responsible for a sequence of lithotypes or at least for distinctness of their primary alternation (Einsele. 1982; Raiswell, 1988). Thus, Alpine cherts sometime contain aptichi (cephalopod mandibles) indicating a post-sedimentary chertification of Aptichus limestones (Hsü, 1976). Lime migration from homogeneous to porous (shelly) limestone varieties produce secondary differences in the relative clay content and hardness (Einsele, 1982). Yet the fossil content of alternating lithologies indicates climatic control of carbonate cycles. In a convincing example of hard/soft marl cyclicity in the Cenomanian platform carbonates of northern Crimea (Krassilov, 1984, 1985), plant fossils – the compressed fern pinnules, conifer shoots and angiosperm leaf fragments – are confined, together with occasional remains of terrestrial insects, to the hard marl member about 0.3-0.4 m thick, abounding in pyritic nodules, inoceramid and ammonite shells that occasionally show a well preserved mother-of-pearl layer. The somewhat thicker soft marl members lack both plant remains and ammonite shells, but contain numerous aptichi, as well as trace-fossils and coprolites attesting to a diverse bottom life (Figs. 15, 16). Since cephalopod shells are aragonitic, whereas the aptichi are calcitic, their respective association with pyrite/plant remains and trace fossils indicates a leading role of oxygen level at the sediment/water interface as a factor of carbonate cyclicity. Influence of dissolved organic matter on deposition and diagenesis of carbonate minerals is related to pH changes and precipitation of ferric compounds. Aragonite dissolution increases with oxidation of organic matter. Conversely, both the depletion of oxygen and precipitation of FeS enhance aragonite saturation (Tribble, 1993).


33

Fig. 15. Carbonate cyclicity: hard marl – soft marl alternation in the Cenomanian of northern Crimea. The hard marl (lower member) is deposited during a humid phase of the cycle with an increased river runoff diluting surface water and inflicting anoxy down the water column, hence the preservation of plant compressions and aragonitic shells, enhanced by H2S. The soft marl (middle member) corresponds to a drier phase of the cycle with water circulation restored over the carbonate platform supporting a diverse bottom life (coprolites); the calcitic aptichi are preserved while the aragonitic shells are dissolved (Krassilov, 1985).

The following model seems accountable for the above taphonomic features: - Carbonate silt accumulates upon the shelf platform populated by a diverse benthic community; - Oxidation of organic material, corrosive to aragonite, prevents preservation of cephalopod shells, except the calcitic aptichi; - At the next stage, an influx of terrestrial organic material indicates an increase in riverine discharge, most probably owing to a more abundant precipitation. - Delution of surface waters decreases vertical mixing; - Incorporation of terrestrial organic material in the carbonate sediment generates anoxy of the boundary layer exterminating benthic life (except inoceramids tolerant of a mild anoxy level). Pyrite precipitation makes the environment less corrosive to aragonite providing for preservation of ammonite shells; - Post-depositional redistribution of lime makes the shelly marls the harder of the cycle.

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c

a

b

d

e

Fig. 16. Carbonate cyclicity in the Cenomanian of northern Crimea: (a-d) soft marl with coprolites, trace fossils and aptichi; (e, f) hard pyritic marl with shell fossils (ammonites, inoceramids) and plant debris (Krassilov, 1985).


35

Since precipitation changes are a major mid-latitudinal climatic effect of precession cycles and since the carbonate cyclicity is of a 104-year scale, the orbital forcing seems a plausible explanation. Orbital forcing has been detected in the Mediterranean platform carbonate sequences, with the meter-thick couplets reflecting the precession cycles, their bundles – the eccentricity cycles (D’Argenio et al., 1994; Buonocunto, 1999; Fiet & Gorin, 2000). With the larger fluctuations of water chemistry extending the limits of carbonate deposition, the 2-member carbonate cycles are transformed into the 4-member sequences of silicites, carbonates, black shales, and phosphorites typical of epeiric basins. Their correlation with pH is shown in Fig. 17. With shallowing of carbonate lysoclines, anoxic carbonate facies give way to black shale facies that typically spread with transgressions owing to an input of terrestrial material from inundated coastal wetlands and ponded estuaries. Ponding events, followed by eutrophication, are inferred in the case of the Permian/Triassic boundary clay with Tympanicysta (IX.2) as well as for the plant-bearing clayey shallow-water marine horizons in the Cenomanian (the Nammoura locality of Lebanon) and Turonian (the Gerofit locality of southern Negev) of Middle East (Krassilov & Baccia, 2000; Krassilov & Dobruskina, 1998) tentatively correlated with Mediterranean black shale levels. The Milankovitch-type cyclicity provides a common basis for correlation of marine and non-marine sequences. Remarkably, the oceanic black shale events are nearly synchronous with the lacustrine ones over a succession of Cretaceous talassocratic/temperature maxima in the Barremian–Aptian, late Albian and late Cenomanian to midTuronian (see Schlanger & Jenkins, 1976; Tissot, 1978; Graciansky et al., 1984 for marine events; the corresponding lacustrine events are represented by the Turginian, Baisian and Daladze levels in Transbaikalia and eastern China, see Krassilov & Bugdaeva,

Fig. 17. Changes in pH (supposedly related to CO2 concentrations) over a series of carbonate-black shale lithologies: (D) Dolomites, (L) Limestones, (PL) Phosphoritic chalk, (BL) Black bituminous limestones, (BS) Black shales, (Si) Chert, (SiL) Siliceous limestones.

36


2000). Such chronological correlations suggest common basic factors. Lacustrine sedimentation is regulated by drainage runoff in turn controlled by sea level (the basis of erosion) and precipitation driven by the precessional cycles. Similarly, marine cyclicity results from sea-level-driven, as well as precipitation-driven fluctuations of terrestrial runoff with impacts on density stratification, biotic productivity of surface waters and oxygen concentrations down the water column (VII.1.1, VII.2.3, IX.3 and elsewhere).

Chapter 3. Palaeoecology

37

III. PALAEOECOLOGY In essence, palaeoecology is reconstruction of past life forms, biotic communities and their habitats based on taphonomic information and present-day examples. Species autecology is inferred from functional morphology, facies occurrences and traces of biotic interactions (such as in situ symbionts, gut contents, etc.). Synpalaeoecology deals with fossil assemblages as a source of inference on source communities and environmental factors. Certain synecological variables can be deduced from geochemistry or isotope ratios without reference to particular fossil assemblages. However, a conversion of isotopic ratios into temperature, salinity or biotic productivity is likewise a matter of inference, the reliability of which depends on soundness of theoretical premises and on robustness of the techniques. The aut- and synecological approaches are certainly interrelated. Yet a search for autecological solutions for long-standing evolutionary problems, such as posed by mass extinctions, might prove futile if the solutions occur at synecological level.

III.1. Life forms Past biota is customary depicted as a display of extinct species, each represented by a few individuals combined in an artistically pleasing manner. Such reconstructions might give a crude idea of flora and fauna as compilations of fossil plant and animal species but fail to reproduce the biota as a functional system of co-adapted life forms. In plant ecology, life forms are considered as morphological expressions of ecological roles, or niches (Drude, 1913). Although a life form unit may correspond to a taxon (the more so as many plant taxa are based on life forms), their one-to-one correspondence is not mandatory. A reconstruction of an extinct life form is a functional interpretation of both morphology and taphonomy. As a functional unit, the biotic community is a set of life forms that is, to a certain extent, predetermined and predictable. Primary producers, consumers of two or more levels and dead-mass destructors are obligatory life form categories, although their species contents vary from one ecosystem to another. When a link in a trophic cascade or a succession of seral forms is missing, there should be a potential space for it conventionally conceived of as a “vacant niche”. By implication it will be filled sooner or later, if not by contemporaneous life forms then by new ones derived for the purpose.

III.1.1. Multifunctionality Life form reconstructions of extinct organisms represented by isolated structures, such as leaves, disseminules, pollen grains, etc., are notoriously ambiguous because of multifunctionality of morphological characters. Samaras, for instance, are thought of

38


as serving for wind disperal, but in certain cases their utility is in adhering flat to the ground. Pappose disseminules are dispersed either by wind or exozoochorously. Saccate pollen grains are usually taken as evidence of wind pollination, although pollen bladders as volume-regulating structures occurred in insect-pollinated plants as well (Krassilov & Rasnitsyn, 1997; Krassilov, 2000a). On the other hand, the morphologically nearly identical asaccate pollen grains are either wind-dispersed (in Ginkgo) or insect-dispersed (in cycads). The latter examples also illustrate substitution of functions over geological time. The adaptive meaning of pollen bladders might have been different in Palaeozoic gymnosperms, for which a protosaccate structure (Fig. 18) might have served primarily for attracting pollinivorous insects, and in their Mesozoic descendants, in which these structures (proto-eusaccate) were disentangled from attraction and, with strobilation, acquired new functions of enhancing pollen catchment by spinning pollen grains in air vortices between cone scales. Ambiguity arises from homeomorphy of different, even contrasting, adaptations, as in the case of leaf trichomes that either reduce transpiration in dry climate or protect leaf surface from excessive vapour in humid climate. Trichomes also prevent insect damage, incidentally by assisting in predator regulation of herbivorous insects, as in the intriguing observations of leaf pubescence reducing population densities of herbivorous mites by encouraging webbing activity in the two-spotted spider mite (Roda et al., 2001). Trichomate bracts commonly occur in xerophytes, but in the early spring “downy plants”, the extremely dense bract trichomes keep primordia warm (Tsukaya & Tsuge, 2001). Stomatal papillae reduce evapotranspiration or protect stomatal pits

Fig. 18. Partly digested Lunatisporites–type pollen grains from a gut compression of a Permian insect Parapsocidium, showing a protosaccate structure of rooted sexinal elements.


39

from water drops. Most morphological characters conventionally related to climate are likewise ambiguous (VII.1.3). Climatic adaptations are those pertaining to temperature and vapour pressure at the leaf/air or root/soil interface. Both root mass and leaf mass (their ratios) are thus affected by climate. But an excessive development of underground organs (both gametophytic and sporophytic) in early land plants and their derived arboreal spore plants (stigmaria of Palaeozoic lepidophytes) might also reflect a less advanced mycorrhizal system of the primitive rhizospheres. Plant growth is limited by the water conductivity of vascular supply, the morphological elaborations of which (bordered pits, perforation plates, etc.) prevent emboly – a plugging of conductive elements by air bubbles and their spread over the system. A risk of emboly increases with water stress, hence the smaller dimensions – higher density of vascular elements, a differentiation of seasonal increments in the seasonally dry/freezing climates, a correlation of early/late xylem ratios with latitudes (Carlquist, 1978). However, since water stress is regulated from above by leaf pumping it eventually depends upon evapotranspiration. Reduction of evapotranspiration in dry climate clashes with the prevention of embolism. Xeromorphism is a morphological expression of this physiological problem. Xeromorphy develops in respect to a deficiency/inaccessibility of soil moisture or exposure to direct sunshine/strong wind pressure, as well as an ineffective vascular system. The latter factor might have had an evolutionary dimension explaining the xeromorphy of early land plants. A restriction of the plant/air interface (dense leaf packing, appressed leaves, thick leaf blades, revolute margins, small leaf area, deciduousness) is the most general solution of the water stress problem. Xeromorphic plants also reduce evapotranspiration by decreasing stomatal conductivity (papillate stomata, wax plugs) or increasing aerodynamic roughness of their leaves to wind pressure (higher in conifers than in broadleaves: Meinzer, 1993). These foliar variables are related to the boundary layer that decouples vapour pressure at leaf surface from that of the ambient air by accumulating the transpired vapour (Fig. 19). Leaf surface micromorphology, in particular the prominent venation network over stomatiferous lower surfaces, scabrate cuticle, striation, cuticular folds, stomatal pits and trichomes pertain to the decoupling efficiency (measured as the decoupling coefficient: Meinzer, 1993) of the boundary layer that tends to increase from conifers to deciduous broadleaves to evergreen broadleaves. In thickly cutinized leaves, the boundary layer is restricted by the areas of densely packed stomata, such as stomatal grooves in cordaites, peltasperms, bennettites, taxaceous and sciadopityaceous conifers, etc. Such leaf microstructures frequently occur in helophytic life forms. Xeromorphy broadly overlaps with scleromorphy, an excessive development of mechanical tissues in relation to distrophy, in such features as an elevated root/shoot ratio, small fibrous leaves, thick cuticles, sunken stomata, etc. (Figs. 20, 21). In palaeoecology, their separation is not always feasible, the more so since in evergreen vegetation dry habitats usually associate with low-nutrient soil types (Hill, 1998).

40


Fig. 19. The excessively trichomate phyllodes in a protognetalean plant Dinophyton from the Triassic redbeds of Arizona (Krassilov & Ash, 1988).

a

c

b

Fig. 20. Scleromorphy in a helophytic bennettite, Otozamites lacustris, from the Early Cretaceous of Mongolia, represented by the shed pinnules of pinnate leaves (Krassiliov, 1982): (a) festooned epidermal cell walls, (b) stomata sunken in the large-celled spongy mesophyll, (c) dense ridged venation.


41

a

b

d f

c

e

g

Fig. 21. Scleromorphy in Mesozoic helophytes: (a) papillate stomatal grooves in Phoenicopsis, (b, c) trichomate lower epidermis and (d-g) papillate stomata in Ginkgoites from the coal-bearing wetland facies of the Bureya Basin, Russian Far East (Krassilov, 1972b).

42


In terms of life-form evolution, scleromorphy is considered as a preadaptation to dry or cool climates (Hill, 1990). Sclerophyllous vegetation might have appeared as a marginal derivate or a seral stage of laurophyllous woodlands adapted to a low fertility soil (Axelrod, 1975). Such seral sclerophylls might have given rise to the chapparal–macchia type xeromorphic communities initially appearing as undershrubs of sclerophyllous woodlands (Axelrod, 1975). Sclerophyllous communities seem also to have been engaged in the altitudinal differentiation of frost-resistant Alpine vegetation belt. Taken literally, xeromorphy as indicator of aridity is a prolific source of erroneous climatic inference (e.g., in the case of a climatological interpretation of the Late Permian peltasperm–conifer assemblages: Krassilov, 2000b). Even in the commonly recognized xeromorphs, such as the Permian callipterids, a scrutiny of epidermal features reveals a helophytic adaptation to excessive waterlogging (Busche et al., 1978). A disambiguating piece of evidence comes primarily from the lithofacies. In particular, a constant association with coal beds suggests a scleromorphous helophytic, rather than xerophytic, life form. Thus, because of its functional ambiguity, morphological evidence should be supplemented by taphonomic data. The modes of dead mass accumulation can even be a better guide to life forms than morphology. Some leaf shapes, in particular those with a long petiole and peltate wind-rough blade, imply deciduousness but, in addition to or irrespective of morphology, seasonal leaf shedding is verified by leaf-mat type taphonomy. Disseminules occur either in the leaf-mats of their source plants or separately. Notably, not all deciduous trees accumulate leaf litter in their stands. Maple stands produce thick leaf mats but lime stands do not. These distinctions in littering correlate with dispersal strategies and sprouting under the canopy of parental plants or lack of such (in lime). In the case of plants that flower before leaves, shed floral parts tend to associate with leaves of other plants. Such associations sometimes lead to erroneous assignments but, when properly recognized, are evidence of deciduousness and early flowering.

III.1.2. Life form inference: examples The following examples illustrate the significance of taphonomic inferences in extinct life form reconstructions. Pleuromeia is a peculiar Triassic lycopsid with a typically unbranched (or forked) stem, bulbous 4-lobed rhyzome, long bifacial leaves and a terminal heterosporous cone. It was originally interpreted as xeromorphic and perhaps psammophytic, a reconstruction that fitted a traditional notion of arid Triassic landscape (Mägdefrau, 1956). I have studied Pleuromeia from three Early Triassic localities, the Ryabinsk on the Volga River, the Russian Island near Vladivostok, and the Olenek River in northern Siberia (Fig. 22). The facial occurrences are both terrestrial and marginal marine. The abundant remains of rhyzomes and aerial stems indicate autochthonous assemblages, in which Pleuromeia


43

Fig. 22. Pleuromeia, a dominant form of the lycopsid wetlands spreading worldwide in the Early Triassic: a cone in the Olenekian marine sandstones of Russian Island, Primorye, Russian Far East (Krassilov & Zakharov, 1975).

was the only macrophyte. It obviously formed pure stands. The most common lithologies are organic-rich shales with plant compressions and occasional marine fossils indicating anoxic estuarine environments. Occasional finds in sandstones with chaotic cross-bedding represent a stand buried by hurricane wave deposits (tempestites). Dispersed sporophylls and spores are also found as allochthonous material in nodules with small ammonoids at the core. Since the ammonoid shells and megasporophylls are similar in shape and dimensions, they might have been transported together by sea currents. In agreement with this suggestion, the sporophylls are boat-shaped, with a solitary sporangium sunken in a spoon-like depression of the upper surface. Both the morphology and taphonomy suggest dispersal by floating sporophylls as a regular feature of reproducitve strategy. A reconstruction emerging from these considerations is that of a vast reed-like coastal wetland extending into river mouths and further inland (Krassilov, 1972a; Krassilov & Zakharov, 1975; compare Retallack, 1975; Wang & Wang, 1982). The dryland Early Triassic vegetation is poorly known because of a strong filtering effect of the ubiquitous Pleuromeia stands. Nilssonia is a widespread Mesozoic gymnosperm that was conventionally assigned to cycads. The genus was based on fossil leaves that are simple, elongate, shortly petiolate, irregularly to regularly lobed, with entire or, rarely, dentate margins. Nilssonia leaves are common in the levee and muddy flood-plains facies where they are scattered

44


over bedding plains. The taphonomy reminds of willows and, indeed, some narrow-leaved Nilssonia are of a willow aspect. The cycad stereotype was shaken by the finds of leaf clusters apparently borne on spur-shoots and of branching long shoots with attached leaves (Krassilov, 1972a; Kimura & Sekido, 1975; Spicer & Herman, 1995). It became obvious that Nilssonia was a profusely branched heteroblastic tree of a growth form comparable with Ginkgo rather than cycads. Although cosmopolitan, as might be expected of an ubiquitous pioneer plant, Nilssonia was far more prominent in the temperate Arcto-Mesozoic zone than in the xeromorphic Meso-Mediterranean zone. The difference might have been partly owing to deciduousness of northern nilssonias while their southern congeners were evergreen (in contrast to angiosperms, dentate nilssonias associate with evergreen, rather than deciduous, vegetation). The thick-stemmed (pachycaul) bennettites of Cycadeoidea type, an extinct Mesozoic life form (Fig. 23), have traditionally been compared with living cycads that grew under the canopy of tropical forests or in open stands of a scrub savanna. However, the in situ silicified bennettitalean stems in growing position are locally abundant in coastal mudflat deposits indicating a quite different life form. In addition, the cycadeoids were cauliflorous, bearing strobilate flowers (flower-like strobili) on stems among the armour of persistent leaf petioles. With hundreds of such flowers, the trunk functioned as a huge inflorescence (infructescence after ripening of the seeds). These structures scarcely have any living equivalents, although the basicaulicarpy is characteristic of the presentday Annonaceae, of which the mangrove apple-tree, Annona glabra, grows in floodplain swamps and along the edge of mangroves on the Atlantic coasts of Africa and South America. In this and other annonaceous species, the basicaulicarpy is functionally linked to dispersal by turtles, alligators and iguanas. Saurochory, or dispersal by reptiles, is certainly an ancient mode, perhaps going back to the bennettites and dinosaurs. Brachyphyllous conifers, widespread in the Mesozoic, with the maximum diversity in the Meso-Mediterranean realm, are often interpreted as representing a xeromorphic dryland vegetation. However, in the Ust-Baley locality on the Angara River, East Siberia, the brachyphyllous genus Elatides is represented by cones and dispersed seeds (Samaropsis), many of which germinate, with seedlings showing tap roots and linear cotyledons (Fig. 24). The seedlings are buried in situ or nearly so, which is characteristic of wetland plants. A Mongolian find confirms such preservation being regular rather than occasional. In the taxonomically related extant Glyptostrobus, a small deciduous marshland tree, the cones ripen in September–October, the seeds are shed in winter, germinate on waterlogged substrate in early spring. The taphonomy thus attests to a wetland habitat of Elatides and supposedly of other morphologically similar scale-leaved conifers, the ecological interpretation of which as dryland xeromorphs has to be radically reconsidered. At the same time, the Ust-Baley locality contains abundant catkin-like cones of the Ixostrobus–Leptostrobus type, belonging to deciduous Czekanowskia and Phoenicopsis (of the extinct order Leptostrobales). In sandy facies, the transported cones


45

a

b

c

Fig. 23. Cycadeoidea, a pachycaule bennettite: (a) transversely cut stem in an armour of leaf petioles, (b. c) multicellular ramenta. The mid-Cretaceous (Albian) of Primorye.

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b c a d

e

h f

g

Fig. 24. A “fossil seed-plot”: the germinant Samaropsis-type seeds of a wetland conifer, Eatides, with seedlings at various developmental stages: (a-g) the Jurassic of Ust-Baley, Siberia, (h) the Cretaceous of Cobi, Mongolia.


47

are even more numerous than leaves (the opposite is more common elsewhere), which suggests their shedding before leaves (and simultaneously with seed germination in Elatides). The Classopollis-producing brachyphylls are also sometimes depicted as dryland xerophytes. However, of all gymnosperm pollen types, Classopollis is most common in offshore marine deposits sometimes amounting to 80% of palynological spectra (e.g., in the mid-Cretaceous deposits of Yamato Rise, the Japan Sea: Markevich, 1994). Such offshore abundances suggest, in addition to a coastal habitat, a winter – early spring flowering. In other seasons, the pollen rain would have been diverted inland by summer monsoons.

III.1.3. Direct and indirect environmental correlates Life form characters are a source of palaeoenvironmental, primarily palaeoclimatic, inference. A correlation of the entire/non-entire leaf margin ratios with temperature is a well-known example. The ratios are believed to be a direct response to the climate. However, a functional meaning of leaf margin structures remains obscure. A neglected aspect of functional plant morphology is its perception by insects. As has been shown experimentally (Brown & Lowton, 1991; Rivero-Lynch et al., 1996), the frequencies of insect visits are differentially affected by the leaf margin morphologies. For example, vine-weevils prefer plants with entire leaves, the flea-beetles – those with non-entire leaves. Leaf margin is formed by the marginal meristem that does not produce vascular tissue. Lateral veins thus tend to attenuate, curve or loop at a certain distance from the margin. However, in leaves with craspedodromous venation, the marginal expansion is arrested except at the vein ends that protrude as marginal serrations. A localized vascularization of the leaf margin can also be induced in association with marginal glands. Glandular teeth had first appeared in the mid-Cretaceous angiosperm morphotypes Protophyllum and Trochodendroides (Fig. 25) while serration is a later development. Leaf glands might repel leaf-cutters by attracting ants while their derived marginal features might provide a visual protection. Pollinivorous insects abundantly represented in the gymnosperm-dominated midCretaceous biotic communities might have been visually guided by conspicuous leaf shapes. Those insects that paid frequent visits to a certain plant species might have also been attracted by other plant species with similar leaves. Hence leaf mimicry might confer not only a protection from folivorous insects but also an attraction for potential pollinators. Leaf convergence might arise not so much in response to climate as owing to leaf mimicry induced by plant/insect interactions. Hence a correlation of leaf characters with climate might have been inderect, via climatic requirements of a co-adapted insect.

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a

f b d

c

e g

h

Fig. 25. Serrate leaf morphology as evidence of plant–insect interaction falling under climatic control, anticipated in a Palaeozoic gigantopterid Gigantonoclea from the latest Permian of South China (c), and further developed in the early angiosperms, starting with marginal glands, as in the platanoid leaves from the Cenomanian of Kazakhstan (d, e), and elaborating over an irregular double crenulation, as in Nammourophyllum from the Cenomanian of Lebanon (a) or Trochodendroides from the Senonian of Sakhalin (g), to a regular one in the younger broadleaves, e.g. Corylites from the Early Palaeocene of Sakhalin (b), and to emarginate types in Ushia from the Late Palaeocene of Cisuralia (f) or Tiliaephyllum from the Early Palaeocene of Tsagajan (h). After Krassilov (1976a, 1979), Krassilov & Shilin (1996), Krassilov et al. (1996a), Krassilov & Baccia (2000).


49

Certain morphological phenomena in extinct plant groups seem also related to leaf mimicry. Thus, cycadophytes constitute a heterogeneous group of Mesozoic gymnosperms including cycads, nilssonias, bennettites and pentoxyleans. These orders, though not closely related, are indistinguishable on the basis of their gross leaf (cataphyll) morphologies (Fig. 26). Formal genera for fossil leaf morphotypes include representatives of

a

c

d

b

e Fig. 26. Leaf convergence in the Mesozoic nilssonias (a, b) and bennettites (c-e) from the Late Jurassic to Early Cretaceous of Bureya Basin, Russian Far East.

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different plant orders that are readily discernible by their epidermal characters. Examples of such cycadophytic genera are:

Morphotype

Epidermal characters

Cycadophytic

Cycadalean

Bennettitalean

Taeniopteris Pterophyllum Cycadites Cycadolepis

Doratophyllum Nilssonia Paracycas Deltolepis

Nilssoniopteris Anomozamites Pseudocycas Cycadolepis

Leaf convergence is by no means a rare phenomenon but such a regular pairwise isomorphism is a unique feature of Mesozoic vegetation. Except the leaf shapes, the sparsely branched pachycaul bennettites had little in common with heteroblastic nilssonias. Such growth form disparities are against a purely climatic interpretation of the pairwise leaf convergence. In the cycadophyte-dominated assemblages, a single bennettitalean species commonly prevails in association with a few subordinate cycadalean species. For example, in the Lower Cretaceous of Primorye (Krassilov, 1967a), a few Doratophyllum leaves came from the same locality as abundant Nilssoniopteris leaves. In the Jurassic Yorkshire localities, Pseudocycas commonly associates with a morphologically similar Paracycas (Harris, 1969). It may thus seem that leaf similarities were advantageous for such paired species or one of them. Since we have direct evidence of insect pollination in cycads involving insect families with fossil representatives in the Cretaceous, and since there is indirect but convincing evidence of insect pollination in bennettites (Crepet, 1974), a plant/insect interaction might seem a plausible cause of the pairwise convergence, with a rare cycad species mimicring on the dominant bennettite species of a plant community.

III.2. Plant communities Although put to doubt time and again as objective biological entities, biotic communities remain the major focus of ecological research as the reproducible supraorganismic systems of life forms (descriptively substituted by their representative species) that are bound by their shared trophic resources, seral developments and environmental constraints. This definition reflects a holistic concept of biotic community as an evolving functional system. The opposite view, still held by some influential ecologists, depicts it as a collection of species that live together not because they are functionally complimentary to one another, but rather owing to their ability to maintain a kind of competitive equilibrium. In plant ecology, these latter views constitute the individualistic species concept


51

(Gleason, 1926; see Bennett, 1997 for a recent review), which is supplemented by the vegetation continuum concept (McIntosh, 1967). However, any biosociological concept has to account for such phenomena as the constancy of life form combinations over space and time and their orderly successions over developmental (seral) stages. Some ecologists argue that plants disperse as individuals forming communities whenever the environmental conditions allow their coexistence. The supportive evidence comes mostly from the post-glacial assemblages of pioneer plants not yet integrated into sustainable communities. Plant species may disperse as individuals but communities disperse by outflow of propagules, a spore–seed rain, that maintains its structural diversity. In particular, the structure of zoochorous communities is maintained by seed dispersers and is affected by their replacements (Huwe & Smallwood, 1982; Christian, 2001).

III.2.1. Community records Insofar as a fossil assemblage is derived from dead mass production of a biotic community, it reflects, with inevitable distortions, at least some systemic features of the latter. But how far the reflection deviates from the original is a matter of palaeoecological analysis. A problem most commonly encountered by plant palaeoecologists pertains to relative contributions of local versus regional sources. Do fossil assemblages represent distinct plant communities or are they a blend of plant communities constituting regional vegetation? It is commonly believed that macrofossils more often represent local sources while regional components are more prominent in spore-pollen assemblages. This being undoubtedly the case for certain types of terrestrial biomes, such as mixed conifer – broadleaved forests, does not exclude a fair representation of local plant communities in the pollen record. For example, different types of wetland communities, such as brackish marshes, cypress swamps, pinelands, mangrove forests, dwarf mangroves, etc., reflecting local fluctuations of hydrologic–edaphic conditions, are recognizable (and can serve as environmental proxies) in the pollen spectra of Everglades (Willard et al., 2001). Direct information on species associations comes from in situ assemblages, such as stump horizons. However, most of them reflect single-species flood-plain stands. Soil assemblages contain autochthonous rhizospheric remains as well as hypoautochthonous debris of aboveground stands. Two examples of palaeosol associations are shown in Figs. 27, 28. In the regressive Umalta sequence, Bureya Basin, Russian Far East, plant-bearing deltaic deposits rest on marine Lower Callovian sandstones with ammonites. The transitional horizon contains limulid remains and microplankton (Tasmanites). A hydrosere palaeosol on top of it is a ferruginous clayey root-bed about 5 mm thick. Plant debris obtained by bulk maceration contains pinnules, sporangia and spores of Osmunda diamensis and Dicksonia burejensis, two dominant fern-marsh species, as well as slender horsetail shoots.

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There are also minute leafy liverworts and other bryophytes, mixed with Selaginella, a ground cover. Well-preserved soil mites came with plant debris (Fig. 27).

d

a

b

c

g

h

i

e j

f

l

k

m

n

Fig. 27. A fern-marsh fossil soil from the regressive Umalta sequence, the Jurassic (Callovian) of Bureya Basin, with macroscopic remanis of a dominant tree fern, Dicksonia burejensis (a). The micro- and mesofossils obtained by bulk maceration represent an autochthonous plant debris from the rhizosphere, ground cover and the aboveground stand: (b) sporangial annulus with attached spores, probably of Dicksonia burejensis, (c) trilete fern spore, (d) detached leaf of Selaginella from a bryophyte–lycopsid community of the ground cover, (e, f) leafy liverwort, Cheirorhiza, from the same community, (g-i) oribatid mites, (j, k) naturally macerated pinnule compressions of Osmunda diamensis showing stomata (Krassilov, 1978). The palaeosol horizon rests on top of littoral shales with(l) microplankton and (m, n) limulid remains.


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Palaeosol from a dinosaur locality in the Maastrichtian of Nemegt Basin, Mongolia is preserved as a thin, light brown with yellow ferruginous spots, clayey film on laminated alluvial siltstone. It contains horizontal roots of aquatic macrophytes and a mycelium of dense hyphal strands. A group of fruiting bodies attached to the mycelium show radially split exoperidia, as in the modern saprophytic gasteromycetes of the earthstar genus Gaestrum (Fig. 28). Mycelial films might have contributed to incipient soil development on detritic levee deposits. Pollen grains in the gut compressions of fossil insects reflect pollen sources within a flight distance, not too vast in large pollinivorous insects. If two or more pollen types are mixed in the guts, they are most likely to have been produced by plants that grew side by side. Pollen extracted from gut compressions of several insects species of the Permian Tschekarda locality, Cisuralia mostly belong to the taeniate types (Fig. 29) Among them, Lunatisporites, produced by Ullmannia and allied coniferoids, is abundant in the guts of two large insects, a hypoperlid Idelopsocus diratiatus, and a book-louse Parapsocidi-

Fig. 28. Fossil soil of a dinosaur locality from the Maastrichtian of Nemegt Basin, Mongolia, with a saprophytic fungus allied to the extant earth-star genus Gaestrum: root-like mycelial strands and a group of fruiting bodies, one showing a radially split exoperidium (arrow and enlarged).

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Valentin A. Krassilov. Terrestrial Palaeoecology b

a

c Fig. 29. Examples of Palaeozoic and Mesozoic phytophagous insects (taxonomic assignments by A.P. Rasnitsyn): (a) Idelopsocus, Hypoperlida and (b) Sojanidelia, Grylloblattida from the Early Permian of Tschekarda, Cisuralia, feeding on taeniate pollen of coniferoids and pteridosperms, (c) Aboilus, Tettigonioidea (archaic katydids) and (d) Brachyphyllophagus, Embioptera (?), from the Late Jurassic of Karatau, Kazakhstan, feeding on pollen grains (Classopollis) and scale-leaves (Brachyphyllum) of hirmerellids, an extinct group of gymnosperms related to gnetophytes.

d

um uralicum. Both also contain Protohaploxipinus pollen grains attributed to pteridosperms (glossopterids in the Gondwana-type assemblages). The Vittatina type, produced by peltasperms (Phylladoderma and allies), is prominent in the gut assemblage of a grylloblattid Sojanidelia floralis where it associates with Lunatisporites and Protohaploxipinus. The occasional Florinites and Potonieisporites of cordaitalean or coniferoid (walchiaceous) origins are admixed to the gut assemblage of Parapsocidium uralicum. The numerical representation of Lunatisporites and Protohaploxipinus in the gut assemblages is reverse of that in the contemporaneous dispersed pollen grain assemblages, in which the former constituted a minor component. Vittatina is relatively rare in


55

the mid-Permian Tschekarda-level dispersed pollen grain assemblages but is abundantly represented in the younger Permian deposits. These findings indicate the ullmanniaceous-pteridosperm community as a major source of edible pollen for all hitherto studied Permian pollinivorous insects. Phylladoderms (Vittatina) might represent a marginal, perhaps seral, component of this community that became more prominent with the Late Permian climate change, while a remote cordaitalean-walchian Florinites–Potonieisporites community was occasionally visited by pollinivores with a wider foraging area (Figs. 30 – 32). In addition, the gut assemblages give some evidence of seasonal phenology. Mixed pollen loads, as in the Permian insects of Cisuralia (Rasnitsyn & Krassilov, 1996; Krassilov & Rasnitsyn, 1997; Krassilov et al., 1999d) indicate overlapping pollination periods in a number of coenotic populations suggesting a long flowering season. In the Early Cretaceous xyelids from Baisa, Transbaikalia, the pollen loads are typically of a single morphotype. However, two species foraged Vitimipollis produced by a proangiospermous Preflosella while in two conspecific individuals of Ceroxyella dolichocera the pollen load, although monospecific in each case, was of two different proangiosperm species. Since xyelids are early spring insects feeding on the most abundant pollen source of the season, their different pollen loads indicate a consecutive, rather than synchronous, shorttime flowering of the forage species.

Fig. 30. A schematic distribution of plants visited by pollinivorous insects over a mid-Permian coastal catena, Tschekarda locality, Cisuralia. Mixed pollen loads indicate (1, 2) an ullmanniaceous–pteridosperm assemblage as a dominant source (Protohaploxipinus and Lunatisporites), with (3) phylladoderms representing a marginal (seral) component, a forerunner of the geologically younger Late Permian vegetation (Vittatina), and (4) a cordaitalean–walchian assemblage as a remote source (Florinite and Potonieisporites). Insects: (I) Idelopsocus diradiatus, (P) Parapsocidium uralicum, (S) Sojanidelia floralis.

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a c b e

d

f

Fig. 31. Permian pollinivores with taeniate pollen grains in the gut compressions: (a) Idelopsocus diradiatus, a hypoperlid insect with a load of (b, c) Lunatisporites coniferous pollen; (d) Sojanidelia floralis, a grylloblattid insect with a mixed pollen load of (e) Lunatisporites, the same species as in Idelopsocus, and (f) Vittatina produced by pteridosperms (Krassilov & Rasnitsyn, 1997).


a

57

b

c

d

e

Fig. 32. A mixed pollen load in the gut compression of (a) a Permian pollinivorous booklouse, Parapsocidium, containing: (b) an occasional grain of a cordaitalean–walchian pollen morphotype, Florinites, (c) a pollen clump with Lunatisporites and Protohaploxypinus (the larger grain), the dominant pollen morphotypes of ullmanniaceous conifers and pteridosperms, respectively, and (d, e) a partly digested Lunatisporites-type pollen grain (Krassilov et al., 1999d).

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A more regular evidence concerning composition of source communities comes from a recurrence of fossil plant assemblages over stratigraphic sequences. Although consecutive plant assemblages are never the same, they may represent the same syntaxonomic unit in terms of species contents and numerical representations. For a source community, the recurrence of fossil assemblages is evidence of persistence reflected in the constancy of dead mass production. The dead mass bears information on diversity and relative abundance of source species. However, the dead mass contributions are not strictly proportionate to ecological significance of source species in terms of numerical dominance and biomass. Disproportionate contributions are most conspicuous at the level of specific dead mass components, such as pollen grains. Species representation in palynological spectra is a function of outcrossing potentials and the modes of dispersal. Wind-pollinated species typically produce more pollen than insect-pollinated ones (except when extra pollen is produced to attract pollinivorous pollinators). Due to such differences palynological spectra can be totally misleading, as in the case of the larch taiga spectra lacking larch pollen. Fagus is likewise systematically underrepresented in the pollen spectra of beech forests while maple and ash are seldom represented at all (Kvavadze & Stuchlik, 1996). In Neotropical forests, wind-pollinated plants constituting 2.5% of tree species produce 27% of palynological spectra (Bush, 1995). Tree line records tend to be displaced owing to a long-distance transport of arboreal pollen production (see Hicks & Tinsley, 2001). Special methods are developed to overcome such distortions (reviewed in Krassilov, 1972; Birks & Gordon, 1985). General considerations bearing on the problem of inadequate representation are the following: (1) Owing to the statistical probabilities of burial (about 1/3 of exported dead mass), fossilization and collecting (about 1/10 of the buried material), species contributing less than 3 per cent of the dead mass tend to be lost from the fossil record. (2) Because of (1), a tail of rare species is commonly cut off. A reflection of species richness in the fossil record thus depends on the shape of a particular density/diversity curve: the longer the tail of rare species, the less adequate the fossil record. Since taxonomically rich communities tend to contain more rare species than poor communities, fossil records of the latter are more adequate in this respect. (3) Since numerical representation of source species is roughly proportionate to their dead mass contribution, the relative abundances are scewed in favour of prolific dead mass producers that may or may not be ecological dominants. (4) Overrepresentation in the fossil plant record relates to plant size and longevity (trees shed more leaves than shrubs or herbs; massive pollen contribution of a few arboreal species can disguise a herbaceous dominance), phenology (deciduous plants are overrepresented in leaf assemblages), reproduction strategy


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(wind-pollinated species are overrepresented in pollen assemblages) and seral status (early successional species tend to be overrepresented owing to their relatively high dead-mass production and their tolerance of unstable habitats at or close deposition sites, such as flood-plains). (6) Preservation potentials are influenced by seasonality of dead mass production and export. Plant remains shed or dispersed during a preservation-favourable season are overrepresented. Thus, leaf litter shed before cold season has a somewhat better preservation potential than that accumulated during dry season. For this reason, summer-green plants are better recorded than winter-green plants. In the temperate amentifers, catkins shed before leaves have a preservational advantage of being transported and buried by high floods (e.g., an exceptional abundance of catkin-like strobili of dominat Jurassic trees in association with germinate Samaropsis seeds in the flood-plain facies of UstBaley locality, III.1.2). In the areas of monsoonal climate, the winter to early spring flowering supplies more pollen to marine deposits owing to the then prevailing offshore wind directions. The Classopollis abundances in offshore facies (III.1.2) might have been due to this factor.

III.2.2. Palaeosyntaxonomy The modern principles of a phytosociological classification have been laid down by Braun-Blanquet (1964 and earlier work) and his school introducing such operational syntaxonomic criteria as fidelity, significance, seral status, etc. (examples in Barkman et al., 1986; Quézel et al., 1992), some of which are applicable in palaeoecology as well. However, no one-to-one correspondence between fossil plant assemblages and their source communities is to be expected even in the case of in situ burial (II.3). We can realistically expect a crude reflection of a distinct source community or, as is more commonly the case, of mixed sources. The basic units of palaeosynecological analysis are the single-bed (bedding plane) assemblages. They are grouped in respect to their facial or stratigraphic occurrences summarily constituting a basinal taphoflora. Although consecutive single-bed assemblages are never identical, they are more or less similar in their floristic composition, species ratios and taphonomic features that bears on synecological interpretation. In other words, fossil assemblages are classifiable in terms of plant syntaxonomy. In cyclic stratigraphic sequences, fossil plant assemblages of a few types commonly recur with their respective lithofacies. Since certain operational criteria of modern plant sociology are lacking or ambiguous in the fossil plant record, the palaeosynecological classifications are relatively crude, of the nature of aspective classifications that are still applied to the present-day vegetation at initial stages of syntaxonomic analysis. An aspect is what is seen from some distance,

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being determined by dominant life forms. Regular aspective changes are brought about by seasonal dynamics (e.g., deciduousness) of the dominant forms. A closer view would distinguish structural components of growth forms (sinusia), mutualistic associations (consortia), etc. Ecological dominance in terms of structural role does not necessarily coincides with abundance. In geobotany, ecological dominants, or ediphicators, are sometimes distinguished from numerical dominants. Likewise, the aspect of a fossil plant assemblage is determined by a single or few conspicuous species. The dominance is quantitatively defined on the basis of frequency (the number of occurrences per total number of localities) and importance (the proportion of dominant occurrences among total occurrences: Krassilov, 1979). While the lowland dominants are revealed by their importance values, the upland dominants tend to be scattered over a number of localities as the high frequency but low importance species. Our guess of their dominant status over a remote source has to be confirmed by their actual dominant representation elsewhere. A scantily represented upland dominant would come forth with an upland–lowland shift of altitudinal vegetation belts. In palaeosynecology, dominant forms can be defined for a sedimentary basin, facies domain or geographic realm. The respective palaeosyntaxa roughly correspond to associations, unions and classes of the currently widely used syntaxonomic system. The nomenclature follows that by Braun-Blanquet school, but reduced to three ranks. Recommendations for formal descriptions of palaeosyntaxa are given in Krassilov (1972a) and Retallack (1981). They include the type locality, dominant species, subordinate species and facies. The names are derived from generic names of dominant species prefixed by -etum, -ion, and -tea, respectively. A co-dominant name can be added for associations (ass.). The following examples are from my studies of Mesozoic fossil plant assemblages of northern Asia, Mongolia, European Russia, the Middle East and elsewhere (Krassilov, 1967a, 1972a, 1979, 1981, etc.). Class Phoenicopsitea includes the mesotemperate deciduous vegetation of northern Eurasia (north of 50°°N) dominated by the arboreal leptostrobalean genus Phoenicopsis, a biomic dominant of high importance/frequency values over the Arcto-Mesozoic realm, forming single-species or mixed leaf-mats and represented also by allochthonous remains in the majority of localities. The co-dominants and facies dominants are ferns Coniopteris (Dicksonia) burejensis, Osmunda (Raphaelia) diamensis, horsetails Equisetum laterale (ferganensis), ginkgoaleans Ginkgoites sibirica, Sphenobaiera czekanowskiana, Pseudotorellia angustifolia, leptostrobaleans Czekanowskia rigida (setacea), and conifers Elatides ovalis (asiatica) and Pityophyllum angustifolia (longifolia). The constituent unions are: Phoenicopsion: a major zonal mesotemperate union of riparian to lower slope woody assemblages, with Phoeni-Sphenobaieretum, Phoeni-Pseudotorellietum and PhoeniPityophylletum facial (edaphic) associations. Coniopteridion (Dicksonion): an interzonal mesotemperate to warm-temperate delta front to riparian facies union of fern-marsh assemblages.


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Osmundion: a zonal mesotemperate peat-bog facies union of fern-marsh assemblages, with Osmundo-Equisetetum (ferganensae) as an edaphic association of waterlogged alluvial soil. Czekanowskion: a major zonal mesotemperate (also interzonal, but rapidly decreasing south of the mesotemperate realm) delta-plain/flood-plain peat-forming union of woody wetland assemblages, with edaphic (waterlogged alluvial to peat-bog facies) CzekanoPseudotorellietum (angustifoliae) and Czekano-Pityophylletum associations. Elatidion: an interzonal mesotemperate or warm-temperate delta-plain back-marsh facies union of a taxodiaceous woodland type, polydominant, admixed by Czekanowskia or boggy cycadophytes, spread over fern marshlands as Elati-Osmundetum (Elati-Todetum in the warm-temperate zone). Pityophyllion: an interzonal mesotemperate to warm-temperate peat-bog facies union of a bog forest type represented by pure Pityophylletum angustifoliae (longifoliae) or by a mixed Pityo-Pseudotorellietum or else in association with cycadophytes and bog ferns, as Pityo-Nilssonietum, Pityo-Osmundetum, etc. Class Ptilophylletea (Cycadeoidetea), includes the Meso-Mediterranean vegetation of a warm-temperate to xerotemperate aspect, taxonomically rich, dominated by the pachycaul cauliflorous bennettites of the pinnate Ptilophyllum, Zamites, Otozamites leaf morphotypes and by brachyphyllous conifers and gnetalean coniferoids of the scaleleaved Brachyphyllum, Pagiophyllum, Frenelopsis, Pseudofrenelopsis morphotypes, with diverse matoniaceous, schizaeaceous and cyatheaceous ferns (Nathorstia, Klukia, Tempskya, Cyathea), and with giant horsetails in the wetlands. The major unions are: Ptilophyllion: a major zonal estuarine–lagoonal to riparian facies union of xeromorphic bennettite assemblages, monodominant (Ptilophylletum) to polydominant with other bennettites or a peltasperm Pachypteris as co-dominants in estuarine associations (Ptilo-Otozamitetum, Ptilo-Pachipteridetum), with a fern-marsh component in the delta fringe associations (Ptilo-Todetum) and with Zamites–Zamiophyllum large-leaved bennettitalen morphotypes prominent in a riparian association Ptilo-Zamitetum. Brachyphyllion: an interzonal but prevailingly xerothermic, fluvial to lacustrine facies union of xeromorphic scale-leaved woodland to shrubland vegetation with diverse bennettitalean, cycadophytic and coniferoid components. The associations are numerous, not yet strictly defined, with Ptilophyllum, Otozamites, Ctenis, Baierella, Czekanowskia as co-dominants. The Late Cretaceous to Early Tertiary examples are given in Krassilov (1976a, 1979) and below.

III.3. Basinal systems The dead mass exported from a local ecosystem is imported and partly recycled by neighbouring ecosystems downslope (as in the case of eutrophication by terrestrial or-

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ganic runoff). All ecosystems of a drainage basin are thus linked by the flow of matter. They form a functional system summarily recorded by the fossilized dead mass of the basinal sedimentary fill. The spatial structure of basinal biotas is determined by geomorphology, ground water levels, soil conditions, microclimate and environmental stability. The biotic communities are sequenced in respect to the physical gradients over orographic slopes forming the pedogenic/vegetational series (catenas) that are superimposed by the mosaics of environmental spots – the bedrock, waterlogging, esposure, insolation and other local heterogeneities. Basinal habitats are also ranked according to their temporal stability, the gradients of which may also be altitudinal, as in the fluctuant flood-plain – progressively constant upslope environments. Fossil assemblages, even the prevailingly in situ ones, contain dead mass imported from more than one source communities, thus providing information on dead mass exchange over the basinal systems. Yet, since the local sources of organic material tend to be grossly overrepresented, a focus on the better-preserved material (e.g., the structurally preserved coal-ball fossils) may blur the general picture. Poor localities are not to be neglected for they summarily provide a more complete evidence of vegetation structure. A replacement of fossil plant assemblages reflects either the basin-wide restructuring or local diversion of the dead mass flows. An observed change over consecutive stratigraphic levels is by no means adequate to evolutionary significance of the event: a most conspicuous replacement may be caused by a minor shift of a zonal pattern and is reversible while a more radical restructuring transpires over a sequence of recurrent replacements as a change in the pattern of the changes. In our analysis of palaeobasinal systems we are guided by predictions based on modern examples. Such extrapolations are justified for persistent basinal structures but they can be misleading in respect to biotic responses that change over time. The Carboniferous wetland vegetation might have been relatively fine-grained (less sensitive to environmental heterogeneities), hence uniform over a wide range of basinal facies. The biotic communities of a Palaeozoic basin would then be less diverse than those of a comparable present-day basin. Or, conversely, the Carboniferous wetland zonation might have been more conspicuous and more stable than in the present-day wetlands, while the uplands were scantily if at all vegetated. A response to environmental change would then be different from what might have been actualistically expected. Substitutions of dominant species within the zonal framework might have more commonly occurred than shifts of the framework itself. Actually, over the Carbonioferous coal-bearing sequences, lepidophyte assemblages alternate with either pteridosperm or fern assemblages forming two- to three-member repeats. Their taxonomic compositions change, but their facial occurrences remain basically the same (Scott, 1977). This type palaeosuccessions projects into the Permian, being, for example, represented by alternation of Viatcheslavia/Phylladoderma assemblages in the Volga Basin. In the latest Permian, Phylladoderma was replaced by an-


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other peltasperm, Tatarina, without an appreciable change in the facial occurrence of pteridosperm belt. Yet intervention of conifer assemblages suggests elaboration of the basinal vegetation structure.

III.3.1. Seral systems Biotic communities are sustainable owing to their ability to recover after disturbances. Recovery starts with a pioneer stage and proceeds, over a series of seral (successional) stages, to an equilibrial climax state. According to the Clementsian concept of plant community (Clements, 1916 and elsewhere) as an equifinal system converging toward the climax whatever the starting point, the developmental pattern is fairly predictable for a given set of climatic variables but it changes with climate producing a sequence of climaxes, a clisere. The idea of climax is based on historical records of persistent vegetation types and is amply confirmed by the fossil record of plant communities that maintained their basic structure over millions of years. An example of such a long-term stability is Phoenicopsion, a widespread deciduous lowland formation (union) that appeared in the mid-Triassic and, after 100 m.y. of dominance, lost its significance in the mid-Cretaceous crises (though Phoenicopsis locally survived to the end of Cretaceous). Its history is paralleled by Ginkgoion, a long-standing upland climax of roughly the same temporal range (Krassilov, 1972c). These examples also support a more recent idea of polyclimax, a regional vegetation with more than one stable states. Several concurrent seral types – hydroseres, mesoseres and xeroseres – develop over a geomorphologically differentiated basinal landscape. The concurrent Ginkgoion and Phoenicopsion seem not to have been in the seral relations with each other, though the former periodically intruded the lowland vegetation during the cooler phases of Mesozoic climate cycles (Krassilov, 1972c). Sere is commonly conceived of as a reparation process taking place each time the climax is disturbed by environmental impacts. In reality, however, consecutive seral stages not only replace each other in time but also co-occur side by side over a gradient of stressful–stable habitats. Seres halted at different developmental stages form a basinal seral system. Seral developments are often argued as being too rapid to be reflected in the fossil record. This may not be true for a high-rate deposition of 5-10 mm per year, as in the case of accumulative flood-plains, and where seasonal increments are distinguishable over a considerable stratigraphic interval. It should be reminded, however, that in palaeoecology we deal with basinal seral systems rather than occasional seral developments. Seral stages are recorded as contemporaneous, as well as sequential, fossil plant assemblages replacing each other over the lateral, as well as vertical, successions of basinal facies. In the Mesozoic coal-bearing deposits of Siberia, the commonest sequence of recurring fossil plant assemblages is represented by Czekanowskia beds (leaf-mats

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of aciculate leaves in the peat-bog facies), replaced by Phoenicopsis beds (leafmats and imbedded ribbon-shaped leaves, their spur-shoot clusters in the alluvial fan to levee facies). Drifted Phoenicopsis leaves regularly admix Czekanowskia mats and vice versa, yet Phoenicopsis is far more common in allochthonous assemblages over a wide spectrum of sedimentary lithologies as a regional climax dominant would be. The relative significance of seral stages changes with basinal morphology. In the graben-like taphrogenic depressions of Transbaikalia, Czekanowskia builds cuticle coals, locally up to 2 m thick. Their patchy distribution suggests an irregular subsidence. In the broader slowly subsiding depressions, such as the intracratonic Kansk Basin of Siberia, Czekanowskia is abundant at the early stages of hydrosere, but the leaf-coals are more commonly dominated by Pseudotorellia, Sphenobaiera or Pityophyllum (Figs. 33, 34), reflecting an advanced seral system of several stages. The Czekanowskion might have been replaced by the Phoenicopsion not only upsere but also over a gradient of disturbed to relatively stable habitats. Seral developments were occasionally retarded by environmental impacts, with the pioneer stages persisting for a considerable time interval. Notably, in the Jurassic Shadaron flora of Transbaikalia, Czekanowskia prevails over a thick volcanoclastic sequence with no evidence of Poenicopsion (Bugdaeva, 1992). Here the seral development might have been halted at an early stage by volcanic activity. It is important for palaeoecological reconstructions that waterside habitats are marginal geomorphologically as well as ecologically, in terms of enironmental stability. They are frequently flooded, abraded or buried under new sediments that are then rapidly recolonized. In effect, waterside habitats are constantly occupied by pioneer seral stag-

Fig. 33. Leaf-coal types from the Middle Jurassic of Kansk Basin, Siberia: (a) Czekanowskia coal of floodplain mire, (b) Baiera–Sphenobaiera coal of a ginkgophyte peat-bog forest.


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Fig. 34. Cuticle of a ginkgophyte Sphenobaiera from the Middle Jurassic coal of Kansk Basin, Siberia.

es that are replaced inland by their successional to climax stages. Seral sequences thus partly coincide with catenic sequences.

III.3.2. Palaeocatena, general The catena, or chain, is basically a sequence of soil types reflecting a directional change in pedogenic processes upslope. Vegetation affects pedogenesis and is, in turn, affected by the latter so that vegetational catenas are closely linked to soil catenas (Greenway et al., 1969). In terrestrial palaeoecology, reconstructions of catenic sequences are based on taphonomic differentiation of proximal to distal (relative to deposition site) dead-mass sources. Catenic distance is inferred from the relative abundances and preservation of fossil material. The commonest plant fossils come from proximal members, such as wetlands. Plant material from distal members should be less abundant not only because of distance but also owing to the filtering effect of waterside vegetation. Quantification of fossil plant assemblages is thus a major source of palaeosynecological inference. Additional evidence of distance comes from preservational, as well as morphological, features related to transportation, such as durability or the prevalence of structures adapted to long-distance dispersal, like winged seeds, saccate pollen grains, etc. Contrary to what may be expected, fragmentation of leaf material is a poor guide to distance since it more commonly results from local redeposition than transport (examples below). In an assortive deposition, organic remains of different size (surface to volume ratios) come to different localities. On the other hand, an association of axial, foliar and floral organs of

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one and the same plant suggests a nearby source. Reconstructions of extinct plants, such as Caytonia (Thomas, 1925), are based on such repeated associations. A proximal source of Caytonia is amply documented by its constant occurrence in estuarine facies (Yorkshire, Kamenka, Peski, etc.: III.3.3). Hence palaeocatena is a deduction based on a sum of direct and indirect evidence, open to correction in light of new or reconsidered evidence. Present-day examples are of a decreasing value back in time. More often than not, such preconceived notions (such as an upland source of xeromorphic fossils) lead to gross misinterpretations, upside down arrangements even, of palaeocatenic sequences. No evidence can be taken for granted because of various taphonomic complications, some of which are listed below. (1) Relative contributions of proximal versus distal sources are determined not only by distance but also by their differential dead-mass production/export potentials. Owing to its prolific dead-mass export, a distal arboreal source typically overwhelms a proximal grassland source. (2) Recognition of “upland” versus “lowland” sources in palaeovegetational studies is highly conventional because an allochthonous material does not necessarily represent far-away vegetation. In the case of narrow, sharply delineated catenic belts, as in the riparian sequence of swamp cypress–palmetto belts of lowland Florida, a palm leaf would come from the distal member, but only a few metres from the deposition site. (3) Preservation quality depends not only on distance but also on a pre-burial decay, transportation media, sedimentation rates, redeposition, etc. Fragmented plant material may be due to redeposition of autochthonous remains, as in the assemblages of leaf cuticles of microbially decayed leaf litter from coastal peatbogs or watterlogged soil. The abundant dispersed cuticles of Debeya pachyderma, a Cretaceous angiospem, in strand deposits of western Sakhalin (Krassilov, 1979) and the taphonomically similar cuticle beds in the Permian of Nedubrovo, European Russia (Krassilov et al., 1999a) illustrate such preservation mode. In both cases, the cuticular structures are xeromorphic but their facial occurrences suggest derivation from a helophytic, perhaps littoral, source rather than a longdistance transport from a dryland source. (4) Rare fossils might represent either rare species of proximal vegetation or common species of a distal source. However, rarity is not a normal feature of proximal habitats, which are unstable, typically supporting pioneer forms of high density and prolific dead mass production. Even such wetland species that are considered rare (endangered) in terms of their geographic ranges are usually abundant in terms of their local population densities (Nelumbo nucifera is a familiar example). (5) Recurring sequences of fossil plant assemblages more often reflect a wetland zonation than a catenic sequence of vegetation belts. Incidentally, an alternation


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of the Late Palaeozoic to Early Triassic lepidophyte/pteridosperm assemblages (III.3) might represent a zonation of fresh-water/brackish wetland types. Finds of lepidophyte remains in pteridosperm-dominated assemblages and vice versa indicate laterally adjacent zones. In distinction, a downslope transport is not reciprocated except by an ascending air transport, which is a special situation in palynology. Harris (1965) has noted that Pachypteris, a Jurassic peltasperm, is either abundant or absent, a taphonomy that probably indicates the lowermost catenic position with little lateral mixing. At the same time, the xeromorphic morphology of Pachypteris and its frequent phytoplanktonic associates (Tasmanites) suggest a wetland, conceivably mangrove-type, habitat. A similar reasoning has made it possible to recognize other Mesozoic examples of littoral catenic belts (Krassilov, 1972a). The Carboniferous coal-measure assemblages have traditionally been interpreted in terms of an upland–lowland differentiation, with the lycopsid, calamite and pecopterid (Psaronious) assemblages representing wetlands, the cordaites, callistophytes, callipterids and coniferoids coming from drylands upslope (Gothan & Remy, 1957; Cridland & Morris, 1963; Rothwell & Mapes, 1988). However, their facies occurrences are more consistent with a microtopographic/edaphic differentiation with little if any contribution from upland sources (Scott, 1977; Pfefferkorn, 1980; Gastaldo, 1983; Collinson & Scott, 1987a, b; DiMichele & Phillips, 1985; Wnuk, 1985). A vegetational zonation, as in the present-day Caribbean example (Phillips & Bustin, 1996), might account for several types of peatland communities, such as the pioneer calamite–pecopterid brackish swamps, freshwater lepidophyte swamps and the raised bogs with a stunted sclerophyllous vegetation represented in the roof-shale assemblages (Scott, 1977), etc. Because of such complications, any catenic reconstruction based on a single or few examples has to be considered as a working hypothesis to be tested by further observation. Sedimentary cyclicity provides an ample opportunity for such testing because catenic successions often recur from one cyclothem to another. Collected over a number of sedimentary basins, such data reveal ecologically and evolutionarily meaningful variations of catenic systems.

III.3.3. Plaeocatenas of regressive sequences Evidence of catenic structures comes, first of all, from regressive sequences recording retreats of epeiric seas. Plants encroach across the emerging land in a sequence reflecting their seral/catenic status. Comparisons of concurrent regressive sequences from different basins reveal regularities in evolution of catenic structures over time and geography. The present-day examples, such as a succession of levee fern beds → Taxodium–Nyssa (Gordonia, Myrica) arborial wetland → palmetto belt several metres

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wide → pine woodland → broadleaved Platanus–Liquidambar river-bench forest of the natural vegetation of Florida are widely used in reconstructions of Cenozoic catenas. A few Mesozoic and Palaeozoic examples are given below. Kamenka. The first reconstruction of a Mesozoic catena was based on a succession of fossil plant assemblages over a sequence of the Bathonian to Callovian shallow marine to freshwater deposits of Kamenka, northeastern periphery of Donetsk Basin (Krassilov, 1972a). Sand bar facies with marine bivalves at the base of the sequence yielded fragmentary pinnae of Coniopteris hymenophylloides, a Mesozoic dwarf tree fern. The back-swamp ferruginous deposits upsection contain both upright and piled stems of Equisetum columnare, a giant horsetail, preserved as moulds or casts. Common in the Equisetum ironstones are leaf fragments of Cladophlebis kamenkensis, an osmundaceous fern, and Ptilophyllum pecten, a pinnate bennettitalean leaf morphotype with revolute pinnae margins, thick cuticles and peltate trichomes. A single-species locality of the latter was found in a gypsiferous clay bed laterally replacing the bog-iron facies. Imbedded in the ironstone are scattered leaves of Podozamites, a broad-leaved riparian conifer, and Czekanowskia, a pioneer wetland gymnosperm with aciculate leaves that are scattered or clustered on deciduous spur-shoots. Fragmentary scale-leaved and needle-leaved conifers are allochthonous remains of the next catenic assemblage. The overlying light-brown silty clay with coaly lenticles and embedded plant debris, penetrated by vertical roots, attests to a bog forest growth over the horse tail mire buried with a crevasse-splay alluvium of a meandering deltaic channel. The fossil plant assemblage, representing the maximal diversity zone of the coastal wetlands, is dominated by Elatides expansa, a taxodiaceous conifer with brachyphyllous shoots and ovoid seedcones, locally in association with Pityophyllum, a needle-leaved pinaceous conifer, Nilssoniopteris, a xeromorphic bennettitalean leaf morphotype with thick lanceolate pubescent marginally revolute leaf-blades, and an Osmunda-like Cladophlebis denticulata. Occasional identifiable fragments belong to Ctenozamites, a xeromorphic cycad, Nilssonia, a heteroblastic cycadophyte with narrow lanceolate leaves, Ginkgoites, a flabellate Ginkgo-like morphotype, and Phoenicopsis, a leptostrobalean gymnosperm with ribbon-shaped leaves. The regressive sequence is topped by white kaolinitic clay with ferruginous impressions of crowded fern leaves. The fern assemblage contains complete leaves of Todites williamsonii, an osmundaceous tree-fern, with subordinate fragmentary Cladophlebis (Todites) haiburnensis, another representative of the family, and Dictylophyllum rugosum of an extinct family Camptopteridaceae. Elatides is fairly common, whereas few other gymnosperms are represented by occasional remains. This system of catenic/mosaic plant communities thus includes, (1) Coniopteridion, an ubiquitous delta fringe fern marshland supplying plant material to offshore deposits; (2) Equisetetum collumnare, a horsetail coastal mire saturated with ferric compounds bound by the decaying plant tissues, laterally replaced by (3) Ptilophylletum pecteni, a littoral, mangrove (?) community, alternating with (4) Podozamitetum/Czekanowski-


69

etum, a pioneer stream-side growths fringing distributary channels, represented by allochthonous material alone, replaced toward the delta interior by (5) Elatididetum expansae, a back-marsh bog forest, of the widest facies range, merging on (6) ElatidiPityophylletum, a coniferous peat-bog assemblage, and (6) Elatidi-Nilssoniopteridetum, a bennettitalean peat-bog shrubland. The allochthonous Phoenicopsis and Ginkgoites represent dryland vegetation. Of special interest is the succession of osmundaceous ferns: Cladophlebis kamenkensis – C. denticulata – Todites (Cladophlebis) williamsonii – Todites (Cladophlebis) haiburnensis, perhaps reflecting a zonation of the fern-marh community partly extending over the horsetail mire and ascending as understorey to the Elatidion. Many uncertainties remained that could only be resolved by comparison with analogous successions in other sedimentary basins. Umalta. In the mid-Jurassic (Callovian) regressive sequence of Umalta, Bureya Basin, Russian Far East, Dicksonia (Coniopteris) burejensis and Osmunda diamensis are dominant in the basal deltaic facies that rest conformably on shallow marine sandstones with ammonite shells and still contain marine phytoplankton, as well as occasional limulid remains. The catenic position of the fern assemblage might have been as at Kamenka (above). Horsetails, much smaller than at Kamenka, patchily occur as an edaphic variant of the coastal marshland. Bulk maceration of a fern bed soil brought fossil mites, diverse bryophytes of the ground cover (Krassilov, 1973c), fern sporangia, as well as scattered scale-leaves and shoot fragments of Elatides ovalis, the latter dominant in the next catenic assemblage upsection, the back-marsh Elatidion. This, however, is less prominent than at Kamenka. Conversely, Czekanowskia (with Leptostrobus-type cupules) is more abundant in the flood-plain deposits. Leaf-mats grading on leaf-coals consist of Czekanowskia accompanied, or locally replaced, by a needleleaved conifer Pityophyllum or a narrow-leaved ginkgophyte Pseudotorellia. Their frequent associates are Osmunda diamensis and Sphenobaiera, a ginkgophyte with deeply dissected leaves. Nilssonia, Podozamites and their reproductive remains cooccur with Czekanowskia in the coarser levee facies. The peatland vegetation is thereby reconstructed as a Pityo-Pseudotorellietum (Sphenobaieretum) mosaic marginally admixed with a riparian Czekanowskietum and with a fern growth as understorey or over the gaps. All plant assemblages of this stage contain allochthonous remains of Phoenicopsis, which comes to dominance over a sequence of the fluvial sandstone/siltstone cyclothems above. Seral links between Czekanowskia, a pioneer form of the delta-plain to flood-plain alluvium, and Phoenicopsis, a lowland climax dominant, are confirmed by repeated observations of their stratigraphic successions and concurrent facies changes in a large number of Siberian localities (II.7.1). Over the Umalta Section, their succession is catenic, as well as seral. The regressive sequence is overlain by a series of coal-bearing cyclothems, in which an avulsion autocyclicity is periodically truncated by a thick lamellate sand-

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stone marking a higher order erosion cycle and containing mass accumulations of Phoenicopsis. Up the fining-upward sequences, Phoenicopsis is regularly joined by osmundaceous ferns, nilssonias, ginkgophytes and Pityophyllum conifers of the peatland mosaic. Ginkgo-like leaves are numerically subordinate but occur as allochthonous material in a wide range of facies. Their dominance in a few Jurassic and Cretaceous localities might have been due to a downslope shift of altitudinal vegetation belts. With accumulation of comparative data, habitats of extinct plants come out as clearly as if they were growing before our eyes. Thus, over the Mesozoic periods, Nilssonia and Podozamites formed willow-like riparian thickets while the narrow-leaved bennettite Nilssoniopteris was a heath-like bog shrub. Certain woody plants commanded a wider range of habitats, as Czekanowskia, a vigorous colonizer of delta-plains, or Elatides, a dominant of coniferous bog forests. Their catenic/seral successors were, depending on the climatic zone, either the deciduous Phoenicopsis forests or the brachyphyll– cycadophyte shrublands that climbed upslope to the luxurious Ginkgo forests, a mesic upland climax. Laibin. A Palaeozoic example comes from the Permian of Laibin Province, South China. In a series of outcrops on the banks of Hongshui River, the Upper Permian (Lopingian) deposits are represented by two sedimentary megacycles corresponding to the Heshan and Talung formations (Jin et al., 1998; Shen et al., 1999). Fossil plants occur in the shallow water marine to paralic facies in the upper members of both megacycles. In a black shale bed of Talung Formation, a few ullmannioid shoots, Steirophyllum, were found in association with abundant ceratitid impressions. In the upper Talung member above, the conifers Hermitia and Steirophyllum are preserved as moulds on the bedding planes of tuffaceous shales with distinct ripple marks and trace fossils. The lowest stand of the Talung Sea is marked by coal-bearing deposits on top of the marine to paralic regressive sequence. The fossil plant horizon is about 50 cm thick consisting, from base up, of a root bed over a lenticular coal, with Paracalamites stems and abundant, though crumpled, pecopterid leaves, overlain by a ferruginous shale with a gigantopterid Gigantonoclea guizhouensis and the numerically subordinate marattialean fern leaves Pectinangium accompanied by few coniferoids. Of these assemblages, the Paracalamo–Pecopteridetum represents a widespread interzonal type of Permian wetlands dominated by arthrophytes and pecopterids. Its replacing Gigantonoclietum is typical of the Cathaysian gigantopterid province, where it occurs both in the coalbearing and redbed facies. Their succession (Fig. 35) is here interpreted as a seral/ catenic sequence following the retreating sea. It started with a pioneer arthrophyte– pecopterid coastal peatland community that was then invaded by a backswamp gigantopterid community. The conifer remains are fragmentary suggesting a distant source upslope.


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a d

b

c

e

Fig. 35. Seral succession in the latest Permian Talung Formation, Penlaitan Section, Laibin, Southern China: (a) pecopterid remains at the base of the regressive sequence above marine deposits, (b) Gigantonoclea guizhouensis, a dominant species of the next successional stage, (c) Gigantotheca, a fertile leaf of Gigantonoclea from the latest Permian of Fujian (courtesy of Prof. Z. Yao), (d, e) a glossopterid leaf or cataphyll from the Gigantonoclea bed of Laibin.

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III.3.4. Facial modifications As the above examples imply, the catena is a universal structural pattern that can be, in principle, recognized in any terrerstrial sedimentary basin. Variation in catenic structures over space and time reflects environmental changes that affect both vegetational patterns and taphonomy. In principle, taphonomic changes are recurrent, whereas structural changes are directional. In practice, both types intermingle. A contribution from a distal source might increase inversely of the filtering effect of coastal vegetation. High tides of a rising sea, as over the present-day mud flats of North Sea shores, would have destroyed coastal vegetation, making inland contributions more evident. However, over the Caspian Sea fluctuations of the 1950’s to 1980’s, the coastal wetlands were reduced during regressive phases but recovered with a transgression– ponding of ground water reservoirs by the rising sea (Glazovski, 1990). Winter monsoons increase the contribution of air-borne particles from distal inland sources, whereas summer monsoons carry them up the maritime catenas. Their relative effects would change with fluctuations of sea surface temperature and other climatic variables. A climatically driven shift of altitudinal vegetation belts would either increase (sometimes dramatically, as in the case of upland–lowland shifts) or decrease (with an expansion of lowland vegetation upslope) the contribution of distal sources. To distinguish such taphonomic effects from evolutionary changes, we must compare vicarious catenas of different facies domains. Thus, deltaic catenas differ from estuarine catenas in the relative development of freshwater/saline marshland belts. At Kamenka (above), the delta fringe belt is represented by a Coniopteridion fern marsh, with Ptilophylletum as a rare edaphic variant. At the contemporaneous Peski locality, Moscow Region to the north, the clayey estuarine clays, pyritic, with compressed plant debris and dispersed cuticles, fill karst cavities of Carboniferous limestones. Here the Coniopteridion is lacking while the Ptilophylletum is a major source of plant compressions. The dominant species Ptilophyllum riparium is similar to P. pecten but with narrower leaves covered with peltate salt glands (Gordenko, 1999). It is joined by Otozamites and xeromorphic Nilssoniopteris (several species, with scattered or intercostal stomata) to represent a salt marsh or mangrove peatlands. In the Jurassic of Eurasia, the cycadophyte diversity decreases northward, yet this trend is countered by the facies. This type facies modifications of catenic sequences was common through the Late Palaeozoic and Triassic sequences where fern assemblages regularly alternated with either lepidophyte assemblages (successively dominated by Lepidodendron, Sigillaria, Vyatcheslavia and Pleuromeia) or pteridosperm–conifer assemblages. In the Permian to Triassic redbed sequences of Volga Basin, freshwater wetlands are represented by stems and megaspores of Vyatcheslavia and other lycopsids. Their antagonists are the xeromorphic peltasperms Phylladoderma, Tatarina and coniferoids, often preserved


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as accumulations of dispersed cuticles. In the transboundary Permian/Triassic Nedubrovo sequence (II.7.2), such cuticle beds contain microplankton and occasional lycopsid megaspores Otynisporites (Krassilov et al., 1999a, 2000). In distinction, over the contemporaneous inland sequences of Tunguska Basin, the proximal catenic communities are invariably dominated by ferns (Sadovnikov, 1997). In the Late Cretaceous paralic sequences of western Sakhalin, deltaic wetlands are represented by the Filici-Nilssonietum (the fern beds of Cladophlebis arctica, Anemia dicksoniana, Cyathea sachalinensis admixed with or occasionally dominated by nilssonias), a marsh assemblage confined to the coal-bearing cyclothems with marine shelly fossils. They also contain drifted remains of Sequoia, platanoid angiosperms and other riparian plants. An alternative type of paralic assemblage is represented by accumulations of xeromorphic Debeya cuticles in black-shale facies representing a salt marsh to mangrove wetland (Krassilov, 1979). In the above examples, the relative development of one or another type of catenic sequences reflects a long-term sea-level trend, the deltic type advancing with regression, the estuarine one – with transgression. However, a similar facies modification can be driven by climate changes, a drier climate inflicting an expansion of salt marshes at the expense of other edaphic varieties. Actually the catenic restructurings reflect an interplay of interdependent sea-level and climatic factors.

III.3.5. Catenic systems over time and geography When a number of vicarious catenas are brought side by side, as in Table 1, we can read in horizontal lines and vertical columns both the environmental modifications and evolutionary innovations. Such spatio-temporal catenic systems form a framework of evolutionary studies (III.3.6). Thus, a comparison of the warm-temperate (Kamenka) and temperate (Umalta) catenas (III.3.3) shows that both deviations and deletions of catenic belts took place (Table 1). The fern-marsh belt is the least affected, with minor replacements of Coniopteris by Dicksonia and of Todites by Osmunda of the same respective families. The horsetail mires differ in stem dimensions mostly. At the same time, the Ptilophyllion is deleted from the temperate catenas while the brachyphyllous Elatidion is considerably reduced. The cosmopolitan Pityophylletum is modified as a Pityo-Pseudotorellietum. The pioneer riparian Czekanowskion and its successive Phoenicopsion are raised to the status of dominant lowland types. The Ginkgoion is a far more important source of transported material than at Kamenka, perhaps reflecting a downslope shift of the upland belt. Conversely, both Brachyphyllion and Ptilophyllion are expanded relative to Kamenka in the xerothermic zone of southern Europe, Mediterranean and Central Asia (e.g., Krassilov, 1982). These examples show that climatically driven restructurings of

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Table 1. Vicarious catenic systems of Jurassic to Early Cretaceous vegetation. Locality

Primorye Far East Gobi Mongolia Bureya Far East Kamenka Ukraine Ust-Baley Siberia Primorye Far East

Age Coastal marsh Con1+ RfOn Bs

NeocomianAlbian NeocomianAptian Late Jurassic- Con1 Aptian Late midCon2 Jurassic Mid-Jurassic Con1 Early Jurassic Nc

Coastal mire OsCl

Catenic FloodPlain Br +Cyc

belts BackPeatBench swamp bog El1+Nl1 Pit+Ns1 Br+ Pz

Slope Gk1

PtOt

Br +Cyc

El1+Nl2 Pit

Br1+ Gk1 Pit

Eq1+ Os

Cz

El2+Nl2 PitPs

Ph+Gk2

Gk3

Eq2+Pt1 +TdCl Ag

Cz

El2+Nl2 Pit

Pz

Gk2

Cz

El2

Ph+Sph

Cz

El2+Nl4

Pit

Pit + Pz

Ph +Gk3

Abbreviations: (Ag) Aegiantetum, a pioneer marsh-type gnetophyte community, (Bs) Baisietum hirsutae, a proangiosperm marsh-type community, (Br) Brachyphylletum, (Con1) Conipteridetum (Dicksonietum) burejensae, a fern marsh, (Con2) Conipteridetum hymenophylloidae, a fern marsh, (Cyc) Cycadeoidetum, a pachycaul vegetation type, (Cz) Czekanowskietum regidae, a pioneer community of flooded riparian habitats, (El1) Elatidetum asiaticae, a taxodiaceous arboreal wetland, (El2) Elatidetum ovali, a taxodiaceous arboreal wetland, (Eq1) Equisetetum sibiricae, a horse-tail marsh, (Eq2) Equisetetum columnarae, a giant horse-tail marsh, (Gk1) Ginkgoitetum pluripartitae, a ginkgoalean forest, (Gk2) Ginkgoitetum digitatae, a ginkgoalean forest (Gk3) Ginkgoitetum sibiricae, a ginkgoalean forest, (Nc) Neocalamitetum, an archaic horsetail marshtype community, (Nl1) Nilssoniopteridetum rhitidorhachisae, a bennettitalean peat-bog community, (Nl2) Nilssoniopteridetum vittatae, a bennettitalean peat-bog community, (Os) Osmundaetum diamensae, an osmundaceous fern mire, (OsCl) Osmundaetum (Cladophlebietum) arcticae, an osmundaceous fern bog or mire, (Ph) Phoenicopsietum, a lowland climax forest, (Pit) Pityophylletum, a coniferous peat-bog forest, (Ps) Pityo-Pseudotorellietum angustifoliae, a coniferous-ginkgoalean peat-bog forest or shrubland, (Pt) Ptilophylletum pectinae, a xeromorphic bennettitalean wetland, (PtOt) Ptilo-Otozamitetum lacustrae, a xeromorphic bennettitalean wetland, (Pz) Podozamitetum, a riparian coniferous shrubland community, (RfOn) Ruffordio-Onychiopteridetum psilotoidae, a fern marsh, (Sph) Sphenobaieretum, a ginkgoalean peat-bog forest or shrubland, (TdCl) Osmundaetum (Cladophlebietum) kamenkensae, an osmundaceous fern mire, (TdCL) Toditetum (Cladophlebidietum) kamenkensae, an osmundaceous fern mire.

catenic systems are more prominent at mid-level belts, with both the lower and upper belts less affected by zonal climates. Tracing the catenic structures over geological ages, we find that evolutionary changes are never uniform over the catenic belts, as some of them evolve more rapidly than others. Like in the geographic variation, the temporal changes are reductions, expansions, deletions, substitutions or altitudinal shifts. Thus, a comparison of the Jurassic and Early Cretaceous catenas of warm-temperate ecotonal zone about 50ºN shows a similar transition from the Coniopteridetum fern marsh to Elatidetum bog forest. However, a new type of fern marsh, Ruffordio-Onychiopteridetum, became more prominent in the Cretaceous. The mid-catenic diversity


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of plant communities increased considerably due to the newly formed brachyphyllous association of shrubby confers (Athrotaxietum) and cycadophytes (Krassilov, 1967a). The Ginkgoion position might have been higher upslope as the remains are rare, but increase in the Albian indicating a downslope shift (Krassilov, 1967a, 1972a). During the mid-Cretaceous, taxonomic replacements took place in all catenic belts, with few conservative communities, such as Filici-Nilssonietum, still maintaining their catenic position. The Coniopteridion was reduced to occasional patches while the Ruffordio-Onychioptridetum was replaced by Anemio-Cladophlebietum arctici, with the dominant species Anemia dicsoniana, a possible phylogenetic successor of Ruffordia goeppertii, and Cladophlebis arctica descending from the Jurassic to Early Cretaceous osmundaceous lineage. In the brachyphyllous communities, the archaic confers Elatides and Athrotaxites were replaced by Sequoia, Geintzia and Parataxodium that rapidly expanded over the ranges of a degrading Phoenicopsion, forming a new type of lowland climax vegetation. With the likewise rapidly rising angiosperms as undergrowth, the Late Cretaceous conifer forests diverged into the temperate Parataxodio-Platanifolietum and the evergreen SequoioLaurofolietum types (Krassilov, 1979; IX.9.1). In these examples, the consecutive dominant species of persistent catenic belts, such as the fern-marsh belt, represent either continuous phylogenetic lineages or different, but morphologically convergent lineages. In either case, morphological gradualism was associated with persistence of vegetation structures. In contrast, the radical catenic changes are accompanied by major morphological innovations. Thus, in the central-eastern Asiatic area of extensive Late Jurassic to Early Cretaceous volcanic activity (Transbaikalia, Mongolia, northeastern China), the Jurassic-type fern marshes (Coniopteridion) were substituted by the newly formed communities involving the pioneer components of both Brachyphyllion and Ptilophyllion. New diminutive grass-like or sedge-like proangiosperms, such as Baisia, Eoantha, Vitimantha, Baisianthus, Preflosella etc. appeared in the process (Krassilov, 1997a; Krassilov & Bugdaeva, 1999, 2000). In this way, a radical restructuring of a catenic system might have procured evolutionary changes.

III.3.6. Evolution over catenic systems It follows from the above examples that a succession of plant communities is meaningful when it is related to the developmental (seral) and spatial (catenic) systems. Vegetation changes typically amount to elaboration, reduction or restructuring of such systems. A seral stage, a catenic belt or a canopy stratum is added or deleted, or shifted to a different position, with modifications by autochthonous evolution or through invasions of new elements. Plant migrations are controlled by pre-existing seral/catenic patterns. At the borders of floristic realms, sequential catenic belts may even have different geo-

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graphic origins, as the “Malaysian” lowland – “Australian” upland belts of New Guinea (Schuster, 1972), an evolutionary perspective interface. In a stable vegetation system, both seral and catenic sequences tend to get longer by an interstitial build-in, such an expansion of the Jurassic brachyphyllous belt upslope and its differentiation into several catenic communities over the Cretaceous (above). With a major environmental change, new belts are added by a terminal build-up. Rising mountain ranges drive the catenas upslope, while a hammada-type community can be added on a rocky ground of the piedmonts. The opposite tendency is manifested by truncation of a sere or by marginal cut-off of catenic belts. Certain members of vegetation sequences, most often end-members, are periodically reduced and then repaired. Where evolutionary innovations arise during the cutoff/repair cycles, they are confined to the end-members. In most cases studied, new plant communities develop either from an early stage of a seral system or from understorey (that can also be seral to the overstorey), or else from marginal catenic variants. Early land plants formed an essentially single-storey vegetation of a small aboveground biomass and a relatively large underground biomass. With the build-up of a multistratal canopy and reallocation of biomass to aboveground organs the slightly modified early communities survived as the marginal/early seral stages. Evolution of the plant body – secondary growth, elaboration of vascular system, planation of photosynthetic organs – was primarily related to elevation of overstorey, with modest reproductive innovations. The overstorey trees of the mid-Devonian–Tournassian forests were freesporing (including the so-called progymnosperms that resemble gymnosperms in their wood anatomy alone), while the precursors of seed-plants (e.g Lenlogia: Krassilov & Zakharova, 1995) appeared in the relatively primitive telomic understorey seral to these forests. The Early Carboniferous evolution of gymnosperms was still confined to the edaphic periphery and/or early seral stages of the free-sporing wetlands (pteridosperms in the roof shales of lepidophytic coals: Scott, 1977). During the Pennsylvanian and Early Permian, gymnosperms gradually invaded all recognizable catenic belts. In the Late Permian regressive sequences, gigantopterids replaced the pecopterid–calamite marshes as a successional phase (III.3.3). Adaptive radiation of Mesozoic gymnosperms was related to differentiation of new catenic systems, such as the scale-leaved Brachyphyllion with cycadophytic helophytes, riparian Phoenicopsion, upland Ginkgoion and the seral/edaphic Czekanowskietum, Pityophylletum, Pseudotorellietum, etc. (III.3.1). Free-sporing dominance was retained in the fern–horsetail marshes, a widespread wetland type, occasionally seral to gymnosperm wetlands. With a decline of fern-marsh communities the belt was invaded by the pioneer elements of the bennettite- gnetophyte wetlands that gave rise to proangiosperm communities (Krassilov, 1997). A mid-Cretaceous wave of catenic/seral replacements brought their descendent early angiosperm communities to the understorey of conifer forests and downslope to hydrophilous vegetation downslope.


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Before the advent of angiopserms, the taxodiaceous rainforests (e.g., the Late Jurassic–Early Cretaceous Elatididion) formed a single-stratum overstorey, with a ground vegetation of bryophytes, small horstails and stunted tree ferns derived from the open fern–horsetail marshes. The cycadophytes were poorly represented, perhaps reflecting an ephemeral seral stage rapidly overshadowed by the coniferous canopy. Angiosperms filled in the vacant niche of shade-tolerant understorey developing a broadleaved, “platanophyllous” habit in the process. The xeromorphic cycadophyte–brachyphyllous shrublands widespread in the midTriassic to mid-Cretaceous Meso-Mediterranean realm might have derived from the symphyllogenetically related sclerophyllous helophytic communities. Their expansion over the Mesozoic drylands coincided chronologically with the appearance of the giant dinosaurian herbivores and was probably enhanced by overbrowsing and deforestation. The thermophilic cycadophyte species were winnowed by the Albian cooling, rendering the xeromorphic shrublands undersaturated, open for the xeromorphic varieties of early angiosperms. The Mesozoic freshwater macrophyte communities consisted of charophytes, mosses and quillwarts, all submerged. A new structural level was added with floating ferns that converged on their antecedent lycopsids by acquiring heterospory with adaptations to long-distance dispersal by water-borne spores (amphisporions in the case of Heroleandra: Krassilov & Golovneva, 1999, 2000). These free-sporing colonizers were soon joined and partly eliminated by floating angiosperms forming a later seral stage. An increasing complexity of floating vegetation reduced light available for submerged plants inflicting a turnover in the rooted macrophyte community that was in effect invaded by aquatic angiosperms. Thus, the ecological and morphological diversity of the newly arising group was rapidly built up with its spread up and down the catenic belts. It came to dominance with further transformation of catenic belts to the end of the Cretaceous. In particular, the appearance of broadleaved riparian vegetation in the Early Palaeocene amounted to a transformation of the Cretaceous lowland forests by a cut-off of the conifer overstorey. The resulting low diversity Platano-Trochodendroidetum is an example of an undersaturated community, the short-term stability of which might have been owing to a highly unstable environments over the Cretaceous/Tertiary transition. Simultaneous innovations at the upper end of the catena are vitnessed by differentiation of the taiga/birch altitudinal belts in the montane areas of the Pacific coast (IV.2.3; IX.9). Most prominent catenic transformations over the Tertiary were owing to aridization of either lowlands or uplands in response to the global climate changes. Sequoiadendron, for instance, was part of a lowland vegetation until an aridization of the lowlands drove it to the slopes of Sierra Nevada (Axelrod, 1986). Similar developments took place in southern Australia where the highland rainforests, with Nothofagus emerging from the understorey (Specht et al. 1992), have appeared in conjunction with the xeric lowland vegetation. In both cases, xeromorphic plant communities might have derived from

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preadapted marginal scleromorphic variants of laurophyllous vegetation (Axelrod, 1975, 1989; Hill, 1990). Dominant forms of the Madrean chaparral, as well as the Tethyan maquis are related to the seral undershrubs of sclerophyllous woodlands which appeared before the late Neogene climate change (Axelrod, 1989).

Chapter 4. Palaeoecogeography

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IV. PALAEOECOGEOGRAPHY Biogeographic studies are focused either on taxonomic composition of floristic/faunistic assemblages reflected in the separate phyto- and zoogeographies, or on life forms and biotic communities as the operational units of ecogeography. While taxonomic biogeography is related to evolutionary histories of plant/animal taxa, the ecogeography conveys a history of ecosystems related to climatic zonation and other ecologically meaningful global variables. The major palaeophytogeographic divisions that were initially recognized on the basis of so-called leaf taphofloras, are biomic rather than floristic, for they are based on dominant leaf morphologies, the primarily life form cheracters. They reflect the aspects of regional plant assemblages rather than taxonomic similarities or the ranks of endemic taxa, as a floristic classification would require (Krassilov, 1972a, 1972b, 1997c, 2000b). The Arcto-Tertiary flora of Engler is the first widely accepted division of the kind. It is equivalent to the temperate broadleaved (nemoral) forest biome that is also recognizable over the Mesozoic (the Arcto-Mesozoic Phoenicopsis realm: Krassilov, 1972b) and further back in time (IV.2.2). Climatically, the nemoral biome corresponds to a humid temperate zone. South of the nemoral zone, there is a no less persistent realm of evergreen and semideciduous arboreal biomes with a prominent xeromorphic element. The climatic conditions were warm-temperate to subtropical, of the Mediterranean type, with a dry summer season. The Neogene precursors of Mediterranean vegetation extended over southern Europe and central Asia to northern China and, as the Madro-Tertiary flora, to western North America (Axelrod, 1958). Their Mesozoic equivalents occupied the same broad latitudinal zone marked by the pachycaul bennettites with pinnate leaves (the MesoMediterranean Cycadeoidea realm: Krassilov, 1972b). On the other hand, the ecological analogues of the present-day tropical rainforests only epizodically appeared in the fossil record, while such biomes as the northern conifer forests (taiga) or tundra are of relatively recent origins (IV.2), with no ecological equivalents in the Mesozoic, but with possible forerunners in the Permian. Therefore, a priory reasoning is not productive in the evolutionary ecogeography. While the uneven distribution of incoming solar energy over the globe grants an ecogeographic differentiation, its pattern depends on a number of factors none of which remains constant over times.

IV.1. Biomes Biomes are the higher order ecogeographic units defined primarily by the physiognomy of their dominant vegetation. The physiognomy is, in turn, determined by the dominant growth forms (conifers, broadleaves, sclerophyllous trees/shrubs, mesic/xeric grass-

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lands, etc.), phenology (evergreen/deciduous, summer green/winter green) and a few structural features, such as patchiness, continuity, density, and canopy structure (closed/ open, single/multistory, etc.). Most of these features are recognizable, with different degrees of confidence, in the fossil record, although classification of past biomes is by necessity more generalized than that of the present-day biomes (e.g., UNESCO, 1973). A palaeovegetation analysis is based on plant life forms rather than taxa. Extant life form characteristis are conventionally extrapolated over the Tertiary taxonomic allies. For a pre-Tertiary vegetation, leaf morphology and phytosociological characters, such as taxonomic diversity and dominance, are of an increasing importance back in time. Fossil leaves rank among the most frequent plant macrofossils providing important life form information. Their morphology relates to structural position (canopy or understorey), habitat, climate, pollination mode, etc. (III.1.3). As will be shown in (IV.3.1), the Permian “geofloras” are not strictly definable in terms of taxonomic distinctions because of morphological parallelisms of their emblematic leaf types that might belong to different, even not closely related, taxa. On the other hand, a biomic interpretation is consistent with the far-reaching convergence of prevailing leaf morphologies in respect to climates, insect symbionts and the “geographic parallelism” supposedly conferred by microorganisms transducing genetic material (Went, 1971). Information on growth forms and seasonality comes also from fossil axial organs, while pollen grain morphologies reflect pollination modes that are related to vegetation structure and climate. Faunistic records may give critical evidence on vegetational structure (e.g., abundant remains of gregarious herbivores indicate open vegetation while a ratio of browsers to grazers among them is suggestive of a woodland/shrubland developments relative to grasslands). A biomic significance of the pre-Tertiary growth forms is deduced from their geographic and facial ranges. Incidentally, the equisetalean stems more than 3 cm thick, as well as the massive tree fern trunks (Tempskia), are found outside the temperate ArctoMesozoic realm while their counterparts in the latter are slender horsetails and dwarf tree ferns. In Asia, the pachycaul cycadeoids are common south of 50÷N, but are lacking further north where the dominant gymnosperms were shedding leafy spur-shoots indicating a leptocaul heteroblastic deciduous life form (Krassilov, 1972a). This type deciduousness commonly occurs in temperate summer-green biomes, whereas the swallen petioles in compound leaves of pachycaul bennettites may indicate leaf shedding in a dry rather than cold season. In angiosperms, several dozens of leaf characters might have been of certain climatic significance, but only few proved useful as distinctions between evergreen and deciduous life forms, such as wind roughness of peltate blades with flexible petioles. Serrate leaf margins are more common in deciduous angiosperms, but they might have a different ecological meaning in the past, more consistently related to plant – insect interections (III.1). Leaves with narrow blades, thick cuticles and protected stomata are relatively frequent in pre-Cenozoic plants. They suggest a greater extent of xeromophytic biomes, but might also be due to a higher than present atmospheric CO2 level or, alternatively, to


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overrepresentation of scleromorphic helophytes from waterlogged habitats. Such uncertainties can only be resolved with pedogenic and taphonomic information. Fossil soil is the best guide to biomic structures and functioning, in particular over the facies realms such as redbeds, that are unfavorable for fossil plant preservation, but commonly contain either in situ caliche-type paleosols or redeposited carbonate rhizocretions or both. However a frequent burial of soil horizons is due to instability of the basinal drainage systems in which full soil profiles scarcely developed, while the aborted hydroseric ones were grossly overrepresented. Seasonality of phenological events can be inferred from taphonomy of leaf assemblages, such as leaf-mats (Fig. 36). Varved sediments containing pollen in their spring increments indicate early flowering, while pollen in offshore deposits suggests flowering during a season of off-shore winds (III.1.2). Leaf roofs – the bedding planes densely covered with overlapping leaf impressions – indicate mass accumulations and rapid burial of leaf litter that normally results from winter deciduousness. Coal is a massive accumulation of

Fig. 36. A leaf-mat of Tiliaephyllum in levee sandstones of Tsagajan Formation, the basal Palaeocene, Amur Province, Russian Far East (Krassilov, 1976a).

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plant litter due to high dead mass production and deposition rates that are controlled by a number of biomic variables, including a processing of primary production over the herbivore and detritivore trophic chains. Coal types are determined by deposition environments to a greater extent than by their source vegetation. Yet some coal-types are confined to certain biomes. Leaf coals are typical of the summer-green deciduous forest zone, e.g., the Czekanowskia coals of Arcto-Mesozoic realm, whereas resinites mostly occur over the ranges of the summer-dry pachycaul Meso-Mediterranean biomes. The reconstructions of palaeobiomes are thus based on constant associations of biotic, pedogenic and taphonomic characters.

III.2. Biome evolution The evolution of biomes is naturally related to the history of terrestrial climates, but the relationships, as the following examples illustrate, are not straightforward ones, but are tightly linked to plant phylogeny and are mediated by plant – animal interactions. Even if biomes are constantly recignized over times, as over the succession of temperate Arcto-Permian, Arcto-Mesozoic and Arcto-Tertiary types, they may not represent a continuous line of discent, but were rather formed by combining elements recruited from different types. Major phylogenetic events tend to occur at an initial stages of biome evolution and the “centres of origin” are the areas of floral/faunal mixing and transformation giving rise to new biomes.

IV.2.1. Tropical rainforests Contrary to a popular view of past vegetation as one global tropical rainforest, palaeobotanical data suggest a discontinuous evolution of the rainforest type biomes. The Carboniferous (Westfalian) type European and North American plant assemblages are commonly considered as representing tropical vegetation. However, more ecologically orientated researchers such as Potonié (1951, 1953) and Jongmans (1958) have argued that extensive swamp forests, like that of the Carboniferous, are atypical of the present-day equatorial zone. Actually, the Westfalian assemblages neither show any rainforest type leaf morphologies nor are comparable with tropical rainforests in terms of taxonomic diversity and dominance. All the overstorey trees were microphyllous freesporing species, whereas the macrophyllous species, both the free-sporing and spermophytic, were small trees or shrubs of understorey or else climbers. This is opposite to the modern tropical rainforest structure, yet resembling leaf size distribution over the storeys of the temperate southern beach or conifer rainforests. The Permian biomic differentiation related to a short-term Asselian glaciation gave rise to a number of extinct biomes (IV.3), of which the Cathaysian gigantopterid flora is


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the nearest approach of the rainforest biome, as we have it today, both in the leaf morphologies, with a strong tendency to fusion of leaf segments, and taxonomic diversity. In its fully developed form, the Cathaysian biome persisted from the Artinskian to Kazanian (Wardian) age, about 30 m.y. The Mesozoic equatorial zone, modern latitudes as well as palaeolatitudes, falls within a broad realm of xeromorphic brachyphyllous vegetation having little in common with the modern-type equatorial rainforests. During the Jurassic, the most extensive salt deposition was confined to the equatorial zone suggesting an arid climate that extended over the larger part of the Cretaceous. Among the extratropical vegetation types, the warm-temperate Late Cretaceous contained a laurophyllous element (Laurophyllum, Magnoliaephyllum and Lindera-like “Araliaephyllum” in the Santonian–Campanian flora of Sakhalin: Krassilov, 1979) as an understorey to the Cupresso–Sequoietum overstorey. With degradation of the latter to the end of the period, the laurophylls segregated as a temperate to subtropical evergreen vegetation prototypic of the modern Canarian-type forests. In the terminal Cretaceous, coal-bearing deposits appeared in equatorial Africa (Petters, 1978; Offodile, 1980), while palynological data indicated a “palm zone” (S. Srivastava, 1983). The Early Palaeocene intertrappean deposits of India contain plants of modern tropical affinities (IV.3.4, VII.8). These developments suggest an incipient rainforest phase over the Cretaceous/Tertiary transition, correlated with a pronounced cooling of the mid-latitudes (VII.8, IX.9). A subsequent warming in the Late Palaeocene to early Middle Eocene gave rise to the so-called paratropical forests that spread to high latitudes along the Atlantic and Pacific coasts. These were essentially laurophyllous, prevailingly of the notophyll leaf class, not matching tropical rainforests in either the leaf physiognomy or diversity. The respective fossil plant assemblages are mono- to oligodominant, confined to redbed facies of seasonally dry climates. The associating dense populations of larger herbivores indicate an open rather than forested landscape. The fossil Dipterocarpaceae, Nypa, Pandanus and other elements of extant tropical vegetation, occasionally recorded in the Palaeogene, need not be taken as evidence of a tropical rainforest because even extant representatives of tropical families, such as Hopea mollissima (Dipterocarapaceae) are found beyond the tropical zone, in highlands with occasional freezing temperatures. Stunted mangroves presently spread far beyond the tropics. Palynological records of rainforests both in southeastern Asia (Muller, 1970) and in the Guinea–Congo area (Maley, 1996) are not earlier than the Oligocene. Remarkably, this was again the period of a pronounced mid-latitude cooling and a spread of continental climates driving the evergreen laurophyllous components to lower latitudes. The origin of the modern-type equatorial rainforests can be considered as a transformation, under the conditions of excessive humidity, of a sparse laurophyllous vegetation into a dense multistoreyed one, for which light became a major limiting factor imposing a selection for shade tolerance.

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Other plant communities for which light poses a major problem thrive in the highlatitude areas of winter darkness. It is commonly taken for granted that a high latitude fossil plant of rainforest affinities must be of a tropical derivation, but it might have been the other way round: polar life forms, preadapted to shading due to their tolerance of prolonged darkness, might find their way under the canopies of primordial rainforests.

IV.2.2. Temperate deciduous vegetation Seasonal deciduousness might have appeared in leafless plants that shed green branchlets, as extant Ephedra species do. The Devonian archaeopterids were shedding leaf-like syntelomic systems that sometimes formed mat-type accumulations (Fig.37) indicating seasonal deciduousness, as in the Late Devonian Tanaitis, a leafless heterosporangial archaeopterid genus akin to Svalbardia, the northernmost (present-day latitudes) representative of the group. The branch-mats occur on bedding planes of a levee-type detrital siltstone containing also abundant fossil wood fragments (Krassilov et al., 1988). The modern-type deciduousness appeared with the Permian Angarian and Gondwana floras of the broadleaved vojnovskyan “cordaites” and glossopterids respectively, forming mats of individually shed leaves and leafy spur-shoots. In the Late Permian, a dominant morphotype of ribbon-shaped leaves with a callous abscission scar was joined by the newly appearing flabellate Ginkgo-like (Rhipidopsis) and aciculate coniferoid morphotypes that became more prominent in the Mesozoic. The mid-Triassic to Early Cretaceous taphofloras of northern Asia were dominated by Czekanowskia, Phoenicopsis, Pseudotorellia, Ginkgoites, Pityophyllum and Nilssonia forming extensive leaf-mats. All these plants were heteroblastic, shedding spur-shoots with leaves (Fig. 38). Shoot dropping was inherited by deciduous conifers, such as Parataxodium, dominant over northern Asia in the Late Cretaceous. They formed mixed leaf mats with broadleaved platanoids and hamamelids (Trochodendroides) that grew as understorey and seral communities of the Platano-Parataxodietum (Krassilov, 1979). With a decline of the latter over the Cretaceous/Tertiary boundary, new broadleaved deciduous forest types were formed on the basis of the seral Platano-Trochodendroidetum, incorporating the deciduous forms taxodiaceous conifers derived from Parataxodium (Metasequoia, Taxodium, Glyptostrobus). They constituted the core of the ArctoTertiary biome mixed forest types, with biserrate betuloid components first recorded as dominant elements from intermontane depressions of circum-Pacific coastal ranges (Krassilov, 1989c), supposedly representing an upland differentiation of the early hamamelid complex. Summer-dry deciduousness is poorly recorded because leaves are shed in a season unfavorable for preservation scarcely forming leaf mats. It might have first appeared in


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Fig. 37. A leaf-mat type taphonomy in a leafless archaeopterid plant, Tanaitis: a bedding plane accumulation of syntelomic branch systems from the Late Devonian of Pavlovsk, Voronezh Region, European Russia (Krassilov et al., 1987).

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a

b

c

d

e

Fig. 38. Ginkgophyte morphotypes forming leaf mats at Mesozoic localities of the temperate Mesozoic deciduous biome: (a) Ginkgoites, (b) Sphenobaiera, (c, d) Baiera, (e) Pseudotorellia. The Late Jurassic to Early Cretaceous of Bureya Basin, Russian Far East (Krassilov, 1972c).


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the Permian coniferoid xeromorphs, often preserved as scattered scale-leaves (e.g., in the Nedubrovo locality: II.7.2). Mesozoic brachyphylls of a highly variable shoot morphologies (Fig. 39), often show basally constricted or obliterate branchlets. Accumulations of detached coriaceous scale-leaves of pro-gnetophytic Dinophyton (Ash, 1970; Krassilov & Ash, 1988) and hirmerellids are examples of the Ephedra-type microphyll deciduousness. In bennettites of Ptilophyllum–Otozamites morphotypes, the swollen petioles of pinnate macrophylls might have been due to insect galling that inflicted leaf shedding that developed into seasonal deciduousness as an adaptation to the wet/dry climate. High frequencies of galled, mined or cut leaves in early angiosperms (Krassilov & Baccia, 2000) suggest a low level of biochemical defense. This invites consideration that angiosperm deciduousness arose in response to herbivory or, with insect pollination, for redirecting insects from leaves to flowers (VIII.1.2).

IV.2.3. Conifer forests The Palaeozoic coniferoids have appeared as sclerophyllous wetland plants derived from helophytic pteridosperms. Ecologically they were perhaps comparable with swamp cypresses of the presently relict coniferous wetlands. The prehistory of the modern-type conifer forests commenced in the mid-Triassic with differentiation of the modern families Araucariaceae, Podocarpaceae, Taxodiaceae and Pinaceae the latter two families linked to the ancestral voltziaceaous stock via the transitional Swedenborgia (Podozamites), Elatides and Schizolepis (Pityophyllum). On taphonomic evidence, all these were riparian trees or shrubs. Both Podozamites and Pityophyllum shed leaves and leafy short-shoots frequently preserved as leaf-mats. Elatides and derived scaleleaved taxodiaceous morphotypes are abundant in fluvial facies. Their germinate seeds and seedlings were found in situ in flood-plain alluvium (Krassilov, 1987, Fig. 24). At the communal level, the Podozamitetum was cosmopolitan as a streem-side community, while the Pityophylletum occurred over a wide range of peatland facies, with the needle-leaved pinaceous conifers forming pure stands or co-dominant in the Pityo– Pseudotorellietum, Pityo–Baieretum, etc., the edaphic variants of temperate wetland forests. The Elatidion formed a distinct catenic belt of riparian vegetation next to fernmarshes (III.3.3) attesting to evolutionary continuity of the arboreal wetland line. At the Cretaceous stage of adaptive radiation, these scale-leaved and needle-leaved wetland communities gave rise to mesic varieties, with morphological innovations in their dominant conifers. The Athrotaxites–Geinitzia–Parasequoia component of Cretaceous shrublands (below) might have derived from the Elatidion, with Sphenolepis or related transitional forms. The mesic derivatives of Pityophyllion were dominant over the highlands of Mongolo-Okhotskian fold-belt, with their wind-borne cone-scales, samaras and occasional leafy spur-shoots reaching deposition sites in rift lakes of Transbaikalia and

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Fig. 39. Diverse shoot morphologies in the conifers and coniferoid gnetophytes dominant in the brachyphyllous Early Cretaceous vegetation of Mongolia (Krassilov, 1982).


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Mongolia (Krassilov, 1982). These allochthonous remains form a morphologic link between the Jurassic Schizolepis–Pityospermum complex and an extant laricoid genus Pseudolarix, and are sometimes assigned to the latter (Fig. 40). Pseudolarix kempferi, a temperate montane species, is the only survivor of the Cretaceous complex of heteroblastic shoot-dropping conifers that, with the mid-Eocene montane glaciation, conceivably gave rise both to the deciduous laricoids with persistent spur-shoots and, by a pedomorphic transformation of heteroblastic shoot system, to the homoblastic piceoids.

c

a

b

e

d

f

Fig. 40. A precursor of mountain taiga, the Early Cretaceous Pseudolarix forest abundantly represented by leafy short-shoots and detached cone-scales in the rift valley deposits of Transbaikalia and Mongolia: (a-c) P. erensis, a modern-type pinaceous conifer growing side by side with (d-f) Swedenborgia, an archaic voltziaceous conifer probably ancestral to the Pinaceae (Krassilov, 1982).

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The conifer rainforests of circum-Pacific mountain ranges still retain a relict taxodiaceous element indicating their origin from the widespread Late Cretaceous to Palaeogene lowland forests with Parataxodium, Parasequoia, Geinitzia, Sequoia, Metasequoia and Glyptostrobus as dominant trees. With the circum-Pacific orogenic belt rising since the mid-Cretaceous, the Mesozoic taxodiaceous forests might have become segregated into the montane and lowland redwood types, perhaps in parallel with segregation of their allopolyploid hybrid complexes into the extant species. In the montane forests, taxodiaceous conifers were gradually replaced by pinaceous conifers. Over the Sikhote Alin ranges, this process is documented by the entries of piceoid and laricoid conifers in the terminal Cretaceous plant assemblages of intermontane depressions (Krassilov, 1989b) and their subsequent increase through the Palaeogene. Fossil wood remains indicate the development of a modern-type montane taiga in the Late Eocene to Early Oligocene, just prior to a widespread temperization of the northern mid-latitudes. The Siziman “fossil forest” assemblage of that age (Blokhina, 1985) includes Metasequoia in association with Laricioxylon xylotypes related to extant larch species of the same mountainous area. Montane taiga is precursive of the lowland boreal forests. Until the Miocene, the northern high latitudes were covered with a mixed Arcto-Tertiary forest. A much more recent segregation of the boreal conifer forests was related to a spread of cold front generated by the Plio-Plistocene continental glaciation. The present-day boreal forest zone occurs between the winter and summer positions of the Arctic atmospheric front dividing cold dry Arctic air from the warmer and moister Pacific air. Solar angle is one of the variables controlling the extent of this zone (Bonan & Shugart, 1989). In the boreal conifers, both leaf surface and leaf mass are larger than in the cool-temperate hardwood forests (Sprugel, 1989). In contrast to the leaf litter of deciduous hardwoods, the needle litter does not inhibit moss growth. In boreal conifer forests, moss cover plays a crucial role in sustaining water balance. Their annual moss production is nearly equal to their arboreal production (Bonan & Shugart, 1989). Since there is evidence of a diverse bryophytic ground cover in coniferous Jurassic forests (Krassilov, 1973c), this aspect of boreal ecology might have already appeared in the Mesozoic wetland forerunners.

IV.2.4. Xeromorphic vegetation Xeromorphic vegetation overlaps with sclerophyllous vegetation (III.1.3) over a broad zone of seasonally dry climates. The present-day xeromorphic plants are mostly angiosperms, though xeromorphy also occurs in lichens, lycopsids, ferns, cycads, conifers and gnetophytes. Cladonia is prominent in the present-day lichen tundra while Selaginella is widespread both in savannas and tundras, as well as over the Pleistocene savannoid (“tundra-steppe”) vegetation.


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The early land plants appear xeromorphic although most of them are found in wetland facies. Owing to their convergent morphologies, xerophytes and helophytes replace each other over both ecological and evolutionary successions. At least in the early plant cover, these variants might have not been as distinct physiologically as at present. It has to be noted here that many present day xerophytes are derived from littoral scleromorphs. No definite xerophytic zone has ever been recognized over the Carbonifeous, yet some precursors of Permian xeromorphs (their taeniate pollen records) have been found in the pre-Cordilleran province (Cridland & Morris, 1963; Read & Mamay, 1964), Tien Shan and elsewhere over the northern mid-latitudes (about 35-40°°N). A more definite xeromorphic zone appeared with the Walchia–Callipteris assemblage rising to dominate the Rotliegenden-type floras in Europe and North America. Its dominant coniferoids and peltasperms might have had their origins among wetland Carboniferous cordaites and callistophytes, respectively. Although commonly interpreted as representing an upland xeric community, Walchia–Callipteris assemblages typically occur in paralic facies (Krassilov, 2000b). In the mid-Permian of the mid-Urals, taeniate pollen grains of coniferoids and peltasperms are often found in the gut compressions of large insects from waterside habitats (Krassilov & Rasnitsyn, 1998). In the transitional Late Permian/Early Triassic sequence of Nedubrovo (II.7.2), a rich, though fragmented, coniferoid–peltasperm assemblage came from the pyritic estuarine facies that typically contain transported debris of riparian vegetation (Krassilov et al., 1999a, 2000). These xeromorphic communities might have formed a helophytic–xerophytic continuum over a range of redbed facies. Likewise, on the southern continents, the Glossopteris assemblages associate with both coal-bearing (in situ Vertebraria-type underground organs) and redbed facies. Remarkably, their palynofacies are also dominated by taeniate pollen grains. Both conioferoid–peltasperm and Glossopteris wetland communities survived over the Permian/Triassic boundary crisis and became a major source of new gymnosperm xeromorphs. The typical xeromorphic epidermal structures are found in Pachypteris, a late representative of Mesozoic peltasperms, as well as in bennettites of the Ptilophyllum–Otozamites group that are abundant in the deltaic and estuarine facies of Yorkshire, Weald, Karatau, Kamenka and elsewhere over the evergreen zone. For them a helophytic habitat is suggested by the stomata sunken in the aerenchymatous spongy mesophyll (in Otozamites lacustris: Krassilov, 1982) and by a dense cover of salt glands (in Ptilophyllum: Watson & Sincock, 1992). A prominent component of the bennettitalean wetland communities are scale-leaved (brachyphyllous) coniferoids (Fig. 39). Their Classopollis-type pollen grains commonly increase with transgressions and are abundant in offshore marine deposits (Markevich, 1994). As in the case of the Permian producers of taeniate pollen grains (above), the proximity of Classopollis-producers to the coastal deposition sites is attested by this pollen preserved in the guts of large insects (Krassilov et al., 1997). Similar but even more diverse cycadophyte–brachyphyllous assemblages came from the inland Mesozoic basins (Krassilov, 1967a). Their macrophyllous component of cyc-

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adophytes Ctenis and bennettites Zamites (Zamiophyllum) is morphologically comparable with the extant cycads, in particular Encephalartos of the African shrub savannas. The brachyphylls are either extinct Mesozoic coniferoids or allied to extant Athrotaxis (fossil Athrotaxites) forming shrublands at drier sites. In the Early Cretaceous of eastern Asia, such assemblages developed in the rain-shadow of the then rising Sikhote Alin ranges (V.9.1). Consistently with a shrubland interpretation, the cycadophyte–brachyphyllous assemblages show a spectacular leaf-shape convergence across taxonomic divisions, as in the present-day scrub–chaparral–maquis formations. In the mid-Cretaceous, the cycadophyte–conifer shrubland communities underwent a major change, with their cycadophyte component replaced by xeromorphic angiosperms of Sapindopsis–Debeya–Dryophyllum–Myriciphyllum leaf morphotypes (Fgs. 41, 42).

a

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Fig. 41. A mixed Mediterranean-type assemblage from the Cenomanian Nammoura locality, Lebanon: (a) Pseudolarix, a conifer needle, (b) Pteris sp., a xeromorphic fern leaf, (c) Pseudotorellia sp., a coriaceous ginkgophyte leaf with intercoastal resin canals, (d) Sapindopsis lebanensis, a compound angiosperm leaf with thick coriaceous leaflets and a prominent stipule, (e) Nammourophyllum altingioides, a serrate leaf of a morphotype typical of deciduous angiosprms, (f) Parvileguminophyllum, a coriaceous leaflet, (g) Nupharanthus cretacea, a nymphaeaceous flower (Krassilov & Bacchia, 2000).


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Fig. 42. A xeromorphic serratoleaved–scaleleaved assemblage from the Cenomanian Silbukhra locality, northern Crimea: (a, b) fern and cycadophyte remains with small coriaceous leaves, (c) charcoal of an angiosperm leaf compression, (d) cuticle of a laurophyllous angiosperm with paracytic stomata, (e, f) Geinitzia, a dominant conifer (Krassilov, 1984).

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The compound to narrowly lobed leaves of Sapindopsis lebanensis with their inconspicuous secondary venation and thick cuticles represent a rather typical leaf morphology (Krassilov & Bacchia, 2000). Its associated leaf morphotypes are either compound and entire-margined or simple, microphyllous, with serrate margins. Fossil wood studies confirm a small tree to shrub habit for most mid-Cretaceous angiosperms (Parrish & Spicer, 1988). Among them, Debeya is often preserved as a single-species accumulation of fragmented compressions or dispersed cuticles (e.g. D. pachyderma: Krassilov, 1979), a taphonomy typical of littoral scleromorphs. Hence a helophyte – xerophyte continuum might have developed at this time level also. With periodic expansions and retreats during the Tertiary, enriched by the marginal sclerophyllous derivatives of laurophyllous vegetation (Axelrod, 1975) and gradually acquiring a familiar aspect of present-day shrublands, xeromorphic communities continuously occupied the Mediterranean zone, in which their precursors first appeared in the Permian.

IV.2.5. Grasslands The history of herbaceous drylands is comparatively short one. As discussed before, fossil fern assemblages, sometimes interpreted as “fern savanna,” were neither herbaceous nor xeromorphic (with the exception of arboreal littoral Weichselia). They represented open freshwater marsh vegetation. With a decline of fern marshes at the Jurassic/Cretaceous boundary, new types of wetland appeared, with herbaceous gnetophytes of grass-like or sedge-like aspects (Krassilov, 1997a). The earliest macrofossil records of graminoid monocots come from the mid-Cretaceous (Krassilov et al., 1983; Krassilov & Dobruskina, 1997; Krassilov & Bacchia, 2000) but the modern-type Arundo–Phragmites wetlands first occurred in the terminal Cretaceous – basal Palaeocene (Krassilov, 1976a), while the xeric herbaceous plant cover dominated by the grass family appeared much later in the Tertiary. These chronological relationships may suggest a derivation of xeric grasslands from grass wetlands, with scleromorphism perhaps initially related to littoral habitats. The terminal Cretaceous appearance of grass wetlands coincided with adaptive radiation of herbivorous mammals. A prominent increase in crown type diversities in mammalian herbivores over the Palaeocene/Eocene transition (VII.8) might reflect a spread of grasslands, but probably of hydroseral types yet, grading on wetlands. Studies of pedogenic processes in the Eocene do not confirm extensive xeric grasslands. Although the origin of grasslands is commonly ascribed to grazing, the historical relationships are less obvious. Grazing developed as an alternative to browsing, a historically much older feeding habit. Logic impels, while geochronology confirms, that grasslands would have appeared before grazing (Retallack, 1982). The incipient Palaeogene grasslands might have been promoted by browsers that destroyed shrublands. Early xeric


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grasslands could have been self-sustainable as a flammable source of bushland fires. Grazing adaptations became advantageous with a stepwise spread of grasslands over the Miocene/Pliocene transition and in the Altithermal (Axelrod, 1985). Grasslands thus appeared as a climatic type of plant community, yet a positive feedback of grazing made them sustainable over a wider range of climatic situations. As a relatively late appearing biome, grasslands invaded the former shrubland zone. “Steppe climate” differs from “scrub climate” mainly in the peak of seasonal precipitation displaced from spring to early summer. In the late Neogene, a “steppe climate” might have developed as a mid-latitudinal response to an expansion of the polar ice caps. The present-day xeric grasslands are sustained in a dynamic equilibrium with shrublands and are rapidly replaced by the latter with climate change and/or overgrazing. Similar relationships occur over the tundra zone that includes the hypo-Arctic shrubland, a dwarfed/creeping shrubland with perennial herbs, meadows and the moss–lichen oligotrophic varieties. Except the coastal meadows, tundra is xeromorphic, with a prominent evergreen element. Considering that the dwarf shrub and creeping components are derived from boreal tree species, the tundras are comparable with savannas differing primarily in the stunted habit of its arboreal component. No typical tundraa has ever been recorded before the Plio-Pleistocene. The Pleistocene “tundra-steppes” had a considerable arboreal component documented not only by palynological data (that might partly reflect a long-distance transport), but also by wood remains, hence more savannoid than the present-day tundras. Remarkably, Selaginella is prominent in both tropical savannas (with dormant megaspores surviving seasonal draughts) and Arctic savannoid (“tundra-steppe”) communities (with draught resistance as a pre-adaptation to cold climate). The Arctic savannoid vegetation supported a rich fauna of larger mammals that disappeared with its decline. The replacing shrub tundra has spread with a drier Holocene climate.

IV.2.6. Wetlands Wetlands play an outstanding role in the global biochemical carbon cycle by storing organic material in peat and peaty soil, as well as by being exporters of dead mass to aquatic ecosystems and regulators of terrestrial runoff. Owing to their prolific dead mass production and proximity to deposition sites, wetlands are the major source of fossil plant assemblages. Wetland assemblages are recognized on the basis of taphonomic criteria, such as autochthonous deposition in alluvial to littoral facies, as well as by helophytic life forms characterized by horizontal roots, aerenchymatic submerged organs and sclerophyllous leaf morphologies. An oxidized zone around the aerenchymatic organs protects them from microbial toxins of the waterlogged anoxic soil (Brändle et al., 1996). Encased with

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ferrous oxides, stumps of wetland trees are often preserved in the living position. Saline ground waters of paralic peatlands enhance mineralization and structural preservation of plant remains. Wetland successions are autogenous, related to peat accumulation inflicting a transition from rheotrophic (fed by ground waters) to ombrotrophic (fed by rains) types, with an edaphic zonation and a widespread sclerophycation in the latter. Exogenous successions are brought about by changing subsidence rates, climate, water table (fluctuating with sea-level and precipitation), penetration of saline waters in the ground aquifers, etc. Freshwater and brackish wetlands replace each other seawards and along the shore over the gradients of the soil, water chemistry and flood intensity, forming catenic series as well mosaic patterns. Many wetland types are seral to dryland vegetation. Even the old-growth mangroves are similar to the early successional stages of adjacent dryland forests in their general aspect and in such structural features as a poorly developed understorey, periodic ground cover of seedlings, canopy gaps, and the loss of apical dominance (Lugo, 1997). Dwarf mangroves are colonizers that can withstand strong hurricanes, with dense roots forming a tough peat. As sere reiterates evolution, many dryland vegetation types might have derived from wetlands. Terrestrial vascular plants most probably appeared in wetlands formed by their algae-like precursors. The Devonian survivors of such primordial wetlands are the peat-forming single-species mats of thickly cutinized thalloid Orestovia, Shuguria, Bitelaria and the like (Istchenko & Istchenko, 1981; Krassilov, 1981b; Krassilov et al., 1987a; Figs. 43 – 45). Their major localities occur at about 50°N (modern latitudes) from New Brunswick via central European Russia (Voronezh) to southern Siberia (Minussinsk Basin) forming a distinct peatland zone. The initial radiation of evolutionary lines leading to bryophytes, rhyniophytes and various pteridophytic groups might have taken place in the Devonian wetlands. A tree habit is advantageous for a frequently flooded vegetation (another option is a submerged habit, as in the Devonian Aglaophyton: Edwards, 1986). Free-sporing heterospory makes sense in conjunction with a dispersal by running water. Vivipary is also a wetland adaptation. In the mid-Devonian, tree habits were acquired in parallel by all groups of the leafless – psilophytic – life forms. Concomitant with the woodiness, there was a parallel tendency toward a heterosporous reproduction, involving the wetland lycopsids, barinophytes and progymnosperms, and further toward endosporangial megagametophyte development giving rise, again in parallel, to seed-like structures. In the Devonian already, these life forms might have diverged in respect to edaphic conditions forming different – protolepidophytic, progymnospermous – types of arboreal wetlands. Ample evidence of further wetland differentiation is provided by the Carboniferous fossil plant records (II.7.1). The lepidophyte wetlands, a major Carboniferous type, slowly degraded through the Permian, with their dominant forms stunted or reduced to pedomorphic aquatic forms which became once more prominent over the latest Permian (Vyatcheslavia) to early Triassic (Pleuromeia), when they formed monodominant coastal marshes.


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Fig. 43. Scanning electron micrographs of cutinized thalloid plants from the Devonian cuticle coals: (a, e) Orestovia, compressed specimens (courtesy of A.R. Ananiev) and a stomata-like structure supposedly derived from a sex organ (Krassilov, 1981), (b, c) gametangial conceptacles in Shuguria, a morphologically similar alga-like form, (d) Bitellaria, bands of hydroid-like fusiform cells traversing the cuticle (Krassilov et al., 1987a).

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Fig. 44. Light microscopy of Shuguria ornata, a structurally preserved thalloid plant from the upper Middle Devonian of Pavlovsk, Voronezh Region, European Russia: (a) shoot fragments from a cuticle coal, (b) cleared cortex tissue showing conceptacles, (c) ripe conceptacle shedding its neck canal cell, (d) perforation at the site of a disintegrated conceptacle.

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Fig. 45. Light microscopy of a structurally preserved thalloid plant, Shuguria ornata, from the upper Middle Devonian of Pavlovsk, Voronezh Region, European Russia: (a, b) intact conceptacles showing gametangia as the radially arranged bulging cells of the basal chamber, topped by a neck canal cell, (c) intact gametangia adjacent to the lower tier of neck cells, (d) conceptacle at a later stage, with empty gametangia.

Chapter 4. Palaeoecogeography a

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Fig. 46. Wetlands, a major source of organic matter in terrestrial runoff: (a) a propagating mangroves on the Little Rock Island, Caribbean (photograph taken during field excursion of the National Park Congress, Venezuela, 1992), (b) a horse-tail bog of the type widespread in the Mesozoic, presently developing over a heavily polluted oxbow of Bitza River, Moscow, (c, d) a swamp cypress stand in the Mississippi delta.

Since the mid-Triassic, both coastal and riparian wetlands were dominated by spore plants forming extensive fern – horsetail marshes. Although their dominant components belonged to the tree-fern families, they might have been of a stunted habit in the temperate zone, as are the present-day coastal Dicksonia growths in New Zealand, but contained large arboreal forms in the taxonomically more diverse evergreen zone. In Asia, representatives of the Schizaeaceae, Matoniaceae and Gleicheniaceae entered the marshlands south of 50°N, while the northern Osmunda and Dicksonia, were replaced by Todites and Cyathea (Krassilov, 1978d; Figs. 47, 48). Only a brackish edaphotype of the predominantly fresh-water fern-marshes survived beyond the Cretaceous as the mangrove fern mires. The horsetail mires are likewise an essentially Mesozoic formation, with the present-day analogues occurring as rare edaphotypes locally replacing Typha swamps in the heavily polluted flood-plain wetland systems (Fig. 46). An alternative type of Mesozoic wetlands was formed by woody gymnosperms derived from several orders, e.g., Pachypteris, a xeromorphic descendant of Triassic peltasperms, as well as some bennettites of the Ptilophyllum–Otozamites group. The latter is common in the evergreen zone (the “Ptilophyllum floras” of Western Europe, Moscow Region, Donetsk, Caucasus, Kazakhstan, India, etc.). Ptilophyllum and Otozamites are the most abundant,

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Fig. 47. A temperate fern-marsh assemblage: a joint occurrence of leaves, sporangia and spores of a dominant fern species, Osmunda diamensis, in the Late Jurassic of Bureya Basin (Krassilov, 1978d).


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Fig. 48. An ecotonal temperate/subtropical fern-marsh assemblage from the lowermost Cretaceous of Tyrma River, 50°N, with termophylic ferns: (a, b, c) Klukia, an extinct schizaeaceous genus, (d, e) Cyathea, an extant tree-fern genus (Krassilov, 1978d).

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sometimes even the only, plant fossils in the carbonate facies of Karatau in Kazakhstan and in the gypsiferous shales of Kamenka, the Donetsk Basin (III.3.3). The Classopollis-producing scale-leaved plants of gnetalean affinities make a strong candidacy for gymnosperm mangroves by their constant association with transgressive facies (III.1.2). The beginning of the Cretaceous was marked by an overturn of the fern-marsh dominance, with the Dicksoniaceae in decline and the schizaeaceous ferns (the Cicatricosisporites spore type) coming to the fore. The arboreal Tempskia–Weichselia wetlands of the southern zone are typically confined to littoral facies perhaps representing fern mangroves. A reduction of fern marshes over the trans-Asiatic Mongolo-Okhotskian rift zone might have played a role in angiosperm origins, with diverse proangiosperms forming an entirely new type of marshland (Krassilov, 1997a). Angiosperms entered the pre-existing wetland types, such as the swamp-cypress mires (Fig. 49), or formed new types of their own. A littoral habitat of the Late Creta-

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b Fig. 49. Nysso-Taxodietum, a persistent wetland community: an early example from the basal Palaeocene of Tsagajan, Amur Province (Krassilov, 1976a): (a) Nyssa bureica, (b-d) Taxodium tinajorum, leafy twig and a cone scale, (e) Metasequoia disticha representing a less flood-tolerant type of taxodiaceous wetland, presently dissociated from this type of wetland plant community.


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ceous Debeya–Dryophyllum communities, conceivably derived from a marginal (seral) laurophyllous element of the broadleaved-coniferous forests, is suggested by the mainly paralic facial occurrences of these xeromorphic angiosperms (IV.2.4). The waterside reed-like communities first appeared in the terminal Cretaceous to Early Palaeocene. These two types might have been ancestral to arboreal and herbaceous Cenozoic wetlands, respectively.

IV.2.6. Aquatic tracheophytes Before the appearance of floating tracheohytes (paradoxically late in the geological history), freshwater ecosystems produced a relatively small indigenous biomass, serving as traps for the dead mass exported from adjacent land. Until the mid-Cretaceous, submerged communities were dominated by charophytes and aquatic bryophytes, with occasional entries of tracheophytic forms (the Devonian Aglaophyton, the Triassic and later Isoetes-like lycopsids) derived from flood-tolerant waterside communities dominated, jointly or sequentially, by zosterophylls, rhyniophytes, lepidophytes, arthrophytes, horsetails, ferns and proangiosperms. Neither of them ever acquired a floating habit. The thallose bryophytes were rhizoidal, forming mats on wet substrates. The floating Riccialike forms have appeared not earlier than the Cretaceous (Krassilov & Schuster, 1983). A great change in aquatic life came in the mid-Cretaceous, starting with a simultaneous appearance of various aquatic ferns, both of modern affinities and extinct, including the Heroleandrales (Fig. 50), a recently found group with megaspores bearing microspores in the pockets of their leasural flanges (Krassilov & Golovneva, 1999, 2000). Such amphisporous dispersal units provided for a close association of mega- and microgametophytes, which is advantageous for colonizers. In fact, the heroleandras, commonly represented by dispersed Balmeosporites–Arcellites type amphisporions, rapidly spread worldwide as pioneers of the previously vacant niche of floating tracheophytes. The diversity of submerged isoetaleans increased appreciably over the Cretaceous, with such short-living genera as Nathorstiana, Limnoniobe and, later, Monilitheca (Krassilov, 1982; Krassilov & Maculbekov, 1995). Their clear-water oligotrophic habitats were invaded by phytomass exported from the neighbouring eutrophic ecosystems, with abundant limnaceous seeds dispersed over the river basins (II.1). Aquatic ferns were closely followed by flowering plants in populating the ecological niche of suspended or rooted macrophytes with floating leaves, of which Quereuxia is a familiar example (VIII.3.3). The short-leaved pioneer forms, such as heroleandras and quereuxias, declined with further build-up of diverse aquatic communities dominated by the modern-type nymphaeaceous genera and later by monocots (Limnobiophyllum: Krassilov, 1976a). It is scarcely a coincidence that floating macrophytes appeared soon after freshwater diatoms and dinoflagellates. Their present allies invade ponds at an advanced

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Fig. 50. Heroleandra, example of a successful aquatic colonizer dispersed by means of amphisporions, the abundantly produced megaspores bearing microspores in the neck and laesural pockets (Krassilov & Golovneva, 1999, 2000).


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stages of plankton/periphyton successions. In the Cretaceous, eutrophication was set up by the newly arising algal groups elevating trophic levels as a precondition for floating macrophytes. These, in turn, greatly enhanced eutrophication and export from aquatic ecosystems, contributing to terrestrial runoff and affecting the aquatic animal life which experienced a major overturn in the mid-Cretaceous (Zherikhin, 1978).

IV.3. Biome chorology Biomes are zonal types of phytogeographic differentiation imposed by atitudinal distribution of solar input defined by a combination of orbital and geographic factors. Precipitation contrasts owing to oceanic and/or orographic effects cause a sectional differentiation over the extent of latitudinal zones, as between the broadleaved and mixed pine-broadleaved forests divided by the Urals or the “dark” and “light” taigas west and east of the Yenisey River in Siberia. The zonal, as well as sectional, pattern of palaeobiomes is a major source of information on land/sea distribution and climates (VI.7). The psilophytic communities were wetlands, interzonal in modern terms. The Middle to Late Devonian Archaeopteris and related genera (Svalbardia, Tanaitis) are reconstructed as deciduous trees with lateral branch systems homologous to macrophylls. Deciduousness was inferred form taphonomic evidence on Tanaitis, a heterosporous archaeopterid from Pavlovsk, central European Russia (Krassilov et al., 1987b), with bedding plane accumulations of lateral syntelomic stomatiferous organs. Archaeopterids were dominant over northern North America and Eurasia, with the southern boundary marked by the relatively rare records in Belgium, Bohemia, Donetsk Basin and Minusinsk. The recently discovered South African localities (Anderson, 1995) suggest a bipolar distribution. The archaeopterid realm overlaps a zone of cuticular coals formed by the alga-like Orestovia–Shuguria at about 50°N (IV.2.6). Scott (1980) recognized the Archaeoptris–Cyclostigma, Rhacophyton and Lepidodendropsis communities, all geographically overlapping. However, the Rhacophyton type is more common over the transitional zone between the northern Archaeopteris and the austral Lepidodendropsis (= southern Archaeopteris) realms, perhaps indicating an incipient latitudinal differentiation. The Early to early Middle Carboniferous vegetation is even less differentiated geographically, partly owing to a taphonomic overrepresentation of interzonal lepidophytic wetlands. However, the size differences of equatorial, mid-latitude and high-latitude (modern latitudes) lepidophyte records (Archangelsky, 1986; see also Thomas & Spicer, 1986) might be owing to a climatic zonation. The Permian geofloras were recognized in the 1930s on the basis of dominant leaf morphologies, with Callipteris, Cordaites, Gigantopteris, and Glossopteris defining the Euramerian, Angarian, Cathaysian, and Gondwana realms, respectively (Jongmans, 1936, 1937, Gothan, 1937; Kryshtofovich, 1937; Meyen, 1970). These realms were considered as floristic divi-

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sions corresponding to separate landmasses. As such they played a significant role in the development of continental drift theory and, later, the plate tectonics paradigm. However, the Permian geofloras are not strictly definable in terms of taxonomic distinctions because of the widespread convergent similarities of leaf taxa. Cordaitestype leaves, for example, occur in the typical cordaites, vojnovskyaleans and glossopterids (Noeggerathiopsis). The phylogenetic kinship of Vojnovskyales to the typical Euramerian cordaites is a matter of discussion (Neuburg, 1965; Meyen, 1984; Krassilov, 1997a). Anyway, the Cordaites morphotype is a poor guide to floristic affinities. The same is true for the Asiatic and North American gigantopterids, the Gondwanic and Angarian representatives of Palaeovittaria– Zamiopteris–Pursongia morphotypes, etc. Glossopteris associates at least with two differernt types of reproductive organs, i.e. cupular, as in Lingetonia, and strobilar, as in Kendostrobus. This leaf morphotype might therefore belong to at least two groups of Palaeozoic gymnosperms, one related to pteridosperms, the other to cycadophytes (Surange & Chandra, 1975; Maheshwari & Srivastava, 1992). Taxonomic affinities of Gigantopteris and its morphological allies are still more intricate. Some of them are related to marattialean ferns (Wan & Basinger, 1992). However, a compressed axis of Gigantonoclea with a partly preserved anatomical structure from the latest Permian of Laibin, southern China shows an araucarioid bordered pitting of metaxylem tracheids similar to that of the Permian conifers and peltasperms (Krassilov, 2000b). This leaf type co-occurs with two different types of reproductive structures, Jiaochengia (Wang, 1999) and Gigantotheca (Fig. 35). At the above-mentioned Late Permian fossil locality (see Shen et al., 1999 for stratigraphic information), the fossil plant bearing beds occurs on top of a marine regressive sequence truncated by a a root bed with arthrophyte stems and abundant pecopterid fronds, a thin coal and a roof shale with Gigantonoclea and fragmentary conifers. This sequence reflects an expansion of arthrophyte–pecopterid peatland after retreating sea, followed by a successional gigantopterid community. The conifers are rare but they also occur as transported material in the marine strata downsection, suggesting a more distant source upslope. Of the two wetland types, the fern-arthrophyte community is found worldwide while the gigantopterid community is characteristic of the Cathaysian realm, substituted by callipterid, cordaite or glossopterid assemblages elsewhere (Fig. 51). Precursory Permian-type assemblages occur sporadically in the Late Carboniferous of both Laurasian and Gondwana landmasses. In the Gondwanic realm, a decline of arboreal lepidophyte wetlands is traditionally related to glaciation. In the Euramerian realm, the coeval vegetational change is ascribed to aridization. However, the humid non-glacial Cathaysia was no exception of the global turnover. Critical to lepidophyte wetlands might have been a drop of ground water level caused by sea-level drop at about the Carboniferous/Permian boundary. On evidence of the better dated tillites, the glaciation was restricted to this transboundary interval. New types of wetlands appeared with a subsequent see-level rise in the Artinskian.


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Fig. 51. A schematic distribution of the dominant mid-Permian plant communities over the vegetation zones: T – temperate, ST – warm-temperate to subtropical, E – tropical (equatorial).

IV.3.1. Permian biomes Using the approach outlined above we can preliminary recognize the climatically defined mesothermic (temperate), xerothermic and humothermic Permian biomes in both the Laurasian and Gondwanic (Noto-Permian) realms. The thermophilic vegetation of southern China belongs to the Permo-Tethyan Realm. Temperate nemoral biome. Permian vegetation of northern Asia, known as the Angarian geoflora or floristic kingdom (Kryshtofovich, 1937; Meyen, 1970), covered Siberia, the Petchora Basin in northern Cisuralia, most of Mongolia and part of China north of 45°N. This flora is typically represented in the Kuznetsk Basin, West Siberia, a major coal-producing area in Asiatic Russia. Cordaites and morphologically similar Rufloria are the dominant leaf genera often preserved as single-species leaf-mats. The Vojnovskyales are reported from a few localities outside northern Asia, such as Midtkap in Greenland and Mount Dall in Alaska (Wagner et al., 1982; Le Page & Pfefferkorn, 1996; Naugolnykh, 2000). Among the spore-plants, the arthrophyte group Tch-

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ernovskiales appears characteristic, but not necessarily endemic, of the Angarian-type floras. The geographic range of “cordaite flora” was thought to coincide with the borders of Angarida. However, the Angarian realm is linked with the Euramerian callipterid realm by a broad zone of mixed floras extending over the Urals and Volga Basin and to Arctic Canada (Fig. 52). Many European taxa, such as Quadrocladus, spread over this transitional zone to Petchora, Taymyr (Sadovnikov, 1997) and further east. Suturing of the seaway between Angarida and Europe seems to have had a minor effect on the position of phytogeographic boundaries which remained nearly the same in the Triassic (Krassilov & Shorokhova, 1975; Krassilov, 1997c). A popular notion of the Angarian “cordaite taiga” is incompatible with the leaf morphology, for the typical taiga is dominated by evergreen needle-leaved trees. In the dry permafrost areas east of the Yenisey River, it is replaced by the larch taiga of deciduous needle-leaved trees shedding needles from persistent spur-shoots. The Cordaites type of broad ribbon-shaped leaves with parallel venation resembles that of monocotyledonous angiosperms, in which they are of a phyllodic origin, derived from inflated petioles with the leaf blade reduced. Phyllody relates to developmental acceleration in turn related to rapid leaf expansion required by a short growth season. In the Permian plants such leaves were shed separately, with clear-cut abscission lines, or clustered on deciduous spur- shoots. They formed thick leaf-mats, the commonest type of fossil plant assemblages in the coal measures. Since a considerable ombrotrophic peat accumulation requires ample precipitations, a combination of morphological and taphonomic data indicates a summer-green broadleaved vegetation ecologically equivalent to the present-day nemoral zone of a humid temperate climate. Warm-temperate ecotone. The Permian flora of Cisuralia and Volga Basin is considered as a broad ecotone of callipterid and cordaite vegetation. It was dominated by Euramerian peltasperms, conifers and their derived forms (Zalessky, 1927; Gomankov & Meyen, 1986; Naugolnykh, 1998). The latest Permian assemblages of this area constituted the so-called Tatarina flora named after a highly polymorphous leaf-genus of peltasperms, the lobate variants of which resemble both European Callipteris and American Supaia. Hence the “Tatariana flora” represents a regional version of the widespread “Callipteris flora”. Here the Angarian ruflorias and tchernovskias formed a significant, yet subordinate, element. Eastward, the ecotonal vegetation extended to Kazakhstan (Karaungir locality that yielded seeds similar to the Euramerian Salpingostoma: Krassilov, 1999b), Central Asia and northern China. On palynological evidence, the warm-temperate vegetation of this type can be traced to Greenland and Arctic Canada (Utting, 1994). In China, the Permian floras of Junggar–Khingan (Hinggan) latitudinal area are considered as belonging to the Angarian realm (Li & Wu, 1996; Li, 1997; Huang & Ding, 1998; Durante, 1992). Actually they are of an ecotonal aspect, with the Volgian Tatarina and Phylladoderma mixed with Cathaysian taeniopterids (Wang & Wang, 1986).

Fig. 52. Mid-Permian biomes and the respective vegetation zones: (1) Temperate nemoral, (2) Warm temperate mesic to xeric, (3) Subtropical xerothermic, (4) Subtropical humothermic, (5) Tethyan tropical and mixed. Symbols of floristic composition: I – Angarian, II – mixed Eurangarian, III – mixed Cathangarian, IV – Eauramerian, V – mixed Eauramerian/Cathamerian, VI – Cathamerian, VII – mixed Eurogondwanic, VIII – Gondwanic. Localities: 1 – Midtcap, 2 – Petchora, 3 – Western Taymyr, 4 – Eastern Taymyr, 5 – Norilsk, 6 – Tunguska, 7 – Kuznetsk, 8 – Verkhoyansk, 9 – Sverdrup, 10 – Soyana, 11 – Kama, 12 – Tchekarda, 13 – Karaungir, 14 – Fergana, 15 – Junggar, 16 – Tabun Tologoy, 17 – North Korea, 18 – Primorye, 19 – Appalachia, 20 – Cantabria, 21 – Orobic Alps, 22 – Turing, 23 – Utah, 24 – Arizona (Hermit Shale), 25 – Kansas, 26 – Guadalcanal, 27 – High Atlas, 28 – Hazro, 29 – Saudi Arabia, 30 – Iran, 31 – Texas, 32 – Mexico, 33 – Shanxi, 34 – Anhui – Henan, 35 – Jiangsi, 36 – Chubut, 37 – Niger, 38 – Gabon, 39 – Kashmir, 40 – Lhasa, 41 – Yunnan, 42 – Phetchabun, 43 – Jambi, 44 – Irian Jaya, 45 – Transvaal, 46 – Handapa, 47 – Perth, 48 – Bowen, 49 – Mount Achernar.

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This zone is traced to southern Mongolia and Primorye in the Russian Far East where the dominant Rufloria–Cordaites morphotypes associate with gigantopterids, a feature shared with contemporaneous floras of North Korea, Anhui and adjacent provinces of eastern China (Mei et al., 1996) and perhaps typical of the Pacific margin. Notably, the Permian faunas of fusulinids and brachiopods from adjacent areas of marine sedimentation are also of a mixed boreal/tropical aspect (Shi & Shen, 1966). Xerothermic biome. The Permian taphofloras of central to Western Europe and eastern North America are dominated by callipterids and conifers accompanied by the highly diverse fern-arthrophyte assemblages. They easily surmounted the boundary with Cathaysia over southern Xinjang and along the northern border of Tarim Basin to the Lesser Khingan (Li, 1997). A widely held view of the Euramerian vegetation as tropical is scarcely compatible with either the deposition environments or plant morphology. The Euramerian plant localities commonly associate with red or variegate deposits containing pedogenic carbonates. (II.7.2). Their dominant peltasperms and conifers were low-stature trees or shrubs with xeromorphic leaves. Their preservational features indicate proximity to deposition sites. The peltasperm-conifer assemblages typically came from estuarine deposits (e.g., the cuticle beds of Nedubrovo: Krassilov et al., 1999a, 2000), whereas the fern-arthrophyte assemblages more commonly occurred in deltaic facies. Their alternation over cyclic stratigraphic sequences might have been related to sea-level fluctuations or wet/ dry climate cycles or both (III.3.4). On sum of evidence, climates of the Permian xeromorphic zone might have come close to the modern Mediterranean type. Humothermic biome. Gigantopterids and emplectopterids are characteristic of Permian plant localities in China, Korea, Japan and Laos constituting the Cathaysian flora. To the southeast they extended to Thailand, Malaysia, Indonesia and New Guinea. Peculiar for the gigantopterid vegetation are also the broadleaved taeniopterids, diverse pecopterids at least partly shared with the xerothermic zone, lobatannularias and other arthrophytes. Their Asiatic ranges extend over a number of tectonic blocks at one time or another separated by seaways. The widely scattered Middle East localities (Wagner, 1962; El Khayal et al., 1980) mark their western extention while the western North American plant assemblages are of Cathaysian affinities owing to Cathaysiopteris and taeniopteroid leaves with marginal ovules (Mamay et al., 1988). On the other hand, Supaia, the emblematic North American genus, has been recently recorded in China (Wang, 1997). Therefore, the so-called Cathaysian flora was actually the Cathamerian. In China, the Cathaysian (Cathamerian) flora is presently subdivided into northern and southern provinces supposedly confined to disjunct terrains at the eastern margin of the Tethys Sea. The northern Cathaysian flora occurs on the Sino-Korean–Tarim block bound from the south by the latitudinal Kunlun–Ginling fault zone traceable to central Japan (Li, 1997). However, the floristic differences involve only few endemic genera, such as Emplectopteridium (Li & Wu, 1996; Li, 1997). More significant are the quantitative distinctions, with Otofolium, Gigantopteris and Rajahia common in the south in


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contrast to the predominantly northern Cathaysiopteris, Emplectopteris, large Taeniopteris and Yuania. Such distinctions suggest a climatic differentiation rather than isolation. Permian flora of Xu-Huai-Yu region in east-central China, at junction of Jiangsu, Anhui and Henan provinces, is considered as a separate floristic subprovince (Mei et al., 1996), but is, in a way, transitional between the northern and southern Cathaysian realms containing both Cathaysiopteris and Gigantopteris as equally common elements. A palaeobotanical standpoint would have the northern and southern Chinese blocks collided in the Devonian already (Li & Wu, 1996; Li, 1997). In the Early Permian, the typical Cathaysian flora of the Shansi (Shanxi) Formation with Lobatannularia and Emplectopteris replaced the essentially lepidophytic flora of the conformably underlying Taiyuan Formation. This rather abrupt floristic change at about the Asselian/Sakmarian boundary (Li, 1997) might reflect a replacement of the cosmopolitan lepidophyte wetland community by the newly appearing arboreal wetlands dominated by taeniopterids, emplectopterids and gigantopterids. In the Lower Shihhotse, these assemblages associate with coal measures that, in the Upper Shihhotse, were replaced by the redbeds with Gigantonoclea (Wang, 1999). In the latest Permian, the northern Cathaysian province was invaded by European conifers and peltasperms (also by Gondwanic glossopterids, see below) while the typical gigantopterid flora survived in southern Cathaysia alone. In the American Southwest, the Proto-Cordilleran Province was differentiated in the Late Pennsylvanian already (Read, 1947). Here, as elsewhere, the fern-arthrophyte marshes expanded at the expense of lepidophyte communities (Read & Mamay, 1964). The Permian plant assemblages of Texas and Oklahoma are of a “Cathaysian” aspect owing to the commonly occurring gigantopterid and taeniopterid morphotypes, the fertile variants of which bore marginal ovules (Mamay et al., 1998). Of the gigantopterids, Cathaysiopteris and Gigantonoclea morphotypes are shared with northern Cathaysia (though possibly convergent) while the other three, Gigantopteridium, Delnortea and Zeilleropteris, are endemic. This gigantopterid province bordered on xeromorphic vegetation of Utah–Arizona with Supaia and Walchia, and with Glenopteris in Kansas (Read & Mamay, 1964). Fossil soil types agree with a dry climate for this area (Kessler et al., 2001). Thus, xeromorphic vegetation penetrated both the Cathaysian and North American parts of gigantopterid realm. Floristic exchanges of the latter with xerothermic realm indicate a climatic continuity. However, the widespread Early Permian coal accumulation suggests a relatively humid climate for Cathaysia. The Dadoxylon-type fossil woods from this area show vaguely marked growth rings (Sze, 1952). Apparently, the northern Cathaysian and Euramerian realms were similar in their temperature conditions but differed in their precipitation regimes like the present-day wet and dry subtropical zones. In the altogether drier Late Permian, their climatic differences became less pronounced. Xeromorph invasions eventually replaced gigantopterid communities by the end of the Permian.

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Gigantopterids and emplectopterids, both prominent in the Cathaysian-type assemblages of eastern Asia and western North America, typically have composite leaves in which the entire or simply pinnate leaf blades were formed by marginal fusion of ultimate and penultimate segments, with the original segmentation betrayed by the conservative venation pattern (Asama, 1976; Krassilov, 1995c, VIII.6). The hierarchical areolate venation of such leaves owes its origin to meristematic fusion zones acting as a plate meristem (Krassilov, 1995c). This leaf morphology is evidence of a prolonged meristematic activity characteristic of megaphyllous life forms in equable climates. An advanced Gigantopteris morphotype, some variants of which show clearly defined driptips (Yao, 1986), occurred in southern Cathaysia. A survival of Carboniferous forms, such as a tree fern Psaronius and arboreal lepidophyte Lepidodendron, also indicates a tropical climate. Noto-Permian vegetation. Classical Gondwana flora came from the Narmada– Son–Mahanadi graben system that transects peninsular India. Boulder beds occur at the base of the basinal sedimentary fill followed by coal measures that contain the bulk of the flora. The Himalayan occurrences are scattered from Kashmir to Assam, with a few outliers in Tibet, Yunnan and the “mixed” floras of Indonesia and New Guinea. A number of well-defined species are shared by Indian and Australian floras, but the other parts of “Gondwanaland” are floristically less similar (Archangelsky, 1990, 1996; Maheshwari, 1992). A primitive glossopterid leaf morphotype Gangamopteris appeared already in the pre-tillite deposits. Early Gangamopteris plants are commonly envisaged as growing in sparse vegetation patches of a glacial landscape. However, fossil plant assemblages found immediately below and above the Dwykka Tillite in South Africa, the Bacchus Marsh Tillite in Australia and the Talchir boulder beds in Peninsular India suggest a less dismal picture, with a number of Carboniferous survivors (Chandra et al., 1992; Pant, 1996). Leaf morphology gives no support to either a “Gangamopteris tundra” or a “Gangamopteris taiga”. Morphologically, Gangamopteris is linked, via Palaeovittaria and Noegerrathiopsis, to the Zamiopteris–Cordaites group of the Arcto-Permian temperate zone. The emblematic Glossopteris morphotype came forth and diversified in the period of a widespread coal accumulation. Some variants of its highly polymorphiic leaves show distinct abscission scars, others had drip-tips characteristic of rainforest foliage. Such a morphological diversity obviously suggests a wide range of habitats (Pant, 1996), which is confirmed by taphonomic evidence (Retallack, 1980). Glossopterids are sometimes preserved as leaf mats containing spur-shoots shed with clusters of intact leaves, as is typical of deciduous riparian trees. A constant association of Glossopteris with Vertebraria, an aerenchymatous underground organ, suggests a peatland habitat. The coriaceous lingulate leaves, as in the modern mangroves, and the uniquely preserved seedlings (Pant & Nautiyal, 1987) are also consistent with a helophyte life form in a majority of Indo-Australian glossopterids. The greatest morphological diversity of Glossopteris


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(42 species from a single locality) is reported (Chandra & Singh, 1992) from the redbed Late Permian of Handapa, Mahanadi Basin bordering on the tropical Tethyan zone. The impression of a floristic integrity of Gondwana is due primarily to the ubiquitous glossopterids that make conspicuous, often numerically dominant fossils. However, the regional species compositions might have been highly endemic (McLoughlin, 1994). Other components show less uniform distribution, with significant differences between the eastern and western Gondwanas (Maheshwari, 1992), or the Notoafroamerican, Australoindian and Palaeoantarctic provinces (Li, 1997). The differentiation primarily pertains to the pecopterid–arthrophyte assemblages of European–Cathaysian affinities. Notably, the Permian floras of Niger in Africa and Chubut in South America have a considerable “Euramerian” component of thermophilic pteridophytes (Archangelsky, 1990, 1996) indicating climatic conditions not too different from those in the northern xerothermic zone of Europe and eastern North America. The fossil woods from Chubut often lack distinct growth rings (Archangelsky, 1996). Notably, the thermophilic South Cathaysian Pectinopteris or a closely allied form occurs in South Africa and elsewhere (Wan & Basinger, 1992). At the same time, a widespread coal accumulation in India and Australia (Retallack, 1980; Balme et al., 1995) suggests a relatively humid variant of a warm temperate to subtropical climate, similar to that of northern Cathaysia. Permian plant assemblages of Antarctica differ from the rest of Gondwana in a lower overall taxonomic diversity and in scarcity of pteridophytes (Archangelsky, 1990; Cúneo et al., 1993). The latter feature was emphasized by Cúneo et al. (1993) as indicating lack of undergrowth. As suggested above, the diverse pteridophyte component associates with the fern-arthrophyte wetlands different from the gymnosperm (peltasperm, glossopterid) wetlands with scarce pteridophytes of a few peculiar species. Incidentally, the Antarctic glossopterid assemblage contains a peculiar fern, Skaaripteris, supposedly aerenchymatous and probably aquatic or semiaquatic (Galtier & Taylor, 1994). Therefore, the Antarctic Permian climates were not too harsh for fern growth in general, but less favorable for pecopterid-arthrophyte communities than the other parts of the Noto-Permian realm. The Antarctic “fossil forest” is a dense growth of small to medium-sized trees with Vertebraria pneumatophores. The Gangamopteris and, later, Glossopteris assemblages associate with coals about 2 m thick containing stems 50 cm in diameter (McLoughlin & Drinnan, 1997). The coal seams contain clastic dikes that might have been formed upon permafrost veins of palsa mires as in the present-day cool temperate areas in Canada (Krull, 1999). The Glossopteris vegetation thus extended from cool temperate to subtropical climatic zones. In the present-day arboreal vegetation, only some riparian plants, such as willows and poplars, have comparable latitudinal ranges. Permo-Tethyan realm. The Tethyan phytogeographic realm is a continuous zone of mixed Cathaysian and Gondwanic assemblages extending from southern Spain and Morocco over the Middle East and Anatolia to Tibet, Yunnan, Thailand (Phetchabun),

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Sumatra (Jambi) and western New Guinea (Irian Jaya). Ecological equivalents of the southern Cathaysian vegetation are yet to be found on the Gulf coast and over the Caribbean. Occasional reports of gigantopterids came from Mexico and Venezuela. Cathaysian gigantopterids occur in the Permian taphofloras of Anatolia, Thailand, southern Malaysia and New Guinea (Wagner, 1962; Kon’no, 1963; Asama et al., 1975), whereas the Gondwanic elements prevail in the mixed floras of Kashmir, southern Tibet (Lhasa) and Yunnan where they are associated with the Cathaysian pecopterids, calamites and sphenophylls (Singh et al., 1982; Pant et al., 1984; Rigby, 1989, 1996; Li & Wu, 1996; Chandra & Sun, 1997). In these areas of floristic mixing, the Cathaysian and Gondwanic elements sometimes come from different localities. If separated by a fault zone, such occurrences are interpreted as of separate terrains (Li & Wu, 1996). Alternatively, a differentiation might have been due to a mosaic vegetation structure, with the glossopterid and gigantopterid communities as distinct ecotypes over a tectonically disturbed landscape. The situation might resemble a juxtaposition of “Malaysian” and “Australian” vegetation as altitudinal belts on New Guinea (Schuster, 1972) Both macrofossil and palynological data show a combination of Euramerian, Cathaysian and Gondwanic elements in the Permian western Tethys extending over Northern Africa, Middle East, Iran and Pakistan (Aassoumi, 1994; Lejal-Nicol, 1990; Broutin et al., 1995; Broutin, 1977; Utting, & Piasecki, 1996; Igbal et al., 1998). In Yunnan, the Gondwanic elements amount to 80% of palynological spectra (Gao, 1998). Glossopterids might have penetrated the Cathaysian realm over Kashmir–Tibet and along the Pacific coast as far north as Primorye, where both Gangamopteris and Glossopteris were found in a single locality (Zimina, 1967). Glossopterids once more became rather common in the Triassic of north-central China (Meng, 1995). In conclusion, the Permian “floristic” divisions could be more consistently interpreted as ecogeographic units, their ranges controlled primarily by climatic zonation. At about the Carboniferous-Permian boundary, the cosmopolitan lepidophyte wetlands were replaced by new types of plant communities dominated by various gymnosperms. This vegetational turnover was related to a global climatic change, regression and, in the wetlands, a drop of ground water level. Temperization of mid-latitudes was accompanied by differentiation of a humothermal tropical zone. The biomes defined on the basis of leaf morphologies and supplementary taphonomic evidence are assigned to temperate nemoral (Angarian and Antarctic realms), xeric subtropical (Eurameria and West Gondwana), humid subtropical (North Cathaysia and East Gondwana) and tropical (Tethyan) zones. Tectonic implications of this scheme are considered in (V.9).

IV.3.2. Mesozoic ecogeographic differentiation Over the Mesozoic periods, a fairly distinct ecogeographic boundary is recorded at about 50°N in Asia gradually ascending to about 60°N in Europe (Krassilov, 1972b,


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1997c). The boundary markers are Phoenicopsis, a deciduous tree dominant in the northern realm (Siberia, northern Kazakhstan, the Urals, Scandinavia, western Canada, Greenland) and Cycadeoidea, a pachycaul bennettite characteristic of the southern realm as far as the Black Hills, Dakota, British Weald, Middle East, Mongolia, and northern Sikhote Alin. Cycadeoidea associates with pinnate bennettitalean leaves of the Ptilophyllum–Zamites–Otozamites morphotypes, neither of which was ever found in the northern realm except at a few ecotonal localities (e.g., in Mongolia or Scotland: Krassilov, 1982; Van der Burgh & Van Konnijnenburg-Van Cittert, 1984). The Phoenicopsis/Cycadeoidea boundary thus marks the northern limit of the Ptilophylletea, as well as of the matoniaceous and schizaeaceous ferns Nathorstia, Klukia, Ruffordia, Stachypteris, Weichselia, the peltasperms Pachypteris and allied genera as well as the araucariaceous conifers. At the same time it divided the ranges of giant (stem diameters over 3 cm) and slender horsetails, Todites and Osmunda among the dominant osmunadaceous ferns, aphlebial and non-aphlebial Coniopteris among the treeferns and their dwarf varieties (Krassilov, 1978d, 1987). These latter distinctions reflect aspective differences of the northern and southern fern-marsh communities (IV.2.6). No less conspicuous are the ecogeographic differences in the zonal distribution of scale-leaved coniferoids (including the hirmerellaceous gnetophytes), both the abundance and diversity of which greatly increased across the boundary giving the southern vegetation a sclerophyllous aspect. However, in northern Asia, Classopollis, a palynological marker of brachyphyllous vegetation, reached to 60°N, justifying a division of the temperate zone into the warm-temperate and cool-temperate subzones. The angiosperm palynological markers, such as Proteacidites, indicate a similar differentiation in the Late Cretaceous (Herngreen & Chlonova, 1981). Deciduousness of the northern Arcto-Mesozoic (Phoenicopsis) vegetation is amply documented by the taphonomy of leaf mats with leafy spur-shoots while their association with coal beds suggests a summer-wet climate. In contrast, leaf mats are rather uncommon in the southern zone. Here the gypsiferous redbeds and lacustrine carbonates are the commonest lithologies while coal accumulation was patchy and insignificant. Both sedimentological and palaeovegetational features suggest a Mediterraneantype climate (Francis, 1984). Dry season is marked by false rings of fossil wood, as well as by narrow sclerites on the fossil fish scales (Krassilov, 1983a, VII.1.1). The distinctness of the boundary is scarcely related to temperature differences alone, the more so that the equatorial to polar temperature gradient was much gentler than now, with the Cretaceous equatorial means nearly as at present against about 0°C at the poles (Sellwood et al., 1994). Seasonality of precipitation might have been critical, with the boundary marking a transition from summer-wet to summer-dry climates. Tethys seas would have sustained a low latitude anticyclone that made the Meso-Mediterranean zone drier than its modern equivalent. On the weight of all available evidence, the Phoenicopsis/Cycadeoidea boundary can be considered as the first order ecogeographic divide separating the nemoral Arcto-

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Mesozoic realm from the xerophytic Meso-Mediterranean one (Figs. 53 – 56). Ecotonal assemblages are known from northern Sikhote Alin (the northernmost occurrence of Cycadeoidea in Asia: Krassilov & Shorokhova, 1989), the Tyrma River northernmost

Fig. 53. A scheme of the Late Jurassic to Early Cretaceous climatic zonation based on plant macrofossil and palynological records: (1) A cool temperate zone lacking Classopollis, (2) temperate zone of deciduous Phoenicopsis forests, (3) summer-dry subtropical zone with thick-stemmed cycadeoids, Ptilophyllum, and diverse brachyphyllous coniferoids, (4) equatorial zone with brachyphyll coniferoids (gnetaleans) and elaterbearing pollen morphotypes, (5) southern temperate zone with diverse ginkgophytes, small-leaved Ptilophyllum and with fossil woods showing seasonal growth increments (after Krassilov, 1985).

Fig. 54. the northernmost occurrence of Cycadeoidea, a pachycaul bennettite in the Albian of northern Sikhote Alin, about 50°N (Krassilov & Shorokhova, 1989).


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Fig. 55. A slab association of Czekanowskia, Leptostrobus and Pseudotorellia, a dominant assemblage of flood-plain peat bogs in the Arcto-Mesozoic realm. Late Jurassic of Bureya Baisn, Russian Far East.

a

b

c

d

Fig. 56. Ust-Baley locality on the Angara River, Irkutsk Region, Siberia, discovered in 1860s as first evidence of temperate Mesozoic vegetation: (a) view of the outcrop, (b) plant-bearing shales intercalated between massive sandstones, (c, d) slabs with remains of ginkgophytes, czekanowskians and gnetophytes (cones of Aegianthus sibirica: Krassilov & Bugdaeva, 1988).

localities of Klukia, Cyathea and Pachypteris (Krassilov, 1987), Transbaikalia, southern Siberia and southern Kazakhstan with joint occurrences of Phoenicopsis and Ptilophyllum (Otozoamites). In northern Europe, the Late Jurassic assemblage of Scotland, with Phoenicopsis and matoniaceous ferns, is of an ecotonal aspect. While the Svalbardian floras are typically Arcto-Mesozoic, those of western Greenland at the same latitude are ecotonal, with Phoenicopsis, Pseudocycas and Nathorstia, an anomalously

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high position of the boundary. In continental North America, the Blairmore flora south of Kalgari has Phoenicopsis (Bell, 1956) that is replaced by Cycadeoidea in the Black Hills of North Dakota. In the Southern Hemisphere, an equivalent boundary would have occurred south of the Patagonian locality of Nathorstia (Halle, 1913), a thermophilic fern known from the MesoMediterranean zone. Ptilophylletea extends to Santa Cruz in Patagonia (Bonetti, 1965; Seoane, 1999) and South Shetland Islands at 60°S (Fuenzalida et al., 1972). Otozamites is recorded from Scope Bay, Antarctica (Gee, 1989). In southern Australia, the deminutive pinnate bennettitalean leaves (Douglas, 1969) suggests stunting at about 40°S while the relative abundance of Ginkgoites signals the proximity of a warm-temperate ecotone. Ginkgoites is also represented in the mid-Cretaceous conifer-dominated flora of Eromang Basin, Queensland (Pole & Douglas, 1999), indicating an extratropical climate. The Cretaceous palynological assemblages of South Africa are ecotonal, with joint occurrences of Classopollis and bisaccate pollen grains (Herngreen & Chlonova, 1981). The Fossil Bluff flora of Alexander Island and the allied Antarctic assemblages (Cantrill, 1996, 1998; Cantrill & Falcon-Lang, 2001; Falcon-Lang et al., 2001) are temperate on account of the relatively frequent ginkgophytes, but the fern-marshes and brachyphyllous conifer communities, with a prominent araucarioid element (Césari et al., 2001; Falcon-Lang & Cantrill, 2001), appear taxonomically more diverse than in the northern temperate realm, being closer to the ecotonal warm-temperate floras of Transbaikalia, Tyrma River, etc. Wood anatomy indicates growth rhythms as in the present-day mid-latitudes (Jefferson, 1982). The Late Cretaceous flora of King George Island, western Antarctica is dominated by cyatheaceous ferns and podocarpaceous conifers, with angiosperms represented by Nothofagus and the broadleaved (platanoid in the opinion of the present author) Sterculiaephyllum (Dutra & Batten, 2000), hence ecologically similar to the warm temperate Late Cretaceous flora of Sakhalin at about 50°N, with diverse conifers and the riparian CyathoPlatanietum (IX.9.1). The most intriguing feature of Mesozoic vegetation is a complete absence of plant assemblages matching the modern rainforest in any of its typical synecological features. The Jurassic to Early Cretaceous macrophossil assemblages from southern Mali are single-species accumulations of scattered (probably shed) coniferoid scale leaves and occasional pollen cones (Krassilov, 1978a) suggesting a taxonomically poor, perhaps deciduous, variant of brachyphyllous vegetation inside the Mesozoic equatorial zone of elatero-plicate-rimulate pollen assemblages (Herngreen & Chlonova, 1981; Dino et al., 1999). Both macro- and microfossil assemblages of this zone are clearly inadequate as a rainforest ecological analogue on account of their scale-leaved morphology, low diversity and the accompanying salt deposits. Elements of tropical rainforest affinities appeared in the palynological “palm assemblages” and as macrofossils concomitantly with coal deposition in the terminal Cretaceous to Early Palaeocene (IV.2.1).


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VI.3.3. Mid-Cretaceous biomes Phoenicopsetea dominated the Arcto-Mesozoic realm from the mid-Triassic to midCretaceous (although Phoenicopsis itself locally survived until the Maastrichtian) and was then replaced by the conifer forests with Parataxodium and platanophyllous angiosperms (the major localities of northern Siberia, Kolymo-Chukotskian region, northern Canada and Alaska). Wood anatomy, with wide growth rings, few false rings, narrow late wood increments suggest a temperate climate with a light seasonality along the northern border (Parrish & Spicer, 1988). The thermophilic producers of Classopollis and, later, Proteacidites pollen types scarcely penetrated this northern zone. However, dinosaur localities from Svalbard, the Koryak Highlands (Nessov & Golovneva, 1990), Arctic Canada and Alaska (Davies, 1987) attest to a circum-polar distribution of this group. Siliceous Cretaceous sediments of the Arctic Ocean indicate an essentially icefree climate. To the south, the deciduous forest type gave way to the semi-deciduous Lauro-Sequoietum with platanophylls as riparian/seral elements (Sakhalin, Amurland, southern Siberia, northern Kazakhstan, northern Europe, northern Rocky Mountains, South Greenland). Within this broad zone, fossil plant localities of the circum-Pacific belt, Atlantic coast of North America and northern Europe are distinctive on account of exceptionally high conifer diversities: 15 genera in the Aptian to early Albian flora of Primorye (Krassilov, 1967a) and a comparable number in the Potomac (Berry, 1911) and Scandinavia (Srinivasan & Friis, 1993), perhaps representing an analogue of the present-day montane coniferous rainforest. In the phytogeographic scheme of Cretaceous biomes that is further considered in VII.2.3, this submeridional zone is extended to southeastern Asia on account of Sequoia finds in Khorat Series of Thailand (Kon’no, 1963). The compound palmato-pinnate Sapindopsis–Debeya leaf morphotypes are common over the warm-temperate zone, gradually increasing to dominance, in parallel with the narrow-leaved laurophylls (“Proteoides”), in the adjacent xerothermic zone extending over central Europe (Kvaèek & Knobloch, 1967), northern Africa (Lebanon: Krassilov & Baccia, 2000) and Caspian region to northern China (Guo, 1985, 1996). The Cretaceous palm records associated with xerothermic zone extend to Vancouver and northern Japan Islands. Notably, the northern limits of warm-water foraminifera reach to Vancouver in the Pacific (McGugan, 1982) and to Orphan Knoll in the Atlantic (Berggren & Hollister, 1977), and that of rudist mollusks (Philip, 1972) nearly coincides with the temperate/subtropical phytogeographic boundary. The mid-Cretaceous floras of Crimea (Krassilov, 1984), Central Asia (Samsonov, 1970), northern Africa, India (Barkar & Chiplonkar, 1973), central China, southeastern Asia (Kon’no, 1963) and southern North America fall into the xerothermic zone but differ in an absolute preponderance of scale-leaved coniferoids. Occasional records of scale-leaved assemblages (single-species, of shed leaves, as described in Krassilov, 1978a)

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extend to the tropical zone on both sides of the equator (the present-day and palaeolatitudinal Cretaceous positions alike). The equatorial biome is thus defined as a taxonomically poor (deciduous?) variant of the Brachyphylletea. The elatero-plicate-rimulate palynoflora of equatorial western Africa – South America (Herngreen & Chlonova, 1981) is traceable to Papua–New Guinea (Batten & Wenben, 1987) indicating the equatorial zone nearly coinciding with the present. Major Cretaceous salt deposits are confined to this zone. The equivalent biomes of the Southern Hemisphere are nearly symmetrical, with analogues of the Meso-Mediterranean vegetation in southern South America and southern Australia. The Late Cretaceous boundary of the nanophyllous subtropical and microphyllous temperate biomes runs at about 45°S (Specht et al., 1992), not far from the present-day position of the temperate/subtropical ecotone, implying a considerably cooler climate than in the Northern Hemisphere. The respective marine climates (the Austral mollusk province) extend to 40°S (Doyle, 1987). The Proteaceae are more diverse in the Senonian of southern Australia than in the contemporaneous floras of New Zealand, with Knightia lacking in the latter. On morphological grounds, the Cretaceous leaf assemblages of New Zealand are comparable with those of the northern temperate biome (Kennedy & Raine, 2001). The temperate Antarcto-Cretaceous zone of Gothan (19011903) is based on fossil wood structures from Antarctic islands and New Zealand (Alexander, Seymour, Snow Hill, etc.). The non-marine Senonian of Antarctic Peninsula contains redbed facies while an ice-rafted material is lacking in the contemporaneous marine deposits (Birkenmayer, 1985). To the end of Cretaceous, a spread of Nothofagus (Nothofagidites pollen grains) to about 40°S in South America and 36°S in Australia indicates temperization, evidently ahead of an adequate climate change in the Northern Hemisphere. Insofar as temperization of the mid-latitudes is related to glaciation of the poles, these data indicate Antarctic glaciation that has been inferred also from sedimentological data (Francis & Frakes, 1988).

IV.3.4. Tertiary precursors of extant biomes Since Tertiary biomes are identified with the extant ones, only a brief summary of their origins (see IV.2) is given here. The laurophyllous vegetation apparently descended from an understorey/seral stage of the Cretaceous Lauro-Sequoietum segregated after the overstorey was cut off. During its maximum development in the Early to early Middle Eocene, it gave rise to the so-called “paratropical” vegetation of the Atlantic and Pacific margins extending to 60-70°N, with edaphic coastal ecotypes evolving into the modern-type mangroves. Northern ecotypes, adapted to winter darkness by physiological shutdown, might then have formed a shade-tolerant strata of more humid laurophyllous types evolving into multistratal rainforests. In the equatorial zone these were foreshadowed by the end-Cretaceous “palm zone”, but the fully developed tropical rainfor-


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ests appeared not earlier than the Oligocene, concomitant with temperization of the midlatitudes. The nemoral forests might have their origins with the platanoid–trochodendroid broadleaved communities arising from the understorey of Platano-Parataxodietum that declined over the Cretaceous/Tertiary boundary. Higher hamamelids might have evolved with altitudinal differentiation of the platanoid–trochodendroid complex, the betuloids first appearing in the Late Cretaceous highlands (records from intermontane depressions). With the upland-lowland migrations over the Cretaceous/Tertiary boundary and in the Oligocene they entered the lowland broadleaved vegetation, giving it a typical Arcto-Tertiary aspect. Coniferous rainforests of circum-Pacific belt are in a direct line of descent with the ecologically equivalent Cretaceous biome. Distinctions in their modern aspect is mainly due to the invasion of pinaceous conifers, still modestly represented in the Late Cretaceous but increasing during the early Palaeogene to give rise, in the Late Eocene already, to the montane taiga with Sequoiadendron, Metasequia and laricoid–piceoid genera of modern affinities. The Palaeogene xeromorphic communities of the Dryophylletum type are closely linked with the Cretaceous forerunners. They periodically expanded and shrank with climatic fluctuations, incorporating scleromorphic ecotypes of laurophyllous vegetation. Grasslands seem to have appeared, on the basis of the Early Palaeocene grass-sedge wetlands, during the Eocene xerothermic phase (on evidence of faunistic records mainly), but their modern aspects and ranges are owing to aridization over the Miocene/ Pliocene boundary and in the Holocene. The Pleistocene high-latitude savannoid vegetation supporting a rich fauna of larger herbivores was replaced by xeric dwarf-shrub tundras over the Late Pleistocene – Holocene transition. These biomic events are part of global change.

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Chapter 5. Tectonic factors of global changes

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V. TECTONIC FACTORS OF GLOBAL CHANGE Changes of the face of the earth are surface manifestations of interior processes. Their impact on the biosphere involves: - endogenous heat flux, - atmospheric effects of mantle outgassing, - effects of volcanic ash/aerosol on climate, stratospheric ozone layer and biota, - tectonically derived sea-level fluctuations, - orographic effect on the global heat transfer, precipitaion, drainage and terrestrial runoff to marine basins, - biological diversification over heterogeneous landscapes of tectonically active areas. Not only certain habitats (mountain ridges, rift valleys, hydrothermal vents, etc.) owe their origin to particular tectonic events, but also the biomass/dead mass production of biosphere as a whole is under direct or indirect control of tectonomagmatic activities. Hence an explanation (prediction) of global change depends on our understanding of the nature and regularities of tectonomagmatic processes.

V.1 Current situation in the geotectonics Despite ca. 700 years of geological research, the driving forces of planetary geological processes remain obscure. Geology is traditionally considered as a regionally orientated, hard fact-backed discipline with an inherent aversion for global generalizations. Global geological models come not so from geology as from geophysics, a technically advanced branch of the geosciences with an opposite tendency of overgeneralizations and a superlative attitude to trivial geological observations. Yet the geodynamics disentangled from geology is fertile ground for unverified assumptions to grow into dogma. The currently dominant plate tectonic theory has replaced the geodynamic pluralism of mid-1900s, with the then rivalling earth contraction, expansion, pulsation, polar wander and continental drift theories, each of them illuminating a selected facet of regional geology rather than providing a global synthesis. In contrast, the new global tectonics, as it was called in the 1970s, emerged from a geophysical compilation of deep-focus earthquake data that showed a limited number of lithospheric plates divided by the focal zones (Isacks et al., 1968; Le Pichon et al., 1973). The plate model was then coupled with the Wegenerian theory of continental drift. However, the rotation forcing of the latter was substituted by a mantle convection forcing, with mantle plumes surfacing as magmatic hot-spots. As a matter of fact, the rotation forcing was a focus of anti-Wegenerian criticism of 1950s. Hence it was essential for the new paradigm to resolutely alienate itself from rotation. Geological evolution was interpreted as resulting from a convective rifting, rafting and consumption (subduction) of lithospheric plates. As the model developed, inter-

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pretations overshadowed the phenomena. So the mid-ocean rises and trenches became “spreading centres” and “subduction zones”. The theory thereby became immune to any factological criticism. Meanwhile the mantle/lithosphere decoupling at the astenospheric zone posed a problem for the assumed convection rafting (Runcorn, 1980). The alternative models of plates being pushed from mid-ocean ridges, pulled into subduction zones or driven by hot spots were proposed as preliminary, so far lacking substantial geodynamic foundation, and were left like this. In the case of the marginal Benioff zone (“subduction zone” of plate tectonics) dipping away from the continent, it is not the denser oceanic crust, but the lighter back-arc crust that is allegedly being subducted, which is contrary to the basic postulates of the model. Such inconsistencies are camouflaged by assertions that both spreading and subduction may occur at any location – they jump, change geometry, replace each other, a game without rules. Seismic data indicate isostatic compensation of the earth crust heterogeneities, with the mantle melting zones depressed under continental swells (Li et al., 1998; LithgowBertelloni & Silver, 1999) and mid-ocean ridges (D.S. Wilson, 1992) and raised beneath oceanic depressions (Wan et al., 2001). The magmatic continental margins (“subduction zones”) are underlain by mantle upwellings, the mid-ocean ridges – by downflows (Keith, 2001), whereas the plate tectonic model postulates the opposite. A geometric puzzle of the simultaneously spreading Atlantic, Arctic, Indian and Pacific oceans, compensated by peripheral subduction in the latter alone, was not resolved by the model. Contrary to the postulates of the model, at least some of the supposed subduction zones are admittedly extensional rather than compressional (Aubouin et al., 1984). An interpretation of mid-ocean ridge crests as conduits of mantle material for seafloor spreading is disproved by the heterogeneity of ridge-crest volcanics and by occasional occurrences of sediments much older than predicted by the model (the Maastrichtian on the equatorial mid-Atlantic ridge: see Rezanov, 1993). A localized – in a series of along the strike knots – magmatic activity of mid-ocean ridges, complicated by a widespread off-ridge volcanism, could not produce, as the original theory holds, the parallelsided stripes of oceanic crust marked by linear magnetic anomalies. The present-day dynamics of the mid-ocean ridges indicates a leaky strike-slip zone, with an extenstional component varying along the axis, impregnated with a sheared ultramafic material, and with volcanism mostly confined to the nodal basins at intersections with transcurrent faults (responsible for the so-called hot spots). The pattern of parallel sea-floor magnetic anomalies can be alternatively (and perhaps more convincingly) ascribed to impregnation of magmatic material in the series of extensional faults flanking mid-ocean ridges. Their apparent off-ridge age progression might reflect a sequence of magnetic signatures acquired with progressive cooling of impregnated material away from the heated axial zone (Krassilov, 1985, 1995a and a similar explanation in Keith, 2001). Actually, the “beaded” structure of magmatic anom-


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alies (Schouten & Klitgord, 1982), with polarity changes over their vertical sections (Lowrie & Alvarez, 1981), better agrees with impregnation than spreading. In other words, the so-called spreading anomalies (the M-series beneath the Cretaceous sediments in the Atlantic) are basically of the same nature as the J-anomaly of the same zone reflecting an exhumed basaltic ridge. Likewise, the “quiet zone” of continental margin is an exhumed mantle impregnation into the marginal fault zone (Witmarsh et al., 2001). As for the age progressions over mid-plate volcanic ridges, such as the Hawaiian Islands, their plate-tectonic interpretation as hot-spot (plume) traces is scarcely compatible with radiometric dating. The Hawaiian Islands and the roughly parallel Line Islands occur on the same plate, but are of the widely disparate chronologies and rates of volcanic progression. Moreover, the diagonal Line lineament overlaps in age with the orthogonal Emperor lineament – a paradoxical situation of two simultaneous hot spots leaving diverging traces on the same plate. No systematic age progression is found over the Cameroon Line, an alleged plume trace in western Africa (Holliday et al., 1990). In the case of the Ninetyeast Ridge, Indian Ocean, the age progression is inconsistent with the hot-spot mechanism (Sclater & Fisher, 1974). The plume theory fails to explain a midplate volcanism along the Cooke – Austral Island Chain, with an inconsistent age progression apparently controlled by lithospheric stress (McNutt et al., 1997). Such features more realistically reflect propagation of tectonic stresses and concurrent volcanic activities over the parallel, but diachronous, shear zones. Insofar as igneous activity over meridional lineaments systematically propagates towards the equator, the Eötvös forcing appears a possibility. Although still afloat owing to the powerful lobby, the plate tectonics theory shows no prospect of further development, which means that it will be soon abandoned. Stones rejected by the builders of the model (essentially geophysical, with little respect to stones) will then be used as cornerstones for new geodynamic theories. But even then plate tectonics will remain in the annals of epistemology as a paradigmatic Kuhnian paradigm, as such deserving a critical analysis. Meanwhile it seems best to seek explanations of global tectonomagmatic developments in the most obvious global processes, such as rotation of the earth. The rotational model is by now the only one accounting for a correlation of geomagnetic, tectonomagmatic, epeiric and climatic processes (Krassilov, 1976d, 1985, 1991, 1995a and below).

V.2. Rotaion tectonics The earth is not the only tectonically active planet, and there is little ground for believing that its tectonic structure is unique in the solar system. Therefore, geodynamic models can be verified by applying them to other planets or their moons. Comparative planetology tells us that only non-synchronous planets (those with widely disparate rotation

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and revolution periods) tend to be tectonically active (such as Jupiter’s satellites Io and Europe, which give definite evidence of volcanic activity: McKinnon, 1997), while the synchronous planets, such as our Moon, are quiescent. Generally, the more non-synchronous the spin and revolution, the more signs of interior heat generation, magnetic field and crustal motions are found. Also, the non-synchronous planets are the only plausible candidates for extraterrestrial life, which arises as an energy preserving substance evolving in response to the day/night alternation. On non-synchronous planets, tidal friction produces an asymmetrical bulge that retards their spin. This drives planetary evolution from non-synchronous to synchronous mode, with respective outer shell deformations and interior mass flows. For Earth, mean tidal deceleration is calculated as about 2.5 milliseconds in 100 years, but instrumental observations show periodic fluctuations indicating a negative feedback between tidal friction, deformation of lithosphere and the offsets of hydrospheric/atmospheric circulation. Still often denied (hence little studied) long-term positive accelerations of the Earth’s spin may have been either gravitational – during the passage over the spiral arms of the Galaxy – or inflicted by impacts of celestial bodies that transfer their momentum to our planet. In distinction from the alleged driving forces of plate tectoncs, rotation is not hypothetic. Since Galileo it has been recognized as a force shaping the globe as well as driving the atmospheric/hydrospheric circulation. According to the Wegenerian theory, it drives the continents as well. Doubts concerning the quantitative aspects of the latter were based on calculations assuming mobility of continents as integral landmasses against the oceanic floor. These calculations are no longer valid because the lithosphere is shown to be split into a great number of individually mobile plates or blocks underlain by the lowdensity asthenospheric layer. Generally, before dealing with quantitative aspects, it may be better to develop a clearer idea of what moves against what. Angular velocity of rotation determines the ratio of polar compression/equatorial expansion. Hence the first postulate of rotation tectonics, based on the physical law, states: variations in the Earth’s rotation rate generate tectonic stresses of opposite sign in the polar and equatorial areas. The postulate predicts distribution of crustal stresses, their related planetary fault network and the pattern of fold belts, as well as the geoidal sea-level fluctuations. The crustal and subcrustal stresses are defined by a physical law describing the density-dependent acceleration under rotation forcing. The Earth’s solid crust is a mosaic of plates of variable densities. Oceanic plates are denser than continental plates (average densities 2.7 g/cm3 and 3.3 g/cm3 respectively), with further differentiation within each of the types. The Earth’s interior consists of concentric spheres with sharp density contrasts across their boundaries. Rotational acceleration acting on both crustal plates and interior spheres gives them differential momenta depending on their densities. Thus, all density boundaries are tectonically stressed, with a potential decoupling and shear. This is the second postulate of rotation tectonics that applies to the shear melt-


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ing and geomagnetic field fluctuations at the interior density boundaries. Isostatic compensation of crustal plates and the consequential deformations of the geoid are related to the density heterogeneities in the rotation stress field. Earth’s rotation rates are sensitive to surface processes, such as changes of atmospheric circulation, sea-level fluctuations or polar ice growth, rendering a feedback of rotational forcing. Thus, tidal deceleration fluctuates with dissipation of tidal energy over the tide flats that expand with geoidal transgressions. The feedbacks of orogeny are owing to a mountain torque accountable for a rapid transfer of momentum from atmosphere to solid earth (Salstein & Rosen, 1994). A transformation of the Tethys into a bulging area over the Eocene as well as the larger orogenies/regressions in the Late Palaeozoic or at the Cretaceous/Tertiary boundary would have a major effect on the Earth rotation. Hence the third postulate: rotational forcing invokes geographical feedbacks. In the following discussion, the present-day tectonic structures are compared with those of the Cretaceous as a period antithetic to the present. Their comparative analysis is intended as a testing ground for the rotation tectonics.

V.3. Geoid The geoid is a figure of the Earth at the mean sea level approximating an oblate ellipsoid of rotation. The present-day geoid highs roughly correspond to the Pacific and Atlantic segments of oceanic crust and the adjacent continental crust (Fig. 57), with the highest bulge near New Guinea (+100 m). A broad geoid low extends from North Siberia to the Indian Ocean, with the deepest depression about the Maldives. The present-day geoid depressions extend over landmasses, with Africa anomalous, although close to the Maldives depression. The configuration of insular landmasses in the Malucca – Banda seas suggests fragmentation and dispersal of continental crust away from the New Guinea bulge. Conversely, the Maldives depression appears as a focus of convergence between the Indian and African landmasses. Notably, the seismically active zones presently approach the geoid mid-level (zero contour), which lends no support to the plate tectonic model of large convective mass transfer at such zones. It is worth noting that the major intracontinental trap provincies of both the Tunguska Basin (Permian/Triassic) and Deccan Plateau (Cretaceous/Tertiary) are confined to a vast geoid depression over a planetary system of orthogonal (submeridional) faults extending from the Obskaya Guba (Ob’ Gulf) in the north to the Maldives in the south. Trap magmatism marks a rise/expansion of cratonic areas, with the controlling fault zones repeatedly extensional at both the end-Permian and end-Cretaceous geological crises (IX.1). The rifted cratonic areas were depressed and inundated by epeiric seas over the Western Siberian and Turanian Plains.

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Fig. 57. Geoid: (1) zero contour, elevated areas hatched, (2) active tectonic zones, (3) highs and lows, (4) Deccan trap province north of the Maldives low (-60), (4) Jurassic dolerites of the Victoria Land graben system near the Antarctic low (-120), (5) giant ophiolite massif of Papua–Solomon Islands near the New Guinea high (+100).

The Antarctic low (-120 m) in a close proximity of the Victoria Land graben system with massive dolerites (Stump et al., 1983) is another evidence of association of the present-day geoid depressions with former extensional zones. In contrast, the New Guinea geoidal rise, presently an extensional feature, was a focus of compressional tectonics forming the ophiolitic Stanley Range – Solomon Islands foldbelt in the Cretaceous. If these relationships stood for past geological periods, then the Cretaceous geoid configurations might have been antithetic to the present one, with a geoid low over the Sunda shelf and a geoid high over the Maldives – Deccan area.

V.4. Rotational stresses Insofar as the geoid is a figure of rotation, all the Earth’s orbital/spin variables affect the stress fields in the outer shell and interior. On a short-term scale, a correlation of global seismicity with rotational variables is confirmed by instrumental observations (Munk & Macdonald, 1960; Shtengelov, 1982; Merriam, 1983; Pariyskiy, 1984). The eccentricity, obliquity of ecliptic and precession cycles are calculated for the last 5 m.y. (Berger, 1978) and are indicated by sedimentological evidence far back in time (Kerr, 1987; Laferriere et al., 1987; VII.2.1). Spin variables (wobble of the axis, fluctuations of angular momentum) appear likewise constant for the period of astronomic


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observations, but might have been different in the recent past owing to geoidal effects of glaciations and tectonic napping. Palaeomagnetic data suggest a far greater polar wander over geological periods than inflicted by the observable wobble of rotation axis. Frictional deceleration in the Earth – Moon system theoretically amounts to about 2 ms per 100 years, with the monthly/ seasonal fluctuations related, causally with feedbacks, to atmospheric, glacial and tectonic (earthquakes) processes (Pariyskiy, 1984). Fragmentary evidence of tidelites and growth increments on skeletal structures either agree with theoretical expectations (Sonett et al., 1966) or indicates a long-term variation (Rosenberg, 1982), yet quality of the data is not always adequate. Increments on shells of marine invertebrates reflect growth periodicities in phase with tidal variations over diurnal, weekly, fortnight and monthly cycles, recording rotation rates as the length of day. An assessment of growth increments on epitheca of fossil corals (Wells, 1963) has been taken as evidence of constant deceleration rates from the Devonian (396 days per year) to present (Runcorn, 1967,1979; however compare the figures for days per lunar month in Wells, 1963, Scrutton, 1964 and Aveni, 1966). Actually Wells (1963) gives a mid-Devonian range of 385 to 410 days per year. In his material, studied under low magnification, ranking of growth increments is scarcely discernible, which makes recognition of diurnal increments problematic. My SEM studies of growth lines on Cretaceous corals suggest that increments of several ranks are preserved, marking not only diurnal, but also weekly/monthly periodicities (Krassilov, 1985). Day lengths can only be deduced from the lower rank increments in the hierarchy of growth lines (Fig. 58). The coarser ridges/furrows are seasonal increments (specimens 4 cm and 2.5 cm long have 9 and 5 such ridges respectively). The second rank growth lines, still discernible at low magnification, correspond to monthly increments (perhaps corresponding to “daily increments” shown in Wells, 1963). They are grouped into zones of narrower (5-6 months) and wider (6-7 months) bands between the annular ridges. The weekly and daily lines are discernible with SEM over the betterpreserved areas of the epitheca. They give about 24 days per lunar month, nearly in agreement with the Late Cretaceous estimates by Khan & Pompea (1978). Despite many complications, such data suggest a fluctuant, rather than linear, change in rotation rates over geological time.

V.5. Planetary fault system Geometry of the earth’s rotational stress field is defined by (1) the polar – equatorial compression – extension gradient, (2) the inertial (Coriolis/Eötvös) forcing, and (3) a disparity of angular momentum over the density heterogeneities of lithospheric structures. In the outer shell, the rotational stresses generate a pattern of faults that divide the lithosphere into a great number of larger and smaller blocks.

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Valentin A. Krassilov. Terrestrial Palaeoecology a

b

c

Fig. 58. Growth increments of fossil corals as evidence of day length: (a) seasonal (ridges and furrows) and monthly (thin striation) increments on epitheca of a coral from the Early Cretaceous of Crimea, (b) SEM of the monthly increments, (c) SEM of diurnal increments marked on the better preserved epitheca areas (Krassilov, 1985).

The major elements of planetary fault network were denoted by Stille (1924) as the orthogonal and diagonal, or O and D systems. Subsequent theoretical and experimental work has shown these fault systems to conform to the geometry of rotational stress field (Vening-Meinesz, 1947; Katterfield & Charushin, 1970; Kogan & Khatarov, 1982; Chebanenko, 1977; Chebanenko & Fedorin, 1983; Melosh, 1977). According to Vening-Meinesz (1947), rotational stress can split a lithosphere of half the present-day thickness. This means that the fault network was laid down early in the earth history when the hard shell was not only thinner but also of a more uniform density. Further developments might have distorted or even obliterated the original fault pattern. However, some presently, or recently, active fault systems can be traced back to the Precambrian. The East-African great rift system is a spectacular example. It is defined by a combination of O/D faults forming S-shaped joints extending 4000 km from Bab el Mandeb to Zambezi River. It is traceable 1500 km further south, to the Orange River, as a fossil rift silled by Great Dyke about 2.5 – 2.8 billion years old (McConnel, 1972). Various


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parts of the rift system were sequentially active during the Archaean Mozambique tectonism, Proterozoic Katanga Cycle and Palaeozoic to early Mesozoic Karroo Cycle. The younger Mesozoic cycle of the terminal Jurassic – Early Cretaceous, about 136 – 123 Ma, commenced with alkalic magmatism in the southern Malawi segment, then spreading northward over Tanzanian and Kenyan segments in the mid-Cretaceous, 101 – 113 Ma, and over the Cretaceous/Palaeogene transition, 60 – 70 Ma (Macintyre, 1977). Cenozoic magmatism was confined to the northern segment. In the Tanganyika Lake zone and Mutigo Fault zone, Kenya, Cenozoic intrusions cut the Archaean ultramafic bodies. The faulted hard shell adjusts to rotational stresses by both radial and tangential shear of fault-bound crustal blocks. Respectively, the block boundaries are of three types: normal-faulted, reverse-faulted (thrust) and strike-slip faulted. In agreement with the second postulate of rotation tectonics (V.2), the most conspicuous fault boundaries separate lithospheric blocks of discrepant densities, such as the continental and oceanic plates that are decoupled in the rotational stress field. Both normal faulting by centrifugal forcing and strike-slip translation driven by disparity of the angular momentum, are to be expected at the block boundaries (Fig. 59).

Fig.59. Segments of orthogonal and diagonal fault systems active in the Cretaceous. The numbered lineaments belong to the transcontinental/transoceanic systems of long-term tectonic activity: (1) Denali fault – Rio Grande fault, (2) Amazon fault system – Gulf of Guinea– Gulf of Aden, faulted coast lines, (3) Guinea Ridge – Cameroon Line – Iskenderun Bay median fault zone, (4) Walvis ridge – Owen fault – Chaman fault, (5) Scoresby Sound (East Greenland) – Viking graben (North Sea) – Vardar rift – Red Sea rift, (6) Pripyat – Dnepr – Donetz – Emba graben system – Karakoram fault – Hanoi rift, (7) Yamalo-Pur rift system (West Siberia) – Ninety-East ridge (Indian Ocean), (8) Godavari – Fitzroy graben system, (9) Lomonosov rise – Sette-Daban fault – Central Sikhote Alin fault – Fossa Magna fault zone (Japan), (10) Mongolo–Okhotsk rift system, (11) Emperor – Line Islands lineament.

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Geological data seem to come up to these expectations. All the continental shelves are bordered by disjunct faults. Examples are the right-lateral circum-Pacific shear fault system extending from Atacama of the Chilean margin to Denali in Alaska and from Fossa Magna of Japan to Alpine Fault of New Zealand. Their impregnated magmatic wedges are responsible for the marginal magnetic anomalies that are most prominent along the Atlantic and Indian margins (Hayes & Rabinovitz, 1975; Rabinovitz et al., 1983). The so-called passive (Atlantic-type) margins are normal-faulted, while the socalled active (Pacific type) margins are reverse-faulted. Both types have a more or less considerable strike-slip component. Geomorphologically, a combination of normal and strike-slip faulting is expressed by straight shelf lines, while reverse faulting generates arcuate margins. Straight shelf lines, as of the Blake – Bahamas Escarpment or the Arctic Shelf, indicate a presently (or recently) active strike-slip faulting. Transcurrent lineaments are offset by such marginal faults. For instance, the Gakkel Ridge of Arctic Ocean is conspicuously offset by the sinistral strike-slip faulting along the Svalbard – Severnaya Zemlya shelf margin (Bogdanov et al., 1995). The familiar continental outlines are defined by intersecting O/D fault systems forming several types of geomorphological patterns. Most common of these are triangular wedges, rectangles and S-figures. Thus, straight arms of the Arctic Siberian – Sverdrup shelves meet at the Mackenzie Delta forming the Beaufort Sea triangle. The wedgeshaped southern extremities of Greenland, Africa, Arabia, India and Australia (with Tasmania) are formed by intersecting D-faults. No less conspicuous configurations of the present-day continental margins are the O-fault bound rectangular gulfs (bays), such as the Gulf of Mexico, Gulf of Honduras, Bay of Biscay, Anadyr Bay (Anadyrskiy Zaliv), Iskenderun Körfezi, Gulf of Guinea, etc., with right-lateral displacement in high to mid-latitudes switched to left-lateral in the low latitudes. The Sea of Okhotsk and the Gulf of Alaska are symmetrical structures formed by combinations of orthogonal (latitudinal) and diagonal faulting. S-figures likewise consist of O and D elements. Notably, an S-shaped system of intracontinental Tanganyika – Rukwa – Malawi rifts is repeated, with parallel orocline bending at the equator, by the North Atlantic – South Atlantic and the Arabian – Central Indian mid-ocean rises. In the Pacific, a similar S-configuration is formed by the Tuamotu – Line islands chain. Symmetry of these planetary structures over equator is spectacular evidence of their rotational origin. In the D-pairs, the northwest striking shear faults are dextral, while their intersecting northeast striking shear faults are sinistral. They thus form a pair of forces that push the fault-bound crustal wedges to the north. India is the most convincing example of northward displacement in whatever dynamic model. Arabia moved to northeast with the opening of Red Sea, while a northward displacement of Greenland is indicated by palaeophytogeography (IV.3.4)


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V.5.1. Mid-ocean rises and transcurrent share As noted above, the mid-ocean rises are S-shaped fault zones consisting of D-elements over mid- to high latitudes and of O-elements over low latitudes. The Atlantic and Indian mid-ocean rises are roughly parallel to each other and to the East-African transcontinental rift system. The latter is flanked by swarms of minor faults and the same must be true for the mid-ocean ridges which arise as leaky median cracks of thin oceanic crust split by a pair of forces generated by marginal shear at the contacts with the much thicker continental crust (Krassilov, 1985). The plate tectonic concept of mid-ocean rises as sea-floor spreading features is based primarily on interpretation of the linear magnetic anomalies as sea-floor increments. An alternative interpretation of these lineaments as fault-bound crustal blocks (Krassilov, 1985) relates the age progression of remnant magnetization and sedimentary cover to a sequential cooling and coeval subsidence of block-faulted crust away from the heated ridge crest. This mechanism explains a correlation of successive magnetization events with deposition rates, yet not precluding sediments older than remnant magnetization of sea-floor basalts. In fact, although pre-Tertiary sediments are mostly assimilated by the ridge volcanics (Saito et al., 1966), they are locally preserved in nodal basins at intersections with transcurrent faults (Bonatti & Crane, 1982). Even Cretaceous coaly shales and breccia with shallow-water fossils were occasionally found in the axial zone of MidAtlantic Rise (Rezanov, 1993). Heat flow over axial zone is evidence of friction heating rather than of extension and passive intrusion of basaltic magma. Actually, a left-lateral shear over the ridge crest is confirmed by geodetic work in Iceland (Vadon & Sigmundsson, 1997). The volcanics of the rises are heterogeneous, varying from tholeiitic to calk-alkaline to rhyolitic, as might be expected of a combined shear-extensionl kinematics, and even on the ridge crest the tholeiites constitute no more than 50% of lavas (predominantly rhyolitic on Hekla in Iceland: Thorarinsson, 1967). An extensional component indicated by tholeiitic petrochemistry is imposed by a pull apart of ridge borders by trancurrent shear. The mid-ocean rises are cut by a series of transcurrent shear faults that displace the ridge crest segments giving them a broken-stick outline. At least some of transcurrent faults extend into the adjacent continental crust (which does not comply with their plate tectonic interpretation as transform faults). Thus, the Romansh Fault is continued, with marginal offsets, to the intracontinental Amazon Rift System (Grabert, 1983) and, to the east, to the inland Cameroon line marked by alkaline volcanics (Machens, 1973). The fault-bound England Seamounts, Walvis Ridge and probably other transcurrent ridges extend over continental margins (Shafter, 1984; Mayhew, 1986; Hurtubise et al., 1987). In the North Atlantic, a transcurrent fault zone defines the southern margin of Biscay and projects as the Main Pyrenean Fault (Vielzeuf & Kornprobst, 1984). In the Pacific, the Tuamotu – Line island chain marks a fossil mid-ocean rise, with the left-lateral Papua – Solomon Island shear zones as its trancrurrent fault. The Fossa Magna of Japan

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and the Alpine Fault Zone of New Zealand are likewise related to transcurrent shear zones of oceanic crust. Ultramafic bodies and their associated metamorphic, volcanic and sedimentary rocks found in oceanic trancurrent fracture zones correspond to the Alpine-type ophiolite assemblage (as defined by Penrose Field Conference, 1972). They typically consist of tiered ultramafic tectonites, cummulative gabbroids, dike swarms with doleritic bodies, and basaltic to rhyolitic lavas interbedded with chert and pelagic limestones. Ophiolites were once interpreted as remains of oceanic crust and upper mantle generated in midocean fissures and obducted upon continental margins. However, both deep-sea drilling and seismic profiles of oceanic lithosphere have shown an only superficial similarity with onland ophiolite assembages (Anderson et al., 1982; Lewis, 1983) inviting a different interpretation. Ultramafic rocks of transcurrent fault zones are emplaced as hot bodies, with contact metamorphism of the surrounding limestones/shales. In the Clarion Fault, the ultramafic tectonites, gabbroid cumulates and the dike swarm – lava complexes are developed as distinct tiers thrust upon each other (Rudnik et al., 1984) as is typically the case in foldbelt ophiolites. On the Gorringe Bank, a dissected transcurrent fault ridge off the Portuguese coast, large ultramafic and gabbroid blocks slide down the terraced slopes and are buried in serpentinite sand – a would-be matrix of ophiolite melanges (La Gabrielle & Auzende, 1981). Transcurrent faulting shows a non-random latitudinal distribution obviously related to the rotational stress-field. The amplitudes of transcurrent shear tend to increase, with the gradient of angular velocities, towards the equator, forming large steps at “critical latitudes”. In both hemispheres, prominent transcurrent fault zones occur at about 40° (Azoro-Gibraltar zone of North Atlantic, Mendocino zone of North Pacific, Kermadek – Hikurangi zone of New Zealand). Such “critical latitudes” of solid crust reflect zonal rotation of the underlying astenospheric layer. Segmentation of Mid-Atlantic Rise makes the pattern quite obvious (V.6). The larger displacements at about 40 – 50°N correspond to the critical latitudes of atmospheric circulation, likewise driven by the earth’s rotation. South of 20°N, the ridge segments are increasingly shifted to the east, most prominently by the equatorial Romansh Fault, while the northern slope of Brazil is shear-faulted, with a sinistral displacement of Amazon cone (Grabert, 1983). The Amazon fault system mediating a sinistral displacement of the Peruvian block (Heki et al., 1983) and the volcanic Cameroon Line extending inland from the Gulf of Guinea are fossil lineaments of the equatorial shear zone. The MidIndian and Mid-Pacific ridges are similarly offset by equatorial faults. This zone of maximal displacement cuts both the Malvinas – Chagos and Tuamotu – Line chains and extends to the Galapagos Fault Zone. The significance of equatorial maximal shear zone (belt) as a rotational feature, coinciding with the equatorial minimum of Coriolis force, has been recognised by Tanner (1964), Krause (1970) and occasionally even by the proponents of plate tectonics (So-


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lomon et al., 1975; Hamilton, 1988). With continental blocks amassed to the north of the equator, the decoupling of hemispheres by a giant shear belt is enhanced by the difference in their crust densities.

V.5.2 Cretaceous O/D faulting and magmatic activity Except the equatorial shear faults of the Amazon system in Brazil – Cameroon Line in Africa, a few presently or recently active fault zones were magmatic in the Cretaceous. Over the African rift system, the Cenozoic and Cretaceous volcanics only locally overlap in the northern segments. On the other hand, many presently quiescent continental and sea floor lineaments are fossil magmatic ridges of Cretaceous age. They belong to both O- and D-systems, of non-random distribution over the globe. The O-faults over high northern latitudes, such as the North Sea Rift System, the Ob’ Bay – Pur Rift of Siberia, the Lomonosov Ridge – Verkhoyansk Range System of the Arctics, etc. (Fig. 59) had controlled the Triassic rifting and trap magmatism of cratonic areas followed by a shallow subsidence of epirift basins in the Cretaceous. Active O-faulting between 30°N and the equator involved the intracontinental Gulf of Aqaba-Jordan rift system and a series of regularly spaced meridional faults traversing the oceanic crust between Madagascar and Andaman Islands, with such prominent linear features as the Chagos – Laccadive Ridge, Indrani, Indra, 86° Ridge, East Indian (90°) Ridge, Investigator, etc. (Fig. 60). Their inland extensions (the Kambay Graben) are marked by the Cretaceous traps at the intersection with the diagonal Narmada – Son rift zone of Deccan Plateau. Fossilized O-faults are recognizable in the Pacific as well. Here they are represented by the submerged meridional Shirshov Ridge and its parallel Emperor Ridge that extends from Aleutian Islands to the Hawaii and is traceable further south to intersection with the latitudinal Marcus-Neccer Ridge. Cretaceous volcanics are found on all these ridges (Jackson & Schlanger, 1976; Saito & Osima, 1977). At the same time, the D-faults were more conspicuously active over the mid-latitudes. The Mongolian (Kerulen) basaltic province and its outliers in Transbaikalia and Heilongjiang are controlled by the NW – SE Mongolo-Okhotskian system, intermittently active since the Precambrian (Buchan et al., 2001). In the Atlantic, the Southeastern Newfoundland Ridge, New England Seamounts, Walvis, Rio Grande, all magmatic in the Cretaceous, are segments of the D-system. An extensive diagonal rift zone is traceable over the Red Sea – Vardar Rift – Rhine Graben (Zytko, 1982). It is intersected by the submerged Owen Ridge extended inland as the Naramada – Son fault system. To the south, a no less extensive oceanic D-system comprises the Kerguelen – Crozet – Broken Ridge – Diamantina fault zones marked by the Cretaceous volcanic ridges (Guilfy, 1973; Silantyev et al., 1983).

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Fig. 60. Fault zones and marginal structures of the Indian Ocean. The Australian margin: (Bo) Bonaparte, (Ca) Canning, (Cr) Carnarvon, (Fr) Fitzroy. The Indian margin: (CP) Cauvery – Palk, (Go) Godavey, (Mh) Mahanadi, (NS) Narmada – Son. Oceanic fault zones (ridges): (AN) Andaman-Nicobar, (BD) Broken – Diamantina, (Ch) Chagos, (La) Laccadive, (In) Indira, (Id) Indrani, (Iv) Investigator, (KC) Kerguelen – Crozet, (Ma) Madagascar, (Ms) Mascarene, (Ow) Owen, (90°) Ninety-East Ridge. (1) Peri-Indian ophiolite belt extending to the Andaman – Nicobar Ridge (AN), Mentawai trough (M) and Timor (T), (2) Volcanic zones of the African margin and the Deccan Plateau, (3) Marginal basins, (4) Deep sea drilling stations, (5) Oceanic ridges, (6) Mid-ocean Ridge.

Such distribution of planetary fault systems is consistent with the modelled fault pattern of a despan planet (Melosh, 1977). Certain regularities emerge also form comparative chronology of volcanic events. In the Kerala zone of southern India, the Precambrian basement is cut by several generations of dykes, ranging from Palaeozoic to the Cretaceous – Tertiary boundary, the latter coeval with Deccan traps of the Narmada – Son zone, as well as with the Cretaceous basaltic province of Madagaskar (Pande et al., 2001). Instead of recording alleged breakup between India and Madagascar, these comparisons indicate synchronous magmatic events over the African and Indian segments of planetary fault system.


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The older volcanics in the diagonal fault zones of Red Sea, Owen Ridge, Narmada – Son, etc. are of the mid-Cretaceous age (the “pre-rift phase” in the Red Sea zone, the earlier traps of Shillong Plateau, Rajmahal Hills and elsewhere: Subbarao & Sukheswala, 1981). The major magmatic phase falls at the Cretaceous/Tertiary boundary (Mahoney et al., 1982). Basalts of the Ninety East Ridge in the Indian Ocean are of the same age. The initial stage of Thulean basaltic province in the North Atlantic is roughly coeval (Noe-Nygaard, 1974), as are the younger volcanics of D-ridges in the central Pacific (V.6). Here the 4500 km long Line Islands Chain forms a giant offset zone extending to the Emperor Ridge in the north (Farror & Dixon, 1981). Over the Cretaceous, this structure might have assumed the role of a mid-ocean rift system, with nephelinites and related alkalic volcanics. The Line Islands Chain is crossed by a series of west-northwest striking ridges, the so-called Cross Trend, also magmatic in the Cretaceous, but of a somewhat diachronous activity (Winterer, 1976; Jackson & Schlanger, 1976). The sequential shift of volcanism from the Cross-Trend to Line to the Emperor Ridge might have reflected a changing orientation of stress field during the Late Cretaceous. Thus, as indicated by the ages of volcanic ridges over oceanic fault zones, the maximal activity of D-system is confined to the early – mid-Cretaceous. It has been intercepted by the O-system in the latest Cretaceous and over the Cretaceous/Tertiary boundary. Re-orientation of the stress field might have been due to expansion of equatorial bulge in turn related to acceleration of the Earth’s rotation rates to the end of the Cretaceous.

V.6. Mobile belts Since the rotational acceleration is density-dependent it generates a strike-slip shear at the boundaries of continental, oceanic and transitional crust (V.2). Thus, cratonic blocks of tectonically active areas are typically bound by the systems of dextral/sinistral faults, as in the case of Tibet Plateau bordered by the Kun Lun (sinistral) and Karakoram – Gangdize (dextral) fault zones with ophiolite assemblages (Xuchang et al., 1980; Girardeau et al., 1984;). Such shear zones tend to acquire annular configuration mediating a rotation of crustal blocks. In agreement with the postulates of rotational tectonic model (above), zones of longterm tectonic activity, or mobile belts, are confined to the major density boundaries, such as between oceanic and continental crust. In fact, all the continental margins are tectonically active (as will be shown later in the chapter, a distinction between the so-called active and passive margins, stressed by plate tectonics, is transient, with reversals related to geoid configurations and the tilt of the continents). Instead of being, as sometimes interpreted, chaotic collages of terrains rafted from various locations and accreted at different times, the marginal mobile belts are flexible buffer zones mediating the rotation

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of oceanic crust against its continental borders. They are cut into independently rotating minor blocks that are sutured by ophiolites at their marginal overthrusts and consolidated by granitoid magmatism, with tectonic developments synchronized all over the belt. A first order planetary structure of this type is the circum-Pacific shear belt mediating anticlockwise rotation of the oceanic plate against its continental surrounding. It can also be considered as a boundary of the continental/oceanic hemispheres. In this respect it is analogous to the equatorial shear belt (V.5.1).

V.6.1. Pacific arcs Morphologically, the circum-Pacific shear zone has a coastal range – backarc basin – island arc – trench structure at the western margin and a simpler coastal range – trench structure in the east. The typical geological structures are the ophiolite assemblages, sedimentary and volcanomictic melanges, paired metamorphic belts of blue schist – glaucophane/eclogite facies, granitoid intrusions and their related andesitic to trachitic volcanics. These structures have repeatedly developed at successive time levels – since the Late Palaeozoic at least, but probably much earlier – as the concentric, partly overlapping belts. Over tectonic cycles, a thickening of marginal crust by napping, granitization and superimposed terrestrial volcanism was followed by fragmentation at the next stage. A counterclockwise rotation of the Pacific floor, postulated by Benioff (1954), Tanner (1964), Allen (1962), and other researchers (see Badgley, 1965; Krassilov, 1985 for review) agrees with a dextral shear over the bordering fault girdle. The American margin from the Atacama Trench off Peru to the Tintina-Kaltag fault system of Alaska is a continuous right-lateral strike-slip zone. Western Pacific is similarly framed by a rightlateral fault system extending from the Alpine Fault Zone of New Zealand to the Tonegava (“Fossa Magna”) Fault of Japan and perhaps continued as the Median Kamchatka Ridge. Together they form an annular shear system. On the continental side of backarc basins it is coupled with the sinistral fault systems, such as the Central Sikhote Alin Fault and parallel fractures in the Sea of Japan. The Sorong-Sarawak Fault Zone of Indonesia may belong to this sinistral system (Katili, 1970). It has been long assumed that geomorphological differences between the western margin with extensional backarc depressions and eastern margin with trenches adjacent to and occasionally overriden by the continent are owing to the rotational Coriolis effect. The difference is not absolute, however, and might not be that obvious in the past (V.6.3). The Queen Charlotte Islands of the eastern margin form an arc with Cretaceous volcanics. Baja California is a strip of marginal belt cut off by an extensional shear zone (below). The better studied examples of island arc – trench evolution (such as Japan Islands on New Zealand) suggest that they initially developed as Andean-type margins, then thrust on the adjacent oceanic floor acquiring an arcuate configuration. Most backarc seas are pull-apart basins created by a combination of along the shore and transcur-


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rent shear. On the western margin, the Coriolis effect adds to the pull-apart tension while counteracting it on the eastern margin. A displacement of the present-day fracture zones of eastern Pacific against their preceding Cretaceous structures marked by the fossil ridges (Carnegie/Tehuantepec, Clipperton/Cocos, Galapagos/Nazca) reflect re-orientation of the stress field that has been also suggested in respect to the Pacific fault-bound island chains (above). An alleged polar wander would inflict much larger displacements, hence improbable. At the same time, as discussed in (V.6.3), the Cocos Ridge is a fossil transcurrent structure that might have extended across the thrust zone of Osa Peninsula to the geophysically similar Beata Ridge in the Caribbean. Similarly, the Tehuantepec Ridge is displaced against its putative Cretaceous position in line with the Motagua thrust zone and its Caribbean extension as the Cayman ridge/trough system. Their parallel Nazca Ridge might have conceivably formed a triple junction with the Andean orocline at about the Peru – Chile border. The offsets are ca. 500 km implying ca. 5° counterclockwise rotation since the Cretaceous.

V.6.2. Trenches and ophiolites Trenches are defined by deep fault zones, expressed as the earthquake focal (Benioff) planes, typically dipping at 30° to about 60 km, then at the an increasingly steeper angles of 40° to 300 km and of 60° to 700 km (Benioff, 1954). This geometry probably reflects a combination of strike-slip/thrust faulting. There are noteworthy variations in the middle member in the Andean and Pacific fault planes correlating with the density contrasts of continental/oceanic and transitional backarc/oceanic crusts (Letnikov, 1997). The transitional backarc crust about 25 km thick is attenuated by subcrustal shear and its related eclogite phase metamorphism, with extensional features expressed as graben-like depressions, such as the Tinro or Deryugin troughs in the Sea of Okhotsk (Gnibidenko, 1983) or the Yamato trough in the Sea of Japan (Ludwig et al., 1971; Honza, 1979). Geological data suggest a continuation of Sikhote-Alin structures over the Sea of Japan (the Triassic granitoids on Yamato Rise: Ludwig et al., 1971), with the Central Sikhote-Alin Fault Zone corresponding to the Titibu Belt of Central Honshu, the coastal volcanic belt – to the Hiroshima magmatic province (Krassilov, 1985). Their separation is owing to an eastward displacement and bending of the arc hinged at the Itoigawa – Fossa Magna Zone of Central Honshu (palaeomagnetic evidence in Sugimura & Uyeda, 1973; Uyeda & Miyashiro, 1974). Volcanism of the submerged Oki Ridge (Hilde & Wageman, 1973) reflects a concomitant backarc extension. On the leading edge, the metamorphic rocks are thrust from beneath the island arc. The thrust sheets are stuck in the inner slope (the Oyashio Thrust: Lomtev & Patrikeev, 1983) or transported beyond the trench. Such thrust sheets are outlined by granitoids

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found on the Zenkevich Ridge and Erimo Seamount outside of the Japan – Nansei trenches (Vassilyev et al., 1978; Vassilyev & Yevlanov, 1982; Lomtev & Patrikeev, 1983). In the Eastern Pacific, the Middle American Trench was considered as a typical subduction zone presently consuming oceanic crust of Miocene age. Instead, the deepsea drilling data of early 1980s shows an extensional graben-like structure filled with the Upper Cretaceous to Miocene hemipelagic/turbidite/volcanic deposits onlapped by ophiolites transported from beneath the continent over a series of reverse faults parallel to the Benioff zone (Aubouin et al., 1984). Recent data, although interpreted as consistent with subduction – underplating of Cocos plate off the Nicoya Peninsula, actually indicate a downfalting of the Mesozoic marginal structures in the Miocene (Vanucchi et al., 2001). These and similar data support interpretation of Benioff planes as shear zones conducting the deep-rooted overthrusts of subcrustal material, accompanied by the shallower overthrusts cutting across the trench slopes. Given an extensional component to the Middle American trench, a spreading shear zone of the Gulf of California can be considered as its continuation cutting off a continental fragment on its way to develop as an island arc. This interpretation would explain tectonic evolution of American margin without resorting to a highly improbable hypothesis of a mid-ocean spreading ridge overriden by the continent. A comprehensive arc-trench model would also account for the Benioff zones dipping away from the continent as in the case of the Manila trench, the current explanation of which involves subduction of an imaginary buoyant plateau, as well as of a mid-ocean ridge (Boutista et al., 2001). Seismic studies of Pacific margins indicate a thick astenospheric wedge (Rodnikov, 1988) supposedly generated by subcrustal shear. During the development of the Pacific-type arc-trench systems, a friction-melted subcrustal material was exhumed from beneath the continental margin and thrust over the Benioff zone and its parallel reverse faults (Fig. 61). Attenuated by subcrustal shear, the marginal crust subsides as a backarc basin. The thrust complexes are then reshuffled by backthrusting resulting in the Manila-type trenches. The Pacific ophiolites have been interpreted as originating in either the mid-ocean ridge (MOR) or backarc or else subduction zone (SZ) environments (Pearce et al., 1985; Edelman, 1988). Actually the equivalents of the Pacific-type onland ophiolites occur in the Tonga, Puerto Rico and Middle American trenches that develop as extensional shear zones (Chase, 1971; Tucholke & Ludwig, 1982; Aubouin et al., 1984). However, the alleged distinctions between these “supra-SZ” ophiolites and those of the “supra-MOR” Bay of California may not be that profound, because both are confined to the marginal shear zones, cutting off a continental fragment in the latter example. Mechanisms of ophiolite emplacement can be sought for in the pulsating strike-slip/ thrust faulting (Fig. 61): (1) The picritic mantle material is squeezed in the leaky fissure, tectonized by shear pressure and intruded by dyke swarms and gabbros,


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Fig. 61. A schematic origin of the circum-Pacific ophiolite nappes: I. Subcrustal material is thrusted from beneath the continent over the oceanic crust that is isostatically elevated. II. The bulging nappe zone is faulted, with subcrustal material penetrating the fault zone; hanging nappes are back-thrusted upon the continental margin; oceanic crust exhumed from under the nappes sinks as a back-arc basin.

(2) The volcanic–sedimentary cover is disentangled and napped upon adjacent oceanic crust to be re-sheared toward the continent in the next phase, (3) The ultramafic bodies of Benioff zone, exhumed from under the nappes, form a median cordillera of the shear belt, shedding ophiolitic olistolites to the residual flysch troughs. The associated metamorphism evolves from the low pressure/high temperature type produced by friction heating to the high pressure/low temperature type beneath the nappe piles, while the lavas show a concomitant progression from the predominantly toleiitic in the lower horizons to andesitic and trachitic above. This model is further discussed in respect to Cretaceous examples.

V.6.3. Cretaceous circum-Pacific structures The circum-Pacific fold belt is formed by consolidation of numerous crustal blocks. The mosaic structure, comprising, from north to south, the well-defined Chukotskian, Kolymian, Omolonian, Okhotskian, Bureyan, Khankian and a number of smaller blocks has been first recognised on the western Pacific margin between Siberian craton and the coastal ranges (Fig. 62). The blocks are cratonic, surrounded by thrust belts emplaced upon their margins. Anticlockwise rotation by a system of sinistral shear is well documented for the Omolonian Massif in the Russian Northeast (Bondarenko, 1996). Such rotating blocks are considered below as the millstones of global fold belts. The Kolymian, Omolonian and Chukotskian blocks are surrounded by complete, though asymmetrical, annular shear zones with Palaeozoic ophiolitic nappes re-sheared during the Late Jurassic/Early Cretaceous phase. Consolidation of the blocks and their accretion to the Siberian craton have been accomplished with the granitoid intrusions (the Kolymian batholites and coeval granitoids elsewhere), after which the tectonic activity has shifted to the east (Archipov, 1984; Natapov & Stavsky, 1985; Alexeev, 1987).

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Fig. 62. Structural zones and the major crustal blocks of northwestern Pacific margin. Blocks: (Bu) Bureyan, (Ch) Chukotskian, (Kh) Khankian, (Ko) Kolymian, (Om) Omolonian, (Wo) West-Okhotoian, (Ya) Yamato. (1) Terrestrial volcanic belt, (2) Foredeep, (3) Cretaceous island arcs, (4) Areas of back-arc extension.


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Similar developments are recorded over the periphery of the Bureyan and Khankian blocks to the south, with their marginal structures extending over the Lesser Hingan Ranges of Amur Province and Sikhote Alin in Primorye. Volcanic belts periodically emerged over the eastern margin of the Khankian block in Primorye since the early Middle Devonian at least, with continental deposition in the rare (Krassilov, 1967a, 1967b). Thick volcanoclastic deposits of Vladivostok series indicate a mid-Permian belt probably extending to the Abo Mountains of Japan (Sugimura & Uyeda, 1973). East of it, the heterogeneous complex of reefal carbonates, tuffaceous turbidites with terrestrial plant remains, olistostromes and basalts represents an active island arc–trench system. After a relatively quiescent phase of a platform carbonate and deltaic deposition in the Triassic and Early Jurassic, an outbreak of andesitic volcanism marked a new phase of arctrench developments over the Lesser Hingan to western Primorye and further north to the Okhotsk margin, partly overlapping the Permian system. The forearc deposits of the eastern Sikhote Alin were overriden by the nappes that involved the Permian to Triassic reefal carbonates and Jurassic cherts. They were locally emplaced upon Neocomian flyschoids with terrestrial plant remains (Krassilov & Parnyakov, 1984). The approximately coeval Early Yanshan tectonism has generated the suture belts of the Anhuj – Zhejiang, Pingxiang and Anfu zones around the Yangtze and Wugongshan blocks (Wang et al., 2001) and elsewhere in eastern China. Synchronous developments took place over the Kamuikotan – Kurosegawa belts of Japan, Talkitna belt of Alaska, Gravina – Nutzotin – Queen Charlotte belt of Canadian margin, Franciscan thrust belt and elsewhere in the circum-Pacicfic girdle (Burk, 1965; Yorath & Chase, 1981; Kienast & Rangin, 1982; Yokoyama, 1987; Kojima, 1989). Geochronologically they belong to the same tectonic phase as the Alpine ophiolites and “schistes lustres” (V.6.5). The Neocomian tectonism of the Pacific phase (Krassilov, 1985) resulted in consolidation of the block mosaic over the Pacific margin, with volcanic activity shifting to the east (Krassilov, 1972d). The terrestrial volcanic belt of eastern Sikhote-Alin has developed, since the Late Albian, over the Neocomian thrust ridges intruded by granitoids and covered by their derivative lavas and ignimbrites (Krassilov, 1989b, Mikhailov, 1989). The Sikhote Alin Ranges are separated by a transform fault zone from the chronologically and structurally equivalent Okhotsko-Chukotskian volcanic belt to the north, while the southern extension is traceable over the inner zone of Honshu to the marginal South China and the Natuna Island, a Cretaceous granitoid massif in the South China Sea. In the Sikhote Alin and elsewhere over the volcanic belt, the most prolific phases of ingimbrite eruptions have occurred at the Turonian/Coniacian and Campanian/Maastrichtian boundaries. Volcanic activity went on over the Maastrichtian/Palaeocene boundary, but the tectonic regime switched from compressional to extensional, with basalts and the coal-bearing molassoids accumulating in the graben-like intermontane depressions (Krassilov, 1992b). The volcanic belts of American margin show roughly synchronous developments (e.g., in the Rocky Mountains: Dickinson et al., 1988). As on the

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Asiatic margin, the volcanic activity migrated to the east (see Hansen & Crosby, 1982 for the Rockies, Katz, 1971 for the Patagonian Andes). A series of coeval ophiolitic thrust belts extends from Sakhalin to the Lesser Kuril Islands, Kamuikotan/Hiddaka zones of Hokkaido, Kitakami zone of eastern Honshu and Simanto zone of Shikoku (Faure & Charvet, 1983; Wakita, 1983; Krassilov et al., 1988; Miyashita & Yoshida, 1988). They are traceable to the south around the East China Sea over Nansei islands, Taiwan and, on the continental side, the Zhejiang – Anhui fold belt (Fig. 63). The Philippine block to the south is framed by the Alpine-type ophiolites of Luzon and the Kyushu – Palau Ridge, a submerged Cretaceous island arc the volcanic and metamorphic rocks of which are dragged from the ridge crest, while the peridotite – gabbroid complex emerges as the Yap Island (Shiraki, 1971). Further south, a sinstral shear zone of Palawan – Sarawak Fault contains the Cretaceous to Palaeogene flysch series and, on Kalimantan, the 3000 m thick ophiolitic Rajang – Danau Group (Liechti et al., 1960; Tjia, 1970). The world largest ophiolite massif, 400 x 40 km, is thrust over the metamorphic rocks of arcuate Owen Stanley Range, Papua New Guinea, and extends along a sinistral shear zone (Krause, 1965; Katili, 1970) to D’Entrecasteaux and Solomon Islands. Its tectonized harzburgites and gabbros are of mid-Jurassic age, emplaced in the Late Cretaceous. The coeval ophiolites occur also in the Kermadek – Tonga trough on the eastern side of the Fiji block (Chase, 1971) and in New Caledonia, the uplifted edge of submerged Norfolk Rise. New Zealand is essentially a raised thrust belt over the Alpine-Taupo dextral fault zone dividing the northern mass of Lord Howe, Norfolk and Three Kings from the southern mass of Chatam-Campbell Plateau (Fig. 64). The southern edge of Norfolk (the Northern Island of New Zealand) is overthrust by ophiolites associated with the Vening Meinesz Fault, a transcurrent shear zone (Brothers & Delaloye, 1982; Schofield, 1983). Forland of the Southern Island is analogous to the Siklhote Aline – Hiroshima zone in being formed of a heterogeneous nappe complex, including the Murichiki and Torless “terraines” as the larger sheets re-sheared during the Late Jurassic/Early Cretaceous Rangitata orogeny (see Fleming, 1969; Waterhouse, 1975, Austin, 1975; Craw, 1979 for a comprehensive structural analysis, although the above interpretation is not exactly after these authors). In distinction from the Alpine-type ophiolites formed in the Red Sea type leaky shear zones and often emplaced with all their tiers intact, most of the Pacific ophiolites are dismembered, with their lava, sheeted dike and gabbroid tiers separated from their ultramafic members (e.g., on the Lesser Kurils: Krassilov et al., 1988). In the Kamchatka – Lesser Kuril segment, there were at least two distinct napping events (coeval with ignimbrite events in the marginal volcanic belt) in the mid-Cretaceous and over the Campanian/Maastrichtian boundary (Shapiro, 1976; Krassilov et al., 1988). Both, and particularly the latter, are recognisable over the Moluccas, the Papua – Solomon Islands – New Caledonia segment and on the Northern Island of New Zealand (Hall, 1987; Cullen & Piggott, 1989; Schofield, 1983).


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Fig. 63. Marginal and oceanic structures of the western Pacific. Island arcs: (IB) Izu-Bonin, (KP) KyushuPalau, (LS) Luzon-Sarawak, (OB) Olyutorsk-Bowers, (RT) Ryukyu-Taiwan, (PC) Papua-New Caledonia, (VK) Valdez-Koryakia, (Y) Yap Island. Oceanic rises, plateaux and basins: (Bo) Borodino, (Em) Emperor, (Er) Erimo, (He) Hess, (Li) Line, (Ma) Magallanes, (Mh) Manihiki, (Mn) Marcus Necker, (Na) Nauru, (Ob) Obrutchev, (Sh) Shatsky, (Zn) Zenkevich, (1) Island arcs active in the Cretaceous, (2) Oceanic structures, (3) Granitoids and their derived volcanites of oceanic rises, (4) Faults (CT – Cross Trend of the Line Islands).

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Fig. 64. Crustal structures of New Zealand and adjacent Pacific rises (plateaux) interpreted as crustal blocks borderd by imbricate foreland/foredeep/islandarc belts over the major fault zones. Blocks: (Ca) Campbell, (Ch) Chatham, (Lh) Lord Howe, (No) Norfolk, (Tk) Three Kings. Ophiolitic thrust belts: (AP) Alpine, (HR) Hikurangi, (KM) Kermadek, (TA) Taupo, (VM) Vening Meinesz. (1) Ophiolites, (2) Flysch trough, (3) Volcanic arc, (4) Granitoid belt (Foreland), (5) Faults, (6) Contours of oceanic structures.

Turning to the eastern margin, we find a similar, although differently interpreted, collage of “terrains”, some equivalent to cratonic blocks of the western margin, the other appearing as giant overthrusts. A once popular idea of their transequatorial drift from various locations in the Southern Hemisphere is presently under a revision (Colpron & Price, 1995; Keppie et al., 2001; Morozov et al., 2001; Stamatakos et al., 2001). New palaeolatitude estimates indicate a Cretaceous location in the Northern Hemisphere midlatitudes (about 53°N for Wrangellia: Stamatakos et al., 2001). The available factual evidence for long-distance transport is far from convincing (see Colpron & Price, 1995 for a critical review). The discrepant magnetic signatures might have been due to tectonic rotation rather than long-distance rafting. However, a counterclockwise rotation might have actually displaced continental fragments north of their original location in the circum-Pacific shear zone. The western and eastern arms of circum-Pacific belt are linked by the tectonic structures of Chukotka, Koryak Highlands, Alaska and the Bering Shelf (Fig. 65). The Triassic – Jurassic geosyncline deposits of northern Alaska are traceable over the – dextrally displaced – Chukotsk Peninsula from Laurentiy Bay to Kolyuchinskaya Guba (Inlet), while the Late Jurassic to Neocomian Koyukuk andesite belt (Globerman et al., 1983) might extend to the coeval Kuryan belt of southern Anyuy (Migachev et al., 1984). A volcanic belt of the same age has developed along the southern margin of Talkitna block. It is overthrust by the Maastrichtian to Eocene Valdis/Orca ophiolitic nappes (Tysdal & Case, 1983) and intruded by the Aleutian-Alaskan batholites. The thrust belts appear continuous form the Kenai – Koryak arc over the Bering Shelf to the Olyutorsk Range of northern Kamchatka, where the coeval nappes constitute the Campanian/Maastrichtian to mid-Eocene sequence of the Olyutorsk, Ukelayat and Koryak complexes (inter-


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Fig. 65. Imbricate structural zones of Alaska (After Burk, 1965) extended to the Asiatic margin: (De) Denali Fault, (K) Kenai zone of ophiolitic nappes, supposedly continued across Bering Shelf to Olyutorskiy zone of Koryak Highlands, (Ko) Koyukuk andesites superimposed upon the Palaeozoic – Early Mesozoic belt, traceable to the southern Anyuy zone of northern Siberia, (Pz J) Palaeozoic to Jurassic nappe belts extending over the Bering Strait to the Laurentiy Bay and across the Chukotski Peninsula and dots- the Alaskan counterparts of the Asiatic terrestrial volcanic belt and foredeep (Fig. 62), respectively.

preted as accreted terrains after the North American example: Harbert et al., 2001). Oblique orientation of the fossil Kenai–Olyutorsk belt against the nearly latitudinal Aleutian Trench may reflect a clockwise rotation of the Beringian block since the Early Palaeogene. Offshore, the ultramafic and metamorphic rocks are thrust over the western slope of the Bowers Ridge (Ludwig et al., 1971), an arcuate structure dextrally displaced in respect to the similar structure of the Shirshov Ridge. The trans-Beringian structures are controlled by a series of arcuate dextral faults, of which the Tintina Fault extends to Yukon Territory as a boundary between the cratonic margin and the supposedly allochtonous southern Alaska and Yukon-Tanana terrains. The Yukonian cataclastic complex of the fault zone consists of thrust metamorphic rocks and ophiolite melanges (Erdmer & Helmstaedt, 1983). The collage of allochtonous complexes comprises several generations of nappes involving the Precambrian to Mesozoic metasedimentary and metavolcanic rocks. In the island zone of British Columbia they are assigned to a number of terrains, the larger of which are Wrangellia, Alexander, Stickine, Quensellia, Ferni and Sonomia (Monger & Irwing, 1980; Ben Avraham & Nur. 1983; Archibald et al., 1983). Their superimposed mid-Jurassic to Neocomian volcanic belts, intruded by the apparently comagmatic granitoids, extend over the fault zone of the Hecate Strait – Queen Charlotte Strait and the pull-apart Gravina Basin (Yorath & Chase, 1981; Gehrels, 2001). In the coastal ranges of southern Oregon and California, the thrust belt assemblage is rather typically represented by the Franciscan complex of metamorphic melanges (Ernst, 1970; Maxwell, 1974; Kienast & Rangin, 1982; Aalto, 1981). Their age assignments vary from Late Jurassic to Late Cretaceous, with three major napping phases in the

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Neocomian, mid-Cretaceous and Campanian. They are overthrust by the contemporaneous Great Valley turbidites deposited in the foredeep of granitoid Sierra Nevada arc (Moxon et al., 1987) that extends south to Baja California as a volcanic Alisitos Arc. Ophiolites of the Franciscan/Great Valley contact zone are thrust from beneath the arc over the metamorphic complex. They are also exposed over the Klamath Mountains to the north. The melange assemblages of various metamorphic grades show evidence of multistage napping and back-thrusting (Aaalto, 1981). Further south, the Cretaceous shear belts seem continuous over the mosaic of Central American blocks including Yukatan, Honduras, Nicaraguan Rise of the Caribbean, as well as Marakaibo of northern South America. The collage is fairly similar to those of the northern Pacific, with the block boundaries marked by ophiolite assemblages. These are exposed over the circum-Caribbean fold belt and the transcurrent sea-floor ridges (Fig. 66). On the western Caribbean border, the system of ophiolitic thrust belts comprises (1) the Polochic – Motagua belt between the cratonic blocks of Guatemala and Honduras (Maya and Chortis) and perhaps extending over the Cayman Ridge across the Caribbean, (2) the Nicoya Arc of southern Nicaragua, Costa Rica and Panama (with large ultramafic bodies of St. Helen Peninsula), (3) the Santa Marta belt, and (4) the metamor-

Fig. 66. Circum-Caribbean structures interpreted as crustal blocks surrounded by imbricate metamorphic/volcanic belts. Crustal blocks and fragments: (A) Aves, (B) Beata, (C) Campeche, (MC) Maracaibo, (N) Nicaraguan, (V) Venezuelan. Metamorphic zones: (CC) Cordillera Central (Hispaniola), (Ca) Cayman Ridge/Trough, (De) Desirade, (Ma) Margarita, (MP) Motagua-Polochic, (Ni) Nicoya, (Pn) Pinar, (Sm) St. Martha. Oceanic structures: (CL) Clipperton fracture zone, (CO) Cocos Ridge, (TP) Tehuantepec.Ridge. (1) Ophiolites, (2) Imbricate metamorphic-volcanic belts, (3) Extensional zones possibly continued as oceanic fault zones.


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phic/volcanic La Riconada/Caracas/Tiara complex related to the dextral Oca-El Pinar fault zone of Venezuellan margin (Shagam, 1960; Wilson, 1974; Dengo, 1975; SchmidtEffing et al., 1980). Rather typical of them is the Nicoya complex of Osa Peninsula (Borrangé & Thorpe, 1988) composed of the sheared Jurassic to Late Cretaceous picritic/andesitic lavas with pillow structures, low-grade metamorphic graywackes, silicites and siliceous limestones, intruded by granitoids. A downfaulted extension of the complex is traced under the Miocene calcarenite breccia in the Middle American Trench (Vannucchi et al., 2001). It shows close affinities to the northern Andean structures and it might have been actually linked via the Romeral zone of southwestrn Columbia (Spadea et al., 1987) to the Cayo-Masuchi complex of Ecuador and northern Peru (Lebas et al., 1987). A structural similarity between the Nicoya arc and the ophiolite complexes of Columbian Andes indicate continuity of the Cretaceous circum-Pacific thrust belt over central America, although land connections (also indicated by faunistic links: Borrangé & Thorpe, 1988) might have been intermittent, interrupted by periodically extensional transcurrent shear zones. Thus, the Motagua – Polochic shear zone of southern Guatemala appears related, with a dextral offset, to the Tehuantepec Ridge marking a transcurrent shear zone of eastern Pacific, presently substituted by the Clipperton fracture zone (Fig. 66). It’s trans-Caribbean continuation over the Cayman ridge/trough system shows a recent extensional component. The analogous transcurrent structures to the south include a fossil shear zone of the Cocos Ridge and its possible trans-Caribbean continuation over the geophysically similar Beata Ridge, separated by a recent extension zone from the Nicaraguan Rise of intermediate crust. The southern segment of circum-Pacific belt over the Drake Passage is fairly similar to the Central American and Caribbean structures. It is represented by the thrust complexes of Patagonian Andes bordering the Magallanes Basin and linked to those of Antarctic Peninsula via the structurally equivalent metamorphic/ophiolitic/volcanic belts of the Tobifera – Navarino (the Yahgan Complex) – South Shetland Islands – Triniti of Alexander Island. Their associated granitoids constitute the Patagonian – Livingston Island batholite (Adie, 1964; Katz & Watters, 1965; Dalziel & Elliot, 1973; Wit et al., 1977; Tanner et al., 1982; Storey & Meneilly, 1983). Like in Central America, the continuity of the southern link is disrupted by transcurrent shear, including the sinistral Magallanes fault zone of Patagonia and its probable continuation to South Georgia (Katz & Watters, 1965). It follows from the above overview that: (1) The transional zones between continental and oceanic domains are the buffer zones of a rotational shear consisting of continental and intermediate blocks surrounded, and marginally overlapped, by annular shear belts. These structures were sequentially accreted to the mainland cratonic areas and consolidated by granitization of their suturing shear belts.

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(2) A similarity of structural style over the marginal structures attests to the integrity of circum-Pacific belt as an annular shear zone since the Cretaceous at least. Fragments of older structures incorporated in the Cretaceous thrust complexes indicate the same style tectonics recurring over a sequence of pre-Cretaceous cycles since the mid-Devonian at least. (3) A synchrony of tectonic phases – the Late Neocomian, mid-Cretaceous, Campanian/Maastrichtian and Maastrichtian/Palaeocene pulses of large-scale thrusting, high rate igneous activity and blueschist – eclogite metamorphism all over the circum-Pacific girdle – is evidence of developmental correlation. An eastward migration of volcanic activity is recorded over both the western and eastern margins (Krassilov, 1972d; Moxon et al., 2001). (4) The Cretaceous tectonic/volcanic structures were continuous over the coastal ranges from Chukotka to Sikhote Alin and the Anhui – Zhejiang zone of eastern China, over island arcs of western Pacific as far as New Zealand, over the Olyutorsk – Kenai arc across the Bering Shelf to the Yukonian and Franciscan complexes of North Americ, over the Nicoya Arc of Central America to the Romeral zone of Columbian Andes, and over the Yahgan – Triniti Arc from Patagonia to Antarctica. (5) The tectonic structures of the western and eastern Pacific were less asymmetrical than now, with volcanic arcs developing along the American coast as well. A widespread backthrusting suggests that Benioff zones might have been dipping off the continent – an exceptional orientation for the present-day Pacific margin. The circum-Pacific structures have been discussed at some length because their interpretation may help understanding of other continental margins and intracontinental fold belts, in particular of the Atlantic “passive margin” and the Tethys.

V.6.4. Cretaceous peri-Atlantic and peri-Arctic belts “Passive margin” is a misleading definition because, at certain stages of its geological development, the Atlantic margin was by no means quiescent. Over the Mesozoic, it has been affected by at least three successive phases of igneous activity (Forland et al., 1971; Macintyre, 1977; Larsen et al., 1983; Hanisch, 1984; review in Krassilov, 1985): (1) The terminal Triassic to Early Jurassic, about 200 – 170 Ma (the Avalon dyke complex of Newfoundland, an older magmatic phase of White Mountains, the great dyke complexes of western Greenland, Liberia and South Africa), (2) The Late Jurassic – Early Cretaceous, about 140 – 110 Ma (a younger phase of White Mountain magmatism, the offshore basalts of Newfoundland Basin, Spur


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Ridge, J-anomaly Ridge, Blake–Bahama escarpment, Rockall Trough, eastern Brazil and shelf, Kaoko volcanic province in Namibia, Benue Trough in Nigeria), (3) The Late Cretaceous, about 95 – 70 Ma (the Newfoundland Basin, Cabu of Brazil, Ombe in Angola, Cameroon Line, etc.). To those who believe that these magmatic structures are related to the alleged opening of the Atlantic it should be reminded that in eastern Brazil (the Paraná – Parnaiba basins) they rest on the Palaeozoic marine limestones and evaporites that are recorded in all the marginal basins from the Alagoas in the north to Rio Grande in the south (Bigarella, 1973). The Late Triassic – Early Jurassic lagoonal salt-bearing deposits extend from the Grand Banks to the Gulf of Mexico (Austin et al., 1989). In other words, the oceanic margin was there in the Palaeozoic and Early Mesozoic. The Triassic magmatism predates, while the Late Cretaceous phase postdates the alleged opening of the central Atlantic propagating poleward. The widespread Late Jurassic – Early Cretaceous Wealden facies suggest downfaulting of marginal basins that was accompanied by volcanic activity over the whole perimeter of the present-day Atlantic. Large basaltic provinces appeared over the marginal basins of eastern Brazil and Namibia. Their coeval basalts occur in Newfoundland Basin, Rockall and elsewhere over the Atlantic perimeter (Fig. 67). Alkalic intrusions are commonly ascribed to continental rifting. However the Brazilian basaltic province is intruded by nepheline syenites of at least two age groups, 133 – 123 Ma and 80 – 60 Ma (Amaral et al., 1967). The latter group postdates the alleged opening of Central Atlantic by about 100 m.y., while the carbonatites of Maio (Cape Verde Islands) are still younger. The coeval alkalic magmatism is widespread over the Appalachians, Canadian magmatic province, southeastern Greenland, Angola and Namibia. During the Late Triassic – Early Jurassic magmatic phases, the oceanic crust might have been decoupled in the rotational stress field from the surrounding continental crust by a marginal leaky shear zone marked by the along-the-shore magnetic anomaly. With consolidation of the marginal zone over the subsequent magmatic phases, the rotational shear pressure was released through a median crack of oceanic crust developing as the Mid-Atlantic Rise. The peripheral magmatism was confined to transcurrent fault zones abutting the continental margin, e.g., the Southeastern Newfoundland Ridge, New England Seamounts, Walvis Ridge and Rio Grande Rise of diagonal fault system, controlling sedimentation in the Cretaceous as well as in the present-day Atlantic Ocean. Like in the circum-Pacific girdle, the periphery of the Atlantic Ocean is a mosaic of larger and smaller blocks. In the northern Atlantic, between Greenland and British Isles, the mosaic includes the Jan Mayen, Iceland, Faeroe, Hatton, Rockall and Porcupine continental fragments separated by fault zones, volcanic in the late Early Cretaceous and Palaeocene (Hanisch, 1984). The Reykjanes zone belongs to this fault system, with a possibility of a still not recorded pre-Miocene volcanic phase. The western periphery

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Fig. 67. Peri-Atlantic and sea-floor structures: (Ag) Agulhas, (Am) Amazon Fault Zone, (Ba) Bahamas Escarpment, (Bl) Blake, (Br) Bermuda, (Bs) Biscay Seamounts, (Ca) Caoko basaltic province, (Cb) Cabo basalts, (Cm) Cameroon Line, (Cn) Canary Islands, (CV) Cape Verde Islands, (FC) Flemish Cap, (FN) Fernando de Noronha, (Gu) Guinea Fault Zone, (J) J-anomaly Ridge, (M) Madeira, (NE) New England Seamounts, (NR) Southeastern Newfoundland Ridge, (OK) Orphan Knoll, (Om) Ombe, (Pa) Parnaiba basaltic province and offshore extension, (Pn) Paraña basaltic province, (R) Rockall Trough, (RG) Rio Grande Rise, (Ro) Romansh Trough, (SL) Sierra Leone Rise, (T) Torre Rise, (W) Walvis Ridge. (1) Axial zone of the mid-Atlantic Rise, (2) Eastern zone of the Tore–Madeira–Fuerteventura (Canary Islands)–Maio (Cape Verde Islands)–Sierra Leone thrust structures, (3) Marginal volcanic belts and ridges, (4) Buried ridge of the western margin, (5) Seamounts.


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comprises the Orphan Knoll, Bermuda, Blake, Bahamas, Demerera, Fernando de Noronha, Rio Grande, Falkland and Scotia blocks of continental and intermediate crust. Their eastern counterparts are the Gorringe, Madeira, Torre, Canary, Cape Verde, Sierra Leone, Guinea, Walvis and Agulhas blocks. They are fringed by submerged thrust structures that, although partly buried and fragmented by transcurrent faulting, still resemble those of the Pacific margin. Thus sheared ultramafic bodies with serpentinite debris occur on the Gorringe Bank off Portugal (La Gabrielle & Auzende, 1981; Verzhbitsky et al., 1989) extending over the Madeira – Tore ridge to the south. They emerge on the western Canary Islands Lanzarote and Fuerteventura in association with the sheared metagraywackes, black shales, silicites, limestones and basalts of at least two generations, the older Late Jurassic – Neocomian and the younger Late Cretaceous – Early Tertiary (Rothe & Schmincke, 1968; Dillon & Sougy, 1974; Robertson & Bernoulli, 1982). A major post-Cretaceous napping event is dated as the Eocene. This volcanic arc – foredeep complex is intruded by a swarm of sheared parallel dykes characteristic of ophiolite assemblages in the Pacific and elsewhere. Notably, the lithosphere back of Canary Islands is of a transitional type (Hayes & Rabinowitz, 1975), like in the backarc basins of the western Pacific. Further south, the shear zone is traceable to Maio, western Cape Verde Islands, where a similar sequence of basalts, cherts, limestones and graywackes is sliced by a series of imbricate thrusts (Robertson & Bernoulli, 1982; Rasnitzyn & Trofimov, 1989). Volcanics of the Sierra Leone Rise associate with the layered mafic-ultramafic rocks of Freetown Massif (Krause, 1963), once considered meteoritic, structurally similar to the Alaskan-type intrusions typical of the Pacific island arc complexes. The Fuerteventura – Maio shear structures of eastern Atlantic find their equivalents in the thrust complexes of Greater Antilles, the imbricate belts of ophiolite/volcanic/ metamorphic nappes involving the Late Jurassic to Eocene deposits and intruded by the Late Cretaceous to Eocene granodiorites (Pyle et al., 1973; Mattson, 1973; Nagle, 1974; Pardo, 1975; Kesler et al., 1977; De Albear et al., 1984). These structures are typically represented on Cuba by the Las Villas ophiolites emplaced upon the Placetas carbonates of the Bahamas margin and successively overthrust by the Sasa-Cabaiguan island-arc andesites (comagmatic to the Camaguey batholite), the forearc flysch and the Juventud – Sierra Maestra metamorphic complex. These zones are traceable over the Blue Mountain zone of Jamaica and the Los Ranchos/Duarte/Cordillera Central zones of Hispaniola to the andesite/chert complex overlying the giant batholite of the Virgin Islands. Their bordering Puerto Rico Trench marks a sinistral shear zone with ophiolites (Tucholke & Ewing, 1974) probably transported from beneath the island arc. The symmetrical thrust belts of Venezuellan margin, including the ophiolites of Santa Rosa fault zone, the volcanics of Tiara arc and the metamorphic complex of Margarita Island (Shagam, 1960; Weyl, 1966) are linked to their northern counterparts by the metamorphic inner zone of the Lesser Antilles locally exposed on Desirade, Leeward Islands (Fink, 1970). Outer arcs of the Lesser Antilles, facing the Central Atlantic, are formed

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of the Palaeocene to Eocene graywacke–olistostrome complex and the overlying tuffs of Barbados (S.N. Davies, 1971). The (long ignored) structural similarity of the Antilles and the Canaries – Cape Verde belt, as well as their asymmetry, is of the same nature as those of the western and eastern Pacific structures (above). The Scotia Block, like the Caribbean, is surrounded by island arcs bridging the thrust complexes of Teirra del Fuego (Navarino Island) to those of Antarctic Peninsula (V.6.3) and projecting to the South Atlantic as the eastern counterpart of the African Agulhas arc. They are dissected by a series of transcurrent faults, including the better known Magallanes, Bransfield and South Sandwich fault zones (Fig. 68). A combination of the left-lateral (the northern fault zones) and right-lateral (the southern periphery) transcurrent faulting might generate the pull-apart basins of the modern Scotia Sea. However there is little evidence of its origin by a back-arc spreading as suggested by Alvarez (1982) and Barker et al. (2001). One can argue that island arcs, if at all present in the Atlantic, are of a much lesser extent than in the Pacific. However, in addition to the emergent arcs, there are submerged marginal structures that might have been active arcs in the past. Incidentally, the Ivory Coast – Ghana marginal ridge marks a shear zone of thrust Cretaceous sediments (Huguen et al., 2001; the alleged relation of marginal shear to continental separation is not supported by geological evidence). On seismic evidence, the marginal evaporitic basins of Sierra Leone, Gabon and Angola are fringed by buried arcuate ridges, e.g., the Longa – Cabo Ledo Ridge fringing the Cuanza Basin (Sheridan et al., 1969; Franks & Nairn, 1973; Cooper, 1978). They form a nearly continuous series traceable to the metamorphic Agulhas Arc over the southern extremity of Africa. Their western counterpart is a series of submerged morphologically and/or magnetically distinct positive structures extending over the shelf margin from Newfoundland to Bahamas (Uchupi & Milliman, 1971). Some of them, such as the J-anomaly Ridge on the northern flank, are topped by the partly subaerial mid-Cretaceous basalts (Tucholke & Ludwig, 1982). This ridge system has been once interpreted as a buried island-arc belt (Watkins & Geddes, 1965), the most plausible of several possible explanations. In the South Atlantic, a similar system of buried arcs borders the shelf basins of Argentina between Colorado to San Jorge (Zambrano & Urien, 1970). The Arctic Ocean consists of the larger Amerasian and Eurasian basins subdivided by a series of parallel ridge/trough structures, with a spreading zone over the Gakkel Ridge belonging to the same fault system as the fossil Alpha-Mendeleev and Lomonosov ridges. The periphery of the Arctic Ocean bears evidence of tectonic activity in the mid-Palaeozoic and the Cretaceous – Palaeogene, in reference to which we can speak of the mobile peri-Arctic belt involving both the Caledonian and Laramid structures. The Caledonian thrust belts of Norway, Novaya Zemlya, Svalbard, Greenland, submerged Lomonosov Rise (with the post-orogenic Devonian molassoids and Cretaceous continental deposits: Herron et al., 1974; Grantz et al., 2001).), Brooks Range and the Parry


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Fig. 68. Scotia structures: (Ma) Magallanes Basin, (Na) Navarino, (SG) South Georgia. Shear zones: (B) Bransfield, (M) Magallanes, (S) Shackleton, (SS) South Sandwich. (1) Cratonic block, (2) Volcanic arc (Tobifera), (3) Foredeep, (4) Mesozoic thrust belts (Yahgan – Cumberland), (5) Palaeozoic thrust belt.

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Archipelago are bordering the larger crustal blocks of the Norwegian, Barents and the Beaufort – Chukchi seas (Fig. 69). The southern margin of Chukchi block is formed by the Anyuy thrust belt probably extending to the eastern De Long Islands in the Laptev Sea (Sokolov et al., 1997) to join the Lomonosov Rise zone as an annular belt of Caledonian tectonism. The Neocomian basalts of eastern Svalbard, Franz Josef Land and Barrow, Alaska (Parker, 1967; Tarakhovsky & Fishman, 1982) attest to a spreading phase synchronized

Fig. 69. Arctic crustal structures interpreted as superposition Palaeozoic and Cretaceous–Eocene patterns of crustal blocks and their bordering thrust belts: (Ba) Barents block, (Bf) Beaufort block, (Ch) Chukchi block, (Ka) Kara block, (Nw) Norwegian block. Thrust belts: (AM) Alpha (Mendeleyev) Ridge, (AB) Anyuy Belt, (BR) Brooks Range, (BINZ) British Isles – Novaya Zemlya belt, (DL) De Longa Rise, (EB) Eureka Sound belt, (GR) Gakkel Ridge, (LB) Lyakhovkiy belt, (LM) Lomonosov Ridge, (WS) West Svalbard belt. Basins: (Md) Mackenzie Delta, (SV) Sverdrup. (1) Caledonian structures, those of Lomonosov Rise and eastern De Longa (thick dashes) supposedly dividing the Palaeozoic blocks, (2) Cretaceous–Tertiary thrust belts, (3) Foredeep basins, (4) Contours of continental shelf and oceanic rises.


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with that of the Atlantic (above), but antithetic to that of the Tethys (V.6.5). The Late Cretaceous – Palaeogene thrust belts are supperimposed, with a dextral offset, upon the Palaeozoic frame, partly re-shuffling the Caledonian nappes (Elswick, 1984). They are exposed on Lyakhovski Islands of the New Siberian Archipelago, western De Long Rise, western Svalbard, the Canadian Archipelago (Eureka Sound) and over the Brooks Range of Alaska. Like in the Atlantic, a major thrusting event occurred in the Late Palaeocene to Early Eocene (Saalman & Tiedig, 2001), involving the Cretaceous volcanic/metamorphic deposits. A large thrust sheet of Cretaceous andesites and silicites is recorded on the Alpha (Mendeleev) Ridge (Herron et al., 1974), a submerged arcuate structure over the western border of the Beaufort – Chukchi block, linking the ophiolitic melanges of the northern Ellesmere Island, Arctic Canada, with those of the Lyakhovski Islands, New Siberian Archipelago (Savostin & Drachev, 1988). The peri-Arctic mobile belt is, thus, strikingly similar to the Caribbean (V.6.3) and, as we shall see, to the Tethys (V.6.5), with the Lomonosov Rise of intermediate crust corresponding to the Nicaraguan Rise and with fossil volcanic arcs represented by the Alpha – Mendeleev Ridge and the Aves Rise respectively. In both the recent extensional zones are controlled by transcurrent fault systems, the Nansen ridge/trough system of the Arctic Ocean having its equivalent in the Cayman ridge/trough system of the Caribbean. Both extended over continental borders, the former to the Moma graben of Yakutia (Volk et al., 1984), the latter – to the Motagua–Polochic fault zone of Guatemala. The above comparisons show that both the Atlantic and Arctic margins might have been active over the Mesozoic. Their mosaic block structures, their thrust belts with ophiolitic melanges are similar to those of the Pacific margin. However, the Atlantic – Arctic phases of marginal expansion in the Neocomian, thrusting over continental margins in the mid-Cretaceous and back-thrusting, supposedly with re-orientation of Benioff zones, in the terminal Cretaceous and Eocene, are antithetic to the roughly synchronous tectonomagmatic events in the circum-Pacific belt.

V.6.5. Cretaceous Tethys In the northern mid-latitudes, the peri-Atlantic and circum-Pacific fold belts are linked by a seismically active zone extending over the Mediterranean and southern Asia. This zone corresponds to the Cretaceous – Tertiary and older belts of folded marine deposits that prompted the idea of Tethys, an extinct seaway between the Laurasian and Gondwanic landmasses. In the classical geotectonics, the Tethys fold belt was ascribed to contraction closing the seaway. Later the emphasis has shifted to the ophiolite assemblages as remnants of oceanic crust obducted on continental blocks (microcontinents) that were rafted across the Tethys Ocean (or the successive Tethys I – Tethys II oceans) to their random collision with each other and with mainland Asia. A brief overview of the Tethys would show this as not having been the case.

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The western Mediterranean is a mosaic of blocks, the better defined of which are the Iberian, Alboran, Corso-Sardinian, Maltian (Pelagian), Adriatic, Serbo-Macedonian and the Rhodopic (Fig. 70). They are sutured by the Gibraltar – Baleares, Penninic, Umbrian, Ligurian, Calabrian, Ionic, Dinarian, Vardar and the less prominent ophiolitic belts (Aubouin, 1963; Hsü, 1971, 1976, 1977; Bosellini & Hsü, 1973; Dewey et al., 1973; Vandenberg, 1979). The ophiolites and associated metamorphics are exhumed from beneath the volcanosedimentary cover and thrust over the Lower Cretaceous foredeep turbidites that are in turn emplaced, with melanges, upon the block margins, as described for the Umbrian, Dinaran and Ionic sutures (Celet, 1977; Abbate et al., 1980; Karamata et al., 1980; Trotet et al., 2000). These structures appear continuous over northern Africa and southern Europe (Fig. 71). The Bettic –Subbetic belts are linked via Gibraltar arc to their equivalent Maroccan Rif and Tellian structures, while the Calabrian arc is allied to the Atlassian ophiolite belts. Whatever the relative motions between European blocks and Africa since the Permian, they have not disrupted the connections. The ophiolites are tectonized before emplacement, which indicates a shear zone environment, such as the Red Sea type leaky strike-slip fault zones injected with peridotite bodies (Styles & Gerdes, 1983). Their associated silicitic turbidites contain terrestrial plant remains (Abbate et al., 1980), which rule out a mid-ocean ridge environment, but are consistent with a ridge – trough structure of a shear fault zone. Instead of representing a great number of local spreading – subduction events, the thrust belts are synchronized in their development all over the western Mediterranean.

Fig. 70. Alpine – Mediterranean crustal blocks of the western Tethys belt: (Ad) Adriatic, (I) Iberian, (M) Moesian, (Ma) Maltian, (P) Pelagian, (Pa) Pannonian, (R) Rhodopian, (T) Transylvanian, (WM) Western Mediterranean, (G) Gibraltar arc.


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Fig. 71. Alpine-Mediterranean ophiolitic nappe zones of Southern Alps, Carpathian, Penninic, Dinaric and Hellenid systems: (Ab) Abruzzi, (Ap) Apulian, (C) Calabrian, (Ch) Chalkidiki, (Io) Ionian, (JA) Julian Alps, (L) Ligurian, (Ly) Lycian, (Ot) Othris, (Mm) Maramaros, (Mu) Mures (Apuseni), (P) Platta, (Pn) Pindus, (R) Rechnitz, (SL) Sesia-Lanzo, (T) Tauern, (To) Toscanian, (U) Umbrian, (V )Vourinos. Faults: (BS) Banat-Srednegorye, (Ga) Gailtal Line, (Gb) Gibraltar, (Gu) – Guadalquivir, (E) Engadine, (In) Insubric Line, (Lu) Lubeniki, (NP) North Pyrenean, (Va) Vardar. (1) Ophiolites, (2) Faults, (3) Nappe belts.

Radiometric dates for the ophiolites are systematically clustered at about 160 Ma (midJurassic). Their associated melanges, recording a thrusting phase, are mostly assigned to the Late Jurassic – Early Cretaceous, which is the approximate age of the Penninic schistes lustres and similar metamorphic melanges elsewhere (Abbate et al., 1980; Dietrich, 1980; Hawkesworth et al., 1972; Karamata et al., 1980; Wiebel & Frisch, 1989). The retrogressive metamorphism is related to back thrusting in the Late Cretaceous and Eocene. The mobile belt extends to the east, over the Vedinian, Zangesur and Sevano-Akerian zones of the Lesser Caucasus (Sokolov, 1977; Aslanyan & Satyan, 1977; Fig. 72). The chronology of tectonic events is markedly different over the peri-Arabic and peri-Indian belts of eastern Tethys. The Taurian block is bordered by the Cyprus-Antakya shear zone containing the well-known Cretaceous ophiolites, Baer-Bassit in Siria and Troodos on Cyprus (Miyashiro, 1973; Juteau, 1980; Pantazias, 1980; Ozkaya, 1982). The radiometric dates for their magmatic phases fall within the mid-Cretaceous – early Late Cretaceous interval, 105-75 Ma, mostly the Late Cenomanian to Turonian. The ophiolites are emplaced upon the Campanian to lower Maastrichtian autochtonous deposits. The thrust zone of Zagros and the roughly contemporaneous nappes of Makran bordering a larger Iranian block Dasht-E-Lût (Fig. 73) are obducted over the Coniacian – Santonian deposits and are overlain by the neoautochtonous Maastrichtian reefal limestones (Berberian & King, 1981). The Samail Ophiolite of Oman might have been con-

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Fig. 72. Ophiolote zones in the Lesser Caucasus segment of the Tethys belt: (SA) Sevan-Akerian, (V) Vedinian, (Z) Zangezurian; (Kf) Kafanian Block.

Fig. 73. Crustal blocks and ophiolite zones of the Middle East segment of the Tethys belt: (At) Antalya, (Bb) Baër Bassit, (Fh) Farâh, (H) Hilmend, (Ht) Hatay, (L) Lût, (Ly) Lycia, (Om) Oman, (Sb) Sabzevar, (T) Taurus, (Zs) Zagros. (1) Ophiolites, (2) Sutures of the crustal blocks.


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tinuous with Zagros, but it was displaced by transcurrent faults of Hormuz and thrust to the south, over the Arabian margin (Smewing, 1980; Tippit et al., 1981; Coleman, 1981; Woodcock & Robertson, 1982). Its basaltic to rhyolitic two-stage lava complex is covered with umbra (Lippard & Rex, 1982) indicating an exposure to hydrothermal weathering. The Hilmend (Central Afghan) block is bordered by the Farâh and Chaman fault zones (Blaise et al., 1978), the latter also controlling the western thrust border of the Indian block. The thrust complexes of Sulaiman Range, Kokhistan – Ladakh arc, TethysHimalayas and Naga Hills (Fig. 74) are similarly structured, with ophiolite assemblages,

Fig. 74. Crustal blocks, faults and ophiolite zones in the Tibet-Himalayan sector of the eastern Tethys belt: (Ar) Arakan, (At) Astin Tag, (Ch) Chaman, (F) Fergana Basin, (Gs) Gangdise, (HK) Hindu Kush, (IZ) IndusZangpo Fault Zone, (K) Kohistan, (Ko) Kokuxili Fault Zone, (Kr) Karakoram Fault Zone, (Kt) Karatau Range, (La) Ladakh, (Lh) Lhasa Block, (MT) Main Thrust Zone, (N) Nujiang (Nu Chiang) Fault Zone, (Ng) Naga Hill Zone, (Sh) Shan Block, (Su) Sulaiman Range, (Th) Thanglha Block, (Tr) Tarim Basin, (Ts) Tsajdam Basin, (Xi) Xigaze. (1) Ophiolites, (2) Granites, (3) Faults, (4) Thrusts, (5) Tarim – Fergana trough.

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melanges, flysch and island arc-type volcanics topped by shallow marine deposits. The radiometric dates obtained for Xigase (Nicolas et al., 1981; Tapponier et al., 1981; Allegre et al., 1984) and stratigraphically equivalent ophiolites indicate their mid-Cretaceous age (Honegger et al., 1982; Thakur & Misra, 1984; Schärer et al., 1984). They form a thick series of nappes emplaced upon the Tethys-Himalayas in the Late Cretaceous and are overthrust by the Trans-Himalayan granitoid belt in the Eocene. The belt extends over the Naga Hills to the Arakan Yoma Range in Burma and over the Andaman – Nikobar ridge to joining the circum-Pacific girdle. The crustal mosaic north of the Himalayas include the Lhasa, Quingtang, Songpan–Ganzi, Yangtze and Qaidam blocks sutured by the (from north to south) Yarlung– Zangbo, Banggonghu–Nujiang, Jinsajing, Longmenshan and Kunlun ophiolitic zones (Zhang, 2000). The Zangbo (Tsangbo) ophiolites of Lhasa (Nicolas et al., 1981) seem closely correlative with those of the Ladakh zone of Himalayas (Thakur & Misra, 1984; Pedersen et al., 2001) indicating a synchronous tectonic development while ruling out the allegedly separate diachronous accretions. On the Tibetan side, the Donqiao ophiolite north of Lhasa comprises a Jurassic magmatic assemblage emplaced during the mid-Cretaceous (Girardeau, 1984). Thrust complexes of this age are typical of the Tanglha and Kukushili shear belts further north. Granitization following the emplacement of ophiolite nappes is likewise synchronous over the peri-Indian belts. As a whole, the Mesozoic Tethys appears to consist of two arms with diachronous tectonic developments. In the northern arm extending from the western Mediterranean over the Lesser Caucasus to Tibet, the mid-Jurassic ophiolites are emplaced in the late Neocomian to mid-Cretaceous. In the Taurian – Iranian – Afgano-Himalayan arm to the south, the ophiolites of mid-Cretaceous age are emplaced in the Late Cretaceous, at about the Campanian – Maastrichtian boundary. The end-Cretaceous to Eocene rethrusting took place over both arms. The timing of tectonic events is roughly synchronized with the circum-Pacific belt in the northern arm and with the peri-Atlantic belt in the southern arm.

V.7. Tectonomagmatic phases Chronological correlation of tectonomagmatic events in the major fold belts indicates global pulses of tectonic activity as postulated by Stille (1924), Umbgrove (1947) and other classical school geologists. Thus, the Late Jurassic/Berriasian, Hauterivian/Barremian, Albian/Cenomanian, Campanian/Maastrichtian and Cretaceous/Tertiary tectonic pulses are felt all over the circum-Pacific, peri-Atlantic and Tethys belts, with resonances over cratonic areas. The Late Jurassic to Neocomian downfaulting of continental margins was accompanied by rifting of the continental crust responsible for the world’s largest basaltic provinces of eastern Brazil and central – eastern Asia (with volcanics up to 600 m thick in the Kerulen Basin, Mongolia: Frikh-Har & Luchitskaya, 1983).


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By the mid-Cretaceous, the volcanic activity has spread over oceanic fault zones forming plateaux and ridges in the Atlantic (the Rockall Plateau, J-anomaly Ridge, Newfoundland Ridge, Walvis Ridge, Rio Grande Rise, etc: Hekinian, 1972; Humphris & Thompson, 1983; Hanisch, 1984), eastern Indian Ocean (the Ninety East Ridge, Broken Ridge, Kerguelen Plateau, etc.: Guilfy, 1973; Hekinian, 1974; Dosso & Vidal, 1979; Mutter & Cande, 1983) and Pacific (the Emperor Ridge, Tuamotu – Line island chain, Neccer Ridge, Shatsky Rise, Hess Rise, Magellan Rise, Manihiki Plateau, etc.: Jackson & Schlanger, 1976; Jenkyns, 1976; Saito & Osima, 1977; Orwig & Kroenke, 1981; Vallier et al., 1981; Thiede et al., 1981; Kogan et al., 1983; Neprochnov et al., 1984). Whatever the position of these sutures relative to alleged plate boundaries, their igneous activity was synchronous with the admittedly “intraplate” volcanism, e.g. of Nauru Basin in the Pacific (Batiza et al., 1980). The mid-Cretaceous oceanic events correlate with ophiolite emplacements in the western Mediterranean, Lesser Caucasus and Tibetan thrust belts, as well as with initiation (in the Albian) of the circum-Pacific terrestrial volcanic belt developing over the allochtonous arc/trench complexes thrust upon continental margins (Krassilov, 1989b). The circum-Pacific thrust belts were granitized by polyphase intrusions, with a distinct Late Cretaceous phase also recorded in the Trans-Himalayan belt (Huang, 1984). The Cretaceous/Tertiary phase is marked by a major intracratonic rifting/trap magmatism event (in particular, by Deccan traps, about 65 Ma: Mahoney et al., 1982; Subbarao, 1999). Most vigorously in the Trans-Himalayan belt, but also in the circum-Pacific, peri-Atlantic and peri-Arctic belts, the nappe piles were re-thrust, with the granitoid basement of volcanic arcs emplaced upon the fore-arc flysch trough and melanges (as in the Sierra Nevada Batholite/Great Valley turbidites/Franciscan melanges of California or the Ladakh Batholite/Kokhistan Arc/Shyok melanges in the Himalayas: V.6.5). Thus thickened, the fold belts were isostatically raised as mountain ranges, with a major orogenic event in the Middle/Late Eocene recorded in the Himalayas, as well as over the Caribbean, the Canary – Cape Verde belt of eastern Atlantic, the Brooks Range in Alaska, and elsewhere (V.6.3 -V.6.5).

V.8. Summary of tectonic factors Spatial distribution of tectonic processes is non-random, governed by the planetary fault system comprising the orthogonal and diagonal subsystems. Regularities of the fault network pertain to the rotation-induced stresses. Distribution of the latter over the globe inflicts antithetic activity of the high-latitude and low-latitude fault systems. Extension over the northern North Atlantic – Arctic Ocean coincides with compression over the Tethys. Although the presently active global lineaments, such as oceanic fracture zones, are appreciably offset against their preceding Cretaceous and older structures, the

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relative stability of the fault network over times rules out a possibility of large-scale polar wander. Rotational forcing is reflected by the series of transcurrent offsets of meridional lineaments, such as mid-ocean ridges, with the amplitudes increasing at critical latitudes. The maximum shear zone is the equatorial. The faulted lithosphere adjusts to rotational acceleration by the lateral, as well as radial, shifts of its constituent blocks mediated by their marginal shear belts. These develop as global fold belts over the density boundaries of continental and oceanic lithosphere. Not only the Pacific and Tethys belts, but also the presently “passive” Atlantic and Arctic margins were active in the Cretaceous, displaying a typical mosaic structure of stable blocks fringed by ophiolitic thrust zones. Tectonomagmatic events are synchronous over the fold belts attesting to their developmental integrrity. The marginal magnetic anomalies are tracks of the leaky shear faults that decouple continents from the surrounding oceanic crust. Their impregnating magmatic wedges are periodically exhumed from under the volcanic – sedimentary cover that is thrust, with high-pressure metamorphism, over the continental margin. The exhumed ultramafic – mafic material is involved in ophiolitic nappes that impose the high-temperature metamorphism, granitization and comagmatic volcanism. The nappe belts, cut off by marginal shear and pulled apart by transcurrent faulting, develop as island-arc systems, with extensional backarc basins and trenches. Asymmetry of the western/eastern marginal structures, as in the present-day Pacific or the Cretaceous Atlantic (with the Greater Antilles opposing to the Canary – Cape Verde arc) betrays rotational forcing, as does the westward drift of volcanic activity (the Coriolis vector) and the equatorward propagation of volcanic bulges (the Eötvös vector). Mid-ocean rises are S-cracks generated by the marginal couple of forces. Impregnation of magmatic material in the crest zone and flanking faults generates a series of magnetic anomalies, with an age progression reflecting a sequential cooling and subsidence. Chronology of tectonmagmatic events over the fold belts and cratonic areas reveals a global cyclicity of tectonic processes. Thus, the Cretaceous Period is bracketed by the major intracratonal rifting events over the lower boundary, the Late Jurassic/Neocomian, giving rise to the basaltic provinces of eastern Brazil, Namibia, Barents Shelf, Mongolia and elsewhere, and over the upper boundary, the Cretaceous/Tertiary, with the even more prominent trap magmatism about 65 Ma terminating the 180 m.y. Mesozoic megacycle. A mid-Cretaceous Austro-Alpine thrusting phase of the western Tethys, circum-Pacific belt and elsewhere, accompanied by basaltic volcanism of oceanic ridges, divides the subordinate Early/Late Cretaceous cycles about 30 m.y. each. These periodicities are typical of the Phanerozoic time scale, which is based on biotic events causally linked to tectonomagmatic evolution (their relations are discussed in IX.3-4). The first order cycles of about 180 m.y. (two such cycles in the Palaeozoic) roughly correspond to revolution of the Solar system around the center of the Galaxy


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(Galactic Year), the 30 m.y. cycles – to its radial fluctuations about the Galactic plane (Rampino & Strothers, 1984). In general terms, these periodicities are owing to gravitational forcing inflicted by the huge mass concentrations (about 500×106 solar masses: Yusef-Zadeh et al., 1990) at the centre of the Galaxy and about the Galactic plane (White & Paradijs, 1996; Stridnar & Touma, 1996).

V.9. Tectonic inferences from taphonomy and palaeoecology Tectonic developments define taphonomic landscapes (II.5), the heterogeneity of which is evidence of tectonic complexity, as in the case of intracratonic basins, volcanic ranges, their foredeeps and island arcs of the western Pacific margin, each zone recognisable as a distinc taphonomic domain (VII.6; IX.9). Certain types of taphonomic environments, such as the shelf zones of carbonate/silicoclastic/ organoclastic cyclothems, the rift lake paper-shale facies, the coal-bearing or redbed molassoids of orogenic belts, are diagnostic for particular phases of tectonic development. The allochthonous molassoid coals of intermontane depressions are different from autochtonous rheotrophic coals of paralic foredeeps and from ombrotrophic coals of intracratonic basins (II.7.1). An increase of kerogen concentration in marine deposits reflects a massive influx of dispersed organic matter with continental runoff trapped in epeiric seas or, with epeirogenic (isostatic) uplift of continental crust, reaching to oceanic basins inflicting stratification of the water column and black shale deposition. Continental runoff is a major source of oil-gas accumulations. Their associated palynofacies vary in their ratios of amorphous to fibrous kerogens and the optical signatures of thermal alteration (examples in Batten, 1982). The fibrous “land plant” kerogens typically originate from organic debris of deltaic facies, while the amorphous “algal” kerogens are supplied by algal blooms in hypertrophic estuarine waters. Thick deltaic deposits typically occur in downfaulted marginal basins (as over the Neocomian Atlantic margins with Wealden facies) owing their origins to a divergence of isostatic compensation levels for oceanic and continental crust (V.2). Conversely, an isostatic convergence of oceanic and continental crust is marked by ponding of river mouths by the rising sea and a spread of estuarine facies. Rapid switches from deltaic to estuarine facies indicate tectonic instability, in particular over the critical boundaries, such as the Permian/Triassic IX.3). Likewise, a tectontic development of continental margins, such as elevation of coastal ranges or emergence of island arcs, affect taphonomy by deflecting or trapping the terrestrial organic runoff. Such events are reflected in the deposition rates and composition of coal as well as in an input of allochtonous material to the fossil plant assemblages. Certain features of tectonomagmatic evolution can be inferred from their related taphonomic events. Incidentally, an input of terrestrial plant material into volcanoclastic forearc deposits (e.g., of the Lesser Kuril Islands at the Cretaceous – Tertiary boundary: Krassilov et al., 1988) attest to the emergence of volcanic island arcs.

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Sedimentary basins involved in orogenic uplift and subsequent denudation, might leave no traces other than redeposited palynomorphs (e.g., the reworked Carboniferous palynomorphs in the areas devoid of Carboniferous deposists: Guy-Ohlson et al., 1987). Likewise, redeposited palynomorphs provide a timing of orogenic uplift of the source areas (e.g., of the Rocky Mountains exhumation from under the Palaeogene peneplain, with the reworked palynomorphs reaching to the Gulf of Mexico). A chronology of orogenic uplifts can also be deduced from the dynamics of altitudinal vegetation belts (Axelrod, 1965, 1997) and a spread of xeromorphic vegetation in the rain shadows, as over the piedmonts of incipient Rocky Mountains in the Late Eocene (Axelrod & Raven, 1985). Thus, both latitudinal and altitudinal differentiation of fossil plant assemblages provide evidence of tectonic reconstructions. Globally, the vegetational zonation reflects a heat transfer controlled by geography in turn controlled by tectonics. Mountain ranges form boundaries of vegetation types, as in the case of the broadleaved/pine forest types divided by the Urals. Similarly, the Angarian and Cathaysian taphofloras in Primorye and Japan were separated by the Vladivostok – Abo system of volcanic ranges (Krassilov, 1972a). As the relationships of the Laurasian and Gondwanic floras suggest, orographic barriers for floristic exchange over the Tethys belt might have been more efficient than marine barriers. Tectonic developments also indirectly affect phytogeography by their impact on climate. Thus, a distinctness of zonal vegetation boundaries increases with a restriction of the meridional heat transfer by latitudinal mountain ranges. Sea-floor uplifts might have a similar effect by restricting deep-water circulation inflicting a monsoon depression and xeromorphism of equatorial vegetation. Conversely, a spread of equatorial rainforests is owing to summer monsoons that are sustained by the tropical upwellings of psychrospheric oceanic circulation in turn related to a subsidence of land barriers and a free passage of deep water through marine gateways.

V.9.1. Tectonics and vegetation of continental margins A tectonic zonation of active continental margins with the coastal volcanic ranges, their foredeeps, island arcs and forearc basins is reflected in the pattern of vegetation zones and taphonomic environments. Changes of the latter are evidence of tectonic restructuring. In the Sea of Japan example (Krassilov, 1975b), differentiation of plant communities reflects both the scale and timing of tectonic events. A vegetational zonation over Japan Islands (Kimura, 2000 and previous work) is traceable to the maritime area of Primorye (Krassilov, 1967a). About one third of Neocomian species are distinctive for the inner (Tetori) zone of Japan as compared with the roughly contemporaneous outer (Rioseki) zone. On account of bennettite diversities, including several Dictyozamites morphotypes, the Tetori flora appears more thermophilic of the two. Since their dividing “median line” is a shear fault zone, the vegetational differences might


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have been owing to juxtaposition of terrains from different latitudinal zones. However, this explanation does not hold for Primorye, where the early Neocomian flora was of a Tetori aspect, with Dictyozamites kawasakii and D. falcatus prominent among the bennettites, while the subsequent Aptian to Albian assemblages varied across the incipient Sikhote Alin Ranges. The lee side assemblages were rich in bennettites and brachyphyllous conifers representing open xeromorphic vegetation. Those of the see-facing slope were of a conifer rainforest type fringed by the coastal fern marshes. Coal accumulation was much thicker in the foredeep than in the inland basins, with resinites characteristic of the latter. The late Albian redbeds are also confined to the inner zone. Thus, both the vegetational and sedimentary features indicate rain shadow over the lee side of the coastal ranges. In the early Neocomian, the rain shadow covered the inner zone of Japan as well as the eastern Primorye. Uplifts that cast this shadow might have developed over the midline fault zone of Japanese Islands that were then part of the continent. A radical change has occurred since the Late Neocomian tectonism, a major thrusting phase of Pacific margin. The Permian to Jurassic reefal limestones and cherts were emplaced over the axial fault zone of the incipient Sikhote Aline ranges (Krassilov & Parnyakov, 1987; Kemkin & Filippov., 2001). The overethrusts might have been rooted in the uplifted landmass east of Primorye, including the Yamato Rise (emergent in the Cretaceous, with Classopollis pollen assemblages coming from an inner zone vegetation: Markevich, 1994) and the cratonic foreland of Honshu. Before the thrusting event, the rain shadow over eastern Primorye was conceivably cast by this landmass. It was removed with a displacement of the Yamato – Honshu landmass over the NNE striking dextral strike-slip fault system. A series of heterochronous pull-apart basins has developed in between, with a stepwise oroclinal bend of Japan arc commencing in the Early Cretaceous, 120 Ma (Uyeda & Miyashiro, 1974) and intermittently continuing until the Miocene (Ito, 2001). Rainforest and wetlands spread over the coastal plains of the arising backarc basin, while the rain shadow was displaced to the inland basins west of the rising orogenic structure over the Neocomian overthrusts (Fig. 75).

V.9.2. Land connections and gateways Floristic similarity is evidence of a nascent or recent land connection. The Angarian lepidophytes occur in the Carboniferous of Brooks Range, Alaska (Thomas & Spicer, 1986). The Cathaysian elements are found in the Permian flora of western North America (IV.3). The Cretaceous – Palaeocene floras of eastern Asia and western North America share virtually all their dominant species of conifers and angiosperms, both terrestrial and aquatic (Taxodium olrikii, Metasequoia occidentalis, Trochodendroides arctica, Nordenskioldia borealis, Platanites raynoldsii, Viburniphyllum finale, Limnobiophyllum scutatum, Quereuxia angulata, etc., see Krassilov, 1976a, 1979; Tiffney, 1985; Manchester, 1999). These floristic data confirm at least intermittent trans-

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Fig. 75. Differentiation of the Early Cretaceous vegetation in South Primorye and Japan (the Japanese localities after Kimura, 1961): fossil plant localities of the inner and outer zones in the Berriasian-Valanginian (filled and open circles, respectively) and in the Barremian-Albian (triangles), showing a displacement of vegetation zones after the initial opening of the Japan Sea basin.

Fig. 76. Possible intercontinental links over tectonic arc and ridge systems : (A) Azores Plateau, (BS) Bering Shelf Arc, (GA) Gibraltar Arc, (MB) Macquarie-Balleny Arc, (MT) Madeira – Tore Rise, (NA) Nicoya Arc, (NR) Southeastern Newfoundland Ridge, (PA) Patagonian – Antarctic Arc, (PC) Papua – New Caledonia Arc, (RG) Rio Grande Rise, (SP) São Paulo Plateau, (S-F) Scoresby – Faeroe Ridge, (SS) Sunda Shelf, (T) Tasmantis, (WR) Walvis Ridge.

Beringian land connections over most of the Phanerozoic times. They obviously contradict the continental reassemblies in which northeastern Asia is widely separated from northwestern North America (e.g., Scotese, 1984, 1991; Scotese & McKerrow, 1990). Yet land bridges (Fig. 76) are prominent in the history of terrestrial biota not only as routes of terrestrial migrations, but also due to their far-reaching effects on climate and


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vegetation. Notably, the Early-Middle Eocene thermal event correlates with the emergence of the North Atlantic land bridge (McKenna, 1975; Tiffney, 1985a, 1985b) that not only provided for floristic exchanges between Eurasia and North America, but also interrupted the psychrospheric circulation in the Atlantic. A closure of gateways for cold water masses resulted in a warmer North Atlantic, allowing the paratropical coastal vegetation to spread beyond 60°N (VII.5). A similar, though shorter-term, vegetation event in the Pacific sector (palms in northern Kamchatka and Alaska) attests to an emergence of the Bering Shelf. This prominent phytogeographic anomaly is thus related to a general upheaval of Arctic land as an epeirogenic consequence of decelerated earth’s spin (VII.1). A subsequent temperization of climate and biota in the Late Eocene, continuing into the Oligocene, correlates with separation of the North Atlantic landmasses due to the opposite tendency of tectonic evolution – a contraction/depression of the polar areas, with consequences for oceanic circulation and climate.

V.9.3.Palaeoecological constrains of continental rearrangements Continental reassemblies are palaeogegraphic inferences based on sea-floor spreading geometry and palaeomagnetology. Since both these sources are hypothetical, the reassemblies have to be warranted by palaeogeography in the broadest sense, including palaeoclimatology and palaeoecology. If an apparently anomalous palaeogeographic feature can be explained otherwise than by a continental rearrangement it need not be used as an incentive for the latter. If, on the other hand, the neutral data are used as a confirmation or the contradictory data are neglected, then we have conspiracy rather than science. For example, the floristic differences of Laurasia and Gondwanaland are commonly used as evidence of a wide continental separation neglecting mixed floras of the Middle East, South China, New Guinea, etc. (IV.3). Floristic uniformity is a venerable argument for assembling the Gondwanaland supercontinent, although a floristic differentiation of the western/eastern Gondwanas became evident far before its alleged break-up in the Jurassic. At the same time, floristic similarities, as between the northern and southern Chinese blocks or over the Palaeozoic Uralian seaway, negate tectonic arguments for vast oceanic barriers. A critical test for continental reassemblies is climatic/vegetational zonation. On the global scale it reflects solar flux, hence latitudinal. In comparison, the cross-latitude vegetation boundaries are owing to such factors as rain shadows, monsoons, substrates (permafrost in the case of the mid-Siberian differentiation), etc., hence regional. In this respect the Palaeozoic and Mesozoic phytogeographies (IV.3) are more consistent with the present-day coordinates in which the planetary boundaries are roughly latitudinal, than with continental reassemblies in which they are differently positioned.

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In the Cretaceous, the mesothermic/xerothermic ecotone extends at about 50°N over continental interiors of Laurasia. (IV.3.2) In the Southern Hemisphere, the equivalent boundary divides the mesothermal assemblages with diverse Proteaceae and evergreen Nothofagus (brassi–type) from those dominated by the deciduous Nothofagus species at about 45°S, modern latitudes, both in Patagonia and New Zealand (Specht et al. 1992), but is totally incongruous on the reassemblies that displace New Zealand to 60°S. Growth rings indicate a warm-temperate Cretaceous climate for the Fossil Bluff wood assemblages of Alexander Island (mean width 9.5 mm) changing to a cool-temperate (mean width 2.3 mm) at James Ross (Jefferson, 1982; Francis, 1986). Such differentiation corresponds to the Northern Hemisphere zonation at about 60°N, whereas on the continental reassemblies it crosses the Southern Pole. With the mean widths of growth rings in gymnosperm trees evolving from 4.5 mm in the Permian (Taylor et al., 1992) to 2.3 mm in the Cretaceous, the palaeoecological data suggest a southward, if any, displacement of Antarctica since the Palaeozoic. A major deviation of zonal boundaries from latitudinal orientation is their northward displacement in the North Atlantic sector, which is systematically recorded since the Carboniferous to the present, hence unrelated to the alleged opening of the Atlantic (instead it may attest to its being open through all these times). Presently the displacement is due to the Atlantic air masses and the same may be true for the past. The northward shift of vegetation zones was maximal at the time of low European land, as in the Early Carboniferous, Late Cretaceous and Early Palaeogene, allowing a deeper penetration of Atlantic air to the continental interiors. However, while the present-day Greenland is cooler than the Arctic lands of northern Europe, fossil plant assemblages suggest the opposite for both the Triassic and midCretaceous (IV.3), indicating an anomalously warm climate for its present latitudinal position. A northward displacement of about 5° since the Cretaceous, mediated by the NW – NE striking fault system between Arctic Canada and Greenland would explain this phyogeographic anomaly.

V.9.4. Permian landmasses and phytogeography In 1920s a uniform Late Palaeozoic Glossopteris flora of India and southern continents prompted the idea of Palaeozoic supercontinent Gondwanaland separated from the Laurasian continents by the Tethys Ocean. A closure of this ocean, or a series of successive closures, gave rise to Pangea or Pangeas. These influential ideas gave rise to the continental drift theory and, fifty years later, were adopted by the plate tectonics. Some reconstructions of Gondwanaland place India near Africa, while others link it to Australia. Either of the fits may seem geometrically plausible because configurations of southern continents are defined by the S-fault systems (V.5), hence congruous. Yet the Pangea poses further problems. Nearly identical structures have to be ascribed to


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convergence in the western part of it and to divergence in the east. It is virtually impossible to reconcile the reconstructions for the Permian/Triassic with those for the Late Triassic without invoking a megashear that is not there (Weil et al., 2001). Despite such geological inconsistencies, the Gondwanaland concept is immensely popular among biogeographers and evolutionists relating phylogenetic divergence to continental separations. In particular, it encouraged separate phylogenies for the “northern” and “southern” plant orders, such as conifers (Florin, 1954; Meyen, 1987). There are also proponents of a pre-Cretaceous divergence of the “Laurasian” and “Gondwanic” angiosperms. The following considerations seem pertinent to the Gondwanaland concept as it stands today: (1) All terrestrial plants are able to surmount water barriers. Some terrestrial plants, such as the coconut palm, are adapted to dispersal by sea currents. If terrestrial plants do not disperse over seaways (as, for instance, the redwood species that are lacking on the offshore Californian islands), an explanation can be more readily related to habitat limitations rather than water barriers. (2) Floristic differentiation, even of the highest phytogeographic rank, as between the African Sahel and Cape provinces, can develop even in the absence of water barriers. The continuity of Glossopteris flora does not prove a continuous landmass. (3) Landmass separation in the Early Carboniferous did not interfere with, even seemed promoting a floristic homogeneity, whereas the floristic divergence of Gondwana and Laurasia developed at the time of landmass aggregation on the way to “Pangea”. (4) The major groups of the present-day “southern” conifers, the Araucariaceae and Podocarpaceae, were widely spread over the northern continents during the Mesozoic and early Tertiary (Krassilov, 1967a, 1976a). Extinct angiosperms related to the presently typically “southern” Proteaceae and Nothofagus have been also found on Laurasian lands (e.g. Ushia: Krassilov et al., 1996b). Their present-day southern distribution is owing to their persistence there rather than the Gondwanaland origins. In much the same way, the “Angaraland” has preserved its climatic and vegetational identity far beyond its supposed isolation as a landmass by the Uralian and MongoloOkhotskian seaways. The ecotonal Eurangarian (“Subangarian”) and Cathangarian provinces over the boundaries of “Angaraland” are evidence of climatic differentiation rather than landmass separation. As discussed in (IV.3.1), the Permian phytogeography reflects a system of cooltemperate to tropical biomes that concur with the present-day latitudinal zones, thus scarcely compatible with the palaeogeographic reconstructions in which Laurasia is placed

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submeridionally, with the Tethys widely gaping to the east (Scotese, 1984, 1991 and elsewhere). Actually the Permian phytogeographic realms do not seem anomalous on the present-day geographic maps. The Angarian realm roughly coincides with the presentday zone of northern conifer forests and tundras (no ecological equivalence is implied; just a congruence of geographic outlines similarly controlled by latitudinal climatic zonation). The western boundary of the Angarian realm is oblique to the modern latitudes, which does not necessarily mean rotation of the continent because the corresponding boundaries are similarly oblique through the Mesozoic (Krassilov, 1972b, 1997c and elsewhere), the Tertiary and even at present. Likewise, the geographic range of Glossopteris is fairly similar to that of the presentday “Gondwanic” family Proteaceae that appeared about 100 m.y. after the supposed break-up of Gondwanaland. Notably, the mixed Cathaysian – Gondwana flora of New Guinea occurs fairly close to the present-day Wallace line between Bali and Lombock. A widely held biogeographic opinion links such present-day geographic features with past distribution of landmasses. But traced back to the Permian they are scarcely compatible with a radical reassembly of the continents. Neither of the Permian phytogeographic realms was restricted to a separate landmass. Their expansion over the Uralian, Tibetan and Ginling sutures is consistent with intermittent seaways rather than vast oceans. Suturing of the Uralian geosyncline rendered a minor effect on the Angarian biome that preserved its climatic and vegetational identity far beyond its supposed isolation as a landmass. In fact, the Triassic Siberian realm still coincided with the “Angaraland” (Krassilov & Shorokhova, 1975) and even the post-Triassic vegetational – climatic boundaries still tracked the familiar outline (Krassilov, 1972b; 1997c) In much the same way, the seaways of central and southern China had little effect on floristic similariry of the northern and southern Cathaysian blocks. Their floristic distinction increased (rather than decreased with a supposed collision of landmasses) in the latest Late Permian owing to expansion of dry subtropical zone that brought a wave of European migrants to northern China. The integrity of the Gondwana realm as a continuous landmass can be put to doubt on both phytogeographic and climatological grounds. If glaciated, the “Gondwanaland” interiors would experience a dry climate, with cyclonic air masses barred from penetration inland by cold atmospheric fronts. Neither coal accumulation nor a permanent ice cover would have been sustainable in such conditions. The apparent floristic homogeneity is due primarily to a broad ecological range of glossopterids spreading from cool-temperate to dry subtropical redbed zone, with a maximal diversity in the latter. All over these ranges, including Antarctica, the arboreal broadleaved glossopterid communities had little in common with the Pleistocene to recent periglacial vegetation, suggesting a totally different character of the Permian glacial events. In the Triassic, the glossopterids survived in a redbed zone extending from India to northern China. Continental reassemblies placing the Permian South Pole near Africa and the Euramerian redbed zone at the equator are scarcely consistent with the Permian phytogeography.


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On both palaeovegetational and lithological evidence, central Antarctica experienced a much cooler Permian climate than South Africa or South America, which implies the Permian South Pole close to its present-day position. From an ecological standpoint, the Euramerian plant assemblages did not match either humid or arid variants of tropical vegetation. They were oligodominant, of xeromorphic aspect and with abundant well preserved conifer remains that obviously came from lowland sources (conifers scarcely grow in tropical lowlands). There were free floristic exchanges between this and the temperate Angarian realms via a warm-temperate ecotone over the Volga Basin and Cisuralia. The Euramerian-Cathamerian-Gondwana “floristic mixing” might have resulted from penetration of interzonal wetland communities (Fig. 51) in a disturbed vegetation of the tectonically active Tethys belt. Equally probable is a taphonomic mixing of plant remains coming from different altitudinal belts. Remarkable in this respect is the present-day mixed vegetation of New Guinea, of Malaysian aspect in the lowlands while of Australian affinities upslope (Schuster, 1972). As shown on the present-day map (Fig. 52), the Tethys corresponds to a great shear belt that extends over the Sulaiman Ranges of Pakistan, the Karokoram-Nujiang fault zone of Tibet, the Arakan Yoma Range in Burma, the Andaman-Nikobar Ridge over the Andaman Sea to the Sunda Trench, Timor and the Owen Stanley Range of New Guinea (reviewed in Krassilov, 1985). Tethyan ophiolites are confined to the shear fault zones between cratonic blocks of the mosaic belt structure over the Mediterranean, Middle East, Himalayas and southeastern Asia. At the time of Tethyan transgression, most of the area was covered with epeiric seas. Yet the Permian floristic mixing is incompatible with the idea of Tethys as a broad ocean between Laurasia and “Gondwanaland”. Floristic links across the ophiolitic zones mean that the Tethyan blocks were parts of a continuous fold belt in the Permian already. Both the Indo-Australian floristic links and a constriction of the Tethys belt over the junction of Tibet and Himalayas are consistent with a large-scale right-lateral shear that displaced continental blocks of southeastern Asia north of their pre-Permian position closer to the Sunda plate. Since in the earth’s rotation stress field, the maximal displacement zone corresponds to the equatorial minimum of Coriolis force, such megashear would also comply with equatorial position of southeastern branch of the Tethys.

V.9.5. Trans-Tethyan biotic links and collision of India to mainland Asia The Tethys has been originally conceived of as a broad seaway marked by flysch and pelagic carbonates that extended over the presently uplifted lands of southern Europe, northern Africa, Middle East, Himalayas and southern China. Since the Alpine type orogenic structures were interpreted as series of nappes, the Tethys belt has been related to a tectonic contraction closing the seaway, with the ophiolites as remains of oceanic crust crushed in collision of continental blocks (microcontinents).

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In this latter model, the Indian plate has collided with the Xizang (Tibet) microplates sutured by ophiolitic thrust belts. Of them, the Lhasa (Thanghla) plate of southern Tibet is a Precambrian cratonic block, the sedimenatry cover of which includes, successively, the Gondwana-type tillites, the non-marine clastic deposits with a mixed Cathaysian/ Gondwana flora, and the Cretaceous (Aptian to Albian) Takena redbeds, folded in the Late Cretaceous and the overlain by the Palaeogene volcanics. The fossil wood from Takena shows distinct growth rings indicating a temperate climate, similar to that over the present-day mid-latitudes (30-40°N). Since at the northern (Bangong-Nujiang, or Nu Chang) thrust border the Takena redbeds spread over the Donqiao ophiolite (Xuchang et al., 1980; Girardeau et al., 1984), the Lhasa plate should have been sutured to the mainland by the mid-Cretaceous. On evidence of the older Cathaysian flora it might have occurred anywhere close to the southern Chinese landmass in the Permian already, while an association of the Gondwanic floristic elements and tillites indicates that Indian plate should also have been not far from its present-day position relative to Tibet. Their suturing Yarlung–Tsangbo thrust belt contains the Albian ophiolite assemblage (the Xigase ophiolite: Allegre et al., 1984; Nicolas et al., 1981; Tapponier et al., 1981) emplaced by the Campanian. In angiosperm assemblage from Xigaze Group, 71% of species have non-entire leaf margins, about 64% are deciduous (Guo, 1996), which indicates a warm-temperate climate. The Tsangbo belt is connected to the Kokhistan – Ladakh arc to the west and to the Naga Hill – Arakan Yoma arc to the east, forming a semicircular suture with a series of nappes emplaced upon the Tethys-Himalaya in the Late Cretaceous to Palaeogene and overthrust by the Trans-Himalayan granitoid belt of about the same age (Honegger et al., 1982; Thakur & Misra, 1984; Schärer et al., 1984). The significance of these age assignments will be clear in relation to the other parts of the Tethyan mosaic (Figs. 70 – 75). Of the better known Tethyan ophiolites, the Troodos of Cypres on the southern border of Taurian block is radiometrically dated as the mid-Cretaceous. Cenomanian foraminifers are found in limestones between lava complexes the upper of which is topped by umbra and overlain by the Maastrichtian to Palaeogene carbonates (Delaloye et al., 1980; Pantazias, 1980). In Zagros, between the Taurian and Iranian (Dasht-E-Lût) blocks, an allochthonous complex is thrust over the Coniacian – Santonian and is overlain by the Maastrichtian reefal limestones (Berberian & King, 1981). The Samail ophiolite of Oman was similarly emplaced in the Campanian to Early Maastrichtian, with the shallow-water Maastrichtian carbonates deposited on a lateritic weathering crust (Smewing, 1980; Tippit et al., 1981; Coleman, 1981; Woodcock & Robinson, 1982). The above chronology of ophiolite emplacements indicate that suturing of all continental fragments in both the Tauro-Omanian and Afgano-Himalayan branches was accomplished in the Campanian to Maastrichtian time. The sutures then emerged above sea level exposing their ophiolites to subaerial weathering.


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Biogeographic data are in agreement with this scenario. The Gondwana/Cathaysian floristic mixing over the Tethys belt indicates a proximity of the Chinese and Indian blocks with at least temporary land connections (IV.3). The Jurassic to mid-Cretaceous floras of India shared their dominant bennettite elements (Ptilophyllum–Otozamites group) with Central Asia, while an extinct gymnosperm order Pentoxylales is shared with Australia and New Zealand. Again the Tethys was not insurmountable for terrestrial plants and even for some dinosaurs, such ankylosaurs, a widespread Laurasian group, occurring in the Turonian Lameta Group of central India (Berman & Jain, 1982). The Indian – Australian floristic links do not necessarily mean a direct contact because the then periodically emerging Sunda Shelf might have served as a land bridge. The chronology of marginal sedimentation suggests separation of Australia from India (or any landmass off the eastern margin) in the Oxfordian – Valanginian (Veevers et al., 1991). Since this is also the age of the older Indonesian ophiolites (V.9.5), the separation might actually reflect fragmentation of continental crust over the New Guinean geoidal bulge (V.3). The Campanian to Maastrichtian tectonic consolidation of mobile blocks between Africa, southwestern Asia and India, forming the Tethys mosaic, coincides with penetration of European and African floristic elements, notably of the trans-Atlantic “Normapolles Province”, to the Middle East and India (Herngreen & Chlonova, 1981; Nandi, 1983; S.K. Srivastava, 1983). By the Maastrichtian, India and northern Africa belonged to one and the same palynological province, with the newly appearing morphotypes, such as Constantinisporites, at the roughly equivalent stratigraphic levels on both continents. The Indian example demonstrates a congruence of tectonic and phytogeographic evidence neither of which is compatible with a simplistic idea of the Indian block being rafted across the vast spaces of Tethys Ocean to its collision with mainland Asia in the Eocene. Instead the chronology of ophiolite emplacements indicates a simultaneous suturing of cratonic blocks all over the southern Tethys by the end of Cretaceous, immediately followed by floristic invasions that attest to land connections over the consolidated fold belt. A subsequent re-shearing and orogenic uplift in the Eocene has enhanced the floristic differentiation over the Tethys.

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Chapter 6. See-level fluctuations

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VI. SEA-LEVEL FLUCTUATIONS The idea of global change has been prompted by remains of marine organisms on land lending support for the Biblical deluge and derivative catastrophic theories. Sealevel fluctuations are still considered as the most important factor of mass extinctions– innovations (Newell, 1967), although their impact is usually relegated to a reduction/ expansion of certain habitats. Yet the biospheric consequences of sea-level change extend far beyond this. In the system of global change, sea-level fluctuations are related to orbital, tectonic and climatic events, with feedbacks. Their biospheric significance lies, first, in a reduction of terrestrial biomass, which means a near proportionate decrease in the total biomass. The areal extent of pedogenesis decreases with it. Since both the standing crop biomass and pedogenic weathering of silicate rocks are major sinks for atmospheric CO2, there should be a positive correlation of transgressions with greenhouse effect (VII.2.4). As they expand, epeiric seas take over a control of oceanic circulation. Oceanic depressions and epeiric seas form a system of estuarine – lagoonal exchange (Thiersten & Berger, 1978). Those with estuarine circulation (inflow of deep water – outflow of surface water) accumulate organic-rich deposits (the black shale events that correlate with transgressions), while carbonate sedimentation prevails in the “lagoonal” (reverse) basins. The amount of terrestrial runoff is, to a considerable extent, controlled by sea level that determines the basis of erosion for drainage systems. Influx of drained material to oceanic basins is further regulated by the extent of epeiric seas which at the peak of global transgressions trap a larger portion of terrestrial runoff. Hence a decrease in organic content – increase in carbonate deposition. Additionally, the filtering effect of coastal vegetation varies with coastal geomorphology, penetration of seawater into ground aquifers, tide amplitudes, etc. Coastal wetlands, a major source of dead mass exported to oceanic basins, expand or shrink with sea-level fluctuations and the relative development of deltaic to estuarine ecosystems, their outflux of oxydised or anoxic waters, their autotrophic versus heterotrophic functioning reflected in the ratios of fibrous to amorphous kerogens in marine deposits (II.I). Sea-level fluctuations are relatively well-documented, yet their origin is a matter of a lingering controversy. Recent fluctuations of an amplitude of about 100 m are considered as glacioeustatic. However such, and even larger, about 300-400 m, fluctuations over the Cretaceous (Sloss & Speed, 1974; Bond, 1976; Vail et al., 1977), essentially a nonglacial period, must have been inflicted by other factors.

VI.1. Sea-level models Although sea-level fluctuations occurred at all times, the modal sea levels differred from one geological epoch to another, defining an alternation of geocratic and thalasso-

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cratic epochs. Relative extents of marine deposition seem to record a decrease of epeiric seas over geological history. However, on evidence of Cretaceous transgressins, this tendency was not sustained over the entire Phanerozoic. It should be noted that the fossil record actually documents shoreline displacements, which are not exactly the same as sea-level fluctuations relative to the geoid mid-level. Shore lines apparently tend to approach the boundary of continental and oceanic crust domains, but are displaced by either eustatic or tectonic events, or both. The following causative types of sea-level changes are distinguished: (1) Eustatic, related to changes in water volume, including the glacioeustatic fluctuations of 102 m scale recorded in the Pleistocene and extrapolated back in time. The commonly recognized glacioeustatic fluctuations are phased by the obliquity cycles, with a period of 41 k.y. However a similar cyclicity is recorded for non-glacial epochs as well. Eustatic events caused by thermal expansion of seawater correlate with low latitude insolation maxima of the orbital eccentricity cycles. Such eccentricity-driven sea-level fluctuations are globally synchronous, of 100 k.y. and 400 k.y. periodicities. The Tethyan transgression cycles of respective periodicities are recorded in the Cretaceous (Stroll & Schrag, 1996). The 400 k.y. is of a biospheric significance, as indicated by a chronology of speciation events and the average species longevity in such rapidly evolving groups as the Permian conodonts or Cretaceous foraminifers and ammonoids. (2) Geodetic, pertaining to deformations – redistributions of elevations and depressions – of the geoid (V.3). A reversal of the present-day geoid highs (+ 100 m near New Guinea) and lows (- 80 m near the Maldives) would generate a sea-level wave of amplitude about 180 m. With deceleration of the earth’s rotation rates, sea level would have been elevated at the high latitudes and depressed at the lower latitudes (Mörner, 1976; 1980), and the opposite with acceleration. Expansion of the astenospheric equatorial bulge would coincide with depression of polar areas. The Boreal to Tethyan transgressive/regressive events of the geodetic sea-level cycles would then be at counter-phase, and the sealevel fluctuations would correlate with magmatic pulses. The chronology of large-scale sea-level events seems consistent with this model. At least, a major mid-Cretaceous Tethyan transgression was accompanied by an emergence of vast land areas in the northern North Atlantic (the Thulean Land), Beringia and elsewhere in the Arctic Basin, as well as in the Tasmantis area of southern Pacific (Krassilov, 1985 and similar examples in Lapkin & Katz, 1990). Conversely, the maximal Late Cretaceous advance of high-latitude epeiric seas over northern European, West Siberian and the interior North America (Fig. 77) coincides with the peaks of low-latitude magmatism over the Tethyan ophiolite belts and oceanic fracture zones (V.6.5, V.7). There are, however, incidents of globally synchronous transgressions and regressions (e.g., in the terminal Permian or terminal Cretaceous), as well as magmatic pulses, that are not explained by the geodetic model, hence of a different origin (e.g., isostatic, below).


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Fig. 77. Epirift subsidence in the Cretaceous, with epeiric seaways extending over the meridional Triassic and older rift zones: (NA) North American Interior Basin, (NE) North Sea – Northern European Basin, (WS) West Siberian and Yenisey-Khatanga basins.

The pole-wander model ascribes sea-level fluctuations to migration of equatorial tidal bulge with displacement of the axis of rotation (Mörner, 1980; Steinberger & O’Connell, 1997; Mound & Mitrovica, 1998). Sea level falls in front of the migrating pole and rises in the rear (Mound & Mitrovica, 1998). According to this model, a sea-level wave would diachronously spread over the globe transverse to the polar path. The effect remains hypothetical however either for the observed short-term Chandler wobble or the alleged large-scale pole wander of geological history. (3) Isostatic, pertaining to relative hypsometric levels of oceanic and continental lithosphere defined by their isostatic compensation at given earth’s rotation rates. A rotation forcing would displace mean hypsometric levels for the isostatically compensated oceanic and continental lithosphere because of their density contrasts resulting in a differential centrifugal acceleration (V.2). Since oceanic lithosphere is denser than continental lithosphere, a change in the earth’s rotation rates would either depress or raise sea floor relative to the continents. Mean hypsometric levels of land and sea floor would converge with deceleration and diverge with acceleration of the earth’s rotation rates. Convergence results in shallow oceans – low continents, hence a global transgression, and the opposite under divergence. This model (Fig. 78) explains alternation of thalassocratic/geocratic epochs, with the lower order sea-level fluctuations imposed by the associated isostatic effects of surface loads (e.g., the glacioisostatic effects of continental ice load and rebound of deglaciated areas appreciable over the periods of vast continental glaciations: Peltier, 1996). In addition, the tidal deceleration of the earth’s rotation rates may appreciably increase with the tidal flat areas during thalassocratic epochs, hence a positive feedback enhancing the trend. In respect to the geochronology, the model predicts coincidence of global transgressions with sea floor magmatism related to an expansion–elevation of oceanic lithosphere. Such a correlation is actually observed during the Cretaceous, a period of vast epicontinental seas and the most voluminous sea-floor basalts (details in V.5.2, V.7, VI.1

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Fig. 78. An interpretation of sea-level fluctuations resulting from divergence/ convergence of mean hypsometric levels for isostatically compensated oceanic and continental domains under centrifugal rotation forcing. Present: shore-lines close to the boundary of isostatically compensated continental and oceanic blocks. Acceleration – divergence: shore-lines retreat with a rise of continents and subsidence of oceanic depressions. Deceleration – convergence: the opposite tendency resulting in low continents and shallow oceanic basins.

and elsewhere). On the other hand, global transgressions would have been synchronized with intracontinental magamtism (traps), which is obviously the case at both the Permian/Triassic and Cretaceous/Palaeogene boundaries. Thus, through application of rotation geodynamics to sea-level fluctuations, we arrive at a causal scheme providing a long-sought explanation of the antithetic pulses of the cratonic sea-floor and continental volcanism correlating with the talassocratic and geocratic tendencies respectively, The intermediate island-arc volcanism increases with marginal shear over the transitional stages.

VI.2. Hypsometric curve The above model of global sea-level changes in respect to the relative hypsometric levels of isostatically compensated oceanic and continental lithosphere can be tested by comparing hypsometric curves for the present geocratic versus mid-Cretaceous thalassocratic epoch. Hypsometric curve shows how much of the earth’s surface (percentages) is maintained at a particular hypsometric level. Hypsometric steps correspond to the larger lithospheric domains of different density isostatically compensated in the earth’s rotation field. The curve thus reflects isostatic situation at a given velocity of rotation.


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Evidence on evolution of the earth’s hypsometric curve over times comes from land/ sea distributions, as well as thickness and composition, in particular the silicoclastic/ carbonate ratios, of sedimentary cover. In non-glacial epochs (or the epochs with insignificant polar ice caps), sea level is controlled by the capacity of oceanic depressions. Shallowing oceans splay large amount of seawater over the continents as epeiric seas. Carbonate sedimentation is correlated with an extent of epeiric seas and, inversely, with an input of terrestrial silicoclastics, insignificant at the time of low continents. Conversely, with isostatic sea-floor subsidence, a capacity of oceanic depressions increases inviting a retreat of eperic seas from continental margins. Deepwater circulation drives the contemporary carbonate compensation levels to shallower depths inflicting hardgrounds over vast sediment-starved areas. Most of the present-day land area occurs at elevations 0-1000 m above the contemporary sea-level while about 60% of oceanic area are depressed to 4000-6000 m depths. These two hypsometric levels correspond to the isostatically compensated domains of continental and oceanic lithosphere. Epeiric seas are mostly restricted to the Arctic and Antarctic shelf areas. In contrast, the Cretaceous epeiric seas inundated cratonic areas covering about 60% of Europe, 45% of North America, 43% of Africa, and 40% of Australia, altogether about 60.106 sq. km. of the present-day land. In addition, about 20.106 sq. km of periodically emerged Arctic shelves, northern North Atlantic, Tasmantis (Lord Howe, Norfolk, Three Kings rises, Campbell and Chatham plateaux adjacent to New Zealand) and a number of smaller oceanic plateaux occurred at about sea level (V.7). The mid-Cretaceous transgression started in the latest Albian and reached a middle peak of about 30% of continental area in the early Cenomanian, only about 500 k.y. later. Such transgression rates, and the subsequent large-scale fluctuations, suggest a vast land area, comparable to the extent of epeiric seas, at elevations from zero to +200 m. Most epeiric seas were about 200 m deep, with maximal depths about 500 m. Summarily, the epeiric seas and the adjacent periodically inundated land amount to about 80,106 sq. km, or 55% of continental area at altitudes from -200 m to +200 m, about twice more than at the same hypsometric interval at present. A correlation of sedimentary rates with elevation of source areas also indicates a relatively low stand of Cretaceous continents. While the basal Cretaceous Wealdentype facies are due to a considerable relief of faulted continental margins, the Late Cretaceous marginal deposits are about four times thinner owing to a general subsidence of continental crust. At the same time, the carbonate content of marine deposits increased dramatically as most or all sedimentation areas occurred above the contemporary carbonate dissolution depth. On sedimentological evidence of coarse-grained volcanoclastic deposits, areas of a relatively high mid-Cretaceous relief were restricted to the volcanic plateaux of centraleastern Asia (Transbaikalia, Mongolia, eastern China), as well as to the circum-Pacific coastal ranges, a source of thick silicoclastic deposition in the foredeeps (e.g., the Great

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Valley trough of California, the Tatar Strait – West Sakhalin trough of Russian Far East, etc. (V.6.3). Floristic comparisons over the coastal ranges indicate a rain shadow that gave a leeside vegetation a mildly xeromorphic aspect, as in the present-day oak woodland of interior Primorye west of Sikhote Alin (Krassilov, 1967a, 1972a; V.9.1), but not precluding peat deposition and of a far lesser desiccation effect than the present-day Rocky Mountains. These mid-Cretaceous ranges were not higher, lower perhaps, than the present-day Sikhote Alin, that is about 800 – 1500 m (solitary peaks up to 2000 m). Depths of Cretaceous oceans can be inferred from lithofacies and marine fossils. Bosellini & Winterer (1975) have postulated a downslope facies sequence of cephalopod limestones (with aragonitic shells, hence above the aragonite lysocline) → Maiolicatype limestones lacking aragonitic shells, but containing calcitic aptichi → chert of ophiolite assemblages, spanning the depths from few hundred to about 2700 m. They also suggested that, during the Cretaceous, carbonate compensation depths (CCD) were shallower than at present. Even in the epeiric seas, carbonate sedimentation might have periodically occurred below the contemporary aragonite lysocline because only calcitic aptichi were preserved. At least we can safely assume that the aragonite lysocline was not deeper than the present-day pteropod silt deposition at 2600 m. On the other hand, a sequence of cephalopod limestones, aptichi limestones and chert, to which we can add black shales, corresponds to a sequential change in pH values. Pure limestone deposition occurs in a nearly neutral environment with pH about 6.8 to 7. Dolomitic deposition requires a more alkalic environment, with pH further decreasing over the sequence of phosphoritic chalk – bituminous carbonates – black shales – silicites. These lithologies, and their transitional types, such as the siliceous/bituminous carbonates, form a frequently recurrent sedimentary cycle indicating flexibility of depositional environments that can be more plausibly ascribed to pH fluctuations than depth fluctuations (II.7.4). The ultramafic members of ophiolite associations have previously been interpreted as remains of oceanic crust formed at spreading axes. Hence the deposition depth of their covering sediments was considered as the depth of the axes. As discussed above (V.6.2), ophiolites are generated at shear faults rather than spreading axes. Their associated sedimentary assemblages, traditionally interpreted as pelagic deposits, are silicified turbidites sometimes containing fossil wood, as in the Ligurid assemblages (Abbate et al., 1980), or even well preserved leaves of terrestrial plants as on the Lesser Kuril Islands (Krassilov et al., 1988). Zooxanthellate corals in the allegedly deep-water carbonates indicate a sedimentary environment within the photic zone. These carbonate facies are laterally substituted by radiolarian cherts and Aptychus beds (Sanantonio et al., 1996). None of these is evidence of deep-water environments. Depths of 600-1000 m are postulated for oceanic black shales and their facial equivalents, such as the Livello Bonarelli (Bonarelli Level) of Apennines (Van Graas et al., 1983). However the organic matter of oceanic black shales is primarily of an estuarine origin (II.7.4). The black shale facies are linked with contemporaneous carbonates, with


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bituminous limestones as intermediates and with the mm-thick lamellae of black shale sequences left from post-sedimentary decalcification. Shallow-water to paralic Cretaceous deposits occur on submerged oceanic plateaux, rises and aseismic ridges. In the North Atlantic, shallow-water mid-Cretaceous facies and subaerially weathered basalts occur on the J-anomaly Ridge, Newfoundland Sea Mounts and elsewhere (Tucholke & Ludwig, 1982; Sullivan, 1983), occasionally even in the axial zone of Mid-Atlantic Rise (Bonatti & Crane, 1982). Bryozoan limestones are dragged from the Bahamas escarpment at a depth of 3100 m (Uchupi & Milliman, 1971). In the South Atlantic, the littoral mid-Cretaceous deposits are recorded on the Walvis Ridge. After a period of non-deposition they were covered by the shallow-water Maastrichtian limestones. Similar deposits also occur on the Rio Grande Rise and São Paulo Plateau (Hekinian, 1972; Humphris & Thompson, 1983; Schaffer, 1984). In the central Indian Ocean, the Late Cretaceous andesites of Ninetyeast Ridge are topped by shallow-water limestones and lignites (Hekinian, 1974; Sclater & Fisher, 1984). On the Broken Ridge and Kerguelen to the south, the Late Cretaceous shoal carbonates are truncated by a pre-Eocene erosion surface (Guilfy, 1973; Dosso et al., 1979; Mutter & Cande, 1983). Shallow-water Cretaceous facies are widespread over the Pacific rises and plateaux, including the Obrutchev Rise in the north, Magellan Rise in the central area and Manihiki Plateau to the south (Kauffman, 1974; Thiede et al., 1981; Pushcharovsky et al., 1984). Actually, all the positive oceanic structures presently submerged to about 2500 m were at shallow depths or above sea level in the Cretaceous. Summarily, these data indicate an about 2500 m subsidence of oceanic floor since the Cretaceous. By inference, the isostatically compensated oceanic crust were at depths of 1500 – 3500 m in the Cretaceous against 4000 – 6000 m at present. Cretaceous basalts are widespread over oceanic depressions piling as ridges along the then active fault zones. The huge amount of basaltic material, mostly of a tholeiitic composition (but not necessarily produced over the mid-ocean ridges, for there are examples of such “intraplate” basalts: Batiza et al., 1980) is in itself evidence of a general upheaval – extension on a sphere – and rifting of oceanic crust. These considerations give us a crude idea of hypsometric steps for a thalassocratic epoch, with a wide area, no less than 55% of the earth’s surface, at altitudes from +200 m to -200 m, a restricted highland area at +500 m to +1500 m, and oceanic depressions at -1500 m to -3500 m. Such a distribution is conveyed by a gently sloping hypsometric curve like that of the Moon, a slow rotating planet (Fig. 79). A thalassocratic situation like in the Cretaceous thus implies a convergence of mean hypsometric levels for the isostatically compensated oceanic and continental lithospheric domains resulting in shallow oceans and low land. A divergence of these levels generates the alternative situation, with deep oceans and elevated continents. Remarkably, rifting of cratonic areas (trap magmatism) marks a culmination of regressive trends at both the Permian/Triassic and Cretaceous/Tertiary boundaries (IX.1).

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Fig. 79. A reconstructed hypsometric curve of the mid-Cretaceous talassocratic epoch (II) compared with the present-day Earth (I) and the Moon (III).

VI.3. Tilt of the continents A comparison of Atlantic and Pacific margins suggests that, whatever the interpretation of marginal processes, raised margins are seismically/volcanically more active than subsided margins. As discussed in (V.6.3-4), marginal activities change over time. The presently relatively quiescent Atlantic margins are fringed by fossil thrust belts with superimposed volcanic structures (V.6.4). Marginal deposits up to 6000 m thick indicate a high relief of the marginal belts at the time of extensive Late Jurassic/Early Cretaceous basaltic volcanism in eastern Brazil and Namibia. The Pacific margins were then relatively quiescent. Reversal took place in the mid-Cretaceous, with epeiric seas transgressing over the Atlantic margins – retreating from over the circum-Pacific belt. The latter soon acquired its present-day structure, with volcanic ranges rising over the thrust complexes emplaced in the late Neocomian and sequentially granitized in the Albian to Maastrichtian (e.g., in the Sikhote Alin and Okhotsko-Chukotskian belts of Asiatic margin: Krassilov & Parnyakov, 1984; Krassilov, 1989b). Similar developments took place in the Tethyan belt after the Campanian to Maastrichtian thrust phase (V.6.5). In the plate tectonic model, such marginal processes are explained by subduction of oceanic lithosphere, although, realistically, underplating of continental crust by a cool dense slab would result in an isostatic subsidence rather than uplift. Alternatively, a marginal activation might have been owing to rotation forcing that induced a divergence (at acceleration) or convergence (at deceleration) of isostatic compensation levels over the boundary of continental and oceanic crust. Schematically (Fig. 78), the rotation forcing cause activation of the marginal shear zones, with subcrustal material transported from beneath the continent onto the oceanic floor and rethrusted over the continental margin. Thickened by the nappes, the marginal


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zone is heated, granitized and elevated giving the continent a tilt away from the active zone. A correlation of vast epeiric seas over northern Eurasia and North America with the Tethyan–Pacific mid-Cretaceous thrust phases might have been owing to the tilt phenomenon. The present-day Eurasia is tilted to the northwest, with the larger rivers flowing into the Atlantic and Arctic oceans. This configuration results from the Cretaceous and subsequent tectonic develoment over the Pacific margin and the Tethys. Conceivably, until the mid-Cretaceous tectonic phase, the continent was oppositely tilted away from the then active peri-Atlantic and peri-Arctic belts, with most of terrestrial runoff directed to the Pacific and Tethys. Possible climatic consequences of such reversals are discussed in the next chapters.

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Chapter 7. Climate change

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VII. CLIMATE CHANGE Climate basically reflects a heating of atmosphere by solar radiation reaching the earth. The amount of heat captured by the atmosphere varies with solar flux and albedo, while distribution of heat over the globe is controlled by a variety of orbital and geographic factors, including sea level, relief, vegetation, etc., with feedbacks. A climatic theory is an explanation of differential heating and its change over time, verified by the records of various climate-bound phenomena. The greenhouse effect is one of such explanations (Budyko et al., 1985, 1986; Berner, 1990, 1991,1997), its predictive value depending on its consistency with the recorded climate change. Although it is well known that climate is a system rather than a sum of quantifiable variables, the present-day climate is traditionally depicted as such. And the same is expected of palaeoclimatic reconstructions. To meet such expectations a palaeoclimatologist must find in the fossil record what is not there, for what is recorded is climate rather than climatic variables.

VII.1. Proxies Our everyday experience relates to weather. Climate is deduced rather than experienced. Likewise, only transient states of atmospheric air or water masses are recorded by whatever palaeoclimatological techniques. However, the processes reflecting a longterm climatic influence, such as vegetation changes, may prove more informative than those under a short-term climatic control, such as isotope fractionaniton. Climatic reconstructions are commonly conceived of as quantifications of past climates based on some recorded phenomena adequately reflecting the non-recorded climatic variables, serving as substitutes, or proxies, of the latter. Such climatic indicators as leaf dimensions or isotope ratios are quantifiable, but their identification with climatic variables can be misleading: none of the “proxies” is unequivocally related to one and only one climatic variable, be it temperature, precipitation or seasonality. For instance, the oxygen isotope ratios in marine carbonates of skeletal structures are widely used in palaeoclimatology as numerical values allegedly convertible into palaeotemperatures. It is well known, however, that biochemical isotope fractionation is controlled by a complex system of environmental variables, including temperature. Hence the advantages of quantification are illusionary.

VII.1.1. Lithologies The widely used sedimentological correlates of palaeoclimates are glaciogenic deposits, coal, redbeds, evaporites, lateritic palaeosols and a few other lithotypes confined

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to certain climatic zones or prevailingly so. However their climatic constraints are not as rigid as palaeoclimatology would have them. Fossil tills (tillites) of lowland glaciers are only quantitatively different from those of montane glaciers. Moreover, certain tectonic mixtites (olistostromes) are similar to tills. Furthermore, special studies are needed for distinguishing fossil tills from glaciomarine deposits, which do not indicate continental ice caps, and from reworked allotillites (a few reliable criteria are listed in Chumakov, 1990). It is also arguable whether tills signify an advance or retreat of a glacier, which in the latter case is either starved because of low precipitation or melted because of rising temperatures. A latitudinal position of Palaeozoic tillites is uncertain. Even on reassembled continents some of their records occur close to the equator (Evans et al., 1997), which can be an artifact of either confused palaeolatitudes or misinterpreted mixtites or else an evidence in favour of tropical galciation (an often meglected possibility discussed in VII.2.2). Peat accumulation is primarily a function of plant dead mass production/degradation rates. Only ombrotrophic peat bogs directly depend on atmospheric humidity, whereas the rheotrophic bogs depend on ground water table that fluctuates not only with precipitation, but also, and locally to a greater extent, with sea-level (basis of erosion) and drainage. The tectonically controlled basinal morphology is a decisive factor of peat accumulation. Thick coal measures are zonal in their geographic distribution, whereas sporadic peat accumulation occurs over a wide range of climatic conditions. Likewise, massive evaporite deposits are zonal, yet salt lakes occur over a broad latitudinal range from the tropics to beyond the polar circles. Black-shale deposition at several distinct stratigraphic levels has prompted the idea of ocean-wide anoxic events (OAE) in the stratified Strangelove oceans. The models of OAE origins involve sea-level fluctuations, climate change, biotic productivity and organic deposition rates (Schlanager & Jenkins, 1976; Tissot, 1978; Graciansky et al., 1984; Arthur et al., 1988; Wignall, 1994; Prokoph et al., 2001; Wilson & Norris, 2001; II7-4). However, the depth estimates for oceanic black shales (Van Graas et al., 1983) seem to have been greatly exaggerated. Most black shales contain at least sparse benthic fossils, such as aucellinids or inoceramids, indicating depths above the contemporary carbonate lysocline (actually, black shale lamination partly results from post-sedimentary decalcification) and a mild anoxy level. Their sedimentation rates far exceed those typical of deep-water pelagic deposits. Among the paradigmatic “OAE“, the Cenomanian/Turonian boundary shales (the Bonarelli Level) are deposited over a carbonate platform intercalating between the Scaglia Variegata and Scaglia Rossa carbonate sequences of northeastern Dolomites, Italy (Luciana & Gobianchi, 1999). They are 1 m thick, marked by an exceptionally high total organic content, a positive shift of δ13C, and by dramatic foraminiferal turnovers (Nederbragt & Fiorentino, 1999; Keller et al., 2001). Other black-shale events over the Cenomanian– Turonian differ in a much lower organic content and in oligotrophic planktic assemblages (Luciana & Gobianchi, 1999) indicating a diversity of black shale environments.


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The Cenomanian–Turonian black shale episodes over the Mediterranean might have been related to a massive input of plant material with terrestrial runoff from Africa in much the same way as the mid-Pleistocene sapropel deposition reflect high flood periods of the Nile (Rossignol-Strick et al., 1998). The coeval plant localities in shallow-water marine facies, e.g., the Cenomanian of Nammoura, Lebanon (Krassilov & Bacchia, 2000) or the Turonian of Gerofit, Israel, are evidence of terrestrial injections, from an enormously productive coastal vegetation in the latter case (Fig. 80). Shallow-water black shale deposition is ascribed either to a high biotic productivity of surface waters or to a density stratification providing for an enhanced preservation of organic matter below the picnocline (Demaison & Moore, 1980; Wignall, 1984; Wignall & Newton, 2001). Black shale facies occur either at the beginning or over the culmination of transgressive sequences (Wignall & Newton, 2001). The origins of the basal black shales (boundary clay) seem to have been related to ponding of estuaries by a rising sea and a concomitant eutrophication by nutrients coming with continental runoff (as in the Permian/Triassic Nedubrovo sequence: Krassilov et al., 1999a, 2000; IX-3). At the maximal transgression phase, the decisive factor might have been a circulation of water masses in the arising deep ocean – epeiric sea system, with an outflow of hypersaline waters from epeiric seas imposing a thermal/density stratification in the ocean. This mechanism explains a constant association of thick black shale sequences with climate warming (Fig. 81): density stratification results in a lesser CO2 uptake by the oceans that counterbalance a sink of atmospheric CO2 to massive organic deposition. The lacustrine black shales owe their origin to high biotic production rates compensated for by a rapid removal of organic matter from epilimnion to anoxic hypolimnion. These conditions preferentially occur in taphrogenic lakes of intracontinental rift zones, such as Tanganyika of the African rift system or the Cretaceous trans-Asiatic rift zone of Transbaikalia, Mongolia and northern China. The climatic control is primarily related to seasonality of precipitation – terrestrial runoff enhancing biotic production and stratification of the water column.

Fig. 80. Gerofit locality, western flank of the Dead Sea rift valley, southern Israel: view of a corbonate/shale sequence and accumulation of plant debris in the Turonian turbidite deposits, a possible source of eutrophication.

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Fig. 81. A correlation of the Cretaceous black shale events (dashed) and temperature fluctuations based on the cycadophyte index (left) for the Early Cretaceous and the entire/non-entire leaf margin ratios (right) for the Late Cretaceous (Krassilov, 1984).

Redbeds are the most ambiguous of sedimentological indicators. Their association with evaporites, black shale facies of stratified lakes, lacustrine and pedogenic carbonates, peaty marshland palaeosols and thin coals suggests a wide range of climatic conditions. An analogy with the present-day red soil of megathermal climate is misleading because redbed sequences usually contain variegate interbeds or patches, in which red colouration develops over a blue-green background. Such lithologies suggest oxidation of ferrous deposits owing to rapid switches from desoxic to oxygenized environments. Fluctuations of ground water table, reflecting seasonality of precipitation, are critical for post-sedimentary alteration, including a pedogenic translocation of iron (Wing & Bown, 1985). Pedogenesis of redbed environments typically produces calcified horizons with abundant carbonate nodules, indicating seasonally dry climates, with variegation and organicrich horizons due to episodic waterlogging. Silicitic rhizolites and authigenic clay minerals are indicative of wet/dry seasonality. However, most redbed palaeosols are either underdeveloped, with residual lamination, or reworked, with secondary concentrations of carbonate nodules (see examples of redeposition and diagenetic red colouration in Boucot & Gray, 1982). Eberth et al. (2000) infer a wide range of modern analogues of the Permian redbeds, ranging from Mediterranean landscapes (which are rather prematurely rejected on “palaeogeographic grounds”, that is in respect to certain continental reassemblies) to African savannas and Brazilian campos. They refer to waterlogged soil and the absence of caliche as evidence of a greater rainfall than in the arid to semiarid tropical climates. Biological climate indicators associating with redbeds are mostly equivocal. Both horizontal and penetrating root systems may occur, and the leaf morphologies may appear xeromorphic, but on a closer inspection relate to waterlogging (as in the Autunian callipterids: Busche et al., 1978). Wood anatomy pertaining to emboly protection is a guide to a fluctuant water supply, as are the false growth rings (Ash & Creber, 1992), yet different climatic variables might have been involved.


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Growth increment (sclerite) patterns on fish scales are of a potential significance for palaeoclimatology as a source of evidence on changing growth rates related to seasonal trophic conditions (Krassilov, 1983b). Fish scales from the Mid-Cretaceous (Albian) of Crimea within the redbed domain show a regular sclerite spacing interrupted by narrow zones of constricted sclerites, but lacking distinct annual rings. Such pattern indicates only a brief period of retarded growth related to seasonal draughts (Fig. 82). In spite of their apparent utility for palaeoclimatology, the climatic constraints of laterite pedogenesis, caliche, bauxites and related products of biochemical weathering, rather than remaining constant over times, change with atmospheric CO2 levels, relief, geological substrates and vegetation. Most redbeds are volcanomictic molassoids confined to post-orogenic phases of tectonomagmatic evolution (in particular, of the Hercynian and Laramid orogenies). The direct sedimentological effects of tectonomagmatic activation are topographic, creating new basinal morphologies, as well as geochemical, providing a source of metallic enrichment (in particular, by iron from mafic minerals). The indirect effects are owing to the associated sea level and climatic events (II.7.2). Acidification of surface waters and rhizosphere by volcanic substrates enhances biochemical weathering promoting redbeds and their related pedogenic lithotypes far beyond their present-day climatic ranges. With peneplanation of volcanic plateaux, the Cretaceous to Eocene bauxite deposition, traditionally taken as evidence of a tropical climate, spread into the temperate vegetation zone over northern Kazakhstan and West

Fig. 82. Growth increments (sclerites) on a fish scale from the Albian of northern Crimea as a climatic indicator (Krassilov, 1983a): the absence of annual rings indicates a round-year growth, with broad zones of regular sclerites attesting to constant growth rates. The intervening zones of about 5-6 narrow irregular sclerites (arrow) reflect a retarded growth over the relatively brief (four – seven weeks) periods of adverse conditions, such as trophic lows during dry seasons.

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Siberian Plains, hence an apparent controversy between sedimentological and vegetational indicators. To summarize, the redbeds are like coals in including various facial types and in being controlled by a variety of factors, both climatic and tectonic. Since a climatic component is always present as a factor of both primary and post-sedimentary origins of redbed deposits, these are of a climatological significance, though in a broader sense than their current interpretation implies. Redbeds mark a wide range of summer-dry climates, from mildly to sharply seasonal and of a different duration of dry season, sometimes of few weeks only. In their latitudinal distribution, the redbed and coal-bearing deposits represent alternative, though marginally overlapping, facies domains (Fig. 83). Over geological times, the redbed zones either extended from the equator to the mid-latitudes or were restricted to narrow subtropical belts. The coal zones are antithetic, mainly temperate, displaced to

Fig. 83. Distribution of Cretaceous coals, redbeds and evaporites. Coals (black squares – Early Cretaceous, open squares – Late Cretaceous): (1) West Greenland, (2) Edmonton, (3) Blairmore, (4) Wyoming, (5) Potomac, (6) Magdalena, (7) Marañón, (8) Dorothea, Patagonia, (9) Western Svalbard, (10) Southern England, (11) Aquitanian, (12) Aaachen – Gosau, (13) Srednegorye, (14) Ust-Urt, (15) Chulymo-Yenisey, (16) Khatanga, (17) Lena (Sangar), (18) Aldan, (19) Zyryanka, (20) Arkagala, (21) Omsukchan, (22) Beringovskoye, (23) Mongolian Altay, (24) Transbaikalia, (25) Bureya, (26) Western Sakhalin, (27) Zhisi, Heilongjiang, (28) South Primorye, (29) Naktong, Korea, (31) Bau, Kalimantan, (32) New Caledonia, (33) Queensland, (34) Gippsland, (35) New Zealand, (36) Colville, (37) Chignik, (38) Yukon, (39) Niger – Benue, (40) Cuanza. Evaporites (black triangles – Early Cretaceous, open triangles – Late Cretaceous): (1) Guatemala, (2) Paranaiba, (3) Sergipe – Alagóas, (4) São Paulo, (5) Gabon, (6) Congo, (7) Cuanza, (8) Gulf of Thailand. Dashed lines – boundaries of the redbed zone.


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high latitudes with expansion of redbeds (as over the Permian/Triassic boundary or in the Late Cretaceous: IV.3.1). An intermittent low-latitude paralic coal zone intervening the redbed realm, as in the terminal Cretaceous (IV.3.5), is evidence of a radical global climate change.

VII.1.2. Isotope ratios Isotopic compositions of carbon and oxygen, expressed as proportions of the heavier species, δ13C, δ18O, are interrelated, reflecting productivity and the dead mass export from terrestrial and marine ecosystems, the variables that fall under climatic control. Their relation to climate is mediated by the influence of temperature and precipitation on photosynthesis, biomass growth rates, continental runoff, organic and carbonate deposition, calcification of rhizolites and skeletal structures, etc. Oxygen isotope ratios of carbonate skeletons fluctuate with water temperature and salinity. The carbonate δ18O tends to decrease with deglaciation and the influx of meltwater, hence a correlation with deep water (in particular, NADW) production and the Heinrich events (e.g., in the Younger Dryas: Smith et al., 1997). Terrestrial biomass preferentially contributes the heavier oxygen and consumes the lighter carbon, thereby inflicting a parallel evolution of δ18O and δ13C. The lighter carbon is locked in the standing crop biomass, humus and the buried dead mass of organic-rich deposits. It is returned to the atmosphere with plant/soil respiration, biomass burning, and the outgassing of buried organic matter. The ratios of terrestrial to marine production are reflected in δ13C fluctuations owing to their different 12C consumption rates (Jasper & Gagosian., 1989). Their relative contributions depend, among other factors, on sea-level fluctuations that are in these way reflected in the δ13C records. Anaerobic metabolism of methane-consuming Archaea is an additional sink for the lighter isotope potentially leaving a signature in the δ13C records (Kuypers et al., 2001; Orphan et al., 2001). However, a recent tendency to ascribe all isotopic anomalies to this factor (e.g., Prokoph et al., 2001) does not seem warranted by the evidence, because a much higher 13C/12C ratio has been found in other chemoautotrophic organisms of the anaerobic ecosystems (Orphan et al., 2001) and the net effect might have been not so prominent. Since the globally recognized level about 7.3 Ma, but probably even earlier, the δ13C is also affected by the changing ratio of C3:C4 photosynthesis that depends on the atmospheric CO2 concentration and climate. In the shallow-water carbonate precipitation systems, isotope fractionation is in equilibrium with that of the organic carbon reservoir. A correlation with terrestrial productivity is confirmed by δ13Ccarb fluctuations with precession cycles (their induced precipitation cycles) and, on a minor scale, with the monsoon index – rainforest production during the El Niño cycles (Siegenthaler, 1990; Siegenthaler & Sermiento, 1993).

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The episodes of massive black shale deposition (the “oceanic anoxic events”) in the Toarcian, Aptian, Albian/Cenomanian and Cenomanian/Turonian are marked by the positive δ13C spikes, which are considered to be a response (with an appreciable time lag) of inorganic carbon reservoirs to a high-rate burial of 13C-depleted organic carbon (Arthur et al., 1988; Prokoph et al., 2001; Wilson & Norris, 2001). These isotopic events typically correlate with transgressions, global warming and extinction of benthic organisms intolerant of low oxygen levels. Inflow of warm saline waters from epeiric seas enhances thermohaline stratification in oceanic depressions and the resulting spread of anoxic conditions (Thierstein & Berger, 1978). The major biotic turnovers, such as at the Permian/Triassic, Cretaceous/Tertiary and the Palaeocene/Eocene boundaries (IX.1-2), are marked by short-term negative δ13C excursions probably reflecting a drop of biotic productivity with a disruption of the oceanic – epeiric circulation systems. Later on, with a spread of highly productive pioneer vegetation over the emerging land and an increase of organic matter in the terrestrial runoff, the anoxic facies reappear accompanied by a rise of δ13C.

VII.1.3. Fossil plants Fossil plant record as a source of palaeoclimatic inference has much the same limitations as the sedimentological and isotope records. Plant life is controlled not by independently quantifiable climatic variables but by their concerted impacts. The temperature tolerances depend on precipitations, and vice versa. Plants are thought to be the most climatically sensitive components of biotic communities, but this may not necessarily be so, e.g., for tree lines primarily controlled by the mycorrhizal tolerances. Plants are linked to climate not only through their tolerances per se (of which we scarcely have sufficient evidence even for extant plants), but also, and perhaps to a larger extent, through their symbiotic organisms, such as mycorrhizal fungi, insect pollinators, vertebrate dispersers, etc. (Bonan & Shugart, 1989; Van der Heiden et al., 1998; Christian, 2001, examples in VIII.1). In other words, what we actually see in the fossil plant record is a vegetation component of co-adapted response to climate of a biotic community as a whole. Plant characters habitually considered as climate correlates might have initially developed in respect to a different kind of causation. Since plants were impacted by animals from the very beginning of terrestrial plant life (VIII.3.1), most plant adaptations are plant-animal co-adaptations. As discussed above (III.1.3), the leaf margin morphologies are related to plant – insect interactions and only secondarily to climate. Flowers and fruits are more conspicuous on leafless twigs (Fig. 84). A shut-down of photosynthesis in respect to pollination and dispersal occurs over a wide range of climates suggesting this to be the primary function of deciduousness, overshadowed by a derived adaptive response to seasonality of water supply.


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Fig. 84. An original meaning of deciduousness: leafless branches of Erythrina humeana with blossoms showy against the green background. Haifa, Israel.

Evergreenness, prematurely taken for granted as a primary condition in terrestrial plants, might have been more directly related to adverse climatic conditions imposing metabolic shut-down, but requiring a rapid switch to photosynthesis, as in tundra vegetation. A high-latitude origin of evergreen plants in relation to short photoperiod would explain their typically scleromorphic habits advantageous in the nutritionally poor habitats of both tropical and boreal zones. This example calls for awareness of limitations imposed by functional changes. Actualistic approach in palaeoclimatology assumes that (1) climatic tolerances of extinct species correspond to those of their extant allies; (2) morphological expression of climatic adaptations are equivalent in the present-day and extinct life forms, and (3) vegetation aspects were defined by climate in much the same way in the past as in the present. Such assumptions are admissible, with reservations, only as null hypotheses of evolutionary palaeoclimatology. Taxonomic inference. The most straightforward actualistic inference (sometimes called “the nearest relative method”) is extrapolation of climatic tolerances from extant species to their phylogenetically related extinct species. When we refer to fossil palms as indicators of megathermal (tropical) climate, we rely on constancy of climatic tolerances, not so much of a taxon as of a life form with vegetative apices sensitive to frost. Yet a few extratropical forms, such as Chamaerops humilis, introduce an element of

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uncertainty in palaeoclimatic inferences based on high latitude records, such as from the Eocene of western Kamchatka (Budantsev, 1983). Evergreen and deciduous species may have either distinctive or similar leaf morphologies, as in Quercus and Nothofagus among the Fagaceae, respectively. Even the upright and prostrate habits, representing quite different climatic environments, may not correlate with leaf morphology (Hill, 1990). It should be kept in mind, that few plant species, if any, explore the full ranges of their climatic tolerances. More often the plant ranges are restricted, even displaced away from a potential climatic optimum, by competitors, pests, etc. Moreover, the palaeoclimatic inferences are often based on relict species of formerly widespread groups. The survivors may represent marginal environments rather than those typical for the group. It would be risky to assume, for example, that, while the only living population of Glyptostrobus pensilis occurs in the subtropical to tropical zones of southeastern China, the climate should have been like this over the vast Arcto-Tertiary range of the genus. In the Palaeogene assemblages, such as those of Tsagajan in Amurland (Krassilov, 1976a), Glyptostrobus frequently associates with Metasequoia, another living fossil that also survived in a single restricted area, but of a quite different climate, with mean annual temperatures about 13°C, enduring 8°C below zero (Fu, 1992). This may be closer to the climate of their joint occurrences over the Tertiary (Krassilov, 1976a). Both plants are pollinated in February or early March – a reminder of their co-adapted reproductive ecologies. Their surviving species might have been quite recently separated in respect to their different flood tolerances. Seed germination is the most temperature sensitive phase of life cycle for a number of relict species allowing for a reconstruction of temperature conditions over their former ranges. However, a survival of relict species primarily depends on sustainability of small populations that, independently of intrinsic climatic tolerances, do better in equable than fluctuant climates. Because of this, equable climates provide refugia for species of a formerly larger population densiy and broader climatic tolerance, as in the case of ArctoTertiary relicts, such as Glyptostrobus pensilis, surviving near the southern margin of their former ranges. Consequently, palaeoclimatic inferences based on relict occurrences are systematically skewed in the direction of a more equable climate. Morphological inference. A justly expected correlation of plant morphology with climate is complicated by inertia of evolutionary conservative characters formed under different climatic conditions or only inderectly related to former climates. In addition, while climax vegetation is climatically equilibrated, the seral/interzonal communities are less so, and just such communities tend to be overrepresented in the fossil record. This consideration pertains to extinct – extant extrapolations over a wide range of morphological characters. Leaf characters. Leaf morphology is related to optimal photosynthetic functioning that is affected by light, temperature, humidity, atmospheric CO2, wind velocities, etc. This factors impose a widespread leaf convergence, e.g., of xeromorphic leaves or those


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with entire/non-entire leaf margins, providing a major source of palaeoclimatic inference However, leaf similarities might have only been indirectly related to climate, as in the case of leaf mimicry owing to plant/insect interactions (III.1.3). Generally, leaf characters reflect adaptive strategies affected by climate rather than climate as such. Actually the causal link is not climate – leaves, but climate – vegetation structure – leaves, with additional complications introduced by fungal and animal symbionts that make it climate – biotic community – vegetation structure – leaves. These considerations pertain to climatic inferences based on leaf dimensions and shapes alike. The size classes of micro-, noto-, meso- and megaphylls, and the leaf indices of microthemal to megathermal climates based on their ratios have been worked out (Raunkiaer, 1934; Webb, 1959) for the present-day vegetation types, such as multistratal canopies of tropical rainforest. Since leaf size is actually related to shade tolerance, wind velocity and CO2 concentration on the blade surface, the leaf size/climate correlations hold for certain types of canopy structures under constant CO2 levels. An indiscriminate application of leaf indices can be misleading (Dolph & Dilcher, 1980). They are of little, if any, meaning for the pre-Pleistocene vegetation types, e.g., the rainforests of a simpler canopy structure. Correlation between leaf area and precipitation is likewise affected by the canopy structure, with evolutionary changes over times (see Wilf et al., 1998). CO2 concentration on leaf surface varies with leaf area and wind velocities (Ågren, 1985) and is affected by the leaf roughness and micromorphological structures (the costal/intercostal differentiation, stomatal pits, trichomes, striation) related to the boundary layer decoupling leaf surface from the ambient air (III.1.1). Since wind velocities are greater over open vegetation than under canopy, this factor might have contributed to morphological divergence of leaf shapes (dimensions) in respect to vegetation structure. Under similar climatic conditions, wind roughness will be better expressed morphologically in riparian woodlands than under the canopy of old-growth forests. The alternative segmentation/fusion tendencies in leaf evolution might have reflected either climate change or fluctuations of atmospheric CO2 levels or evolution of vegetation structures. In particular, marginal fusion of leaflets, parallel in the initially compound leaves of Permian gigantopterids, Triassic peltasperms or Cretaceous platanoids, resulting in a broad entire blade, is correlated with species richness and, by inference, with an increasing complexity of contemporary vegetation structures. Both leaf venation and marginal characters convey an indirect response to seasonality. In secondary veins of pinnate venation, the angle of departure increases with compactness of leaf folding, which is inversely related to unfolding rates that fall under climatic control (tightly folded leaf primordia take longer to expand, hence correlation with a long growing season: Kobayashi et al., 1998). Yet under similar climatic conditions the unfolding rates vary with edaphic conditions, seral status, pollination strategy, etc. Rapid leafing is advantageous in temperate climates of a short growing season, but may also be disadvantageous in deciduous plants that flower before leaves.

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A higher proportion of non-entire leaf margin morphologies in temperate vegetation relative to tropical types reflects an arrested marginal expansion owing to a shorter growing season. At the same time, leaf margin morphologies play a role in the plantinsect interactions by either attracting pollinators or repelling folivores. In deciduous plants, a visual defence compensates for the relatively weakly developed biochemical defence, hence serrate leaves. On the other hand, entomophily is more common in tropical plants that have to encourage insect visits, hence entire margins. These leaf margin morphologies thus correlate with climate, but in an indirect way, via the climate/defence mode and the climate/pollination mode correlations. Some modern leaf/climate correlations are valid for the pre-Cretaceous non-angiospermous plants as well. Thus, an association of narrowly pinnate compound leaves with the Mediterranean-type climates holds for both the present-day angiosperms and the Mesozoic bennettites of Ptilophyllum, Zamites and Otozamites morphotypes. Cycadophyte contribution to the total floristic diversity (the cycadophyte index) tends to decrease with latitude, partly owing to a loss of pinnate morphotypes. A significance of cycadophyte index as a proxy of climatic evolution is revealed by its parallel changes over the timeseries of Meso-Mediterranean (Primorye) and Arcto-Mesozoic (Bureya) assemblages, with the climatic optima in the Berriassian and Aptian confimed by independent evidence (Krassilov, 1973b; Fig. 85). On the other hand, an adaptive meaning of the non-entire leaf margin morphology might have been different in angiosperms and cycadophytes. For instance, the serratoleaved Nilssonia, an extinct cycadophyte, occurs preferentially in thermophilic (Kamenka), rather than temperate (Bureya), assemblages (III.3.3). Serrate leaves are virtually lacking in

Fig. 85. Cycadophyte index for two basinal taphofloras of different climatic zones indicating parallel and synchronous climatic changes. (Cal) Callovian, (O-V) Oxfordian-Volgian, (Ber) Berriasian, (Val) Valanginian, (H-Bar) Hauterivian-Barremian, (Ap) Aptian, (Ab) – Albian.


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Palaeozoic plants, first appearing in the latest Permian gigantopterids (Gigantonoclea). Over the Mesozoic, serration appeared in all the macrophyllous groups, including ferns (Cladophlebis denticulata), cycadophytes (Nilssonia serrotina), and bennettites (Neozamites). In the early angiosperms, marginal teeth are typically glandular. Double serration in its incipient form of irregular tooth shapes and spacing first appeared in the Cenomanian (Krassilov & Bacchia, 2000), Even at this stage marginal galls seem more frequent on entire than on serrate leaves (Fig. 86), suggesting a role of margin morphologies in the leaf defence. The regularly biserrate morphotypes Corylites, Viburniphyllum, Tiliaephyllum, etc. became prominent towards the end of Cretaceous, their frequencies increasing with both latitude and altitude, indicating a priority of climatic control (IV.2.2). It follows from the foregoing discussion that leaf margin morphologies are linked to climate by two kinds of causation: directly, through developmental acceleration in seasonal climate, and indirectly, through defensive mimicry related to pollination strategies which are in turn related to climate. Their respective contributions vary from one vegetation type to another. A meaningful comparison can only be made between fossil plant assemblages belonging to one and the same vegetation type.

a

b

Fig. 86. Traces of insect damage on leaves from the Cenomanian of Nammoura, Lebanon (Krassilov & Baccia, 2000): (a) traumatic resin ducts and a transverse mine on Pseudotorellia, (b) marginal gall on a leaflet of Parvileguminophyllum.

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This consideration was taken into account in my studies of leaf-margin ratios in the Cretaceous assemblages of Sakhalin (Krassilov, 1975a) that came from successive horizons of the facially uniform deltaic to shallow marine deposits. Counting leaf specimens, rather than species eliminated a potential taxonomic bias. The results indicate a steady rise in the entire/non-entire leaf ratio from the Coniacian to Campanian, followed by a gradual decrease in the late Maastrichtian and over the Cretaceous/Tertiary boundary (Krassilov, 1975a; Fig. 87). In a similar analysis of the mid-Cretaceous plant assemblages from deltaic to shallow-water carbonate facies of Crimea, a low of the entire/non-entire ratios falls at the Albian/Cenomanian boundary, with a rapid rise over the Cenomanian (Krassilov, 1984). As in the case of cycadophyte index (above), a climatic significance of the leaf margin ratios was revealed by their parallel changes over time-series of successive plant assemblages in two distinct vegetation realms. An elaborate multi-character method of leaf analysis has been introduced by Wolfe (1992) and extended back in time by Spicer et al. (1996). The method allegedly provides for a much more detailed climatic inference than any single-character approach. However, the multiple leaf character/climate correlations are obtained for dominant zonal types of extant vegetation. Edaphic and seral communities, overrepresented in the fossil record, tend to develop leaf morphologies different from the bulk of regional vegetation. Moreover, unlike the extant forms used for comparison, the early angiosperms scarcely were the canopy trees. Climatic significance of their leaf characters might have changed with their status in the contemporary vegetation structures. Leaf characters may persist despite of climate change due to their preadaptational value, as in the case of transition from sclerophylly to xerophylly (III.3.1). Since leaves

Fig. 87. Leaf margin as a climate indicator: ratios (%) of entire-margined leaves in consecutive Cretaceous assemblages, western Sakhalin. (Cn) Coniacian, (St) Santonian, (Cmp) Campanian, (Mst) Maastrichtian, (Dt) Early Palaeocene (Danian).


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are evolutionary conservative organs, the more leaf characters are indiscriminately used as past climate indicators, the greater a potential there is for an error to occur as a result of such misleading similarities. This consideration casts some doubt on the alleged advantage of using many poorly understood characters instead of few better known ones. Reproductive characters. Some reproductive features, such as volume-regulating (harmomegathy) structures in pollen grains, are commonly interpreted as climatic adaptations. However, an importance of harmomegathy primarily depends on pollination mode, with air-borne pollen grains normally exposed to desiccation for a shorter time than those carried by insects. Generally, over the evolutionary history of reproductive structures, the climatic effects are indirect, related to reproductive strategies. Thus, wind pollination prevails in temperate climate because of relatively low biological diversity, high population densities and the seasonal leaf shedding (canopies in full leaf screen the pollen rain). Entomophily is more frequent in tropical climates for the opposite reasons. Fossil evidence of pollination ecology is an important, although as yet little explored, source of palaeoclimatological inference. For example, the giant male strobili of Cycandra, a Jurassic cycadophyte (Fig. 88), produced thick-walled sporangia with narrow apical beaks, from which pollen grains, several hundreds per sporangium, were released in small portions. In contrast to instantaneous dispersal of wind-born pollen, this peculiar mechanism of parsimonious exposure offered a persistent pollen source for diachronously developing ovulate cones implying an extended pollination period that in turn implies a weak seasonality or none at all, with pollinators available round the year (Krassilov, 2000a). More than one pollen type in gut compressions of fossil insects is evidence of overlapping pollination periods in the forage species. Such mixed pollen loads are rather typical of the mid-Permian insects of Cisuralia (Krassilov & Rasnitsyn, 1997; Krassilov et al., 1999d). The pollen types of their gut assemblages are also abundant in contemporaneous dispersed assemblages indicating the dominant pollen producers. This type plant – insect interactions indicates seasonality, with synchronous pollination probably confined to the drier season. In contrast, all hitherto studied Early Cretaceous xyelids from Baisa, Transbaikalia, had single-species pollen loads, with pollen types produced by relatively rare plant species. Thus the feeding habits were highly selective and probably adapted to diachronous pollen sources, as in the case of Ceroxyella dolichocera that fed either on saccate pollen produced by proangiospermous Preflosella or on cryptosaccate pollen of an unknown relative of the latter (VIII.3.3). Since xyelids are early spring insects, the succession of flowering should have been rapid as in ephemeroids of temperate climates. Like fossil pollen, fossil disseminules are as yet a poorly explored source of palaeoclimatic inference. The seed/fruit size is a function of (1) growth rates, with larger disseminules facilitating a faster seedling growth; (2) seed dormancy period correlated with small endosperm and thick coats; (3) seral stage, with disseminules typically smaller in pioneer species and colonizers (Wilson, 1993), (4) dispersal mode, in particular, the body

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Fig. 88. Cycandra, a giant male cone, producing thick-walled sporangia with apical beaks (bottom left), through which the pollen grains were released in small portions providing a persistent pollen source. The Middle Jurassic of Georgia, Caucasus (Krassilov et al., 1996a).


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size of dispersers (Mack, 1993). Climate pertains to all these factors, but the effects are modified by vegetation structure, as well as by plant – animal interactions. Thus, a preponderance of small disseminules may indicate a spread of pioneer species or extermination of larger dispersers, with a climate change at the background. The Devonian preovules were relatively small, indicating pioneer habits. The larger Palaeozoic seeds (Lagenostoma) appeared with a build-up of morphological diversity in respect to a longterm climatic stability. Small Mesozoic disseminules under 1 mm first appeared in association with polyspermic cupules, endozoochrous in Caytonia (found in coprolites) and probably also in Leptostrobus. Some of the latter type cupules were produced by Czekanowskia, a pioneer of temperate peatlands. Small seeds, often pappate as in Problematospermum, Basia or Typhaera, further developed in relation to wind-dispersal in proangiosperm-dominated reed-like wetlands of the mesotemperate/xerothermic ecotone, an unstable vegetation type disturbed by volcanic activity as well as periodic dryness (Krassilov, 1997a, IV.3.2). Seed/fruit shapes are primarily a function of dispersal mode indirectly affected by climate. Thus, eliasomes occur in myrmecochorous disseminules of all climatic zones, but are more common in the ground cover of temperate broadleaved forests. Tufts of hairs (pappuses) may serve for either exozoochorous or anemochorous (parachute) dispersal, both characteristic of open vegetation. The parachute mechanism requires dry air (for the hairs are hygroscopic), moist soil for adhesion or open water surface for flotation. It is most efficient in the helophytic communities of a seasonally dry climate. On the other hand, fleshy endozoochorous disseminules more commonly occur in humid tropical to temperate climates. Their first appearances, simultaneous in Caytonia, Ginkgo and cycads over the mid-Jurassic might have signalled endozoochorous dispersal by early birds or mammals. Wings are multifunctional structures serving for wind dispersal as well as for anchoring (by screwing or clinging to soil or snow). The morphology of samaras relates to their flight performance, e.g. gliding or rotation, the rates of descent and, correspondingly, the flight distance (Azama & Yasuda. 1989). One or another function prevails in respect to vegetation structure, seasonality, snow cover, etc. Wing-like appendages also serve for rolling by wind or, in areas of abundant seasonal rainfall, by torrentious flows. Because of their multifunctionality, winged disseminules occur in various types of vegetation and climates, but are more frequent in those combining their different functions, such as the cool-temperate birch-conifer woodlands (wind/snow cover) or the Mediterranean-type woodlands (wind/torrential flows). Wind-dispersal is inefficient in canopy trees, but efficient, due to a much higher wind velocity above the canopy, in emergent trees, such as the dominant rainforest dipterocarps (Osada et al., 2001). The following events over the evolutionary history of seed/fruit dispersal might have been (inderectly) related to climate change: In the mid-Devonian pre-ovules with free integumental lobes, the integumental and cupular appendages might have served for protection of pollen chambers. In the derived

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forms, apical appendages were modified as pollen-trapping structures (Gnetopsis), whereas in the Late Permian Lazarospermum the integumental lobes were more typically glochidial (Krassilov, 1999b), suggesting an exozoochorous dispersal, conceivably gaining in importance with a spread of open xerophytic vegetation and the fur/feather-clad theriodonts. The parachute dispersal by an apical pappus, another modification of integumental lobes (hair), as in Problematospermum, appeared with the Mesozoic cycadophyte shrublands, an open xerophytic to helophytic vegetation type, and became widespread in proangiospermous marshland assemblages (Krassilov, 1997a). A smilarity of pappate disseminules in proangiosperms and the much later appearing marshland monocots is an obvious case of convergent homeomorphy. This type disseminules were simultaneously acquired by arboreal forms of riparian vegetation, such as Platanus, in which they are derived (Maslova, 2001), in relation to a climate/vegetation change over the Cretaceous/ Tertiary boundary. The bilaterally winged samaras might serve for flotation or sticking to waterlogged soil in the Carboniferous arboreal wetlands. They are found in similar facies over the Mesozoic (germinate Samaropsis-type seeds are abundant in the silty flood-plain deposits: III.1.2, Fig. 56). This function is retained in some modern conifers (Sequoia), but is modified for wind-dispersal over sparsely vegetated drylands (Welwitchia) or over snow cover, as in the cool-temperate birch woodlands. Of the screw-type samaras, those with one-sided wing first appeared in the Late Permian temperate (Angarian) peatland assemblages (Sylvella). They became abundant in the Mesozoic bog-forest assemblages of the same climatic realm (Pityospermum) and were then inherited by their derived boreal conifer forests (IV.2.3). In contrast, the propeller-like four-winged samaras with bracteate (bracteolate) wings first appeared in Dinophyton, a xeromorphic proto-gnetalean plant of redbed facies realm (Ash, 1970; Krassilov & Ash, 1988; Fig. 19). They presently occur in dry arboreal woodlands and, in a derived form, in emergent trees of tropical rainforests. Both the geography and chronology of dispersal mechanisms suggest a considerable climatic effect. The above-mentioned innovations took place at the time of prominent global climate change in the Pennsylvanian (winged seeds), latest Permian (glochidia), Late Triassic (propeller-type samaras) and mid-Jurassic (pappate seeds, fleshy seeds/ cupules). These chronological levels are conspicuous also over the palynomorphological evolution, marked by the changing ratios of catasulcate/anasulcate, taeniate/non-taeniate and saccate/asaccate morphotypes (VIII.7). Vegetation types. Of the present-day vegetation types, the equatorial rainforests occur over a humid area with two annual peaks of rainfall. The adjacent summer-wet tropical areas with a pronounced dry period are covered by savannas, scrub and patches of dry evergreen to deciduous forest. The laurophyllous evergreen forests/woodlands of southern Europe and northern Africa thrive under warm-temperate to subtropical climate with a distinct winter precipitation


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maximum. With a shift to a brief spring maximum, they give way to ephemeroid semidesert. A broad zone of summer-wet climate with mean cold month temperatures below zero is forested at mean annual precipitation above 500 mm, supporting the nemoral broadleaved deciduous (annual precipitation means 1000 to 2000 mm) to mixed semideciduous forests. The montane conifer rainforests with deciduous components develop on margins of this zone with round-year precipitation above 1000 mm, whereas the relatively dry East Siberian permafrost area is covered with deciduous larch taiga. Evergreen fir-spruce taiga dominates the boreal zone within the reach of polar atmospheric fronts (Fig. 89). Yet normally under precipitation rates below 500 mm, temperate forests are replaced by open vegetation types among which the steppe (steppoid) grasslands occupy vast areas of continental climate bordering on spring-wet semideserts in the south and on tundras in the north. These vegetation types convey climates in a more holistic way than any quantifiable variables. Their respective climate types are: (1) Broadleaved rainforest (selva) climate, frostless, equable, fairly humid, with two seasonal peaks of rainfall. (2) Savanna climate, frostless, summer-wet, with a dry period. (3) “Mediterranean” climate, frostless to slightly frosty, winter-wet, with a dry summer period. (4) Scrub to ephemeroid semidesert climate, frostless or frosty for a short period, spring-wet, dry in other seasons. (5) Steppe climate, mildly frosty, summer-wet, moderately dry (mean precipitation below 500 mm). (6) Needle-leaved rainforest climate, mildly frosty, equable, fairly humid round year. (7) Nemoral climate, mildly frosty, summer-wet, moderately humid. (8) Taiga climate, frosty with thick snow cover, affected by cold polar fronts, moderately dry over the continental interiors. (9) Tundra climate, frosty, radiation balance typically less than 20 kilocalories per sq. cm per year, moderately dry round year. Such vegetation/climate correlations are broad enough to assume their validity back in time, at least for the ice-cap type geoghraphy. The tundra climate is periglacial, the taiga climate is controlled by polar atmospheric fronts while the temperate nemoral zone is owing to occasional surges of cold polar air. The selva climate owes its origin to a high monsoon intensities related to deep-water oceanic circulation. The steppe climate is evidence of continentality. These climatotypes are relatively short-leaved, reflecting an advance of continental glaciation, while a more typical situation of a milder polar glaciation is reflected in a spread of nemoral climate, persistent over most of the Phanerozoic

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Valentin A. Krassilov. Terrestrial Palaeoecology A

B

Fig. 89. Glacial versus non-glacial climatic/vegetation zonation of Eurasia as controlled by the annual means and seasonality of precipitation: purple – equatorial two-peak precipitation rainforest zone, light green – winter-dry sparse tropical woodlands, savannah, yellow – deserts, pink – Mediterranean summer-dry xeromorphic shrubland to woodland, light pink – spring-wet xeric grassland, dark green – summer-wet deciduous and evergreen forest, blue – wet round-year temperate rainforest and tundra. (A) Present-day distribution, (B) Distribution inferred for non-glacial climate (based on the Cretaceous example), lacking the equatorial rainforest zone, but with broader temperate rainforest and Mediterranean-type xeromorphic zones.


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Fig. 90. The relative extent of the major biomes over the Carboniferous to Neogene: (AT) Arctic tundra, (DF) Deciduous forest, (RF) Rainforest, (TF) Boreal forest (taiga), (XS) Xeric shrublands to grasslands.

history (Fig. 90). The Mediterranean-type climate is likewise persistent, formerly supported by the vast extent of the Tethys seas. Its occasional excursions to high latitudes (as in the Early Eocene) correspond to minimal ice caps.

VII.2. Causes and effects Climate, as a complex system, comprises a number of subsystems with their own dynamics. Climate affects, and is feedback affected by circulation of oceanic water masses, carbonate deposition, burial of organic matter, weathering of geological substrates, pedogenesis, vegetation changes, etc. In each subsystem, climatic evolution with a positive feedback is directional (humidity promotes forestation that increases humidity) while negative feedbacks generate cyclicity (CO2 greenhouse enhances biotic production that reduces atmospheric CO2). Secular climates arise at the interference of directional and cyclic processes and are by necessity less predictable than any of these. Critical for the understanding of climatic cyclicity

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are the astronomic and geological factors that control either the amount of solar energy involved in climatic processes or its distribution over the globe, or both. These the solar constant, eccentricity of the earth’s orbit, obliquity of the ecliptic, wobble of the earth’s rotational axis and the spin angular velocities, the interconnected variables of restricted amplitudes and regular 104-105-year scale periodicities. Climatic effects of astronomic variables are documented for all geological epochs (II.7). However, their significance varies form one epoch to another depending on superimposed geographic variables. Climatic evolution through the Pleistocene and recent times indicates that the orbital (Milankovitch) factors trigger off atmospheric and oceanic events of a great climatic effect, with the polar ice-caps, atmospheric CO2 and other albedo factors acting as amplifiers. Because a climate change results from interaction of many driving forces, a prediction by mere extrapolation of any observed tendencies is scarcely reliable. As is usually the case with complex interactions, any single-factor model (CO2, solar input, orbital forcing, etc.) is plausible but not sufficient. Improving our understanding of multiple climate-generating systems is the only way to a meaningful climatological prediction.

VII.2.1. Orbital forcing The rotational forcing of climate change is a relatively well-studied phenomenon. Effects of orbital cycles on global climate are evident in both historic and geological records as the so-called Milankovitch cycles (E1 and E2 cycles of 100 and 400 k.y., O cycle 41 k.y., P cycle 21 k.y., including two cycles about 19 and 23 k.y.). Notably, the atmospheric CO2 levels and their correlated temperature fluctuations calculated for trapped air from Greenland and Antarctic ice cores correspond to the Milankovitch periodicities (Lorius et al., 1990). The Plio-Pleistocene sequences show evidence of 23, 41, 100 and 404 k.y. cycles that are phase-locked, respectively, to the Pand O-driven insolation cycles modulated by the E-cycles (Clemens, 1997). Recent examples of rotation-forced cyclicity are provided by the biogenic silica records form Lake Baikal, revealing the 23, 41 and 100 k.y. periodicities (Colman et al., 1995). Orbital cycles are the major regulators of the mesocyclic lithological and associated taphonomic alternations. They are recognizable, for example, in the fluctuations of planktic microfossil abundances and the ratios of planktic/benthic forms (Sprovieri et al., 1996; Conti, 1997), as well as in the Cretaceous carbonate cyclicity (Krassilov, 1985; Park et al., 1993, II-7) and magnetic susceptibility cycles, etc. The estimates of surface water temperature (high planktic δ18O), biotic production (benthic δ13C), carbonate dissolution depth, kaolinite/illite ratios, metallic enrichment and turnover of foraminiferal assemblages in the Maastrichtian concertedly show the Pdriven fluctuations of terrestrial runoff (MacLeod et al., 2001). In the fossil plant records, an alternation of different wetland types, such as fresh-water marshes and salt marshes or mangroves, reflects the precession-scale precipitation cycles (III.3.4).


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The alternation of glacial–intergalcial phases over the Pleistocene is commonly ascribed to the Milankovitch exocycles combined with glacial autocycles. Large glaciers are self-sustainable owing to their high albedo and atmospheric feedbacks. Polar ice caps increase vorticity of polar air, with effects on both tropospheric and stratospheric temperatures (and on the stratospheric ozone layer). Yet, as continental glaciers are fed by massive snowfalls that are deflected with an expansion of polar atmospheric cells, these interactions generate a succession of the ice growth → cold front advance → ice retreat cycles. Glacial autocycles produce a profound effect on mid-latitude precipitation patterns (Mayewski et al., 1993) either enhancing or reducing the precession-driven fluctuations. Droughts in California and Patagonia during the Middle Age warming phase are ascribed to reorientation of mid-latitude storm tracks in response to contraction of circumpolar vortices (Stine, 1994). Similar atmospheric phenomena are inferred back in time from vegetational changes, lake level fluctuations, coral fluorescent bands (marking periodic influx of terrestrial organic matter), etc. The Altithermal and subsequent warm phases correlate with a drier North America but a relatively humid North Africa, Middle East and Peninsular India (Kellog, 1979; Hansen et al., 1981; Klein et al., 1990). The rotation effects can be summarized as follows: (1) The obliquity-driven 41 k.y. and the eccentricity modulated 100 k.y. polar insolation cycles trigger off ice volume fluctuations (e.g. the Pleistocene glacial phases that were initiated at the minimal eccentricity/low obliquity phases: Emiliani, 1966, 1993 and elsewhere). The associated pulses of deep water production (in particular, NADW production) regulate CO2 circulation in the system of polar downwellings – tropical upwellings, with combined effects on atmospheric CO2 concentration and monsoon intensities (Saltzman & Verbitsky, 1994; Goldstein et al., 2001). (3) Among the secondary effects of obliquity cycles, the iceberg-discharge (Heinrich) events decrease sea surface temperature (SST) and suppress thermocline circulation through the input of meltwater, as reflected in the δ18O fluctuations (Raymo et al., 1998; Stauffer et al., 1998). The concurrent changes of biotic production are detected by the carbon isotopic (δ13C) ratios. The Heinrich effect is thought to have been responsible for a sharp temperature drop and the negatively correlated anomalous δ13C over the Alleröd/ Younger Dryas transition (Björck et al., 1996). On a short-term scale, δ13Ccarb fluctuatuions correlate with tropical SST over the ENSO (Cole et al., 1993) (4) Both the obliquity-driven deep-water production cycle and the 400 k.y. eccentricity-driven summer insolation cycle (Berger, 1978; Kutzbach et al., 1982; Conti, 1997; Rossignol-Strick et al., 1998) alter the ocean-to-continent pressure gradient thereby affecting monsoons, with multiple secondary effects, such as an increase in precipitation to about 12% over the Sahel and Sahara 12 – 6 k.y. BP

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(Kutzbach et al., 1996). Fluctuations of tropical rainforest biomass with monsoon intensity in their turn contribute to cyclic changes of atmospheric CO2 levels, with feedbacks (VII.2.4). (5) The precession-driven mid-latitude precipitation cycles are of multiple consequences for the climate-generating systems, including erosion and chemical weathering on land, their contribution to terrestrial runoff, surface water salinity, biotic production and desoxygenation down water column. The uptake of CO2 and release of CH4 by wetlands is enhanced at precipitation maxima over the precession cycles. These phenomena are reflected in the sedimentary (e.g. correlation of Mediterranean sapropels with high floods of the Nile: Rossignol-Strick, 1998 and elsewhere), as well as isotopic and magnetic susceptibility records. (6) Trade wind velocities vary with the angular momentum of atmospheric circulation (Lau et al., 1989) reflecting fluctuations of the earth’s rotation rates. Their effects on SST (El Niños) and monsoon intensities are discussed below. Relative contribution of orbital components varies over time indicating an interdependence of orbital and geographic forcing. Thus, the δ13C cycles over the Pliocene – Early Pleistocene reflect the 41 k.y. obliquity fluctuations that switch, with an increase in ice volume about 0.9 Ma, to the 100 k.y. eccentricity-driven periodicity (Raymo et al., 1990) which is further modified by the millennial Dansgaard – Oeschger cycles since 0.25 Ma. In general terms, non-linearity of climatic response to orbital fluctuations (Elkibbi & Rial, 2001) is evidence of internal instability of climatic systems.

VII.2.2. Orogenic effects Net orogenic effect comprises an orographic rain-shadow effect, the albedo effect of rocky slopes and nival crests, a greenhouse effect of increased CO2 consumption rates by weathering of silicate minerals, as well as the derived effects of stratospheric cooling, ozone depletion, etc. (Raymo et al., 1988, 1998; Manabe & Broccoli, 1990; Carslaw et al., 1998). According to Riebe et al. (2001), the tectonic (orogenic) control of weathering rates prevails over climate in a long run. Burial of organic debris in the coal-bearing molassoids of orogenic foredeeps is a major sink of carbon cycle depressing the atmospheric pCO2 (as postulated for the Taconic orogeny – the Late Ordovician glaciation: Kump et al., 1995). The rain-shadow effect is recognizable over times by its sedimentological and vegetational consequences. With Sikhote Alin and Okhotsko -Chokotskian ranges rising in the Albian and over the Late Cretaceous, coal accumulation ceased in the interior sedimentary basins of northern Asia (Krassilov, 1985, 1992b). The climatic affect was enhanced by igneous activity. A short-time volcanic effect, such as “the year without a summer” after a massive Tambora eruption in 1816, is primarily a result of aerosoles


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with a residence time in the atmosphere of about two – three years. Explosive ignimbrite volcanism at the end of Cretaceous might contribute to the climate change over the critical Cretaceous – Tertiary boundary (IX.2). In this and other examples (including the explosive volcanism over the Kamchatka – Aleutian chain at the beginning of Late Pliocene glaciation: Prueher & Rea, 2001), the long-term volcanic effects are scarcely separable from orogenic effects. Likewise, the intracontinental topographic highs impose a large-scale climatic/vegetational differentiation, e.g., over West Siberian Plains and East Siberian Highlands supporting different types of boreal conifer forests, the “dark” evergreen spruce-fir taiga and the “light” deciduous larch taiga respectively. The vegetational differences are due to a combined effect of relief, permafrost and a persistent anticyclonic gyre over the Eastern Highlands. The broadleaved forests are typically developed west of the Urals, within the reach of Atlantic cyclones. The Uralian boundary has been in evidence since the Permian, dividing the Angara-type broadleaved deciduous vegetation from the European-type xeromorphic evergreen vegetation, with a broad Eurangarian (“Subangarian”) ecotone over the Cisuralia and the Volga Basin (IV.3). In the Cretaceous, this submeridional boundary separated the palynological Normapolles and Aquilapollenites provinces (Vakhrameev et al., 1978), with redbeds in the former and thick coals in the latter. Mountain ranges on the way of tropical air masses restrict the atmospheric export of heat and water vapour out of the equatorial zone thus increasing the meridional gradients of both temperature and precipitation. A rapid cooling of ascending tropical air brings abundant rainfalls to the windward slopes and adjacent lowlands. On the lee side, the vapour-starved descending air masses absorb evapotranspiration from vegetation and soil imposing sharp climatic contrasts (Fig. 91). At present, the equatorial nival patches at about 5000 m above sea level are insignificant in comparison with the Arctic – Antarctic nival areas and are scarcely perceived as an incipient equatorial glaciation. However, with the snow line descending to 3000 m, about one half of Africa between the equator and the tropic of Capricorn would have been permanently covered with snow caps. Such extensive nival areas on the way of tropical air masses would inflict a hemispheric cooling entailing a sequence of climatic events with unipolar or bipolar glaciation at the end. Remarkably, a rise of the African plateaux in the late Middle Eocene about 45 Ma coincides with the onset of a planetary cooling trend. The Late Proterozoic tillites of Australia and the Transvaal Supergroup of South Africa occur at equatorial (5°S) palaeolatitudes (Evans et al., 1997). The Late Ordovician glaciation at least briefly overlapped the palaeoequator (Kump et al., 1996). The Lower Permian tillite records extend into the tropical zone in both the present-day and reassembled Gondwana continents. At least in the latter case, a montane glaciation appears more probable than a spread of polar ice-cap debris which is incompatible with the coeval broadleaved Gangamopteris-Glossopteris lowland vegetaion (V.9.4). Williams et al. (1998) have argued that, with the earth’s obliquity greater than 54° (against 23,5° at present), the equator would have been the coldest place on the planet.

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Fig. 91. Global cooling effect of low-latitude orogeny through a decrease in the meridional heat transfer and an increase in the albedo.

However, with such a large-scale increase in obliquity, the latitudinal climatic zonation would have been smeared, which is not the case, at least not for the Permian (IV.3). Contrary to a widely held opinion, the montane equatorial glaciations, rather than being enforced by global cooling, appear and expand in advance of the latter, at the time of ice-free polar climates. This is only to be expected, because in an ice-free climate the lower troposphere would have been overheated, the stratosphere respectively cooler, the altitudinal temperature gradient steeper (Shindell et al., 1998), shifting a potential snow line downslope. Hence a low start for montane glaciers. Evidence for a leading role of montane glaciations comes from the vegetation changes that suggest a sharp downslope shift of tree line in advance of the lowland cooling.


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Since tree line is controlled by soil freezing, its low position indicates a downslope displacement of respective snow line. In the Sikhote Alin Ranges, Russian Far East, the birch woodland belt, now next to the tree line, descended to river valleys, now more than 600 m lower, before the climate-driven vegetation turnover in the lowlands at the Cretaceous/Tertiary boundary (VII.6, IX.9). The Middle Eocene (45 Ma) taphoflora of Idaho Plateau represents a sparse conifer woodland, with the tree line about 550 m lower than at present corresponding to a cooling of about -3°C (Axelrod, 1990, 1992) when the lowland vegetation still retained a paratropical aspect. Altitudinal shifts of tree lines in the Andes over the Pleistocene glacial/interglacial cycles indicate an amplitude of temperature fluctuations of about 8°C (Hooghiemstra & Ran, 1994 and a similar information for other low-latitude regions in Lin & Colinvaux, 1985; Newsome & Flenley, 1988) against the lowland estimates not exceeding 3°C. A downslope expansion of the upland conifer belts during the Pleistocene glacial maxima is palynologically recorded both in subtropical and tropical zones (Heusse & Sirocko, 1997; Taylor et al., 2001), suggesting an increase in the altitudinal temperature gradient disproportionate to the lowland cooling there. The extensive glaciations in the Vendian, Ordovician, Late Carboniferous/Early Permian and the Neogene closely followed the main orogenic pulses of the Transvaalian, Taconian, Hercynian and Alpine tectonic epochs, respectively. The turning point of Cenozoic evolution is marked by a major mid-Eocene orogenic event in the Tethys belt, with multistage overthrusts and high-rate elevation over the bulging area of Tibet–Himalayas (Burg et al., 1984; Hirn et al., 1984), and the concomitant rise of the central African plateaux. Climatic consequences were immediately felt over the vast areas of northern Asia, up to the South Urals and Kazakhstan, where the laurophyllous Eocene floras were rapidly replaced, first by the serratophyllous broadleaved and then by the mixed conifer-broadleaved Arcto-Tertiary assemblages. Such a large-scale temperization signalled an expansion of the North Pole ice-cap. Cyclic fluctuations over the temperization trend, with the successively deeper pessima in the terminal Oligocene to basal Miocene, late Middle Miocene and the Pleistocene, were correlated with orogenic pulses causing a stepwise shrinking of residual Tethys seaways. The Cenozoic glaciation was thus linked to the low-latitude orogeny. The low-latitude mountain ranges, elevated above the potential snow line, were covered with montane glaciers generating cool air masses that spread to the poles. The atmospheric heat export from the overheated tropical zone was barred by orogenic barrier enhancing the extratropical cooling trend. With its southern margin uplifted by the orogeny, Eurasia has acquired a northward tilt that brought most of its runoff to the northern North Atlantic and Arctic basins, diluting surface waters and promoting water-borne glaciers. With their growth, surges of marine ice stirred cyclonic activity bringing heavy snowfalls to northern lands, feeding the continental ice sheets. This scenario is further considered below.

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VII.2.3. Oceanic effects A net effect of sea-level fluctuations on climate combines, (1) an albedo effect arising from redistribution of sea/land areas and the related effects on atmospheric circulation, cloudiness, CO2 levels, and (2) an oceanic circulation effect of thermocline (carbonate lysocline) fluctuations and of land bridges/marine gateways. A correlation between sea level and climate is regularly detected in palaeogeographic studies but, with a great number of variables involved, the mainstream cause and effect relations are difficult to discern. In the most general form, the oceanic effects are related to divergence of continental and oceanic temperatures, most emphatic over the millennial Dansgaard–Oeschger continental ice volume – marine temperature cycles and the 4-year ENSO, with consequences for deepwater circulation, monsoon intensities, biotic production and the atmospheric CO2 levels. A climatic asymmetry and diachrony of the hemispheres seem to have been related to their different land/sea ratios (Blunier et al., 1998; Sachs et al., 2001). Oceanographic barriers and gateways. Over the Cretaceous, the elevated, intermittently emergent Walvis – Rio Grande ridge system prevented the Antarctic water masses from entering the central Atlantic. Both the anoxic sedimentary environments and endemic fauna (Fleming, 1969; Van Andel et al., 1977; Zinsmeister, 1982) indicate an effective isolation of the Austral Province south of the ridges. Insofar as teleconnection depends on deepwater circulation, such situations might result in a thermal divergence of hemispheric climates and asymmetric, even unipolar, glaciations. The Gulf Stream penetrating the Norwegian Sea contributes to the NADW and brings snowfalls to Greenland feeding the largest northern continental ice-sheet. Over the Pleistocene, the Gulf Stream – NADW circulation might have been affected by glacioeustatic sea-level fluctuations over the Faeroe–Iceland Ridge. A closer of the Isthmus of Panama would render a fairly similar climatic effect by increasing Gulf Stream (Haug & Tiedemann, 1998). A fairly popular sea-level/climate scenario ascribes climatic changes to opening of Drake Passage as a gateway for the circum-Antarctic current inflicting thermal isolation of Antarctica in the Late Eocene. Yet a northward spread of Nothofagidites in the Late Cretaceous (IV.4.1) suggests a cooling trend at least 30 m.y. before the alleged opening of Drake Passage. The better documented emergence of trans-North Atlantic land bridge in the Early to Middle Eocene might have a wider climatic effect by interrupting heat transfer between the Atlantic and Arctic oceans (VII.5), and the subsequent cooling trend commenced with a breach of the land connection in the Late Eocene Isostatic and geodetic events. In the barrier/gateway scenarios, climatic changes appear as mere coincidences of certain geographic configurations. Yet a role of barriers/ gateways in relation to the meridional heat transfer or thermal isolation can only be understood in the context of a general geodynamic model. It is scarcely a coincidence that trans-Atlantic land connection and warming in the Eocene were chronologically related to a major thrusting phase over the Arctic fold belts (Saalman & Thiedig, 2001).


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Isostatic elevation of oceanic lithosphere results in shallower oceans, with deep-water passages closed, shallow-water epeiric seaways (notably, over the West Siberian plains and North American interiors) open, and the reverse under isostatic subsidence of oceanic floor relative to continents (Fig. 92). The superimposed sea-level fluctuations driven by the geoid waves (VI.1) are antithetic in the Boreal and Tethyan realms. The boreal transgressions (presently covering the Arctic shelves) inflict cool continental climates, whereas the Tethyan transgressions typically lead to a global warming. Interference of isostatic fluctuations and geoid waves

Fig. 92. A causal scheme of rotation-driven sea-level fluctuations with effects on the temperature – CO2 system.

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would have resulted in a less straightforward correlation, with a periodic mixing of the Boreal and Tethys waters over transcontinental seaways, such as the Turgay Strait. A climatic effect of isostatic sea-level fluctuations is mediated by deep-water circulation. With a convergence of isostatically compensated oceanic and continental lithosphere (shallow oceans), the thermocline is depressed by influx of warm waters from epeiric seas while the psychrospheric circulation is further restricted by oceanographic barriers. Conversely, with their divergence (deep oceans), the psychrospheric circulation is enhanced, bringing cold deep-water masses to sea surface at the equator. Downwelling–upwelling system. Deep-water circulation is a major heat transfer agent coupled with atmospheric circulation via its effects on SST, water vapour and wind velocities. The oceanic downwelling/upwelling system is also a major regulator of biotic productivity and atmospheric CO2. Hence all factors affecting deep water production, such as the Milankovitch cycles of polar insolation (VII.2.1), their related ice volume fluctuations, Heinrich events and dilution of surface waters by meltwaters (presently reducing NADW: Mikolajewicz et al., 1997) or heat transfer by oceanic currents (in particular, the Gulf Stream, presently increased by warm water passing through the Isthmus of Panama: Haug & Tiedemann, 1998) are of profound climatic significance. Deep-water production correlates with equatorial upwellings, in turn enhancing biotic productivity of surface waters. A concomitant increase in monsoon index stimulates terrestrial productivity and contributes to fertilization of oceanic surface waters by the nutrients fluxed in with the terrestrial runoff. So a positive feedback loop is set bringing about rapid climatic changes. A decrease in deep water circulation would entail a reverse sequence of climate-generating events, with the tropical monsoons weakened, terrestrial biotic production falling, and an additional amount of CO2 released to the atmosphere. A coupling of continental ice volume and deep-water circulation is manifestly expressed over the millennial – about 1500 years – Dansggaard–Oeschger (D-O) cycles (Hansen & Takahashi, 1984; Dansgaard et al., 1993) and their bundles known as the Bond cycles about 5–10 k.y. (Bond et al., 1993) terminated by the ice surge (Heinrich) events (HE). The cycle is prominent over the fluctuations of the East Greenland ice sheet, the concomitant planktic δ18O events, and the epizodic deposition of ice-rafted debris marking the HE. Temperature fluctuations over the D-O cycles in the northern North Atlantic greatly exceed their planetary effects (Bond et al., 2001; Shindell et al., 2001). The cycle starts with a rapid warming followed by a moderate cooling and then by an abrupt cooling. NADW production is shut off, or slowed down, by a meltwater injection into the East Greenland Current to be rapidly reactivated after the HE. The D-O cool stadials correspond to the North Pacific warm phases while the East Asian monsoons are intensified during the D-O warm phases. Such antithetic correlation indicates a prevalence of oceanic teleconnection, via the NADW reduction–upwelling turnoff, over the atmospheric


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teleconnection that tends to synchronize climatic events over the Northern Hemisphere (Kiefer et al., 2001; Garidel-Thoron & Beaufort, 2001). A correlation of the terrestrial and oceanic events is confirmed by the SST fluctuations over the Medieval Warm Period – Little Ice Age transition (Keigwin, 1996). Monsoons. The monsoon intensities are of primary importance for the tropical biotic production providing a major sink for the atmospheric CO2. The controlling factors for monsoons are summer insolation (astronomically driven by precession cycles: Rossignol-Strick et al., 1998; Mommersteeg et al., 1995) and the tropical SST (a decrease in the monsoon index associates with strong El Niños, below), as well as the trade wind velocities. Monsoon fluctuations are thus linked to deep-water circulation, in particular, NADW production as a contributor to equatorial upwellings. While deep-water circulation establishes a teleconnection between the polar and tropical marine climates, the tropical monsoons couple them to the terrestrial climates. Incidentally, the world-wide effect of NADW is indicated by a negative correlation of East Asian monsoons with Dansgaard–Oeschger stadials (Garidel-Thoron & Beaufort, 2001). The millennial vegetation changes in southern California, inflicted by fluctuations of winter monsoon (Heusser & Sirocko, 1997), seem related to this coupling mechanism. Conversely, a reduced NADW production, a shoaling of oceanic heat conveyor with sea-level or a barrier for deepwater circulation would have a decoupling effect. Climatic modelling for the last glacial maximum, 21 k.y., gives the tropical SST about 2.2°C lower than at present (Weaver et al., 1998) while the continents were cooler by 3 -7°C (Behling et al., 2001). Over the warming trend, the SST might lag behind the land temperatures. Actually, a weakening of monsoons during the glacial maxima – strengthening during the early interglacials and Holocene are inferred from both sedimentological and palynological data (Faibridge, 1986; Hooghiemstra, 1988; Dupont & Hooghiemstra, 1989; Maxwell, 2001). Yet the rainforest dynamics over the glacial/interglacial cycles was affected by a combination of factors, including a displacement of trade-wind belts. Incidentally, both a drop of terrestrial productivity (recorded by the contemporary pollen rain) and a decrease in clastic runoff (recorded by a lower clay content in the contemporary marine sediments) at the glacial maxima might have been due to a diversion of moist easterlies by strong monsoons. Offshore palynology indicates a negative correlation of the African monsoon intensity and the northeast trade wind velocities over the glacial cycles (Hooghiemstra, 1988). A latitudinal shift of the jet stream is also inferred in relation to vegetation changes in the winter-wet zone (southern California: Heusser & Sirocko, 1997), where the relatively humid phases of conifer expansion correlate with glacial maxima. El Niño. Perhaps a not fully adequate but still instructive short-term model of longterm oceanic/climate change is provided by El Niño, a quasiperiodic rise of SST caused by the wind-driven surface current fluctuations in turn related to oscillations of atmospheric pressure, the “southern oscillations”, in the Pacific and Austral-Asiatic region

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(ENSO). El Niño damps tropical upwellings, thereby correlating with a low surfacewater biotic productivity to the detriments to coastal fishery. The stronger El Niño events are felt globally causing large climatic anomalies with biotic and demographic consequences, such as severe draughts in northern Africa and southern Asia. Their recent periodicities differ by a year or two from the seven-year Biblical famine cycles. The SST fluctuations are inversely correlated with the rates of CO2 dissolution and uptake by biomass production (Sarmiento & Le Quéré, 1996). On land, a weakening of tropical monsoons reduces the rainforest productivity while enhancing seasonal draughts inflicting fires that devastate large areas of tropical vegetation (Siegert et al., 2001). Altogether, the biotic effects of El Niño are large enough for depressing δ13C (Siegenthaler, 1990). The prominence of El Niño events depends on other superimposed cyclic processes, with feedbacks (Gu & Philandes, 1997). Since El Niños are causally related to the trade wind velocities driven by the angular momentum of the atmosphere (Lau et al., 1989), there should be a correlation with the earth’s rotation rates. Indeed, the stronger El Niños regularly fall at the turning point from acceleration to deceleration of the earth’s interannual rotation rates (Zheng et al., 1990).

VII.2.4. Greenhouse effect At the turn of 20th and 21st centuries, for various mostly non-scientific reasons, the international environmental policy has become a tribute to the theory of industrial climate change owing to technogenic emissions of greenhouse gases. A net increase in the greenhouse gases, C02, CH4, CFCs, NO and O3, has been estimated as about 40% over their pre-industrial levels (Jones et al., 1998). At the emission level of the 1990s, a global warming by 1.4 – 5.8°C to the end of 21st century is predicted by the Intergovernmental Panel of Climate Change (IPCC). The C02 theory of global warming has generated an extensive experimental work and vast literature (only sketchily reviewed below). The idea of C02 being a major, if not sole, regulator of global climate has been extrapolated over the long-term climatic evolution (Budyko et al., 1985; Berner, 1990 and subsequent work) rapidly developing into a palaeoclimatological paradigm. Yet the industrial warming is related not only to technogenic gas emissions, but also to deforestation, soil degradation, pollution, etc., thus making a greenhouse contribution difficult to assess. A warming tendency of the last decades is partly owing to volcanic aerosols (Mann et al., 1998) and strong El Niños. Anyway, globally, the 1980s were only about 0.5°C warmer than the 1900s, only very slightly warmer than the 1940s and considerably cooler than the pre-Little Ice Age Medieval (Kerr, 1988). “Pre-industrial level” is a confusing reference point for there were different atmospheric CO2 levels in the pre-industrial times. A level of 300 ppmv (instead of 270 ppmv


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of ice core estimates) is inferred for the beginning of Holocene on the basis of stomatal frequencies (Wagner et al., 1999). The present-day level thus does not seem highly anomalous for the post-glacial times. A level of 300-450 ppmv was sustained during considerable parts of the Tertiary (Roger et al., 2001). As for the large-scale climatic changes of remote geological past, a role of atmospheric CO2 is as yet poorly understood. The following considerations seem pertinent to further discussion of greenhouse theory: (1) The ongoing temperate changes are disproportionate to the net CO2 emissions from technogenic and natural sources. There are no grounds for suspecting them to be proportionate in the past. (2) The greenhouse gases constitute a component of the earth’s atmospheric albedo interacting with its other components (such as claudiness), the net effect being grossly different from their own contribution. (3) In the absence of technogenic emissions, CO2 fluctuations had to be inflicted, with feedbacks, by natural phenomena, such as fluctuations of tropospheric/ sea-surface temperatures, the commonly held cause and effect relations between CO2 and temperature being thus reversed (Krassilov, 1994a). (4) The present-day technogenic emission is a reallocation of organic carbon from its lithospheric reservoir (fossil fuel) to the atmosphere (greenhouse gases), with a sizeable part going to the biota (human population) in the process. This is not the only reallocation pathway. Its significance can only be assessed in relation to the naturally ongoing process of this kind (e.g., outgassing of methane hydrates), as well as the gas exchanges of atmosphere with the ocean, standing biomass and soil, the net effect of which is potentially much larger than the technogenic one. A realistic greenhouse model would thus account for interaction of greenhouse gases with other components of the atmospheric albedo, as well as for the multiple links between the atmosphere and the geospheric, hydrospheric and biotic compartments of the carbon cycle. Sources, sinks and feedbacks. Atmospheric feedbacks of elevated CO2 concentration include, first, an increase in water vapour and convectivity. These in turn enhance cloudiness, an albedo factor. Additionally, the wind-driven salt particles and aerosols contribute to the atmospheric albedo by further enhancing the vapour condensation and cloudiness, as well as by their own optical effects (Prospero et al., 1991). The same factors facilitate the growth of glaciers, a component of the land surface albedo. An increase of only 5% in the polar ice-cap albedo would compensate for the greenhouse effect of all technogenic emissions. The net effect may, therefore, be negative. The stratospheric feedbacks, though potentially considerable, are insufficiently studied. The stratosphere is cooled with a tropospheric greenhouse warming, increasing the

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rates of ozone destruction by heterogeneous reactions with chlorine on stratospheric ice clouds. An enhanced vorticity of polar air may have a similar effect (Shindell et al., 1998). The feedbacks – an increase in ultraviolet radiation inhibiting photosynthesis of oceanic phytoplankton (Neale et al., 1998) while enhancing decomposition rates of dissolved organic matter – contribute to the atmospheric CO2. The geospheric source of CO2 is the outgassing of mantle and lithosphere by magmatic, metamorphic and weathering processes. The amount of CO2 released to the atmosphere varies with pulses of tectonic and volcanic activity. A large terrestrial volcanic source, such as Etna, emits about 25×106 t CO2 yearly. Tanner et al. (2001) estimate a rise in CO2 level by about 250 ppmv in consequence of flood basalt eruptions over the Triassic/Jurassic boundary. The Permian/Triassic (Asiatic traps), Jurassic/Cretaceous (the Brazilian, Barents shelf and Central Asian igneous provinces), mid-Cretaceous (the circum-Pacific volcanic belt) and the Cretaceous/Tertiary (Deccan traps) volcanic events were of a comparable intensity. It is not known how lasting the effects were. A sink of the system is a consumption of atmospheric CO2 by chemical weathering of silicate and carbonate minerals. The Ca2+ and Mg2+ of these exothermic reactions are washed off to the oceans to become deposited as carbonate rocks binding about 100×106 t CO2 yearly. A feedback loop arises from a positive effect of climate warming and the related vegetational/pedogenic processes (Berner, 1998 and critical review in Boucot & Grey, 2001) on chemical weathering that enhances a sink of CO2 to carbonate deposition. In the geological record, a massive carbonate deposition is associated with climate warming, yet, upon the feedback scheme, it might eventually disrupt the balance of atmospheric CO2, with a switch to cooling. Since the World Ocean stores about 80 times more CO2 than the atmosphere, the oceanic source potentially prevails in regulation of CO2 levels (Van Valen, 1971; Moore & Bolin, 1986; Sarmiento et al., 1988; Covey, 1991; Siegenthaler & Sarmiento, 1993). However, oceanic contribution depends on deep-water circulation and other factors. Presently, the North Atlantic deep water (NADW) formation is considered as a major sink of atmospheric CO2 returned by equatorial upwellings circulating about 50×109 t CO2. The 14C age of dissolved organic matter in deep water masses is about 6 k.y., which has been taken as evidence of a fairly slow circulation, but the measurements might have been distorted by an admixture by ancient carbon emitted from bottom deposits. Anyway, a time lag of oceanic response many times exceeds that of standing biomass and soil. Changes in deep-water circulation might have been caused by a system of interconnected variables, such as SST, polar ice surges, carbonate dissolution depth, eutrophication, etc. Each of the variables potentially generates multiple, often counteracting, effects. A warming of surface waters decreases CO2 dissolution (a positive feedback, reinforcing the warming trend), but it also damps a release of CO2 from upwellings (a negative feedback countering the warming trend).


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By some calculations (Stocker & Schnittner, 1997), doubling of the present-day atmospheric pCO2 would halt oceanic circulation thereby decreasing CO2 uptake by the ocean by 30%. However, the interaction of oceanic circulation, SST, ice volume and atmospheric pCO2 are complicated by a number of side effects. A positive loop of elevated atmospheric CO2 → rising SST → melting of polar ice → damping of oceanic heat conveyor by meltwater → reduced CO2 uptake by the ocean (Saltzman & Verbitsky, 1994) might switch to the negative mode by an enhanced vapour transfer to continental glaciers (Oppenheimer, 1998). Land biota stores about 830×109 Gt carbon, 90% of it in the standing crop of which 60% in the fast growing tropical rainforests, a major biospheric sink of atmospheric CO2. Yet the role of the other biomes might prove to be grossly underrated. Seasonal atmospheric CO2 fluctuations, with the mid-latitude highs/lows in April and August respectively, attest to significance of deciduous vegetation (Woodwell et al., 1978; Watson et al., 1991). The boreal forests store less in the arboreal phytomass, but more in the moss layer, the annual production of which can be equal to that of the trees. Moreover, while the tropical evergreen biomes store carbon in their standing crop biomass, the temperate biomes shed much more to peat and soil, which in the long run are more durable carbon sinks (D’Arrigo et al., 1987). Humus contains four times more carbon than vegetation, most of it in the temperate forests and grasslands (Woodwell et al., 1978; Bolin, 1987). The mid-latitude forest uptake of CO2 is estimated as 3.7± 0.7 tons per ha per year, that of grasslands about 11.1 tons per ha per year which is much more than previously expected (Wofsy et al., 1993). The detrital pool contains three times more carbon than living plants. Its role as a sink of atmospheric CO2 depends on the balance of accumulation and decomposition rates. The latter are controlled both by climate and litter quality (Meentemeyer, 1978; Aerts, 1997). While climate (the variables affecting evapotranspiration) is a leading factor for leaf litter, the Ca level and the C:N ratio (relatively high in conifer needles showing the slowest decomposition rates) are more important for the roots (Silver & Mija, 2001). Hence the long-term variation of litter decomposition rates are related to both climate and evolution of higher plant biochemistry. In marine ecosystems, zooplankton sinks carbon with fecal pellets to bottom deposits where it is stored in anoxic sedimentary environments. This mechanism gains in importance with a spread of near-bottom anoxy, partly correlated with an influx of fresh waters. At the same time, planktic production increases with riverine discharge bringing nitrates and iron. During glacial phases, iron comes also with dust from desert areas (Kumar et al., 1995). An active circulation of oceanic waters enhances planktic production by preventing bacterial denitrification of anoxic subsurface waters (Ganeshram et al., 1995). Since plankton is also a source of dimetilsulphate aerosols that increase cloudiness, the net albedo effect of an increased oceanic productivity is much greater than its CO2 effect.

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Land biota plays primarily a stabilizing role maintaining the balance of atmospheric composition through the multiple negative feedbacks of pCO2 fluctuations. From 1/2 to 2/3 of CO2 emitted to the atmosphere sinks to biota (Cao & Woodward, 1998; Schimel et al., 2001). The biotic uptake fluctuates on a decadal scale, with an increase in the 1900’s owing mainly to physiological feedbacks (a longer growing season, fertilization by CO2), fire prevention, forest management and reclamation of abandoned lands over the northern extratropical areas (Schimel et al., 2001). Total phytomass production tends to increase with atmospheric pCO2, in particular over the mid- to high latitudes (Adams et al., 1990). It seems possible to at least partly compensate for the industrial CO2 by afforestation (Fang et al., 2001). With a doubling of atmospheric CO2, agricultural crop production, if not limited by a deficit of soil nitrates and phosphates, would potentially increase by about 10% binding an appreciable amount of carbon. Plant species widely differ in their responses to atmospheric CO2 levels (Wilsey, 2001), an elevation of which impose a selection pressure for acceleration of growth rates. However, an enhancing effect of additional CO2 on photosynthesis is partly negated by an increase in respiration from leaf canopies (Specht et al., 1992) and soil (Cao & Woodward, 1998). Experimental data suggest a rapid adaptation of annual plants to the current atmospheric CO2 concentration level, with a negative effect of further elevation on their physiological efficiency (Bunce, 2001). Anyway, an acceleration in growth rates under an elevated CO2 concentration (Heil et al., 1988) may not be a long-lasting effect, for it is slowed down by a side effect on nitrogen turnover (Centritto et al., 1999; Marriott et al., 2001). More nitrogen is consumed by the rapidly growing plants and less is released from their litter which has a relatively high C:N ratio (Tateno & Chapin, 1997; Cao & Woodward, 1998). An increase in the C:N ratio in foliage stimulates herbivory, in particular root feeding (Wilsey, 2001), with adverse effects on CO2 consumption by vegetation. Greenhouse not only increases plant growth rates but also accelerates reproduction, shifting flowering to an earlier developmental stage/smaller size (Morse & Bazzaz, 1994). An elevated CO2 concentration promotes early flowering typical of early seral stages that expand at the expense of the later stages (Roden et al., 1997), and reverse at a depressed CO2 level. Owing to the parallelism of seral and evolutionary sequences (VIII.4), the latter effect might have occurred during plant evolution as well. The carbon content of lignified plant tissues tends to increase with evolutionary advance. In the Carboniferous wetland forests, it increases from tree ferns (Psaronius) to lepidophytes to medullosan gymnosperms (Baker & DiMichele, 1997). A replacement of lepidophyte wetlands by gymnosperm wetlands at about the Permian/Triassic boundary would have then had a CO2 depressing effect similar to that of the hardwood/softwood replacements over the Late Cretaceous and Cenozoic. A specific physiological response is provided by the CCMs, or CO2 concentrating photosynthetic mechanisms, such as the Crassulacean Acid or C4 Dicarboxylic Acid


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mechanisms (CAM, C4) that bind more carbon than the conventional C3 photosynthesis. The C4 plants are recognized by the characteristic Kranz anatomy of the concentric bundle sheath and palisade cells once thought obligatory for this type of CCM (it has been recently shown, however, that C4 photosynthesis may occur even without Kranz anatomy: Voznesenskaya et al., 2001). CCMs might have appeared repeatedly, from as far back as 15-20 Ma or even earlier, in response to environmental factors imposing limitations on the vegetational carbon uptake, such as extreme temperatures, draught, soil salinity, etc. An increase in CAM and C4 plants resulting from a drop of atmospheric CO2 below the level optimal for C3 photosynthesis is postulated for tropical grasslands on the basis of carbon isotope ratios in collagens and tooth enamels of grazing animals (Morgan et al., 1994; Cerling et al., 1997). However, a direct response of photosynthetic mechanisms to atmospheric pCO2 levels is a debated issue (Prell et al., 1992; Moore 1982; Ehlinger et al., 1991; Ku et al., 1996; Huang et al., 2001). A correlation of high C4: C3 ratios with atmospheric CO2 fluctuations appears plausible for 7.3 Ma and a few subsequent time levels over the Neogene (Kleiner & Stresker, 2001) but not universally so (Morgan et al., 1994). Moreover, the C4 grasses might increase relative to C3 grasses owing to a spread of waterstressed environments rather than as a direct response to the atmospheric pCO2. Sequestration of carbon by forests depends on growth rates that correlate with structural differences between the old-growth and second-growth stands. CO2 uptake decreases with age approaching zero in old-growth forests. Hence, the carbon sink and flux vary with disturbance rates imposed by both natural and human impacts (Dixon et al., 1994; Keeling et al., 1996; Rayner et al., 1999). Extrapolated over the long-term deforestation/afforestation cycles, this means an elevation of atmospheric CO2 with a retreat of forest biomes, a sharp drop at the early afforestation stage (maximal uptake) and stabilization with the new forests coming to maturity. At least the Maastrichtian – Early Palaeocene climate changes seem consistent with this scheme (VII.6). Critical for the terrestrial biotic responses is a global change in precipitation patterns owing to atmospheric CO2 fluctuations. A reduction of tropical precipitation and rainforest production with a doubling the present CO2 interposes a positive feedback (a net reduction of carbon sink to biota) counterbalanced, to an unknown extent, through an advance of biomass to high latitudes (VIII.2.2). Water uptake in terrestrial plants decreases with elevated atmospheric/soil pCO2. Therefore, under greenhouse conditions, terrestrial plants are more tolerant of water deficit. In effect, dense vegetation expands over dry areas, toward the poles and up the mountains, thereby increasing the total terrestrial biomass. Interrelated with vegetation changes are pedogenic processes mobilizing the huge soil reservoir of organic carbon Since CO2 concentration in the air of soil pores is in a dynamic equilibrium with that of the ambient air, an elevatation of atmospheric pCO2 level intensifies pedogenic weathering of silicate minerals (Brady & Carroll, 1994). An influence of elevated CO2 on the rhizosphere includes a decrease in root colonization

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rates by ectomycorrhizal symbionts in turn affecting growth rates and carbon allocation in the forest trees (Fransson et al., 2001; Oren et al., 2001). At an elevated atmospheric CO2 level, an increase in the soil carbon concentration is accompanied by a rise of cellulose decomposers among the soil fungi, as well as collembolans and other fungivorous arthropods (Jones et al., 1998). Generally, an elevated CO2 uptake by vegetation correlates positively with the biological diversity (Naeem et al., 1994) and is considered as a factor of the latter. Temperature – CO2 – temperature system. As it follows from the above discussion, atmospheric CO2 fluctuations, both recent and over geological times, depend on the dynamic interactions of the lithospheric, hydrospheric and biotic reservoirs and a balance of the (mostly not as yet quantified) positive and negative feedbacks. Tapping any of the major reservoirs, be it by a magmatic–metamorphic outgassing, spread of warm surface waters, deforestation, degradation of soil, burning of fossil fuel or biomass, leads to carbon reallocations with potential climatic effects. However, the present-day technogenic “experiment” shows that even a massive emission of greenhouse gases is rapidly compensated for, as about three times less CO2 resides than is emitted. The existing models of greenhouse-generating CO2 emissions assumes a linear response of CO2-regulating systems, ignoring autocyclicity, an inevitable outcome of regulation with negative feedbacks, as well as a possibility of overcompensation. The latter arises due to the inertia of negative feedbacks. Overcompensation potentially occurs in the CO2 – silicate weathering, CO2 – carbonate deposition and CO2 – glaciation systems (above), each of which may cause a reversion of the climate trend. The CO2-centered models of climate change are inadequate not because of an imprecision of the basic estimates, a deficiency surmountable with accumulation of data and sophistication of techniques, but rather because of a reductionist approach to the immensely complicated system that does not come to any simplistic expectations. A prediction (of a rapid doubling the atmospheric CO2) based on an extrapolation of the recent technogenic emission rates would inevitably fail on account of the negative feedbacks that, with elevation of pCO2, would alter not only the rates but possibly the direction of the process as well. Eventually the process would come under control of natural regulation. Thus, over the warming trend, more CO2 is spent on chemical weathering of silicate minerals and is deposited with carbonates to be returned to the atmosphere with their dissolution or thermic decomposition. In this geochemical circuit, the relative rates of CO2 consumption and release depend on the air temperature that affects the weathering rates both directly, as an enhancer of chemical reactions, and indirectly, through its effects on mycorrhiza (Smith & Read, 1997), decomposition of organic matter and related rhizospheric processes (reviewed in Boucot & Grey, 2001). Air temperature and the related SST are engaged in regulation of carbonate deposition (lysocline fluctuations relative to thermocline). The geochemical cycle thus falls under control of the temperature cycles induced by the factors (orbital, tectonic, eustatic, etc) other than CO2.


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The oceanic deep-water circulation is regulated by the SST and by freshwater – saline injections, which in turn fall under control of the air temperature, either directly or through a derived effect of precipitation on the input from rivers and melting ice. In the biological cycle, CO2 is consumed by photosynthesis, the products of which are stored in the biomass, soil, peat, sapropel and the derived organolites. It is returned to the atmosphere with respiration, decomposition or burning of biomass and the fossil dead mass (fossil fuel). Here also the air temperature is effective in controlling, directly or with the aid of related precipitation variables, the ratios of arboreal to non-arboreal vegetation types, growth rates, lignification, C3:C4 photosynthesis, seasonality of dead mass accumulation and burial, as well as soil respiration and the dead mass decomposition rates. With warming, an outgassing of marshgas condensates, silled by permafrost, adds to the greenhous effect. Thus air temperature defines the rates of CO2 consumption/release over the geochemical, oceanic and biotic cycles, thus controlling their net effect. The atmospheric composition is therefore related to climate upon the temperature → CO2 → temperature scheme, rather than a simplistic CO2 → temperature scheme. CO2 is a positive correlate of air temperature in any case but the cause and effect relations are reversible rather than unidirectional. The positive feedback scheme is consistent with the observed atmospheric CO2 fluctuations associated with El Niño (above). The nascent atmospheric CO2 level initially drops with El Niño, perhaps owing to a lesser CO2 contribution from equatorial upwellings damped by the anomalously warm surface waters, then rises in conjunction with a decrease in tropical monsoons and the ensuing drop of tropical biomass (Siegenthaler, 1990). The amplitudes are larger than the sum total of technogenic emissions (Freeman & Hayes, 1992). Over the cycle, the atmospheric CO2 levels lag behind the sea surface temperature by four months and behind the air temperature by one month (Marston et al., 1991). These observations lend support to the regulation scheme, in which CO2 elevation is a consequence, rather than cause, of global warming due to the atmospheric phenomena (the rotationally regulated trade wind velocities in the case) unrelated to atmospheric composition. The greenhouse contribution then sustains the warming trend. At the onset of glaciation, a cooling of about 8°C over the high to mid-latitudes and about 6°C within the tropics greatly exceeds the calculated effect of reduced CO2 (Webb et al., 1997). The Antarctic (Vostok) ice core data suggest a net contribution of CO2 and CH4 (released from wetlands) about 2.2°C, less than 50% of the average 4.5°C postglacial warming (Raynaud et al., 1988). The Holocene and Pleistocene atmospheric CO2 levels based on fossil air samples from Greenlandic and Antarctic ice cores (Lorius et al., 1990; Jouzel et al., 1993; Raynaud et al., 1993) indicate the full glacial to interglacial (and Holocene) amplitudes about 20-30 % (200 and 280 ppmv respectively), correlated with the isotopic palaeotemperatures. This is commonly taken as evidence of CO2 → temperature regulation. However, the fluctuations show a statistically significant 20 k.y. periodicity of the precession cycle, which is definitely in favour of the temperature → CO2 scheme.

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Moreover, a lag between the temperature and subsequent CO2 fluctuations has been actually recorded in some of the fossil air studies (Raynaud et al., 1993). Incidentally, an appreciable lowering of CO2 level has been recorded only 6 k.y. after a cooling event over the 5.5/5.4 isotopic stage boundary (Fischer et al., 1999), the time lag being perhaps related to the inertia of such CO2-regulating oceanic systems as the vertical displacements of carbonate lysoclines (Hodell et al., 2001). The greenhouse records. Modelling of CO2 changes over geological times has started with a seminal work by Budyko et al. (1985, 1986). In their model, the atmospheric CO2 concentration is determined by the ratio of coeval volcanic (source) and carbonate (sink) deposits (based on quantitative estimates in Ronov et al., 1980 and elsewhere). Over times, atmospheric CO2 allegedly decreases with a fading-out of volcanic activity. While bringing to view the geospheric factors that potentially may determine atmospheric comoposition in a very long run, the model fails to account for either reversibility (e.g., a sink of CO2 to chemical weathering of volcanic rocks or its release with thermic decomposition/dissolution of carbonates) or reallocation of CO2 between the atmospheric, oceanic and biotic reservoirs that operationally regulate the greenhouse effect (above). As for the general trend, the postulated decrease in volcanic activity with dissipation of radioactive heat is scarcely substantiated by either theoretical considerations or geological data. The model disregards kinetic heat production by shear at the interior density boundaries within the lithosphere and over the lithosphere/mantle transition (V.2). Moreover, the rock volume assessments neglect the most voluminous terrestrial volcanism of the Permian/Triassic (Siberian traps), Jurassic/Cretaceous (the East Brazilian and Mongolian basaltic provinces), mid-Cretaceous (the circum-Pacific volcanic belt), and the Cretaceous/Palaeogene (Deccan traps) events. Neither does it account for the pulses of oceanic magmatism, as in the Cretaceous. Worse of all, the model is in a poor correspondence with palaeoclimatic records. A postulated Early Permian peak of atmospheric pCO2 falls at the time of glaciation, the next, mid-Triassic, peak is timed to a major cooling, the Jurassic ascending track passes over the prominent Bathonian/Callovian cooling, and the Oligocene low indicates a cooler climate than at present. Contrary to the model by Budyko et al. (1985), the most prominent volcanic events associate with global cooling rather than warming (Briffa et al., 1998). The greenhouse effect of volcanic emissions is grossly outweighed by the volcanic ash/aerosol albedo effects, in particular the backscatter of solar radiation by SO42- plus a depletion of stratospheric ozone. A further elaboration by Berner (1990 and elsewhere) involves a great number of geological and biotic variables, such as land area, river runoff, biotic production, carbonate dissolution depth, etc. Atmospheric pCO2 levels are calculated as a reverse function of the difference between the CO2 spent to weathering of silicate rocks (the rates of which are inferred from the ratios of carbonate to organic deposition) and those emitted form the contemporaneous volcanic/metamorphic sources. A correlation of the lowest


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calculated atmospheric pCO2 levels with the Late Paleozoic and Neogene glaciations is taken as a verification of the model. However, the outgassing rates are derived from the supposed sea floor spreading rates in turn inferred from the observed sea-level fluctuations. In effect, Berner’s curve actually reflects sea-level fluctuations that are climatically significant with or without mediation of CO2 (VII.2.3). Instead of being relegated to mere evidence of alleged sea-floor spreading rates, sea-level fluctuations has to be considered as a major force under greenhouse events due to their effect on oceanic circulation as well as through the land area – terrestrial biomass – pedogenic weathering correlations. The factual evidence of CO2 fluctuations over geological times is controversial. The carbon isotope ratios defined for marine and pedogenic carbonates (that reflect photosynthetic carbon isotope fractionation related to the atmospheric CO2 concentrations) indicate a high CO2 level in the Ordovician, actually an icehouse period (Cerling, 1991; Freeman & Hayes, 1992). The Late Palaeozoic to Mesozoic estimates (e.g., Ghosh et al., 2001) fall within a very broad uncertainty margin (except for the Jurassic), whereas the Cenozoic ones are totally incongruous. The stomatal method is based on the response of stomatal frequencies to CO2 concentrations on the leaf surface that depends on the atmospheric pCO2 (which is no longer a limiting factor at 200% of the present-day level), leaf area and wind velocity (Ågren, 1985). The method is verified by comparison of stomata per cell ratios (stomatal index) in pre-industrial herbarium specimens and living plants of the same species. Objectives and limitations of the method are discussed in Kürschner (1996) and Roger (2001). Inverse response of stomatal frequency to the ambient CO2 levels has been found in 36% of hitherto studied species. CO2 fluctuations of about 30% over the last glacial/ interglacial transition correspond to about 17% drop of stomatal index in Pinus flexilis and Salix cenerea (Van der Water et al., 1994). The estimates for Quercus petreaea, a long-ranging Mediterranean species, seem congruous with other indicators of the Neogene climate change (Van der Burgh et al., 1993). Other limitations aside, the stomatal method requires long stratigraphic records that are rare in fossil plant species. The longliving genera, such as Ginkgo, seem promising (Chen et al., 2001), although interspecies comparisons are scarcely justified by the method. Moreover, the Mesozoic species of Ginkgo (Ginkgoites) differ from extant G. biloba in the densely trichomate foliage leaves, similar to juvenile leaves of the living species, suggesting a different ecology. It must be made clear that the stomatal index method is based on species-specific response to CO2 concentration in the ambient air determined by the genome. When the comparisons transgress species boundaries, we are no longer dealing with the genomespecific responses, but rather rely on the convergent ones. Hence, it makes little diffrence whether the selsected species are phyllogenetically related or not. Mean stomatal length is another epidermal variable correlated with atmospheric pCO2 (Kürschner, 1996) that is more likely to show a convergent response in a group of con-

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temporaneous species. Bennettites, an extinct group of gymnosperms widespread from Late Triassic to Early Cretaceous, have paracytic leaf stomata with dimensions typically constant for a leaf population but varying over stratigraphic sequences (Krassilov, 1967a; Watson & Sincock, 1992), therefore suitable for this kind of analysis. The bennettitalean stomatal length measurements for several Mesozoic taphofloras are presented in Table 2. The means change from moderate, 27-37 ì m, in the Late Triassic/Early Jurassic (Scoresby, Western Greenland) and mid-Jurassic (Yorkshire, England) to a low of 35 ì m in the late Middle–early Late Jurassic (Karatau, southern Kazakhstan), and to a peak of 44–45 ì m in the Late Jurassic to early Neocomian (Wealden of southern England). Further in the Cretaceous, the mean values drop to 24 ì m in the Mongolian bennettitalean assemblage presently dated as the mid- to late Neocomian (Krassilov & Bugdaeva, 2000) and rise again to 39 ì m in the Aptian of Primorye, Russian Far East (Fig. 93). Since all the localities belong to one and the same phytogeographic zone (IV.3.4), a geographic component of the stomatal length variation should have been insignificant. A plausibility of an atmospheric pCO2 control over the mean stomatal length fluctuation is

Fig. 93. Stomatal length index (mean guard cell length) for bennettitalean leaves from Mesozoic taphofloras: (TJ) Late Triassic/Early Jurassic (Scoresby, Western Greenland), (Jbb) Mid-Jurassic of Yorkshire, England, (Jc) Late Middle (Callovian) to early Late Jurassic of Karatau, southern Kazakhstan, (Jk) Late Jurassic (Kimmeridgian) of France, (JK) Terminal Jurassic to early Neocomian (Berriassian) of southern England, (Kn) Early Cretaceous (Neocomian) of Mongolia, (Ka) Mid-Cretaceous (Aptian) of Primorye, Russian Far East.

Mean 27.0

P.subaequale 30 Otozamites sp. 22

P.xiphipterum 16

Mean 37.3

Z.nicolae 40 Z.tatianae 46,25 Z.notokenensis 56,25

O.venosus 40 O.simpsonii 42 O.marginatus 50 Zamites gigas 48,7 Nilssoniopteris major 34,4 N. pristis 37,5 Anomozamites thomasii 35 Pterophyllum cycadites 30 P.fossum 40 P.thomasii 47

P.ptilum 20 P.zygotacticum 20

Z.corderi 56,25

O.gramineus 36

Pterophyllum hanesianum 30 P.astartense 25-30 P.schenki 28 P.rosenkrantzi 20 P.pinnatifidum 40 P.kochi 20

Mean 44.2

Z.distractus 52

Mean 44.9

Z.wendyellisae 43,75

O.beanii 35

Mean 35.0

A.nitida 35

England Mongolia Berriassian Barremian Pseudocycas saportae Nilssoniopteris 61 denticulata 21,3 P.roemeri 63,75 Otozamites lacustris 28,4 P.lesleyae 60 Neozamites verchojanicus 18 Otozamites titaniae Pterophyllum cf. 53,75 angulatum 28,3 Zamites carruthersii 36,25 Z.manoniae 38,75 Mean 24.0 Z.dowellii 40

O.leckenbyi 32 O.thomasii 32,5

Z.feneonis 52

Z.dubisensis 48

Z.carpentieri 29

Zamites pumilio 48

France Kimmeridgian Otozamites reglei 36

A.minor 25 A.amdrupiana 30

Karatau Callovian Sphenozamites sphenozamioides 37,5 Anomozamites kornilovae 30 Nilssoniopteris aff vittata 65 N.boroldaica 50,6

Otozamites penna 32 N.karataviensis 28,5

P. pectinoides 45

Yorkshire Bajocian Dictyozamites hawellii 25 Ptilophyllum pecten 37,5 P.hirsutum 30

Anomozamites marginatus 40 A.hartzi 25

T.ajorpokensis 30

Greenland Rhaetian Taeniozamites groenlandicus 22 T.jourdeyi 30

Mean 39.0

Otozamites Dictyozamit grossinervis Pterophyllum sutschanens Cycadolepis pterophylloi

Ptilophyllum 30 Zamiophyllu buchianum Zamites bor

Primo Apti Nilssoniopte rhitidorachi N.robusta 30

Table 2. Stomatal lengths in the Triassic to Early Cretaceous bennettitalean leaves (Harris, 1932, 1969; Krassilov, 1967a, 1982; Doludenko & Orlovskaya, 1976; Barale, 1981; Watson & Sincock, 1992).

Chapter 7. Climate change 229

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confirmed by the parallel variation of the cycadophyte index, a proxy of temperature changes over the Mesozoic (Krassilov, 1973b; Fig. 85). Dispersion of stomatal measurements has been calculated upon the assumption that a low level of atmospheric CO2 would have caused a wider variation. The maximal dispersion was obtained for the Karatau assemblage but the next low in the mid-Neocomian coincides with a relatively narrow dispersion. Perhaps a more meaningful correlation will emerge with accumulation of data. A qualitative phytogeographic approach (Krassilov, 1997c) is based on a correlation of atmospheric/soil pCO2 with plant tolerances in respect to other limiting factors. In particular, a water uptake by photosynthesis typically decreases with elevation of atmospheric pCO2. In effect, the greenhouse conditions would have rendered plant species more tolerant of water deficit. Their geographic ranges would have been less dependent on precipitation. Since phytogeographic zonation is controlled both by temperature and precipitations, a reduced water uptake would make zonal boundaries more transparent, resulting in a mixing of zonal elements. An azonal plant distribution is thus a phytogeographic testimony of greenhouse climate. In the fossil plant record, the distinctness of the boundary between the mesotemperate deciduous and xerothermic evergreen zones varies with climate fluctuations. Smearing of the boundary is recorded for the relatively short “azonal” intervals in the Early Carboniferous (Visean), Early Triassic (Olenekian) and the Early-Middle Eocene, the supposed greenhouse episodes (VII.4). CO2 trends and feedbacks. As discussed above, the atmospheric CO2 levels depend not so much on the total amount of carbon released from the earth’ interior, as on its reallocation between the atmospheric, lithospheric, oceanic and biotic reservoirs. A directional change of atmospheric CO2 concentration over times would have occurred with directional evolution of these reservoirs or any of them. Net CO2 fluctuations result from a disbalance of multiple sinks and sources, their interaction being affected, among other variables, by the air temperatures and the SST. The time-average atmospheric CO2 levels are sustained, with short-term fluctuations, by the negative feedbacks upon the temperature → CO2 → temperature scheme. There is little doubt about the role of outgassing of the earth’s interior in the origin and regulation of atmosphere before the biospheric regulating mechanisms came to action and took over, of which the Proterozoic banded ore formations might have been an early evidence. Sea-level fluctuations, with an uncertain directional component, but probably of a deminishing amplitude since stabilization of continental crust over the Precambrian fold belts, administer a control over the sink rates of atmospheric CO2 to deep-water circulation, terrestrial biomass and carbonate deposition. During the global regressions, both terrestrial biomass and chemical weathering increased with land area, consuming more CO2 while, simultaneously, deep-water circulation was enhanced by shutting-down an influx of warm saline waters from epeiric seas. Transgressions inflicted the opposite trends (Fig. 94).


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Fig. 94. Sea-level curve (after Berner, 1990) compared with contemporaneous climatic curve (the same as in Fig.100) and a tentative atmospheric pCO2 curve in respect to sea-level and biotic regulation, sustaining a conservative mid-level of about 380 ppmv through the Mesozoic and Cenozoic, with occasional peaks reflecting the greenhouse episodes.

Hence, notwithstanding their first-rate significance in affecting CO2 fluctuations, partly through their effects on biota, the lithospheric and oceanic sources scarcely induced a directional change of atmospheric composition since the Precambrian. The biotic reservoir is one that evolved to a larger capacity over Phanerozoic times. As discussed later in the chapter, the buld-up of terrestrial biomass was accomplished during the Palaeozoic, with the later changes fluctual rather than directional. We, therefore, may expect larger CO2 fluctuation during the Paleozoic than later, which is at least

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partly confirmed by the records. As major fluctuations of terrestrial biomass were inflicted by sea level and continental glaciations, there should be distinct CO2 levels corresponding to the talassocratic and geocratic epochs, and the glaciated/ice-free continents, with the quantitative estimates for the Pleistocene reflecting the former, and the presentday values approaching the latter norm. A CO2 curve through the Phanerozoic times would have, thus, four modal levels, in descending order, the Palaeozoic level, with a directional change over large-scale fluctuations, the Meso-Cenozoic talassocratic level, considerably higher than the present one, but quantitatively poorly defined, the Meso-Cenozoic geocratic level for ice-free continents approached by the present-day values, and the glacial maximum level roughly corresponding to that of the last glaciation. These levels are tentatively shown in Fig. 94. The biotic responses are related to greenhouse gases and temperature as their negative feedbacks, thus having a stabilizing effect. Elevation of atmospheric CO2 level enhances both the biological production and pedogenic dissociation of silicate minerals by organic acids (reviewed in Brady & Carroll, 1994; Boucot & Grey, 2001). The water uptake by photosynthesis decreases allowing plants to colonise dry habitats. Likewise, a lowering of atmospheric pCO2 level invokes compensatory resoponses, such as the C4 photosynthesis and an increased water uptake enhancing the chances of biomass burning that returns CO2 back to atmosphere. Fluctuations of atmospheric pCO2 are further compensated for by soil respiration (Curl & Truelove, 1986). We arrive at a scheme in which an input of greenhouse gases form a source evoking no or insignificant feedback, such as volcanic activity, sea-level fluctuations or technogenic emission, are compensated for by biota, first of all the land biota, producing a stabilizing effect on climate. The biotic responses are more rapid than either the lithospheric or oceanic (we have to delay our judgement of the human responses). Therefore, on a short-term scale at least, the biota is a major regulator of atmospheric composition and climate, compensating for fluxes of greenhouse gases from non-biotic sources, while the large-scale atmospheric fluctuations may be owing to a disruption of the biotic stabilizing mechanisms. Indeed, over the Phanerozoic history, prominent climate changes correlate with biospheric crises (IX.1-2). The recent elevation of atmospheric CO2 with technogenic emissions may or may not result in an appreciable climatic change depending on the extent of deforestation, soil erosion and other human impacts disrupting the homeostasis of natural ecosystems.

VII.2.5. Stratospheric effects Stratospheric ozone, systematically monitored since 1978, is presently depleted by about 3% in the mid- to low latitudes, which means a nearly adequate increase in UV radiation at the DNA-damaging wavelengths. At high latitudes, the seasonal ozone de-


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pletions are much greater, up to 50% over Antarctica at altitudes 20-50 km. This phenomenon, first observed in 1985, became known as the “ozone hole”. It is now considered as a major environmental hazard. Episodic attenuation of the ozone layer is caused by solar flares, as well as by cold winters (Müller et al., 1997). The ongoing ozone depletion is ascribed to industrial emissions of chlorofluorocarbons (CFCs), a source of chlorine that is engaged in heterogeneous reactions on the stratospheric ice clouds. Naturally such compounds are produced by decaying red algae (Wever, 1988). The ozone-destructive aerosols are metabolically produced by terrestrial vegetation (Kavoures et al., 1998) and by phytoplankton. Bacterial nitrification is a substantial source of atmospheric nitrous oxide, enhanced by warm SST (incidentally by El Niño) bolstering a dominance of nitrogen-fixing cyanobacteria (Dore et al., 1998). Volcanic eruptions are a source of chlorides far exceeding the present-day industrial emissions. Concomitantly, the ozone-destroying reactions are enhanced by injections of sulfur acid aerosols (Arnold et al., 1990; Brasseur et al., 1990; Brasseur, 1991). Critical for the stratospheric ozone layer are the tropospheric processes affecting circulation of ozone-destructive compounds (Fig. 95). These are conducted by ascending air, primarily at the tropopause and through the polar vortices. Thus, aerosols from

Fig. 95. A Watson circle scheme (Coughlan & Armour, 1992) of variables affecting the stratospheric ozone layer, including chlorofluorocarbons (CFC), their derived chlorine, nitrous oxides, sulfur acid aerosols, water vapour, hydroxyl, temperature and the vorticity of the polar area (right margin symbol). The relative importance of these factors is reflected in the number of departing arrows, maximal for the vorticity of polar air.

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the low-latitude volcanic eruptions (El Chichón, Mount Pinatoba) enter stratosphere at the tropical tropopause, spreading to the poles (Boering et al., 1996; Volk et al., 1996). The intensity of tropospheric impact on the stratospheric ozone layer thus depends on convectivity of tropospheric air that is stirred by the greenhouse warming. Redistribution of heat by the tropospheric concentration of greenhouse gases results in a cooler stratosphere, hence an acceleration of ozone-destroying reactions on stratospheric ice clouds (Wang et al., 1991; Shindell et al., 1998). However, the effects are non-linear owing to superposition of other atmospheric phenomena, such as a contraction of polar vortices that isolate polar air from ozone-destroying substances borne by the mid-latitude air masses (Stine, 1994; Robinson, 1998). A combination of positive and negative effects gives the 2 – 4 year ozone cycles that correlate with oceanic – tropospheric events, such as El Niño. On account of the volcanic, as well as the tropospheric, greenhouse effects, fluctuations of the ozone layer are bound to geological and climatic processes as an integral part of global change. A depletion of stratospheric ozone is potentially damaging to the biosphere because of an increased influx of UV radiation. This effect is presently mitigated by a concomitant increase in tropospheric ozone excessively produced by photochemical decomposition of technogenic nitrogen dioxide (Penkert, 1989). The photochemical fog of modern cities, though damaging to respiratory systems, shields the urban life, also the adjacent rural life, from UV radiation. In pre-industrial time, a reduction of stratospheric ozone by natural agents, such as volcanic aerosols, might have had even more dramatic consequences. The biotic turnovers of the fossil record correlate with huge magmatic events, such as the Siberian and Deccan traps at, respectively, the Permian/Triassic and Cretaceus/ Tertiary boundaries. A massive discharge of volcanic aerosols to stratosphere inevitably resulted in at least temporary reduction of the ozone layer. At such critical moments, the mutagenic effect of UV radiation might have contributed both to the mass extinctions and innovations. Evidence of UV damage comes from plant cuticular studies. Thus, an exceptionally high frequency of epidermal anomalies, such as an irregular differentiation of epidermal zones, interrupted stomatal rows, underdeveloped and/or contiguous stomata, disorganized stomatal morphologies, etc. (Fig. 96), are recorded in peltasperms and conifers from the Permian/Triassic boundary clay of Nedubrovo, central European Russia, a smectitic horizon correlated with the eruptive phase of Siberian and Petchorian traps (Krassilov et al., 1999a, 2000).

VII.2.6.Vegetation effects Terrestrial vegetation bears on the albedo of the earth’s surface (about 35% over deserts, 9 -12% over forested areas: Sellers et al., 1997), air temperature (decreasing by 3-5 K in forested areas), evaporation, wind velocity, air convection, meridional heat transport (Hadley


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a

b

c

d

Fig. 96. Epidermal anomalies in gymnosperms from the Permian/Triassic “boundary clay” of Nedubrovo as evidence of UF damage enhanced by loss of stratospheric ozone at the time of massive volcanic eruptions (Krassilov et al., 1999a, 2000): (a) an irregular epidermal topography with linear and concentric configurations of compressed cells in a peltasperm, Tatarina antiqua, reflecting erratic local bursts of meristematic activity, (b) contiguous stomata and compact groups of small thick-walled cells derived from abnormally developed stomatal initials (arrow) in a conifer, Ullmannia cf. bronnii, (c) an irregular configuration of subsidiary stomatal cells distorted by intervention of ordinary cells in the same species, (d) a profuse division of encircling cells of an aborted stoma in Tatarina antiqua.

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cell geometry) and the rainfall (Potter et al., 1975; Sellers et al., 1997). Moreover, vegetation introduces an indirect effect on climate by trapping rainwater and snow, controlling floods and riverine discharge and, consequently, the input of terrestrial organic matter to the oceans. Crucial role of vegetation in the carbon cycling makes it a major regulator of atmospheric CO2 (VII.2.4). Humic acids of the rhizosphere greatly enhance mobility of Ca, Mg and Fe cations released by chemical weathering. Their export to marine ecosystems is a significant factor of marine biotic production (Lloyd, 1991). Evidence of pedogenesis by at least 50 m.y. precedes the advent of higher land plants (VIII.I.I). Plant debris (in particular, the wind-borne spores) might have been initially transported from marginal marine environments to land supporting the pre-Devonian microbial-fungal saprophytic life. Even such plantless terrestrial ecosystems might have induced chemical weathering of silicate minerals (Boucot & Grey, 2001). Early land plants might have entered the primordial terrestrial ecosystems owing to symbiotic interactions with fungi forming a mycorrhizal rhizosphere. With this, the terrestrial life became a major sink of atmospheric CO2 enhancing both the oxidation of atmosphere and accumulation of the stratospheric ozone layer. The mainstream transport of organic matter was redirected from land to sea, increasingly prevailing over the opposite flow. Along with a decrease in UV radiation reaching the earth, plants acquired leaf blades and, as a result of light competition, as well as under the pressure of arthropod herbivory, sequentially developed the perennial life forms, arboreal habits, lignified vascular tissues and the multistratal canopies. The build-up of biological diversity, including the beta and gamma components, allowed plant communities to spread over the globe. Seasonal adaptations and the increasing complexity of seral stages provided for developmental homeostasis enhancing the stabilizing vegetational effect on climate. It is commonly assumed, and is at least partly justified by the fossil record, that prior to the build-up of terrestrial biomass the climatic fluctuations were larger (extreme fluctuations are indicated, for instance, by alternation of warm-water carbonates with glaciogenic deposits in the Ordovician). The stabilizing effect of the biosphere on climate might have conceivably increased over geological times with a rise of terrestrial biomass and its capacity as a reservoir of organic carbon. Major steps in this direction were: - The advent of the early land plants and the rhizosphere, with detritivorous arthropods, in the Late Silurian – Early Devonian, - The onset of peat accumulation, first by the heavily cutinized thalloid macrophytes of the Shuguria – Orestovia group (Figs. 43 – 45) in the Middle Devonian, later replaced by the tracheophyte peatland communities, - The appearance, in the mid-Devonian, of arboreal growth form, with lignification typically increasing over the Palaeozoic history; - The rise, since the Late Devonian, of forest biomes with multistratal canopies and the ever increasing diversity of life forms,


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- The development, since the Late Carboniferous, of seasonal deciduousness that increased the carbon sink with a shed leaf mass, - An increase of trophic cascades with elaboration of detritvorous and herbivorous adaptations the diversity of which seems to have been saturated by the Permian. Thus, a major increase of terrestrial biomass occurred in the Palaeozoic, with subsequent developments switching to a fluctuation mode (one can argue that biotic capacity have increased with the relatively recent appearance of such overproductive vegetation types as the tropical rainforests; yet, as discussed above, the rainforests allocate more carbon to their standing crop biomass, but less to their dead mass production). Fluctuations potentially affecting the stabilizing role of terrestrial vegetation in respect to atmospheric composition and climate are induced by sea level and climate through their effect on the area of plant cover, as well as the productivity of terrestrial biomes. Over the post-Palaeozoic history, these variables fluctuate with: - An extent of vegetated land changing with sea-level fluctuations (the consequences of Cretaceous transgressions are discussed in VII.2), desertification and continental glaciations, - The ratios of arboreal to non-arboreal vegetation changing with the periodic afforestation/deforestation waves, with afforestation maxima in the mid-Jurassic, early Palaeocene and Oligocene (VII.5), affecting the net carbon uptake by the biota, - An extent of peatlands, maximal in the Carboniferous, mid-Jurassic, mid-Cretaceous and mid-Eocene affecting both CO2 and CH4 levels, - The ratios of evergreen to deciduous vegetation, with a spread of deciduousness in the mid-Jurassic, Early Palaeocene and Oligocene (IV.2.2), affecting the dead-mass accumulation/decomposition rates, - The ratios of C3 to CAM and C4 photosynthetic types, typically correlated with deforestation (Moore, 1982; Ehleringer et al., 1991; Ku et al., 1996). Land plant photosynthesis depletes the atmospheric carbon reservoir of the lighter isotope that is then released with forest fires to contribute to the Suess effect, a decrease in carbonate δ13C that is in equilibrium with the atmospheric 12C /13C ratio. The δ13C fluctuations are partly ascribed to the changing ratios of C3 to C4 photosynthetic types (Morgan et al., 1994; Cerling et al., 1997), the latter presumably appearing about 15-20 Ma, with peaks at 7.3 Ma and a few subsequent levels over the Neogene (VII.2.4). However, the anatomical syndrome of C4 photosynthesis (the Kranz anatomy) has been recorded in the Mesozoic Pseudofrenelopsis already (Bocherens et al., 1993) suggesting a much longer evolutionary history.

VII.3. Climate trends This chapter summarises the discussion of climate change in VII.1-2. A popular climatological theory holds that, with dissipation of the earth’s internal heat, the global

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climate is getting progressively cooler over geological times (Budyko et al., 1986). In this model, radioactivity is assumed to be the sole source of internal heat released with igneous activity. However, a far prevailing source of igneous activity is friction heating generating magmatic chambers at the geophysically defined density boundaries between and within the earth’s mantle and lithosphere. Contrary to Budyko et al. (1986), the igneous activity, rather than showing uniderectional trend over times, fluctuates with other geological variables related to the earth’s rotation rates. The alleged factors of directional cooling also include a decrease in solar radiation, as well as a sink of atmospheric CO2 to a buildup of terrestrial biomass and pedogenesis (Berner, 1990, 1997). The Pleistocene glaciation has been taken as evidence in favour of the cooling theory, yet the preceding Proterozoic, Ordovician and Late Palaeozoic glaciations were comparable, if not larger. The vegetational zonation does not indicate a general trend over the sequential climatic optima from the mid-Devonian to Early Eocene (VII.4). Lack of a long-term correlation between climatic change and the solar constant indicates a prevailing role of albedo, the proportion of reflected (scattered) light that eventually determines how much of the incoming heat resides in the atmosphere. The earth surface albedo depends on geography, while the atmospheric albedo is primarily regulated by water vapour and the gaseous carbon compounds, hence the CO2 theory of climate change. However, before the technocratic humans, CO2 fluctuations were the result of a disbalance of atmospheric concentrations, on one side, and the lithospheric, oceanic and biospheric reservoirs, on the other (VII.2.4). A reallocation of carbon from these reservoirs to atmosphere and back is caused by tectonic, eustatic and biotic events (and by climatic feedbacks), essentially the same factors as those affecting other components of atmospheric albedo (water vapour, sulfide aerosols, etc.) as well as the earth’s surface albedo. Therefore, it is virtually impossible to single out CO2 from the net albedo effect. As discussed before, the geological record of past climates can be more objectively interpreted in terms of climatic gradients rather than global means. Certain palaeovegetational and/or sedimentological data may suggest that, in the distant past, polar areas were warmer than at present or, conversely, that continental glaciers left their marks far beyond their present-day range. Climatological generalizations from such data, such as an assumed warming/cooling in terms of mean global temperatures, are scarcely warranted by the nature of the evidence. It remains to be learned whether a “greenhouse” was globally warmer than an “icehouse” or the same amount of heat was just differently distributed. Such designations flow in respect to cause and effect since ice-free conditions are not necessarily the consequence of greenhouse effect. At the same time, redistribution of heat and precipitation over the globe is amply documented by the fossil record of climatic zonation. The evidence comes from zonal occurrences of certain rock types (tillites, evaporites, redbeds, coal, etc.) and vegetation types (arboreal/non-arboreal, deciduous/evergreen, mesophytic/xerophytic, etc.). Zonation is either fairly prominent or effaced, indicating a phytogeographic effect of elevated CO2 level (above). Individual zones spread at the expense of their neighbouring zones,


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with reversals coming up later on. What we actually know of past climate changes is a restructuring of zonal patterns, from which global climatic variables can be inferred with a varied degree of confidence. Major climate changes are indicated by alternation of the following zonal patterns: Humid/dry equatorial zone. The Late Cenozoic and Early Permian are the times of a predominantly humid equatorial zone supporting rainforests of macrophyllous angiosperms and gigantopterids respectively. The mid-Mesozoic is the time of a predominantly arid or semiarid equatorial zone supporting a xeromorphic vegetation of microphyllous (brachyphyllous) gymnosperms (IV.4). Dry equatorial zone signifies a pattern of global atmospheric (and, by inference, oceanic) circulation considerably different from the present-day norm, but corresponding to the present-day short-lived anomalous situations, such as desiccation of equatorial rainforests caused by El Niño (VII.2.3). Over the Cenozoic, as well as over the Permian, equatorial humidity is apparently correlated with the advances of continental glaciers. Evergreen/deciduous temperate zones. The ratios of evergreen to deciduous vegetation types vary with seasonality, extreme temperatures, precipitation, atmospheric/soil pore CO2 levels and the length of growing season. In areas of ample precipitation, deciduousness is due to a prolonged period of soil freezing inflicted by a spread of cold air generated by polar ice caps. Cold winters make temperate zones unfavourable for broadleaved evergreens, but, in combination with heavy snowfalls, they promote the needle-leaved evergreens. Dwarf-shrub evergreens are adapted to a short growing season (that need not be wasted on unfolding the leaves). Temperate deciduousness appears with polar ice caps and spreads to lower mid-latitudes with their expansion, replacing both the temperate and xerothermic evergreens. In the fossil record, deciduousness is recognized on the basis of leaf morphology, taphonomy and associated lithologies, among which the intracontinental coal measures furnish evidence of summer-wet climate. Ice-capped/ice-free poles. Glaciogenic deposits are direct evidence of past glaciations, but continental tillites are scarcely distinguishable from glaciomarine deposits and allotillites (Chumakov, 1990), let alone a possible confusion with taphrogenic mixtites. More reliable evidence comes from palaeophytogeography. Vast continental ice sheets generate a periglacial dry zone that reaches to the mid-latitudes. The well-defined temperate deciduous zones correlate with moderate polar ice caps, as at present. On evidence of mid-latitude temperization, polar ice caps might have existed over most of the mid-Palaeozoic to Cenozoic times. The alternative situation of ill-defined temperate zones, with a significant proportion of arborial evergreens spreading over the polar circles, is recorded over relatively short, about 6-10 m.y., intervals in the Carboniferous (Visean), Early Triassic (Olenekian), Late Cretaceous (Campanian) and Early-Middle Eocene (Sparnacian to Lutetian). It is widely held that glaciations are exceptional, occurring over no more than 10% of the Phanerozoic time, while the normal situation is ice-free (Lloyd, 1983). Hence recent

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climate is not a clue to the past or most of it. However, with mid-latitude temperization as a major criterion, polar ice caps come to be considered as the Phanerozoic norm and the ice-free climate as exceptional. Glaciated/ice-free tropical zone. The present is an example of a reduced equatorial ice at about 5000 m above sea level. In contrast, the Late Proterozoic tillites of Australia and the Transvaal Supergroup in South Africa occur at equatorial (about 5°S) palaeolatitudes (Evans et al., 1997). The Late Palaeozoic tillites are recorded at many low (palaeo)latitude sites over the so-called Gondwanaland. Their interpretation as an extensive South Pole ice-cap debris over the reassembled continents is contradicted by Glossopteris flora the broadleaved aspect of which negates a broad periglacial dry zone (IV.3). A vast montane glaciation seems more consistent with the Palaeozoic fossil plant record. Desertified/desert-free continents. Deserts are generated by circulating air masses that adsorb, rather than shed, water vapour. The most extensive deserts are due to desiccation by vapour-starved air masses descending on the flanks of the humid equatorial zone. Desertification is enhanced by mountain barriers in the path of tropical air masses, hence the correlation with the Tethyan orogenesis (VII.2.2). Basically, there are two alternative situations recurring over geological times: (1) Extensive polar ice caps, dry periglacial zones, broad temperate zones of seasonal deciduousness, and the humid equatorial zone bordering on arid/semiarid areas; (2) Reduced polar ice-caps, humid periglacial zones, irregular temperate zones of mixed deciduous/evergreen vegetation, and a broad summer-dry low latitude zone, with precipitation decreasing towards the equator. Typical of (2) is a low temperature gradient and the wider than now Hedley cells of the meridional heat circulation (D-Hondt & Arthur, 1996). Alternation of these climatic situations poses a major problem for historical climatology. Insofar as we are dealing with redistribution of heat/precipitation over latitudinal zones, the geographic factors – land/sea distribution, relief, glaciation, etc. – play a decisive role owing to their profound effect on the meridional heat transfer. However, the above climatic situations correlate not so much with a selected geographical variable as with combinations of interdependent variables. Thus, situation (1) repeatedly occurs during the geocratic epochs with deep oceanic depressions, high terrestrial relief and the mountaneous Tethys belt, whereas situation (2) is typical for the thalassocratic epochs with shallow oceans, low continents and the Tethys seaways. The mid-Cretaceous climate, typical for situation (2), is characterized by the narrow dry equatorial zone (xeromorphic vegetation, evaporites), broader subtropical zones of a seasonally dry climate with a predominantly evergreen xeromorphic vegetation and redbed facies extending to 50°N/45°S over continental interiors and to 60° on continental mar-


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gins, temperate zones of mixed evergreen-deciduous vegetation and the forested polar areas of high biological productivity (IV.3.5). The widespread mm-scale cyclicity of lacustrine and marine black shale sequences indicates a high susceptibility of mid-latitude Cretaceous climates to the rotationally induced (Milankovitch) precipitation cycles. Climatic fluctuations over the Cretaceous, with the temperature maxima in the Berriasian, Aptian, mid-Cenomanian – Turonian and Campanian, and the minima in the Hauterivian, Albian/Cenomanian, Coniacian and over the Maastrichtian/Palaeocene boundary (palaeobotanical evidence in Krassilov, 1973b, 1975a, 1984), correlate with sea-level fluctuations and the related orogenic, volcanic and biotic events. The mid-Cretaceous and Campanian maxima coincide with the larger transgressions, whereas the minima correspond to the orogenic–volcanic pulses of, successively, the circum-Pacific, AustroAlpine, Mediterranean and Laramid tectonism (Krassilov, 1985). Palaeobotanical and sedimentological evidence of polar glaciations, more definitive in the southern high latitudes, appear no earlier than in the late Maastrichtian. The climatic asymmetry of the hemispheres indicates a leading role of geographic factors, such as land/sea distribution, in which they primarily differ, dominating over the atmospheric heat budget (greenhouse) factors. Contrary to the greenhouse theory of global climate change, the black-shale events, a massive carbon sink potentially decreasing the atmospheric CO2 level, overlap the warming phases while the major volcanic events, a source of atmospheric CO2, coincide with the cooling episodes. In the preceding chapters, the concerted geographic changes have been related to the rotation forcing that inflicts a divergence (with acceleration of rotation rates) or convergence (with deceleration) of hypsometric levels for the isostatically compensated oceanic and continental domains. Elevation of the continents is accompanied by marginal faulting and cratonic volcanism, while the rise of sea level is correlated with extensive marginal shear and sea-floor basalts. The greenhouse component of climate change is bound to sea level (VII.2.4) that concomitantly affects the albedo of the earth’s surface (increases with land area), as well as the oceanic heat transfer (interrupted by land barriers) contributing to the net effect. This explains the recorded transgression/warming – regression/cooling correlations. The following climatic scenarios (summarizing VII.2.1-6) are here suggested for each of the trends (Fig. 97). (1) The convergence trend: - The continental crust is inundates by the rising sea, - The thermohaline stratification in the shallow oceans is sustained by the water exchange with epeiric seas, - The deep-water circulation is further restricted by the emerging oceanic rises, with anoxic conditions spreading over oceanic depressions, - The epeiric seas decrease the net earth’s surface albedo, - The epeiric seaways connecting the Tethys with Arctic basins enhance the meridional heat transfer,

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Fig. 97. A schematic interpretation of climate change through the rotation-driven epeiric, orogenic and biotic effects.

- A broad anticyclonic gear generated by the Tethys seas sustains a summer-dry climate over the tropical – lower middle latitudes, - The tropical monsoons weaken due to a high STT and a low ocean to continent temperature gradient, - The atmospheric CO2 level is elevated as the net effect of a decreased land area, reduced terrestrial biomass/pedogenic weathering and the reduced deep-water circulation,


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- The stratospheric ozone is reduced in consequence of a stratospheric cooling related to the tropospheric greenhouse effect. (2) The divergence trend: - The sea-level approaches the boundary of oceanic and continental crust, - The epeiric seaways retreat with consequences for the albedo and the temperature gradients, - The psychrospheric oceanic circulation is set on, with the tropical upwellings enhancing the monsoons bolstering the equatorial biomass production, - The atmospheric CO2 level drops with SST, deep-water circulation and the rise of terrestrial biomass, - An elevation of continental crust over the low-latitude bulge generates a high relief with ridge crests rising above the potential snow line, - The mean global temperatures decrease with montane glaciations. A restricted meridional heat transfer enhances the cooling trend over extratropical areas, - A tilt of the northern continents away from the low-latitude orogenic zone redirects the continental runoff to the Arctic Ocean, diluting surface water and enhancing sea surface freezing, - Polar ice caps appear (spread) with global cooling set on by montane glaciation and an increased freshwater input to the Arctic Ocean, - The ice cap generated polar air vortices affect the mid-latitude climates (temperization) and stratosphere (ozone depletion). The mechanism for glacial autocycles is set up. Essential for this scenario is the understanding that glaciation is a global, rather than polar, phenomenon. Global cooling is initiated by tropical montane glaciations at the time of an equable lowland climate and a relatively high atmospheric pCO2 level. Thus, after the Late Palaeocene to Eocene thermal maximum, the Cenozoic cooling trend starts with the mid-Eocene uplift of the Tethys thrust belt, closely followed by temperization of mid-latitudes in the Late Eocene to Oligocene. A succession of low-latitude orogenic pulses/global cooling events mark each of the evolutionary significant chronostratigraphic levels at about 36, 23, 15, 6, 2.5 and 1.6 Ma, culminating in the dramatic continental glaciation.

VII.4. Displacements of zonal boundaries as evidence of long-term climate change Palaeoecological indicators mostly reflect transient climatic situations that are extrapolated over geological times. In effect, long-term climatic trends can be lost in the noise of minor fluctuations. Long-term trends are better indicated by climatically controlled processes of a comparable time-scale, such as the displacement of vegetation zones.

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The global pattern of vegetation zones is a function of climatic zonation. Zonal climates are generated by the distribution of incoming solar energy over the globe, therefore latitudinal, with deviations owing to geographic effects. The major areas of presentday deciduous and evergreen forests occur north of the O°C and south of the 8°C January isotherms respectively, with a broad (or locally narrow, as in central Asia) ecotone in between (Fig. 98). Their boundaries are defined also by the annual means (above 500 mm for the temperate forest types, except the larch forests of permafrost areas) and the seasonality of precipitation. While the temperature changes with latitude are gradual, in particular during thalassocratic periods of a relatively equable global climate, the precipitation contrasts can be much sharper, accentuating the distinctness of phytogeographical boundaries. In particular, the boundary between the mesotemperate deciduous and xerothermic evergreen vegetation zones is most distinct over the divides of summer-wet and summer-dry climates. This latter boundary is recognized for most of the mid-Palaeozoic to Cenozoic periods on the basis of multiple ecological and taphonomic criteria (II.7, III.1.2), including: - The pachycaul to leptocaul growth form ratio, increasing in the evergreen zone, - Branching habits, such as the heteroblastic with deciduous spur-shoots, - Growth increments, more regular in the deciduous perennials, - The ratios of simple to compound leaves, higher in the evergreen to semideciduous zones, unless decreased by a tendency to leaflet fusion in non-seasonal climates, - Scleromorphic leaf structures, characteristic of evergreen zone,

Fig. 98. Present-day distribution of broadleaved forests against the January isotherms 0o and 8o reflecting an extent of cold atmospheric fronts and the boundary between the equable and seasonal precipitation zones.


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- Foliage wind-roughness associating with deciduousness, - Taphonomic evidence of reproductive structures developing before leaves, - Alpha and beta diversities, both higher in the evergreen zone, - Dominance types, grading from prevailingly mono- and oligodominant in deciduous zone to polydominant in evergreen zone, - Plant-animal interactions, with specialized entomophily more typical of evergreen zone, - Plant litter accumulation types, with leaf mats and leaf coals confined to deciduous zone, - Sedimentological evidence of seasonal input of plant dead mass, such as distribution of organic debris over lacustrine varvites. The zonation was distinct over most of the Phanerozoic except for the relatively brief azonal episodes in the Early Carboniferous, Early Triassic and Early-Middle Eocene.

VII.4.1. Deciduous/evergreen zonal boundary through time Since temperate deciduousness is reflected by both palaeoecological and taphonomic criteria ( the distinctness of growing season increments, wind roughness of foliage, shoot dropping, low diversity, mono- to oligodominance, pollination before leaves, leaf mats), the boundaries of deciduous and evergreen biomes are most reliably documented by the fossil record, even in case of extinct vegetation types. The Middle to Late Devonian archaeopterids shed lateral branch systems (homologous to macrophylls of the later appearing gymnosperms) preserved as the leaf mat-type bedding plane accumulations (Krassilov et al., 1987b). They are dominant over northern North America and Eurasia, with their southern boundary marked by the relatively rare occurrences in Belgium, Bohemia, Donetsk Basin and Minusinsk, about 50°N (modern latitudes). Cuticle coals formed by the thalloid Orestovia-Shuguria type plants are confined to this zone. Over the Late Carboniferous and Permian, the deciduous zone is represented by the Angara-type assemblages, the leaf-mats of Rufloria (“Cordaites”)–Zamiopteris morphotypes and their associated Vojnovskya–Gaussia ovulate heads (IV.3.1). A warmtemperate ecotone extends over northern Kazakhstan, Central Asia, Junggar-Hinggan zone of northern China to southern Primorye, bordering on evergreen vegetation of the Cathaysian realm. The gigantopterid records mark a western outlier of this zone into Asia Minor and Middle East. Over the Late Mesozoic, the boundary is marked by the dominance of heteroblastic taeniophyllous, spur-shoot dropping Phoenicopsis to the north – the pachycaul macrophyllous cauliflorous Cycadeoidea to the south (Krassilov, 1972b). Ginkgophytes are more diverse and numerically prominent in the Phoenicopsis zone while brachyphyllous coniferoids prevail south of it. The Cycadeoidea zone assemblages are more diverse

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and frequently polydominant. Sedimentologically this zone is dominated by redbeds while leaf mats and leaf coals are almost exclusively confined to the Phoenicopsis zone. The Phoenicopsis/Cycadeoidea boundary marks the northern limit of bennettitalean morphotypes Ptilophyllum–Zamites–Otozamites (the Ptilophyllion of syntaxonomic classification). Stem diameters of Equisetum (Equisetites) appreciably decrease north of the boundary (from over 25 mm to less than 10 mm). The thermophilic matoniaceous, schizaeaceous and cyatheaceous ferns Nathorstia, Klukia, Ruffordia, Stachypteris, Cyathea are restricted to the southern zone, rarely penetrating the ecotonal assemblages, as do also the peltasperms (Pachypteris) and araucariaceous conifers. Among the dominant osmundaceous genera, Todites is replaced by Osmunda in the northern realm. Coniopteris species with aphlebial pinnae occur south, those with decurrent pinnae north, of the boundary. The deciduous czekanowskialean-ginkgoalean-coniferous vegetation (the Phoenicopsion) dominated the mesotemperate Arcto-Mesozoic realm from mid-Triassic to mid-Cretaceous when it was replaced by the deciduous conifer rainforest (the Parataxodietum). The deciduous hamamelid broadleaves, Trochodendroides, platanophylls and betuloid serratophylls, supposedly developing as seral stages and understorey of Parataxodium forests, emerged as dominants of a progenitorial Arcto-Tertiary vegetation over the Cretaceous/Tertiary boundary. In the Palaeocene, the Trochodendroidion extended as far south as southern Gobi at about 43°N (Makulbekov, 1988), but retreated to the north with the Eocene warming. The thermophilic Eocene flora with laurophyllous evergreens spread to 60°N over the mainland Eurasia, with marginal extensions to 70°N (VII.5). A reversal of climatic trend with the rise of the Tibet – Himalayas and the central African plateaux in the Late Eocene brought with it a stepwise displacement of the evergreen/deciduous boundary to the south, continued, with fluctuations, over the Oligocene and Neogene. Since the late Neogene, the boundary has remained close to its present-day position over the southern limits of the Euxinian beech-oak forests and the Mandschurian oak forests at about 35°N.

VII.4.2. Climatic interpretation Displacements of vegetation zones might have been the result of polar wander, continental drift or climate change. The maximal amplitudes of displacement are recorded for the Early Palaeogene, with the deciduous/evergreen boundary shifting from about 43°N in the Early Palaeocene to 60°N in the Early Eocene and back again in the Late Eocene to Oligocene. Neither polar wander nor continental drift of comparable magnitudes has ever been inferred for this time interval. Likewise, the pre-Cenozoic shifts of phyogeographical zonation do not correlate with the alleged directions and magnitudes of polar wander or continental drift. All over the Late Palaeozoic and Mesozoic, the boundaries remained roughly parallel to the modern latitudes, which makes the idea of their displacement by continental drift (rotation on sphere) unlikely.


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On the other hand, the appreciable shifts of zonal boundaries, in particular, the better documented boundary of deciduous and evergreen zones, correlates well with the independent evidence for climate change, such as the cycadophyte index for the Early Cretaceous (Krassilov, 1973b) or the entire/non-entire leaf margin ratios for the Late Cretaceous and Paleocene (Krassilov, 1975a). Taphonomic evidence of altitudinal shifts (an increase in upland element in fossil plant assemblages), in particular, a downslope migration of altitudinal belts over the Cretaceous/Tertiary boundary (VII.6, IX.9), agrees with climatically driven latitudinal displacements. The deciduous/evergreen zone boundary is more distinct, and more consistently latitudinal, over continental interiors where its position is less affected by heat exchange with the oceans. For this reason, the displacements of the boundary over geological times can be better assessed for the interior of the larger Asiatic land mass. For this the southern limits of sequential deciduous biomes are shown in Fig 99: Mandschurian – Euxinian, extant – 35°N Arcto-Tertiary, mid- Miocene – 35°N Arcto-Tertiary, mid- Eocene – 55°N a

c

b

d

Fig. 99. The Eurasian temperate deciduous biomes (shaded) over the Devonian to Miocene: a – Archaeopteridion, Late Devonian, b – Vojnovskyion, mid-Permian (P) and Phoenicopsion, mid-Triassic (T), c – Phoenicopsion, mid-Jurassic (J) and Early Cretaceous (K1), d – Parataxodion, Late Cretaceous (K2) and Trochodendroidion, Early Palaeocene (P1), e – polydominant Arcto-Tertiary vegetation, Early-Middle Eocene (P2) and mid-Miocene (N) (after Krassilov, 1997c).

e

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Trochodendroidion, Palaeocene – 43°N Parataxodion, Mid-Cretaceous – 50°N Phoenicopsion, Mid-Jurassic – 35°N Phoenicopsion, Late Triassic – 42°N Ruflorion, Late Permian – 38°N Archaeopteridion, Late Devonian -54°N The resulting curve (Fig. 100) fluctuates between the latitudes of 35°-54° north. The corresponding meridional amplitude amounts to about 2220 km (111 km per 1° of meridian over the mid-latitudes). Since vegetation zones are jointly controlled by temperature and precipitation gradients, their shifts indicate a general direction of climate change, such as temperization or tropicalization, rather than individual climatic variables. Deciduous vegetation spreads with a temperization of the mid-latitudes caused by the surges of polar air generated by the polar ice caps (IV.2.2). Hence the deciduous zone is fairly prominent over the Late Palaeozoic, mid-Jurassic and Late Cenozoic, but shrinks in the essentially non-glacial periods. These are also periods of relatively mild climatic contrasts, or equability, inducing tropicalization of the mid-latitudes (Axelrod, 1992). A penetration of thermophilic elements into the temperate realm is due not only to a lower temperature gradient, but also to a decrease in water uptake with the rise of atmospheric CO2 level rendering plant species less dependent on precipitation (VII.2.4). An interpretation in terms of temperature changes is no more than a crude approximation, warranted in the case of the nemoral zone by a correlation of deciduousness with the cold month component of the mean annual temperatures, which increase by about 1°C per 3° of meridian. A recalculation of latitudes into temperatures gives a cooling effect of about 6°C from the mid-Devonian to present. A prominent Cretaceous

Fig. 100. Climatic curve based on the latitudinal positions of the southern limit for temperate deciduous biomes in Central Asia (Fig. 99). Open peaks correspond to episodes of obliterate zonation supposedly reflecting a greenhouse climate (Krassilov, 1997c).


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peak is only about 1.5°C lower than the Devonian. The mid-Jurassic (Bathonian) trough is about 1°C lower than the Late Permian and of the same magnitude as the mid-Miocene, which is consistent with the glacial mid-Jurassic deposits in northern Asia (Epstein, 1978), as well as a low taxonomic diversity of mid-Jurassic macrofossil assemblages and a drop of thermophilic elements documented by palynological data (Hubbard & Boulter, 1987). The recurrent positions of the boundary at about 50°N (Devonian, Cretaceous, midEocene), 40°N (Permian, Late Triassic, Palaeocene) and 35°N (mid-Jurassic, mid-Miocene, present) correspond to “warm”, “temperate” and “cool” climates respectively. Open peaks in Fig.100 convey the intervals of obscure zonation in the Early Carboniferous (Visean), Early Triassic (Olenekian), Late Cretaceous (Campanian), and Early to early Middle Eocene (Sparnacian – Lutetian), with thermophilic arboreal lycopsids, semiarboreal Pleuromeia, palms and the broadleaved evergreens sequentially spreading over the Arctic Circle. If facilitated by a decrease of water uptake, these anomalous occurrences may suggest a greenhouse climate.

VII. 5. Early Tertiary climates and afforestation–deforestation cycles A culmination of temperization trend at about the Cretaceous/Tertiary boundary (KTB) is indicated by the prevalence of leaf-mat taphonomy and serrate leaf morphology. The northern nemoral zone extended to the southern Gobi (VII.4). Thick conglomeratic sequences, as in the Tsagajan Formation (Krassilov, 1976a), reflect a major erosion cycle driven by a drop of sea level. The poorly sorted coarse-grained deposits, with clay inclusions, suggest slumping of waterlogged debris transported by melt water from a highland snow cover. A steep altitudinal temperature gradient is also indicated by differentiation of vegetation belts upslope, with a succession of polydominant conifer forest to monodominant birch woodland as in the present-day cool-temperate zone. A downslope migration of the birch belt in the Early Palaeocene of Sikhote Alin (Krassilov, 1989c; VII.6) is evidence of an upland cooling ahead of the lowland temperization. Vegetation change had commenced in the Maastrichtian already with the entries of numerically subordinate Palaeocene newcomers, such as the small-leaved Trochodendroides, the biserrate Corylites and other betuloid morphotypes (Krassilov, 1979; Golovneva, 1996; Herman, 1999), supposedly invading the pioneer to early successional stages of disturbed riparian vegetation. These vegetation events coincided with an increase in tectonic/volcanic activity. In eastern Asia, elevation of the marginal volcanic belt in conjunction with ignimbritic eruptions and the emergence of volcanic island arcs might have contributed to the climate change. However, a coeval spread of Nothofagus (Nothofagidites pollen grains) to the southern mid-latitudes (Doktor et al., 1996) indicates the global extent of temperization toward the end of Cretaceous. The early Palaeocene assemblages reflect a taxonomically poor monodominant vegetation (Krassilov, 1976).

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The subsequent fossil plant records indicate a warming trend accompanied by an increase in both α and β diversity until the Late Palaeocene (Thanetian) warm-temperate forest climax has been reached across the continental interiors (e.g., the Late Palaeocene flora of the Volga–Uralian Region dominated by a protofagoid genus Ushia, related to Nothofagus: Krassilov et al., 1996b), with a slightly higher content of evergreen elements along continental margins (e.g. in the contemporaneous floras of Sakhalin: Krassilov, 1979). The so-called Late Palaeocene Thermal Maximum (LPTM) about 1 m.y. before the Palaeocene/Eocene boundary, PEB (the Thanetian/Ypresian boundary in Europe) has been inferred on the basis of a prominent carbon isotopic event that was provisionally ascribed to a thermal dissociation of marine methane hydrates indicating a short-lived greenhouse episode (Dickens et al., 1995; Beerling and Jolley, 1998). However, the evidence of thermal dissociation is meagre if any at all (Katz et al., 2001). Uncorking of gasohydrates might have been due to tectonic instability of continental margins and metamorphic outgassing at about the PEB (e.g., the Alaskan margin: Hudson & Magoon, 2002), of a longer-term climatic effect. The “LPTM” does not correlate with any significant biotic event, unless it is confused with the long-term warming trend over which the traditional PEB was marked by a brief cooling recorded in both marine and non-marine sequences (Pardo et al., 1999; Wing et al., 2000; Kelly et al., 2001). A warming trend over the Early to early Middle Eocene is most evident in Western Europe and North America (Chateauneuf, 1980; Plaziat, 1981; Boulter, 1984, Wolfe, 1994, Wing et al., 1991; Wing, 1998; Collinson, 2000; Collinson et al., 2000b; Huber et al., 2000; Sloan & Hueber, 2001; Sloan et al., 2001). The terrestrial vegetation of this interval is described as a tropical or paratropical multistratal rainforest that extended to 60°N, with evergreen components spreading to 70-75°N (Wolfe, 1985; Collinson, 2000). The rainforest interpretation of the “megathermal” assemblages is to a considerable extent based on their mangrove components, the geographic ranges of which are typically restricted by winter sea surface temperatures of 20°C. Yet mangal communities extend to northern Florida at 30°N and to southern Australia at 38°S, sometimes fringing open landscapes of continental interiors. A spread of mangroves, in particular over the Atlantic coasts, indicates a sea surface warming (reflecting a peculiar oceanographic situation, below) confirmed by the oxygen isotope ratios (Shakleton & Boersma, 1981; Kobashi et al., 2001). In turn, a high flux of methane from mangal ecosystems (Westlake, Kvet & SzczepaDski, 1998; Lu et al., 1999) might have contributed to the climate warming. The freshwater arboreal wetlands were ubiquitously represented over the Late Palaeocene to Early Eocene by the Taxodium–Nyssa type assemblages that are prevailingly extratropical. Conifer remains, including Sequoia and Metasequoia, are relatively common. The angiosperm leaves are relatively small, of the microphyll–nanophyll size classes, xeromorphic, with thick cuticles (Krumbiegel et al., 1983). In association with such morphologies, the apical extensions can be interpreted as prickly tips rather than driptips. In the dominant angiosperm groups, such as Cornales, presently including both shrubs


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and tall trees (Mastixia of montane rainforests), the fruit–seed remains prevail over leaf remains, a taphonomy indicating shrubland rather than rainforest. At the same time, the mammalian turnover, with a modification of the crown types in the larger herbivores over the PEB (Janis, 2000; Bowen et a., 2992), indicates a change to a more open Eocene vegetation. According to a recent palynological analysis by Harrington (2001), the vegetation change was insignificant, with the Early Eocene communities slightly more diverse and “more successional”. However, a drop of Tricolpites, a dominant palynotype of the Palaeocene nemoral biome, and a rise of palms as well as herbaceous monocots that increased by about 80% indicate a change to more open vegetation. The carbonate rhizocretions of thick red palaeosols indicate a seasonality of precipitation. In the Bighorn Basin, Wyoming (Fig. 101) the reworked palaeosols contain silicified conifer wood. A subtropical climate with pronounced seasonal draughts is inferred also for the Sparnacian of Central Europe on the basis of oxygen and strontium isotope records (Schmiz & Andereasson, 2001). The palynological, taphonomic and sedimentological data are thus complementary in suggesting a major deforestation event. In a broader phytogeographic context, the anomalous advance of thermophilic vegetation to high latitudes is a regional trans-Atlantic, rather than global, phenomenon. Chro-

Fig. 101. Redbed sequence over the Palaeocene– Eocene boundary, Bighorn Basin, Wyoming (photograph taken during field trip of the International Conference on “Climate and Biota of the Early Paleocene”, Powell, 2001).

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nologically, it coincides with a major thrusting phase in the peri-Atlantic fold belt, dated as the Late Palaeocene to Early Eocene on Svalbard (Saalman & Thiedig, 2001) and elsewhere, and the appearance of the trans-Atlantic land bridge (McKenna, 1975; Tiffney, 1985a, 1985b) that prevented penetration of cold Arctic waters. In consequence, the Atlantic surface waters would have been overheated suppressing monsoonal circulation, hence dry continental interiors, with xerophytic/helophytic evergreens, but lending a little chance for rainforests. With break-up of the bridge later in the Eocene, a psychrospheric circulation was rapidly restored entailing a retreat of evergreen vegetation. In the western Pacific sector, the mid-latitude early Paleogene assemblages reflect warm-temperate forest vegetation, prevailingly deciduous, with a subordinate evergreen element occurring in the coal-bearing facies of the Amur Basin, western Primorye and Sakhalin, but vanishing further north. Thus, the Early Eocene floras of Sakhalin at 50°N contained Magnoliaephyllum as a prominent laurophyllous element while the contemporaneous floras of northern Kamchatka at about 60°N were dominated by the Metasequoia–Trochodendroides–Corylites assemblage with entire-margined morphotypes either lacking or constituting no more than 10 per cent of leaf counts (Krassilov, 1976c). The assemblage is relatively large-leaved, containing a considerable proportion of compound leaves of Aesculus and Carya types, with broad leaflets. On western coast, palm leaves were found in a fairly restricted stratigraphic interval (Budantsev, 1983) perhaps reflecting a brief penetration of palm thickets as a narrow riparian belt, like in the warmtemperate vegetation of northern Florida. This might have been due to intermittently continuous Bering Land Bridge, related through a geodetic regression ((VI.1) and synchronous to the trans-Atlantic land bridge (above), with a similar, though less prominent, climatic effect. An alternative possibility is a translation of continental fragments with thermophilic plant assemblages from south over the system of sinistral strike-slip faults controlling the block structure of northwestern Kamchatka (Harbert et al., 2001). In contrast to evergreen expansion on the margins, the continental interior shows a more regular pattern of vegetation zones. In India, the Intertrappean Early Palaeocene flora is of a mixed aspect, comprising both tropical and temperate (Ginkgo, Taxodioxylon) elements (Meher-Homji, 1980), the latter perhaps representing a riparian vegetation. The Late Palaeocene to Eocene floras are of a more tropical aspect, with mangroves (Guleria, 199; Mehrotra et al., 1998). A dipterocarp-like wood has been found in the Intertrappean deposits and above, yet the earliest authentic appearance of dipterocarps is recorded from the Oligocene of Borneo (Muller, 1970). The laurophyllous vegetation of southern China was confined to approximately the same latitudes as today (about 25°N). The dry subtropical belt with Palibinia extended over central China at about 35°N (Guo, 1985). The humid to seasonally dry subtropical belt with Sabalites and Dryophyllum was displaced to about 40°, about 10°N of its present-day position over the Yangtze River. The Raychikhinsk flora of Amur Basin at about 50°N, and the contemporaneous Uglovskaya flora of Primorye, represented a warm-temperate broadleaved forest with subtropical elements (Akhmetiev et al., 1978).


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The Early to early Middle Eocene Tastakh flora of northern Siberia (Kryshtofovich, 1958; age based on palynological and dinoflagellate records: Kulkova, 1987; Grinenko, 1989) is broadleaved platanophyllous, of leaf dimensions larger than average for this type assemblages, reflecting a temperate climate of high growing season humidity. Evergreen morphologies are lacking among leaf remains, yet a few palynological records attest to occasional thermophilic forms (Kulkova, 1987). All in all, the Early Palaeocene climate change is not reducible to a greenhouse effect while an increase in global temperatures might have been less dramatic than inferred from the regional European – North American records. A relatively mild climatic amelioration over continental interiors (also in the Southern Hemisphere: Wilf et al., 2001) suggests that a greenhouse effect might have been less significant than, and perhaps consequential to, the regional geographic (oceanographic) effects. In particular, an emergence of the trans-North Atlantic land bridge resulted in anomalously high sea surface temperatures allowing for an expansion of mangroves to high latitudes while inflicting a seasonally dry climate over continental interiors. Open vegetation is suggested by both plant and vertebrate records. Yet at the same time the western Pacific margin remained densely forested, supporting a substantial coal accumulation. An increase in average leaf size and the prominence of compound pinnate leaves with large leaflets indicate an ample precipitation during the growing season implying a high summer monsoon intensity. A comparison of contemporaneous fossil plant assemblages bears on the problem of teleconnection between the Atlantic and Pacific climates. For the ice-free Eocene the teleconnection scheme might have been different from that of the recent past or the Pleistocene (Kiefer et al., 2001). In particular, over the Dansgaard-Oeschger cycles of Greenlandic ice-sheet fluctuations, an increase in eastern Asiatic monsoon corresponded to cold Atlantic phases (Garidel-Thoron & Beaufort, 2001). In contrast, the above correlations suggest a coincidence of summer-dry North Atlantic phases with peaks of monsoon intensities in the western Pacific during the Eocene worming. Later on, the broadleaved deciduous forests of eastern Asia expanded to continental interiors with temperization over the Eocene/Oligocene boundary, almost regaining their Early Palaeocene ranges (VII.4). The following waves of a warmer drier climate caused a recurrence of thermophilic laurophylls, as in the Miocene (Burdigalian – Llanghian) of Central Europe. A teleconnection of these mid-latitude events with tropical climates might have been prevailingly atmospheric, rather than oceanic as over the Pleistocene. The Maastrichtian developments in the Niger – Benue Basin and the lower Intertrappean plant assemblages of India suggest an early rainforest stage corresponding to the major mid-latitude temperization. Next temperization stage concurred with the appearance, in the Oligocene, of fully developed tropical rainforests with dipterocarps (IV.2.1). While induced by climate change, the afforestation/deforestation trends might in turn contribute to it through their effects on the carbon cycle. Seral stages consume more CO2 than old-growth forests, hence an initial afforestation would inflict a drop of atmospheric pCO2 level recovered at culmination of the trend and raised above the time-

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average with deforestation. A significance of these correlations may only be revealed with a more exact chronology of the Tertiary cycles.

VII.6. Mid-latitude Tertiary cyclicity in eastern Asia To account for both temperature and precipitation variables of climate changes, vegetation responses inferred from fossil plant assemblages have to be added by depositional and taphonomic evidence. It is also desirable to restrict the analysis by the less ambiguous responses. In practice, the following set of palaeoclimatic criteria proved useful in distinguishing the recurring combinations of temperature and precipitation over the Tertiary climatic cycles (Table 3). The stratigraphic sequence of plant assemblages is informative when related to a catenic succession, as in the case of Fagus replacing Castanea upslope and upsection (below).

VII.6.1. Succession of plant assemblages Geological setting. The Russian Far East east of the Siberian Platform is a mosaic of cratonic blocks welded by ophiolitic fold belts. The pre-Cretaceous sedimentary baTable 3. Indicators of Tertiary climate types in the Russian Far East. Climate

Warm humid, WH

Warm dry, WD

Cool humid, CH

Cool dry, CD

Diversity

High

Moderate

Moderate

Low

Dominance

Poly dominant

Poly dominant

Olygo dominant

Oligo/mono dominant

Non-arboreal element

Rare

Common

Rare

Common

Taxodiaceae

Common

Rare

Common

Rare

Pinaceae

Rare

Rare

Common

Common

Typical leaf morphology

Simple

Compound

Simple/ compound

Simple

Mean leaf width

c. 60 mm

c. 35 mm

c. 45 mm

c. 30 mm

Entire/serrate leaf margin ratio

High

Moderate

Low

Low

Leaf mats

Lacking/rare

Lacking

Common

Rare

Thick coal

Present

Lacking

Present

Lacking


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sins are confined to cratonic margins or residual troughs over their sutures. A new structural pattern of concurrent tectonic zones emerged in the mid-Cretaceous and persisted over the Tertiary (Fig. 102). The volcanic ranges extended from Chukotka to Primorye and further south to Korea and China. Thick volcanomictic molassoids accumulated in intermontane depressions over the transcurrent fault zones. A foredeep of thick miogeosyncline deposition is traceable from the Anadyr Bay in the north to the Tatar Strait and western Sakhalin. The island arc/trench belt of eastern Kamchatka, eastern Sakhalin and the Kuril Islands is marked by overthrusts and ophiolitic melanges. The eastern flank of the foredeep is now raised as the coastal ranges of western Sakhalin. The trough is filled with deltaic clastics derived from Okhotsk Land in the north while periodically inundated by shallow seas from the south. A zone of interfingering marine and non-marine deposits at 50°N contains the most instructive Cretaceous – Tertiary sequences, including the Khoinjo Section. Khoinjo Section. For most of the Far Eastern Tertiary basins we have composite stratigraphic sections based on scattered outcrops or coal quarries. The coastal Khoinjo Section of

Fig. 102. Tectonic setting of the Cretaceous–Tertiary plant localities in the Russian Far East.

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northwestern Sakhalin (Figs. 103, 104) is exceptional in spanning all the stages, from the Maastrichtian to the mid-Miocene, in a condensed monoclinal sequence continuously exposed in steep coastal cliffs extending about 20 km over the Khoinjo Point. It contains the stratotypes of several lithostratigraphic units (suites, or formations) described by Kryshtofovich (1936), who recognized the Eocene and Miocene fossil plant assemblages. A number of floristic horizons have been added by subsequent studies (Akhmetyev, 1978; Krassilov, 1979; Krassilov & Kundyshev, 1982). The following analysis is based on the joint palaeofloristic and magnetostratigraphic work by Krassilov, Schmidt and Remisovskiy (1986c). Schematically, the Khoinjo Section includes: (1) Cretaceous marine shales with inoceramid shells, more than 20 m. (2) Sandstones with thin coal lenses, 4 m, unconformable on (1), containing the Metasequoia–Trochodendroides assemblage. (3) Conglomerates, gravels and coarse detrital sandstones (Conglomerate Formation), about 100 m, with several floristic horizons representing the Dryophyllum– Trochodendroides assemblage. (4) Coarse to fine-grained sandstones, 25-30 m (Lower Member of the Lower Due Formation), with the Magnoliaephyllum–Dryophyllum assemblage.

Fig. 103. View of the Khoindjo Secton, Cretaceous to Miocene, Tatar Straight coast of western Sakhalin (1981).


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Fig. 104. Schematic drawing of the Khoindjo (I) and Chekhovka (II) sections, Sakahalin, with fossil plant and invertebrate localities. PK, P3P2, NP, N12N11, N13N12 – the Cretaceous/Palaeogene, Palaeocene/Eocene, Palaeogene/Neogene, Early/Middle Miocene and Middle/Late Miocene boundaries, respectively. CD, CH, WH, WD – climate phases (explanations in the text).

(5) Coal beds, 15 m (Upper Member of the Lower Due Formation), with three major coal seams separated by ferruginous shales yielding three successive assemblages: (5a) Platanus–Aesculus–Platycarya, (5b) Alnus onorica–Populus celastrophylla, and (5c) Metasequoia–Zelkova ungeri–Trochodendroides. (6) Marine black shales (Gennoishi Formation) with sanidine crystals (Gennoishi of Japanese researchers), containing Oligocene invertebrate assemblages, about 80 m. (7) Basalts with pillow structures, tuffs (Khoinjo Formation), about 60 m, overlain by the coarse-grained tuffaceous sandstones, 3m. (8) Black shales with conglomerate interbeds, 25 m, containing the Picea–Betula– Ulmus protojaponica arboreal assemblage and grass remains. This is either included in the Khoindjo Formation or considered as the basal member of the Upper Due Formation above. (9) Sandstone/shale alternation, about 80 m, with thin coal beds and with seven successive plant-bearing horizons of the Acer–Ulmus–Carya broadleaved assemblage. (10) Ferruginous sandstones and shales, 12 m, with the Quercus furuhjelmii–Castanea miomollissima assemblage. (11) Coal beds, more than 50 m, with brackish Corbicula shales above the uppermost coal seam and with numerous floristic horizons containing the arboreal Fagus antipovii–Acer–Byttneriophyllum and aquatic Nelumbo–Hemitrapa assemblages.

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The Khoinjo section thus represents two large non-marine sedimentary cycles divided by the marine Oligocene horizon. There are gaps in the floristic succession at the unconformable Cretaceous/Tertiary boundary and over the non-fossiliferous volcanic rocks at the base of the second cycle requiring supplementary data from other sections. Cretaceous/Tertiary boundary (KTB). The basin-wide unconformity over the KTB is caused by tectonic uplift and regression over the miogeosyncline trough of western Sakhalin, with a few continuous transboundary sequences preserved in residual depressions. One such site is located 100 km south of Khoinjo Point, near the Boshnyakovo Village, where the Maastrichtian to Palaeocene deposits crop out on steep banks of the Augustovka River. The black shales of the Pachydiscus gollevilensis–P. subcompressus zone, Upper Maastrichtian, are overlain by the paralic sandstones with the Paralaxodium–Trochodendroides assemblage of a transitional Cretaceous – Tertiary aspect, followed by the fine-grained tuffs with the Metasequoia– Corylites–Trochodendroides assemblage still retaining Cladophlebis arctica, Nilssonia and some other Cretaceous relicts (Krassilov, 1979). Further south, the tuffaceous non-marine deposits are replaced by the shallow marine Sinegorsk Horizon of the basal Palaeocene foraminiferal zone Rzehakina epigona (Kalishevich et al., 1981). This level is missing in the Khoindjo Section. On the Lesser Kuril Islands to the east, the Maastrichtian turbidite deposits with inoceramids (Inoceramus baltica, I. shikotanensis) are overlain by the island-arc volcanics cut by swarms of parallel dykes and intruded by plagiogranites and gabbroids of a dismembered ophiolite assemblage. On the Yuri Island north of Hokkaido, the turbidites are intruded by basaltic sills and covered with basalts, tuffobreccias and agglomerates. A shale horizon at the base of the island-arc complex contains foraminifers and macrobenthos, as well as plant macrofossils and palynological assemblages of a transitional Cretaceous/ Tertiary aspect (Krassilov et al., 1988). The foraminifers belong to the Rzehakina epigona zone. The palynoflora is dominated by Gothanipollis santaloides, Ulmoideipites spp., Tricolpopollenites, Myricites, etc. associating with the diverse Aquilapollenites spp., Orbiculapollis, Cranwellia and other Cretaceous elements amounting to about 11 % of the spectrum. The plant macrofossils constitute the Sequoia reichenbachii – Corylites protoinsignis assemblage, comparable to the Augustovkian taphoflora of Sakhalin (above). Fossil wood of Palaeogene conifers is found in the overlying tuffobreccias. In the cratonic basins west of the marginal volcanic belts, a relatively complete nonmarine KTB sequence is known from the Tsagajan Formation of Amur Province, an alluvial conglomerate/sand/clay megacycle resting on the lacustrine Upper Cretaceous shales (Krassilov, 1976a; Bugdaeva, 2001). The lower coarse-grained member contains dinosaur remains accompanied by the Maastrichtian palynological assemblage with Orbiculipollis globosus. The middle clayey member contains the transitional macrofossil Taxodium–Trochodendroides–Nordenskioldia and microfossil Tricolpiles–Ulmoideipites assemblages, with Protophyllum, Triatriopolleniles, Triprojectacites, Orbiculapollis and other Cretaceous survivors, correlated to the Augustovkian taphoflora of Sakha-


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lin. Leaf-mats of the upper sandy member represent a less diverse Trochodendroides– Tiliaephyllum assemblage, correlated with the Khoinjo unit (2). As a whole, the early Palaeocene is characterized by the leaf-mat type taphonomy, high proportions of conifers, in particular the Taxodiaceae, a moderate to large leaf size, and a high ratio, up to 75 %, of serrate leaves. These features indicate a CH climate phase (Table 3). Late Palaeocene – Eocene assemblages. The coarse-grained member (3) of Khoinjo Section consists of poorly sorted fan conglomerates and mudflow deposits cut by well-sorted channel sandstones, with residual oxbow facies of avulsion cycles. These latter contain plant assemblages of scattered leaf remains, dominated by the narrow serrate Dryophyllum and laurophylls (Cinnamomophyllum) while conifers are rare. In accord with a sclerophyllous aspect of the plant assemblage, the debris flows indicate scantily vegetated slopes. A WD phase is inferred from these data. The unit (4) assemblage is relatively macrophyllous (mean leaf width ca. 60 cm) and contains a more prominent lauroid element (including Magnoliaephyllum) thus recording a relatively humid phase of the warming trend, culminated by a thick coal accumulation. Parallel floristic successions are encountered over the Palaeogene sequences of Snatol Basin, western Kamchatka, with rich Palaeocene to Eocene plant localities of Napan and Snatol Formations (Krassilov, 1976c). The Metasequoia–Trochodendroides– Corylites assemblage of the lower Napan member is enriched upsection by the entries of Parrotiopsis, Acer, Pterocarya, Aesculus, Nyssa and a few magnolioid morphotypes representing a WH phase. Similar developments took place in the intracratonic depressions of the Amuro-Zeyan and Khankian blocks over the Russian/Chinese border, where the WH phase is represented by the clayey Raychikhinsk and Uglovskaya formations respectively, with minable coals and polydominant floristic assemblages containing a significant lauraceous component (Akhmetiev et al., 1978). Eocene – Oligocene transition. Magnetostratigraphically, unit (4) represents an interval of reversed polarity with brief episodes of normal polarity (Fig. 105), correlated with magnetochrones C15/C13 (Ness et al. 1980). The lower coal seam of unit (5) above falls in the relatively long interval of normal polarity supposedly corresponding to a magnetic anomaly at the top of C13. The Eocene/Oligocene boundary is placed between the anomalies 15 and 13 at about 36.5 Ma (Poore et al. 1982). A major floristic change occurs above the lower coal seam of unit (5), from the roof of which came a broadleaved assemblage dominated by compound leaf morphotypes with double-serrate leaflets, as in Aesculus magnificum and Platycarya hokkaidoana. Preponderance of the latter is evidence of a switch to a cooling trend. Thick coals (two adjacent coal seams 2.5-3 m and 5-6.5 m thick) would have required an ample year-round precipitation. Osmunda beds in the roof and splits of the second coal seam reflect a fern growth over the open mire. The arboreal remains represent a bordering vegetation, floristically like those of the preceding stage, though less diverse and with the dominant species relatively small-leaved (mean leaf width ca. 40 cm).

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Above the second coal seam, the normal polarity zone C12 (35.5-32.5 Ma) contains two successive plant assemblages. The lower Alnus omorica–Populus celastrophylla assemblage is characterized by the minimal leaf size for the whole sequence, representing a CD episode terminating peat accumulation. The Metasequoia–Zelkova ungeri– Trochodendroides assemblage above is considered mesothermic on account of zelkova and relatively humid on account of conifers. This brief climatic fluctuation preceded the mid-Oligocene transgression (Gennoishi Formation). Oligocene-Miocene transition. The non-marine equivalents of the Gennoishi Formation are poorly represented in the miogeosyncline trough of western Sakhalin that was inundated by the mid-Oligocene sea. The Oligocene to Lower Miocene sequences are turbiditic, with volcanomictic material in the upper members. Profuse eruptions, like those of Khoinjo

Fig. 105. Palaeomagnetic zonation over the Eocene–Oligocene transition, Khoindjo Section, Tatar Strait coast, western Sakhalin (Krassilov et al., 1986c).


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Formation, are recorded also in eastern Sakhalin (Tchekhovskaya Formation) as well as in eastern Kamchatka (Uspenskaya Formation) and on the Greater Kuril Islands, indicating a major phase of igneous activity controlled by a series of transcurrent faults. In the Krynka Section of eastern Sakhalin, the volcanoclastic Tchekhovskaya Formation on top of Oligocene turbidites is stratigraphically equivalent to the Khoinjo basalts (unit 7). A rich plant assemblage of white tuffs in the lower part of the Tchekowskaya sequence (Krassilov et al., 1984) contains abundant ferns (Osmunda sachalinensis, Woodwardia arctica), conifers (Glyptostrobus, Pityostrobus) and angiosperms (Ceratophyllum, Myrica, Fagus, Coryliles, Nyssa, Fraxinus, etc.). Marine bivalves are found below and above the plant-bearing horizon. Dominant elements of the assemblage represent a modern type bald cypress – black gum coastal wetland (of which green ash and wax myrtle are constant associates), with fern marshes and a pioneer hazel growth over clearings. Expansion of the Coryletum might have been owing to frequent ashfalls and fires. The plant-bearing laminated white tuffs fall in a reversed polarity interval while the massive tuffs and tuffaceous conglomerates above (60 m) represent a mixed polarity interval identified as the standard polarity zone 21 (Theyer & Hammond, 1974). It is followed by a long normal interval in the lower part of the coal-bearing sequence (Fig. 105) assigned to the standard polarity zone 19. This zone is detectable in most of the regional Miocene sections. It comprises several floristic horizons corresponding to the Khoinjo units (8) – (11). Abundant Pityostrobus from the lowermost horizon reflects a spread of a spruce bog-forest during a brief CH episode better represented by the Khoinjo unit (8). Miocene cycle. At the Khoinjo Point, the black shale member with intervening conglomerates (8), resting on basalts and tuffs, was deposited in a stratified rift lake of a gaping transcurrent fracture zone. The plant remains are scattered all over the sequence rather than forming distinct plant beds. Small betuloid and ulmoid leaves predominate in association with spruce and fir samaras as well as grass achenes. A high proportion of wind-borne debris suggests an open vegetation, savanna or a sparse woodland, with a conifer forest on adjacent slopes. This assemblage appears polydominant owing to plant remains coming from various sources. On the basis of prevailing leaf morphologies and a high proportion of non-arboreal elements it is assigned to a CD phase, perhaps corresponding to a 1 m.y. long Aquitanian cooling (Alvinerie, 1980). The subsequent plant assemblages indicate a recurring warming trend culminating with the narrow-leaved Quercus–Castanea assemblage of unit (10). It is followed by a series of mixed taxodiaceous–broadleaved assemblages over the Upper Due coal measures, with Fagus among the dominant hardwood genera. In the contemporaneous Miocene deposits of Primorye, Castanea is replaced by polymorphous Fagus chankensis in the Khanka Lake tuffaceous deposits (Krassilov & Alexeyenko, 1977). Since in the present-day succession of altitudinal vegetation belts (e.g. on Honshu and elsewhere in the mountain range of southwestern Pacific) chestnut forests are systematically replaced by beech forests upslope, such replacements over stratigraphic sequences signify a downslope shift of respective altitudinal belts.

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The Late Miocene sequences (the Razdolnaya Formation in Primorye: Krassilov & Alexeyenko, 1977) show a decline of lowland beech forests, with climatic conditions approaching a CD phase. For comparison, over the Paratethys realm, the Miocene warming culminated in the late Burdigalian–Llanghian (Navarotti & Bicchi, 1996), with a subsequent cooling in the Serravallian to Tortonian and an increasingly dry climate in the Messinian.

VII.6.2. Four-phase climate change The vegetational/climatic succession based on the Khoinjo and supplementary sequences is summarized in Table 4. Four mid-latitude cycles of about 30, 23 and 7 Ma correspond to three typical periodicities of the Phanerozoic scale (V.7). Only the larger cycle is recognized in the condensed lower part of the sequence, with resolution increasing upsection. Nevertheless, the CH (TH)-CD-WD-WH sequence recurs over each of the cycles (a missing WD of the second cycle corresponds to a regional transgression). Each phase is in turn a succession of subordinate phases the relative prominence of which defines its general aspect. This pattern can be traced to a k.y. scale (e.g., over the terminal Pleistocene and Holocene: the Younger Dryas CD, Boreal WD, Atlantic WH, Subboreal CH to CD, Subatlantic CD to WD) and, on a minor scale, for the last two millennia (Table 5).

Table 4. Succession of fossil plant assemblages reflecting cyclic climate change in the Tertiary of Russian Far East. Climatic stages as in Table 3. Razdolnaya, Alnetum CD Khoinjo 11, Khanka, Acero-Fagetum CH ___________________________________________Mid-Miocene,15 Ma Khoinjo 10, Castaneo-Quercetum WH Khoinjo 9, Ulmo-Aceretum WD Khoinjo 8, Piceo-Ulmetum CD Krynka, monodominant Piceetum CH ___________________________________________Oligo/Miocene, 22.5 Ma Krynka, Glypto-Coryletum WH Stage missing WD? Khoinjo 5, Trocho-Zelkovetum TH Khoinjo 5, Trocho-Alnetum CD Khoinjo 5, Platano- Aesculetum TH ____________________________________________Eo/Oligocene, 36.5 Ma Khoinjo 4, Dryo-Magnoliaephylletum WH Khoinjo 3, Trocho-Dryophylletum WD Khoinjo 2, Trochodendroidetum CD Augustovka, Trocho-Parataxodietum CH _____________________________________________KTB, 66 Ma


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Table 5. Four-phase climate cycles of different scales. Climate stages as in Table 3. Phase CH

Ma 1.8

Ka

Holocene

Century

Historic

4.3 6 9.5 13 16

Neogene Piacenzian/ Calabrian Zanclean Messinian Tortonian Serravallian Llangian

WH WD CD CH WH

0.7 3.5 4.5

Sub-Atlantic Sub-Boreal Atlantic

XXI-XXII XIX-XXI XIII-XIX IX-XII

Near future Modern II Little Ice Age Viking Warming

WD CD

19.5 21

Burdigalian Aquitanian

6.8 11

Boreal Younger Dryas

IV-VIII I-III

Great Migrations I Little Ice Age

The present-day warming repeats that of the 1940’s with a brief cooling event in between – the minor fluctuations over the CD/WD transitional interval. The millennial dynamics suggests that we have at least about two centuries of mid-latitude WD before us. The following interpretation of a lower-rank four-stage climatic cyclicity (Fig. 106) for a glacial climate (as argued in VII.3, the climate was mildly glacial over most of the Phanerozoic history) is derived from the theory of atmospheric disturbance as a leading factor of climate change (Lorenz, 1967; Palmen & Newton, 1969; Wallace & Blackmon, 1983). WH: Glacioeustatic sea-level rise and melting of sea-borne ice enhance the maritime polar front surges, resulting in an intense mid-latitude cyclogenesis. The meridional heat transfer is intensified decreasing the temperature gradient. The atmospheric pCO2 level rises with SST and the expanding thermocline. CH: Polar ice caps expand owing to snowfalls brought by cyclonic air masses to Arctic lands. The polar atmospheric fronts advance with the snow cover. The Arctic air surges push cyclonic vortices to the lower mid-latitudes entailing an increase in precipitation and expansion of forests. The atmospheric pCO2 level drops with SST and afforestation. CD: Meridional heat transfer is restricted by the advancing polar fronts. The tropical/ temperate temperatures diverge: an overheating of the tropical zone is accompanied by a cooling of the temperate zones. Later in the phase, the ice caps shrink in consequence of a reduced snowfall. WD: Outbreaks of overheated tropical air shift the subtropical highs to the mid-latitudes where the persistent meridional anticyclone/cyclone couples are formed (Wallace & Blackmon, 1983). Their blocking of the westerlies (as observed in the northern midlatitudes during El-Nño: Davidson et al., 1990) results in deforestation of continental interiors and a build-up of atmospheric CO2 (also recorded during the El-Niño). The four-phase model implies an autocyclic restriction of meridional heat transfer resulting in overheating of the tropical zone – cooling over the mid-latitudes, with consequences for land biota and the carbon cycle. A restricted heat exchange generates zonal temperature anomalies that are resolved by atmospheric disturbances stirring an interzonal heat transfer. This mechanism explains the abruptness of climate changes, a phe-

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Fig. 106. A schematic interpretation of four-phase climate cycle over tne northern mid-latitudes related to the shifting positions of polar front (PF), southern front (SF), cyclonic/anticyclonic air masses, and the alternative roles of the ocean as sink or source of atmospheric CO2.


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nomenon brought to focus by Lorenz (1967), Lamb (1977), Zubakov (1990) and elsewhere. Atmospheric CO2 fluctuations are considered as an integral part – an enhancer – of the processes (Fig. 106).

VII.7. Summary of climate change Since climate affects both biotic and sedimentological aspects of the biospheric carbon cycles, the proxies of climate change are countless, yet all of them potentially change their meaning with evolution, hence ambiguous. Direct inference from indirect correlations, as between climate and isotope fractionation, produce mythic thermal maxima and minima (e.g., the “LPTM”: VII.5). Likewise, too much emphasis on quantification of palaeoclimatic variables is scarcely productive. Mean global temperatures are as informative of climatic situations as mean temperatures for a hospital are of the patient’s conditions. Anyway, robustness of interpretation is preferable over numerical detail. Longterm climatic evolution is better reflected by the processes of comparable rates, such as displacements of vegetation zones (VII.4). The geological record of past climates can be more objectively interpreted in terms of climatic gradients rather than global means. What is recorded as a global climate change is primarily a redistribution of heat over latitudes, with contrasting situations of (1) gentle equatorial to polar temperature gradient and a likewise low precipitation gradient of the opposite sign or (2) overheating of tropical zone – freezing of polar areas, with precipitation reallocated to the polar (ice caps) and equatorial (rainforests) zones. The latter situation is brought about by orographic barriers, such as mountain ranges over the elevated Tethys belt. Glaciation is a global, rather than polar, phenomenon. At a high stand of the continents correlated through rotational mechanisms (V.8) with a low-latitude orogeny, glaciation might have commenced in the tropical highlands later spreading to the polar lowlands. Over the earth’s history, the first-order climatic cycles correspond to alternations of the Tethys seaway geography with the Tethys mountain range geography. The latter is consequential to the major pulses of the Taconian, Hercynian and Alpine orogenies correlated with the Late Ordovician, Late Carboniferous–Early Permian, and the Plio-Pleistocene glaciations. Their ca.180 m.y periodicities correspond to the Galactic Year (Steiner & Grillmair, 1973). The subordinate climate cycles, down to 4-year scale, are triggered by the earth’s rotational and orbital forcing and are amplified by the causally related geographic and oceanographic effects. Polar glaciations inflict temperization of mid-latitudinal climates promoting a spread of deciduous biomes. Insofar as extensive zones of temperate deciduousness are recorded for the greater part of geological history since the mid-Devonian (VII.4), a glacial (mildly glacial, as at present) climate should be considered as a norm while the icefree situations are atypical.

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The fossil record furnishes ample evidence of hierarchial climatic cycles while a directional change over times is less convincingly documented. In theoretical models it is commonly related to a decrease of atmospheric CO2 concentrations. However the thermal capacity of atmosphere depends not so much on the total amount of CO2 emitted as on its reallocation to the oceanic and lithospheric reservoirs mediated by the biota (VII.2.4). It is theoretically conceivable and partly supported by geological evidence that period to the build-up of terrestrial biomass in the mid-Palaeozoic, the climates were more fluctuant globally, while the latitudinal climatic zonation was less distinct than later in the earth’s history. Over the Meso-Cenozoic times, the larger climatic fluctuations were related to episodes of global tectonic instability and large-scale sea-level fluctuations inflicting a reallocation of terrestrial biomass, as over the afforestation–deforestation cycles (VII.5), superimposed by autocyclicity (VII.6). What we now perceive as a directional change is by all probabilities a small segment of a cyclic trajectory. Human activities have a two-fold impact on climate, first, through technogenic emissions of greenhouse gases, which add to their release from the naturally exhumed hydrocarbon reservoirs, and, second, by removal of biomass which is superimposed upon the natural afforestation–deforestation trends. Human impact can only be assessed in conjunction with the natural climate-generating processes. In any case, the industrial emission effect is but transient, counteracted by a complex system of biospheric sources and sinks, while a removal of biomass may have the longer-term consequences. Our long-term climatic predictions must take into consideration both the heat-generating and heat-transfer factors, with the latter having a substantial priority. The idea of deglaciation due to greenhouse effect has so far obtained no supporting evidence from palaeoclimatology. Rather, the glacial/interglacial cyclicity would terminate with the currently observed tectonic activation of the Tethys and circum-Pacific belts potentially inflicting a radical change in geography, such as a subsidence of continental margins or re-opening of the Tethys seaway.

Chapter 8. Ecosystem evolution

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VIII. ECOSYSTEM EVOLUTION When we speak of evolution (of organisms, human personality, ideas, etc.) as different from a mere change, we seem to imply a succession of changes bringing about an essentially novel state of things or affairs. Biological evolution involves both structures and processes at the hierachial biospheric, ecosystemic, coenotic, populational, organismic and genomic levels. Human interests in evolution have traditionally been focused at morphological variation, the adaptational aspects of which have been explored since Aristotle. This line of investigation culminated in the Lamarckian theory of adaptive change. Charles Darwin introduced a population approach that soon jostled the organismic studies (in particular, the developmental studies, the focus of evolutionary research in the early 1800’s) out of fashion. A new dimension was added to population studies by the Mendelian genetics that was incorporated into the Darwinian (now neo-Darwinian, or synthetic) paradigm. Each of the single-level models was in turn proclaimed as the only possible theory of evolution. In effect, the most significant evolutionary phenomena, the origins, the extinctions, the progress of life, the build-up of biomass, diversity and structural complexity, remained unexplained, or they were explained away as either non-existent or occasional, or else owing to intrinsic laws that defy explanation. Ecological theories of climax vegetation as a coenotic equilibrium (Clements, 1916), the ecosystem as a coherence of interacting organisms and their physical environments (Tansley, 1935; recent review in Golley, 1993), and the biosphere as a global system bound by the biogeochemical cycling of matter and energy (Vernadsky, 1967 and earlier publications) were scarcely considered as relevant in respect to biological evolution. Admittedly, each of the single-level systems is capable of a self-organizing development (Kauffmann, 1993). Organism develops over a series of ontogenic stages to a genetically determined condition changed by mutagenic impacts. Population strives for the Hardy-Weinberg equilibrium disrupted by genetic drift or selection. The biotic community evolves, over a succession of pioneer–seral stages, to a potential climax, at which the spontaneous development ends unless the climax is disrupted by environmental impacts. An external impulse is likewise required for each single-level system to advance beyond the limits of self-organization. Hence the single-level theories deal with short-term processes, a mere extrapolation of which fails to explain the long-term evolution (macroevolution) of geological times. Macroevolution differs from microevolution not only in scale but also in being an outcome of interactions between the structural levels. To explain macroevolution, a theory must consider the origin and spread of evolutionary impulses in multilevel systems. Classical biology scarcely recognized supraorganismic systems at all, which did not help evolution to find ecology (Haila, 1989). Charles Darwin (1859) considered the possibility of accelerated evolution rates at times of low sea level, but only in relation to the alleged imperfection of fossil record. Meanwhile his contemporary A.R. Wallace (1870)

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put forth a seminal, but at the time not fully appreciated, idea of a causal link between biological diversity and environmental stability. This was, in fact, a basic postulate of ecosystem evolution theory (see Krassilov, 1995a). Ecosystem includes both living and non-living components that are driven thermodynamically up the negentropic and entropic tracks, respectively (Glandsdorff & Prigogine, 1971; Prigogine, 1980; Kondepudi & Prigogine, 1999). In most general terms, evolution is due to a disbalance of these opposite tendencies, with negentropic processes prevailing over progressive developments interrupted by entropic episodes of backsliding. Evolution impulses are thus generated at the ecosystem level then spreading to the populational–organismic–genomic levels. My attempt at advancing an ecosystemic evolutionary model started with a definition of coherent (in stable ecosystems, encouraging co-adaptation and structural complexity) versus non-coherent (in disturbed ecosystems, inflicting a tendency of ecological expansion and developmental acceleration) evolution modes (Krassilov, 1969). Ecosystem disturbances, by the tectonomagmatic, eustatic and climatic impacts globally, enhance non-coherent processes culminating in ecological crises. A decrease in stabilizing selection over the non-coherent phases releases latent variability disrupting the developmental homeostasis and providing a raw material for adaptive evolution at the next coherent phase. The model accounts for both directional and cyclic processes (Krassilov, 1973a, 1978b, 1985, 1994b, 1992a, 1995a). Further, ecosystem evolution was analysed in terms of few critical variables, the biomass, dead mass and diversity. A build-up of biological diversity was interpreted as an optimisation process, increasing the efficiency of ecosystem as the biomass-producing machine (Krassilov, 1992a, 1995a and elsewhere). Biomass fluctuation and reallocation in response to global change were related to biological diversity through the respective adjustments of population strategies and the top-down regulation of organismic and genomic developments (Krassilov, 1995a, 1997b, 1998a). Premises of the theory are as follows: – Biospheric recycling of matter and energy is a goal-orientated (telic) process producing living matter from non-living matter. Photosynthetic, consumptive and reproductive structures evolve for the purpose of efficient biological production. – Ecosystem as a living matter-producing unit tends to increase its standing biomass (B), at the same time decreasing its wastes, the dead mass (M). The ratio of B to M production rates is a measure of ecosystem’s efficiency directly correlated with its structural complexity conveyed by the biological diversity (S). Progressive ecosystem evolution means that for a time intervals t1 – t2, Bt2/Mt2 > Bt1/Mt1 and, proportionately, St2 > St1. – Upon diversification trend, the demographic strategy evolves from more to less redundant which allows for a decrease of niche overlap and non-competitive co-existence at the population and (in humans) individual levels. – Morphological complexity of dominant life forms is related to the ecosystem complexity through co-adaptations that develop in stable environments. A build-up of structural


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complexity inflicts a tendency of decelerate development at both the ontogenic and seral levels. Conversely, a loss of structural complexity results in accelerate development. Genome evolution is affected by these processes through the developmental genes. – Discordant fluctuations of B/M and S are imposed by environmental disturbances inflicted by the concerted tectonomagmatic, sea-level and climatic responses to the planetary rotation/revolution cycles. Biological developments are thus bound to geological processes as an integral part of planetary evolution. – With evolution, the biota increasingly confers a stabilizing effect on their physical environments. This tendency would potentially culminate in the human care of the earth. Evolution thus makes sense as an increasing potency of ecosystem (biosphere) survival. In pre-Darwinian time, evolution was thought of as unwinding of a preconceived succession of life forms. Indeed, ecosystem evolution, as a telic process, directs its components to a preconceived state of optimal cooperation that provides for effectiveness and sustainability. Hence the ecosystem evolution is predictable.

VIII.1. Biotic communities as evolution units As discussed in (II.2), the biotic communities are bound by their persistent trophic, spatial and developmental patterns (the seral/catenic in the case of plant communities), co-adaptations and the concerted responses to environmental change. Ecological roles (niches) are fairly predictable over both the trophic cascades and seral sequences. The vacant niche concept means that the ecosystem structure is preconceived.

VIII.1.1. Anticipation Organismic evolution is preconceived because adaptations are anticipative. In essence, adaptation is an improvement of biotic/environmental interactions. Since any interaction affects both sides, all adaptations are actually co-adaptations the history of which goes back to primeval ecosystems of prebiotic molecules. In microbial mats already, cyanophytic photosynthesis generates H2 and CO gases that affect both chemotrophic and heterotrophic members of the community (Hoehler et al., 2001). Such basic co-adaptations have been transmitted to all subsequent structural levels. Plants as primary producers were preceded by microbial autotrophs that coevolved with microbial heterotrophs upon the lines of predation, saprophagy and mutualistic integration. Endosymbiosis appeared at this stage giving rise to the functionally integrated systems of cell organelles (Raven, 1980). Plants entered the pre-existing autotroph–heterotroph systems, initially with little effect on their structure or functioning, converging upon the bacterial autotrophs. Even the present-day lichenicolous plants (green algae) are similar to the lichenicolous bacteria

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(cyanophytes) in their symbiotic roles and habits. Both phylogenetic analysis and palaeontological evidence indicate great antiquity of lichens that gave rise to various nonlichenized groups of fungi (Lutzoni et al., 2001). Fungi emerged as dominant herbivores and detritivores long before animals (see Taylor, 1990 on various forms of plant – fungi interaction in the Devonian). In symbiotic systems, fungi facilitate plant propagation as animals do. The Devonian fungi developed a tree habit (Prototaxites: Hueber, 2001) in parallel to, or even in advance of, the higher plants. Soil appeared long before the higher plants. Traces of pedogenesis go back to the Precambrian (Retallack, 1986a; Retallack et al., 1984). The Ordovician palaeosols lack plant roots but contain mycelial tubes as well as deep burrows (Retallack, 1985a; Feakes & Retallack, 1988) suggesting their origin as the fungal–arthropod systems. The Late Silurian records of diverse littoral communities of higher algae and algae-like plants (Fig. 107), with probable ancestral forms of higher plants among them, suggest that air-borne

Fig. 107. Before the conquest of land: A littoral assemblage from the Late Silurian of Podolie, Ukraine; (a) a limestone quarry near the Kudrintsy Village; the girl stands on a dark organicrich layer with algal and conodont remains, (b) slab covered with Caulerpatype fronds resembling an early land plant Baragwanathia, (c) cleared frond showing a filamentous structure, (d) a verticillate-type frond from the same assemblage.

a

b

c

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spores might have been produced by emerging sporangia of some transitional semiaquatic forms. With plant material appearing as a rain of air-borne spores, saprophytic adaptations might have developed before the advent of land plants in the dispersed spores – fungal systems. In the same way, the plant – arthropod interactions might have commenced with the detritivory later involving living plant tissues. Massive rhizomes of early land plants intruded the pre-existing fungal–arthropod soil ecosystems to give rise to the Devonian rhizosphere (an early example in Boucot et al., 1982). As a rule, the endophyte–host interactions are more advanced and more often systemic (rather than local) in underground than aboveground plant organs (Bogle et al. 2001). In early land plants, the feebly developed photosynthetic organs were compensated for by the relatively advanced underground symbiotic systems. Mycorrhizal structures are known in the Early Devonian helophytic Aglaophyton already (Remy et al., 1994). Since that time, mycorrhizal systems seem to have played key roles in plant adaptation to terrestrial environments (Van der Heijden et al., 1998).

VIII.1.2. Co-adaptation The still popular individualistic concept of biotic community holds that organisms of different species live together because their environmental requirements happen to overlap (Gleason, 1926). In contrast, the holistic concept claims pre-eminence of co-adaptation as a reciprocal adaptive response of co-existing species to each other conditioned by their mutual response to their physical environments. The post-glacial history of vegetation in Europe is commonly thought to illustrate individualistic species behavior, with a succession of species invasions being determined by their migration rates alone. However, this theory has recently been put to serious doubt (Tinner & Lotter, 2001). The early Holocene abundance of Corylus avelana represents an early successional phase reflecting unstable climatic conditions with frequent droughts. The subsequent invasions of Abies and Fagus signal the development of climax communities reflecting a climatic stabilization. Such successions are fairly predictable on the basis of seral sequences. Upon migration routs, structural patterns are maintained by differential seed dispersal, in which the ecology of seed dispersers plays a leading role, their replacements inflicting structural changes (Huwe & Smallwood, 1982). This mechanism provides for expansion of plant communities as co-adapted systems rather than loose aggregates of independently migrating species. Actually, immigration of a plant species means an invasion of symbiotic system, including the rhizospheric components, with constraints imposed not only by this species tolerance per se, but also by that of its mutualists (Christian, 2001). A division of ecospace into the tolerance niches of individual species is superimposed by their interaction niches. A plant species and its pollinator each have their own toler-

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ance ranges, yet their interaction at pollination is a sensitive stage of their coordinate life cycles more stringently limited by certain trophic and climatic factors. Therefore, the interaction niche is narrower than either of the species niches per se. Hence, diversity of a co-adapted system is determined primarily by interaction niches (e.g. the mycorrhizal niches or pollination niches that actually control distribution of plant species). The diversity in Palaeozoic plants with edible pollen grains is demonstrably higher than in other contemporaneous species, and the interactions with xyelids appear as a factor of proangiosperm diversity (VII.1.3). The individualistic species concept has been recently coupled with the concept of genetic neutrality into a “unified neutral theory” to explain what appears as random distribution of species, in particular, of plant species in tropical rainforests (Hubbell, 2001). Instead, the apparent randomness may prove to be evidence of exceptionally strong intracommunity ties. Since population is controlled by a number of factors from below as well as from above, a prevalence of one or another factor decouples it from other limitations. Thus, in insect-pollinated plants, population density is controlled by pollen loads on stigmas (Molano-Fiores & Hendrix, 1999), which depends on pollination efficiency, while the trophic constrains by light, phosphorus, nitrogen, etc., are rendered increasingly less important. Both the genetic polymorphisms and physiological adaptations related to these factors are thereby neutralized. Does the “unified neutral biogeography” reflect such a decoupling situation rather than intrinsic neutrality? With co-adaptation, the intraspecies competition is becoming to prevail over the interspecies one as a factor of plant distribution. Consequently, the species tend to minimize their population densities while entering an increasing number of interspecies configurations. This seems to be a situation in the tropical rainforests of high species richness, low population densities and high tolerances of other species simulating random distribution. Climatic adaptations are based on the communal, rather than individual, responses. For many arboreal plants climatic constrains pertain to the requirements of their mycorrhizal symbionts. Tree lines in the polar and montane areas are determined primarily by the limits for the rhizosphere as a symbiotic structure. The notion of zonal vegetation as an assemblage of plant species with similar climatic requirements is a reductionist view of a higher rank co-adaptive system of plant species, their symbionts and the retinue of phytophages. Red flowers of tropical plants conform to the visual preferences of vertebrate pollinators, whereas purple flowers of temperate plants are entomophilous. The causal link is not climate – colour, but climate – pollinator – colour. In the same way, the production of carbon-based allelochemicals, notably the tannins typical of deciduous plants (Karban, 1989), varies with pest pressure, and only indirectly with climate. Climatic impacts are mediated by biotic interactions. Thus, cool wet summers invigorate the boreal conifer forests by inhibiting insect outbreaks rather than directly (Bonan & Shugart, 1989). Trophic strategy is likewise determined by biotic interactions. In a stable plant/herbivore system, plant growth is restrained by herbivores before it is restrained by a deficit


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of nutrients. An excessive supply of nutrients is detrimental owing to its destabilising effect on the plant/herbivore system: while the nutritional quality of plant tissues increases with nitrogen enrichment, the herbivores burst and crash the plant population. Erratic biomass fluctuations occur over a series of overgrazing events until the system comes to a stable state (Rosenzweig, 1971; DeAngelis et al., 1989). Fossil plant record documents a long-term persistence of structural patterns, as in the recurrent successions of Ceskanowskietum → Phoenicposietum in the Triassic to Early Cretaceous (ca. 100 m.y.) catenic/seral sequences (III.3.2). Major evolutionary changes of plant communities are owing to additions, deletions or substitutions of structural levels. These modes were illustrated in the preceding chapters, e.g. the loss of lycopsid wetlands in the Permian and their reappearance as Pleuromeia wetlands in the Triassic (IV.2.6); a substitution of Jurassic fern marshes by proangiosperm communities over the Jurassic/ Cretaceous boundary (III.3.6); a segregation of progenitorial Arcto-Tertiary elements at both the lowest (the riparian Trochodendridion) and the highest (the montane Betulion) catenic levels in the Early Palaeocene (IX.9); a replacement of scale-leaved helophyte communities by the scleromorphic Debeyo-Dryophylletum and the invasion of palm thickets at the same catenic level in the mid-Cretaceous (III.3.6), etc. Such events are part and parcel of biospheric turnovers inflicting a restructuring of vegetation systems.

VIII.1.3. Herbivory and plant evolution A straightforward plant reaction to herbivory is defensive, such as prevention of herbivore access to palatable plant tissues (spines, thorns, deterrent exudates), visual determent (pubescence, serrate leaf margins) and the reduction of palatability or nutritional value by cutinization, lignification or concentration of secondary metabolites that interfere with digestion. With evolving mutualistic relationships, more sophisticated responses have appeared, such as an excessive production of palatable material (e.g., pollen grains), a compensatory growth (most common in grasses: McNaughton, 1979) or even an increase in fitness after injury (as in Pinus strobus: Puettman & Saunders, 2001). Plant–herbivore co-evolution is a generator of new ecological niches, hence an enhancer of biological diversity, with positive feedbacks. In diverse communities, co-evolution acquires a “diffuse” mode of a multilateral, rather than pairwise, interaction of plant/ herbivore species, and the results are essentially different (Janzen, 1980; Fox, 1981; Pilson, 1996; Inouye & Stinchcombe, 2001). The top-down regulation of herbivores by predators is becoming even more essential for plant life than the bottom-up regulation by plant competition for trophic resources (as proved by ecological disorder in the predatorfree communities: Terborgh et al., 2001), a plausible explanation of “random” configurations of plant species (above). Since plant debris produced by aquatic plants entered terrestrial ecosystems before the advent of land plants, detritivory might have appeared before herbivory. At least in

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the early terrestrial ecosystems, detritivory prevailed over herbivory, which is consistent with their relatively high dead mass production rates. In the later history of terrestrial biomes, a significance of detritivory varied with detritus to biomass ratios of carbon storage, in turn dependent on plant growth rates, as well as on the biotic turnover rates (Cerrbián & Duarte, 1995). In the importer-type ecosystems (II.1), such as estuaries, a contribution of detritivory to carbon cycle is proportionate to the ratio of inorganic to organic nutrients in the terrestrial runoff (Obviate et al., 1986; Smith & Hollibaugh, 1993; Kemp et al., 1997; Lin et al., 2001). In mangal ecosystems, detritivory prevails over herbivory because of a high chlorine concentration in living plant tissues. Over evolutionary history, this factor might have constrained herbivory in the coastal Carboniferous wetlands as well. Yet about 20% of mangrove biomass is consumed by insect herbivores. Mangrove roots branch at damages, co-adaptation of a kind. The detritivorous communities of mites, collembolans and millipedes seem to have but little changed since their first appearance in the Palaeozoic (Llabandeira et al., 1988; Scott & Paterson, 1994; Shear & Kukalová-Peck, 1990). Their relative significance for carbon cycle fluctuates with the soil CO2 concentration that is dynamically equilibrated to the atmospheric CO2. Coprolites with spores are abundant in the Late Silurian and over the Devonian (Edwards et al., 1995) attesting to the role of sporivory (presently rare, detected in weevils feeding on tree ferns: Shear & Kukalová-Peck, 1990) as an initial type of arthropod herbivory, conceivably related to an exceptionally high ratio of spore production to vegetative body mass in early land plants, decreasing later on. Defensive responses to sporivory included a perennial habit, an aboveground elevation of sporangia (upright habit), their shift from terminal to the better protected lateral (axillar) position, spines on shoots and on spores, thick sporoderms, etc. Yet at this early stage already, spines with terminal digits (Fig. 108) suggest adaptation to spore dispersal by sporivorous insects. Heterospory might have also appeared in relation to sporivory, securing a few thickwalled megaspores at the expense of abundant microspores. The free-sporing heterosporous plants has been fairly successful and diverse until the piercing of megaspores was invented later in the Palaeozoic (Sharov, 1973; Scott & Patterson, 1994). With an increasing loss of free megaspores to arthropod herbivory, several lineages of spore-plants switched in parallel to endosporangial germination. With this, protection of megasporangia became a priority, involving sterile structures, as in the telomic Lenlogia, the sporangial clusters of which were protected by whorls of terminal branchlets (Krassilov & Zakharova, 1995). As the next step, such structures evolved into the spiny integuments or cupules of early seed plants. In this round of plant–arthropod co-evolution, reproductive structures remained a major target, with both pollinivory (Krassilov & Rasnitsyn, 1998) and seed boring (of Trigonocarpus: Scott, 1992) developing during the Palaeozoic. A new stage of co-evolution commenced with the seed plants discovering the benefits of arthropod visits. The protective devices then became rapidly transformed into the

Chapter 8. Ecosystem evolution a

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Fig. 108. Arthropod attacks on early land plants: chewing (a, b) and piercing (c, d) injuries in Psilophyton from the Early Devonian of Gaspé (courtesy of Harland Banks).

attractive ones. Free integumental lobes protecting pollen chambers from insect damage were fused into micropyles exuding sweat pollination drop while an extra pollen mass was produced as a reward for pollinivorous pollinators (Krassilov & Rasnitsyn, 1998). Likewise, the secreting glands, abundantly developed in early seed plants as repelling structures, were modified as nectaries in the insect-pollinated bennettites and gnetophytes (extant species of the latter are visited by the nectar-collecting short-horn flies: Ren, 1998). Copious resin ducts of gnetophytic Heerala, Eoantha or Gurvanella (Fig. 109) might have alternatively performed a repelling or attracting function in response to different insect species or at consecutive stages of reproductive cycle. To decrease the risk of reproductive failure by damage from a massive arthropod attack, perennial reproduction has developed in all lineages of free-sporing Devonian plants and was later transmitted to spermophytes. A consequential increase in both photosynthetic and mechanical tissues has radically changed the ratio of spore (ovule) production to body mass. With this, arthropod herbivory gradually extended to the less nutritious but more abundant vegetative tissues, as attested by Palaeozoic examples of leaf cutting (Scott & Taylor, 1983). Leaf (petiole) galls first appeared in Palaeozoic pteridophytes (Van Amerom, 1973; Scott et al., 1994; Llabandeira & Phillips, 1996) and evolved into typical zoocecids in Mesozoic gymnosperms (Ash, 1999). Biochemical impacts of

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Fig. 109. Resiniferous pollen cone of a gnetalean proangiosperm Loricanthus from the Early Cretaceous of Baisa, Transbaikalia: (a, b) resin bodies in pelate sporangiophores, (c, d) in situ pollen grains.

insects on their host plants is reciprocated in gymnosperms by production of juvenile hormones preventing maturation (invoking paedogenesis?) in herbivorous insects (Shear & Kukalová-Peck, 1990). In the Carboniferous pteridosperms, ovules developed on little if at all modified leaves. Yet with a spread of leaf cutting (marked on the Pennsylvanian Neuropteris already: Scott & Taylor, 1983, but more common in the Triassic: Ash, 1997, 1999), reproductive structures were dissociated from foliage, with a corresponding divergence of their insect retinues (Fig. 110). Over the Palaeozoic to Mesozoic transition, both male and female reproductive structures were simultaneously strobilized in all major seed plant lineages. While protecting sporangia and ovules from indiscriminate leaf/sporophyll cutting, strobili provide shelter as a new stimulus for insects to co-operate. Early tetrapod herbivory entailed elaboration of food webs (Olsen, 1973). The plant– arthropod systems have further developed as the plant–arthropod–tetrapod systems, with derived co-adaptations, such as reciprocal leaf–insect mimicry (Odontopteris pinnule – cockroach wing: Scott & Taylor, 1983, butterfly wings – Vitimantha pre-flower bracts: Fig. 111). While arthropod herbivory encouraged strobilation, the opposite destrobilation tendency was inflicted, likewise simultaneously, by endozoochory in the ances-


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Fig. 110. Leaf-eating in Mesozoic biotic communities: (a) Phasmomimoides (Phasmatoptera), a folivorous insect feeding on brachyphyllous coniferoids: (b, d) cuticle fragments from compressed gut contents showing characteristic stomata, and (c) Classopollis pollen gain of the same plant adhering to the leaf cuticle. The Late Jurassic of Karatau, Kazakhstan (Rasnitsyn & Krassilov, 2000).

tral lines of Ginkgo, Cycas and flowering plants. These developments have also affected the pollination strategy (VIII.7). Selective pressure of tetrapods on plant communities increases with body size, endothermy and the diversity of dental types. Dense populations of larger herbivores depend on resilience and high growth rates of their forage plant communities, hence for them an open vegetation is preferable to closed forests. Large herbivores in turn propagate plant communities that conform to their trophic requirements. Thus, browsers promote open pioneer vegetation by destroying forest canopies. In this way herbivory generates new vegetation structures. As a rule (the Jarman–Bell rule), the dietary tolerances are wider in larger than in smaller herbivores. Driven by a greater demand for food, larger browsers consume a

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Fig. 111. Mimicry in an early lepidopteran insect, Undoplerix sukatchevae (top), and a gnetophytic proangiosperm, Vitimantha, showing the wing-like perianth bracts. The Early Cretaceous of Transbaikalia.

lower quality leaf mass at higher canopy levels, pruning leafy shoots, while smaller browsers pluck the better quality leaves below. A vertical browsing zonation arises in this way (Woolnough, & du Tait, 2001), enhancing the adaptive divergence of size classes. In effect, over times, populations of larger browsers, such as sauropods, might have suffered from total reduction of leaf mass with replacement of vegetation types, while the smaller guild was more sensitive to quality changes brought about by evolution of the forage plant species. These considerations seem pertinent to diachronous extinctions of herbivorous dinosaurs and mammals over the Cretaceous – Tertiary boundary. Oral processing of tough plant tissues is recorded in the Late Tatarian anomodonts (Rybczynski & Reisz, 2001) attesting to a rapid advance of plant–herbivore systems prior to the Permian/Triassic crisis. In the Jurassic and Early Cretaceous, a highly productive open vegetation of fern-marshes and cycadophyte shrublands was sustained by co-adaptive interactions with herbivorous sauropods (Russel et al., 1980). New feeding habits appeared with, and contributed to, the mid-Cretaceous vegetation change. A re-


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duction of leaf/xylem ratio with extinction of macrophyllous cycadophytes induced dental innovations in the Late Cretaceous ornithopods, their posterior occlusive dentition and toothless beaks increasing the range of palatable plant tissues (Norman & Weishampel, 1985; Wing & Tiffney, 1987). Semenivory and frugivory might have evolved in respect to the ornithopod feeding habits that were reciprocated by endozoochorous adaptations in plants in turn preparing the herbivore niches for the early mammals (e.g. a hard seed gnawing in the mid-Eocene: Collinson & Hooker, 2000) Plant–animal interactions are commonly recognized as a leading factor in the origin of flowering plants and the build-up of their morphological and biochemical diversity. With flowering plants, pollination ecology has become a focal point of ecosystem evolution. Multistratal canopies of angiosperm communities have greatly enhanced animal biodiversity while, on the other hand, certain angiosperm-dominated biomes, in particular the grasslands, are sustained by herbivory. With the rise of flowering plants, the herbivore–vegetation systems (such as the grazer–grassland systems) took control over plant evolution intercepting the climate–vegetation systems. Yet, by the appearance of flowering plants, all the essential modes of plant–arthropod–tetrapod interactions were there, ready for the newcomers to pick up. In this way, diversification of a new group is prearranged by the antecedent co-adaptations.

VIII.1.4. Coenotic gene pool An as yet not fully appreciated aspect of co-evolution is related to the role of exogenous nuclear acids, conceivable since the fundamental discovery by Gershenson (1939). The genomes of multicellular organisms are not impenetrable to genetic material scattered over biotic community in the form of DNA and RNA molecules stored in soil and transduced by various microorganisms (Kordum, 1982). Recent genetic research has revealed the ubiquity of microbial gene transduction as a powerful mutagenic factor (Syvanen & Kado, 1998; Vorontsov, 2000). Gene pools of co-existing populations are thereby included in the higher order, coenotic, gene pool of the biotic community. The coenotic gene pool concept may prove useful in elucidating the origins of symbiotic systems. Thus, plant–fungi interactions involve genetic transduction as a method of symbiotic integration. Plants having ectomycorrhizal symbionts produce elicitors of fungal growth that comprise chitinases of fungal cell walls (Salzer et al., 1997), a biochemical convergence owing to a one way or possibly reciprocated flow of genetic material. Since physiological responses induced by symbionts are more flexible and less costly than constitutional changes, symbiosis plays a leading role in plant evolution. Plant – fungi symbiosis is the most ancient one, with primitive plants penetrating the mycelia and vice versa. Certain plant – fungi interactions, in particular, with an involvement of haploid spores and pollen grains, are fairly persistent over times (Fig. 112). The developing fun-

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Fig. 112. Saccate pollen grains with zoosporangia of chitrid fungi. The Early Cretaceous of Mongolia (Krassilov, 1982).

gal sporangia consume the genetic material of a spore cell nucleus, potentially trunsducing it with propagation of zoospores. In the Late Silurian to Early Devonian, the early land plants appeared in the midst of microbial communities of the rootless soil and were rooted there by way of mutualistic co-adaptation (VIII.1.1) Mycorrhizal symbiosis had already developed at this early evolutionary stage. Extant mycorrhizal fungi still induce lignification of host tissues, a reaction going back to the origin of terrestrial plant habits, with plant lignins owing to a fungal transduction of genes eliciting a benign plant response (as in the case of chitinases: Salzer et al., 1997). Lignification then might have evolved as a defence against arthropod damage eventually acquiring a mechanical function associated with arboreal habit. A widespread phenomenon conceivably related to interspecies gene transfer is the so-called geographic parallelism, a spread of peculiar, selectionally neutral, characters over a range of not closely related species within a geographic realm (Went, 1970). In the geological record this phenomenon acquires a time dimension as the chronogeographic parallelism over a geographic realm/stratigraphic interval. Thus, the Cordaitestype linear-lanceolate leaves with subparallel venation prevail in the Late Palaeozoic fossil plant assemblages of Euramerian, Angarian and Cathaysian realms. On evidence of associated reproductive organs, this leaf morphotype occurred in different gymnosperm orders, such as the Cordaitales, Vojnovskyales and Peltaspermales. Glossopteris, a lingulate leaf morphotype with reticulate venation widespread in the Gondwana continents, appeared not only in the nominate order Glossopteridales but also, simultaneously, in the co-existing cycadophytes (IV.3). The Permian Gigantopteris and allied morphotypes, with composite leaf blades of marginally fused segments and areolate venation, belonged to different gymnosperm and marattialean taxa of the Cathaysian (Cathamerian) realm. The Mesozoic Taeniopteris-type leaf morphology is shared by at least three gymnosperm orders, the Cycadales, Bennettitales and Pentoxylales.


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Even more spectacular are the palynomorphological “fashions”, such as the taeniate protosaccate pollen grains in the Permian or the tricolpate reticulate morphotypes in the mid-Cretaceous (Fig. 113). The taeniate morphotype was produced by glossopterids, peltasperms and coniferoids while the tricolpates simultaneously appeared in the platanoid and ranunculoid lines of early angiosperms (Krassilov & Golovneva, 2001). In both these examples, the exinal characters might have either volume-regulating or signalling functions, therefore not selectively neutral. Still it remains puzzling why, of all the functionally equivalent structures, just these were independently selected in several seed plant lineages at a particular time level.

Fig. 113. Seed plant phylogeny as a succession of evolutionary grades (Krassilov, 1997a). Parallel developments (right margin) across the lineages are marked with horizontal dash lines.

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A possibility of horizontal gene transfer as an explanation of chronogeographic parallelisms and accumulation of innovative traits has been discussed in relation to angiosperm origins (Krassilov, 1997a and earlier, Syvanen, 1994). It has been further suggested (Krassilov, 1998b) that a role of horizontal transfer might have increased in situations of ecological crises. With disruption of coherent community structures the stabilizing selection is suspended thereby increasing the chance of a successful transfer. Moreover, for a crisis survivor species, an invasion of exogeneous genetic material is potentially advantageous, rather than detrimental, because, by extending the limits of genetic variation, it facilitates ecological expansion. Herbivory assists in maintaining the coenotic gene pool by transmissting microbial pathogens. Such transmissions are either occasional or they are required for mutualistic interactions, as in the case of stem-boring weevils transmitting rust fungi (Frideli & Bacher, 2001). Pollinivory is a special mode of such interactions. Horizontal, over a number of parallel lineages, spread of morphological characters, in particular the monogenic palynomorphological traits of haploid generation, has been shown to associate with pollinivory (Krassilov & Rasnitsyn, 1998). On evidence of mixed pollen loads, a horizontal spread of such traits as the taeniate exine involved the phylogenetically not closely related plant species that grew in close proximity to each other as members of persistent biotic communities. These findings point to a potential role of horizontal gene transfer as a factor of co-adaptation, owing to which the biotic community evolves as a unit.

VIII.2. Biomass Biomass is the product of ecosystem, a testament of its functional efficiency. In equilibrial ecosystems, biomass is sustainable within the limits of autocyclic fluctuations. Data on biomass production in the present-day vegetation types, measured as net dry mass, are available from a number of special studies and compilations (Basilevich & Rodin, 1969; Duvigneau, 1971; Whittaker, 1975a; Songwe et al., 1988; Oliver & Larson, 1990; Basilevich, 1993; Lawson, 1994; Westlake et al., 1998). Biomass increases with body size of dominant growth forms. In tropical forests it is about 10 times larger than in tropical shrublands (460 – 500 t ha-1 in the Ivory Coast semideciduous forests against 6.8 t ha-1 in the shrublands: Lawson, 1994). Yet a large size is only advantageous to a certain limit set up by a combination of various factors, such as the efficiency of carbon fixation by photosynthesis, light competition, longevity, reproductive strategy, etc. Among them, the efficiency of carbon fixation tends to decrease with size owing to accumulation of non-photosynthetic material in axial organs, roots and shaded branches that turn into an energy drain. To reduce energy losses, large trees regularly shed a considerable dead mass as spurshoots, lower branches, water-stressed shoots cut off at the basal constriction zone (Sprugel et al., 1991), bark, etc. As mentioned before, shoot shedding occurs in a number of


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pioneer riparian trees, such as poplars. It was fairly common in the Mesozoic gymnosperm forests while the bark-droppers prevailed in the Carboniferous lepidophyte forests. In the ecologically marginal environments (near the tree-line), large plants tend to degrade to a lower habit (Stevens & Fox, 1991) with biomass reallocated to underground parts. In grasslands, the underground biomass amounts to about 25-58% of the total (Titlyanova et al., 1996). It might have been even larger in the rhizomatic early land plant communities.

VIII.2.1. Biomass build-up and fluctuations The evolution of ecosystems, and of the biosphere as a whole, is directed toward the goal of efficiency in producing living matter from non-living matter with a minimal waste in the form of dead mass. Hence, the biomass to dead mass ratio tends to increase with evolution. This theoretical expectation is supported by the fossil record. The Early Proterozoic microbial mats of a few cyanobacterial species produced a disproportionately large dead mass preserved as shungites. They also contributed to accumulation of banded iron ores amounting to about 20% of contemporaneous sedimentation. The subsequent Late Proterozoic mat communities were much more diverse, yet they left a relatively modest dead mass (Hofmann, 1976; Bauld, 1981). The appearance of terrestrial plant life entailed a multifold increase in the total biospheric biomass that was further elevated with the terrestrial vegetation expanding from wetlands over drylands. Large trees appeared already in the Late Devonian but they were telomic rather than leafy, thus producing a much smaller above-ground biomass than modern trees of a comparable size. On evidence of telomic mats, the arboreal archaeopterids were shoot-droppers (IV.3). In addition, most Devonian plants were rhizomatic, and even in the Carboniferous arboreal wetlands a contribution of positive geotropism, forming massive stigmaria, was considerably greater than in the extant ecological analogues. Fossil fuels and kerogen-rich deposits are accumulations of buried dead mass. The biomass of extinct biotic communities can be approximately assessed by analogy with their extant ecological equivalents (as in the Cretaceous example, VIII.2.3). When an ecological equivalence of extinct and extant communities is problematic, as in the case of Carboniferous peatland vegetation, comparisons can be made on the basis of arboreal to non-arboreal plant ratios, tree size and diversity as the biomass correlates. Thus, the Carboniferous wetland forests were sources of huge coal measures, indicating high dead mass production rates, while their extant analogues accumulate a relatively small dead mass. They were inferior to the present-day tropical rainforests both in the tree/herb ratio (about 50% against 70%) and tree size (up to 40 m against 60 m), indicating an altogether smaller biomass. The Carboniferous assemblages from a single locality, typically accumulating plant debris from about 1 hectare of tree stands (Krassilov, 1972a), are oligodominant with no more than 10 tree species, whereas the extant

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rainforest stands of a comparable area are typically polydominant (occasionally monodominant in stressful environments: Hart et al., 1989) normally containing about 40100, sometimes over 300, tree species. Alongside with a general increase, the terrestrial biomass is progressively reallocated from underground to above-ground components, from trunks to crowns (with leptocaul trees taking over the pachycaul ones), and from crowns of few large leaves to those with copious small leaves (dominant in the Carboniferous, the macrophyllous life forms were relegated over the Mesozoic to a lower status of smaller trees or shrubs). These vegetation events are reflected in the biomass fluctuations at the higher trophic levels. Giant herbivores declined with a decreasing leaf to xylem ratio following the cycadophyte extinction in the mid-Cretaceous. They disappeared at the Cretaceous/ Tertiary boundary marked by a major afforestation event (Krassilov, 1981a and elsewhere). Their subsequent recovery in the mammalian faunas associated with a spread of herbaceous wetlands in the early Palaeogene. In general terms, global biomass fluctuations are imposed by environmental impacts, most ostensibly by climatic changes affecting the ratios of vegetated areas to deserts, closed to open vegetation, arboreal to non-arboreal forms and the evergreen to deciduous habits. The total terrestrial biomass increases with both latitudinal and altitudinal advances of tree lines and with expansion of evergreen biomes at the expense of deciduous ones (the latter allocating relatively more carbon in the rhizosphere but less in the standing biomass). The biomass to dead mass production rates, measured as a ratio of dry weight per ha production to litter fall, is several times greater in evergreen vegetation than in deciduous counterparts (Whittaker, 1975a; Lawson, 1994). This ratio decreased with temperization of the mid-latitudes in response to expanding polar ice caps (VII.4.1). In effect, massive peat accumulation was displaced from low latitudes to mid-latitudes over the Permian, Jurassic and Oligocene as compared with the Carboniferous, Cretaceous and Eocene. On the other hand, in non-glacial climates the advance of arboreal vegetation beyond the polar circles, over the areas of pronounced light seasonality, provides an abundant source of dead mass, evidence of which are the huge mid-Cretaceous coal reserves of northern Siberia (Zyryansk Basin) and the northern slope of Alaska. In freshwater ecosystems, charophytes were dominant macrophytic biomass producers over the Late Silurian to Early Devonian when they were joined by submerged or partly submerged tracheophytes (Aglaophyton) ecologically equivalent to extant quillworts. The Devonian algae-like thalloid forms of the Shuguria–Orestovia grade became the first producers of substantial cuticle coals. However, the diversity of aquatic macrophytes remained low, and the lacustrine ecosystems were predominantly oligotrophic, until the Cretaceous when the aquatic biomass production was greatly enhanced by the invasion of both floating ferns and angiosperms. With them, the macrophyte biomass was reallocated to the surface causing the plankton blooms and eutrophication. A high rate of dead mass production over the Cretaceous lacustrine ecosystems is attested by


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the black shale facies spreading over the trans-Asiatic rift system (II.7.3). The lacustrine black shale deposition correlates with marine black shales, with eutrophication exported via estuarine ecosystems (II.7.4). The most prominent Cretaceous black shale intervals in the deep sea drilling sections and their onland equivalents (the Bonarelli Level shales of Apennines, with about 16% organic carbon: Van Graas et al., 1983) are confined to the late Barremian-Aptian and the late Cenomanian-Turonian transgressive episodes (Schlanger & Jenkyns, 1976; Tissot, 1978; Graciansky et al., 1984). Their correlation with the coal deposition maxima in eastern Asia, western North America, Greenland and Svalbard (Krassilov, 1985) indicate either a direct or an indirect link between the terrestrial dead mass and oceanic biomass production. Marine biomass fluctuations are related to oceanic circulation patterns. Tropical upwelling of cold deep waters is commonly considered as a major enhancer of oceanic productivity that decreases with sea surface warming (as during El-Niño events that were initially recognized by their adverse consequences for fishery). Inasmuch as the tropical monsoon intensities also depend on SST, such oceanic biomass fluctuations are in phase with those of the terrestrial biomass. A direct link between the two might have been provided by an influx of organic debris and dissolved compounds with terrestrial runoff causing both dilution and eutrophication of oceanic surface waters. The ensuing combination of high surface productivity, near-bottom anoxy and decalcification of the bottom deposits resulted in a widespread black-shale event. The amount of terrestrial dead mass exported with continental runoff is controlled by climate, vegetation, pedogenic processes (silicate weathering, of which CO2 concentration in soil pores is a critical factor) and sea level. Marine hydrocarbon sources contain kerogens of terrestrial origin coming from estuaries and coastal peatlands, in particular from intertidal mangrove swamps that spread during the high sea-level stand/high SST phases of eustatic cycles. A greenhouse climate would have enhanced both terrestrial and marine biomass production owing to a denser plant cover, as well as by increasing chemical weathering as a source of nutrients exported with terrestrial runoff (VII.2.4). Indeed, the larger accumulations of hydrocarbons as well as coal (e.g. in the mid-Cretaceous: Klemme, 1975; Krassilov, 1985, 1992b) typically correlate with sea-level rise and global warming. Sea-level fluctuations and the related climate change also inflict a large-scale redistribution of terrestrial biomass over the globe.

VIII.2.2. Biomass redistribution Presently, the largest concentration of terrestrial biomass is confined to the equatorial zone bordered by the arid zones of a comparatively small biomass. Secondary maxima fall in the temperate forest zones, from where the biomass is reduced again toward the treeless circumpolar zones. This pattern of latitudinal biomass distribution has developed

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with the Late Eocene and subsequent orogenic events in the Tethys belt and a consequential stepwise expansion of the polar ice caps (VII.2.2). Notably, palynological data suggest the appearance of rainforests both in southeastern Asia (Muller, 1970) and the Guinea-Congo area (Maley, 1996) not earlier than the Oligocene. The equatorial concentration of terrestrial biomass is owing, in the first place, to a steep meridional temperature gradient imposed by mountain ranges in the path of tropical air masses and the advancing polar atmospheric fronts that jointly reduce a heat export from the tropical zone. In front of the mountains, the ascending air cools rapidly shedding rain over the tropical zone. Beyond this, the vapour-starved descending air masses inflict desiccation over a vast area of tropical deserts. At the same time, the nival deserts and lowland tundras develop within the polar atmospheric cells, and the taiga forests thrive within the reach of the polar fronts. Temperization of the mid-latitudes by surges of polar air results in a spread of deciduous vegetation. The ensuing two-peak distribution of terrestrial biomass is typical of a mildly glacial climate as at present. A different pattern would appear with deglaciation and a decrease in climatic gradients. The equatorial zone would secure less heat while precipitation would drop with a decrease in the tropical monsoon intensities. Simultaneously, a warmer SST would stir cyclogenesis at the middle to high latitudes thereby shifting the peak of terrestrial biomass as in the Cretaceous example below. Climate change also affects the latitudinal pattern of dead mass production. In the tropical zone, peat accumulation is mostly confined to coastal wetlands, such as mangroves, that also provide a major source of organic matter exported to oceanic basins, whereas in the temperate zone more peat is deposited in the inland basins. Accordingly, the paralic coals spread with tropical climates, as in the Early to Middle Carboniferous, whereas the limnic (inland) coal basins gain in importance with temperization, as in the Late Carboniferous to Early Permian Angarian and Gondwana basins, the mid-Jurassic Irkutsk and Kansk-Atchinsk basins of Siberia or the Miocene browncoal basins of central Europe. The distribution of terrestrial vertebrate biomass presently shows a single peak over the xerothermic zone of tropical savannas and woodlands.. The extinct mammoth fauna indicates a different distribution in the Pleistocene, with a secondary maximum in the high-latitude savannoid zone. Likewise, the relatively frequent Cretaceous dinosaur occurrences in the warm-temperate ecotonal zone about 50°N and over the circumpolar zone (Nessov & Golovneva, 1990) seem to indicate a two-peak pattern.

VIII.2.3. Assessment of Cretaceous terrestrial biomass The following operations are proposed (Krassilov, 1999a) for an approximate assessment of terrestrial biomass over times: (1) Reconstruction of palaeobiomes on the basis of plant morphologies and the relevant taphonomic information (IV.3-4).


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(2) Mapping of palaeobiomes and assessment of their ranges (modern latitudes preferable over hypothetical palaeolatitudes by reasons given in V.9.3). (3) Selecting the extant ecological analogues for which the unit biomass production per area is available from the literature (Basilevich & Rodin, 1967; Basilevich, 1993; Basilevich et al., 1996; Duvigneau, 1971; Whittaker, 1975a; Songwe et al., 1988; Lawson, 1994; Adams et al., 1990, etc.). These values are extrapolated to the respective palaeobiomes. (4) Calculation of biomass production for the palaeobiomes and total for the terrestrial biosphere, based on (2) and (3). (5) Calculation of respective atmospheric CO2 gains/losses with the biomass fluctuations. Extrapolation of unit biomass production from extant biomes to their extinct analogues is justified on account of the closely matching values for the ecologically similar, though taxonomically disparate, vegetation types, such as the North American, Mediterranean and Australian scrub communities, the European and East Asian laurophyllous forests, the South American, African and southeast Asian tropical rainforests, etc. A schematic phytogeography of mid-Cretaceous biomes (IV.4.1, Fig. 114) was obtained by analysing ca. 50 Aptian to Turonian fossil floras. Six major palaeobiomes were recognized on the basis of their leading plant assemblages, dominant growth forms, species richness, taphonomic evidence of phenology, and sedimentological data (Table 6). The summer-green mixed forest biome, typified by the Cenomanian assemblages of northeastern Asia and Alaska, was dominated by taxodiaceous conifers, heteroblastic Table 6. Cretaceous biomes.

Biomes Mixed deciduous

Plant community Platano-Parataxodietum,

Coniferous evergreen

Cupresso-Sequoietum

Mixed deciduousevergreen

Lauro-Sequoietum (Geinitzietum)

Debeyo-Dryophylletum Mixed evergreen with deciduous elements Xeromorphic Debeyo- Brachyphylletum serrato-scale leaved Xeromorphic Ephedro- Brachyphylletum scale-leaved

Aspect Coniferbroadleaved forest, woodland Conifer rainforest

Diversity Moderate

Phenology Summer-green

Moderate to high

Evergreen with summer-green components Summer- green with evergreen components

Coniferbroadleaved forest with laurophyllous element Mediterraneantype scerophyllous forest, woodland

High

Moderate to high

Evergreen with winter-green components

Dry open forest, woodland

Moderate

Scrub, dry woodland

Low to single-species

Evergreen with winter-green components Evergreen to summer-green

Fig. 114. Mid-Cretaceous biomes: (1) Xeromorphic scale-leaved, (2) Xeromorphic serrato-scale leaved, (3) Mixed evergreen (lauro-coniferous of Mediterranean type), (4) Mixed deciduous with evergreen elements (sequoio–platanoid with laurophylls), (5) Coniferous of temperate rainforest type with Sequoia and Cupressinocladus, (6) Mixed deciduous (parataxodio-platanoid). Typical localities: 1 – Anadyr, 2 – Sikhote Alin, 3 – Vilyuy, 4 – Kem, Yenisey basin, 5 – Nemegt, Gobi, 6 – Western Kazakhstan, 7 – Voronezh, 8 – Crimea, 9 – Bohemia, 10 – South Scania, 11 – Portugal, 12 – Bali (Guangze), 13 – Xigaze, Tibet, 14 – Jabalpur, 15 – Katch, 16 – Lebanon, 17 – Sudan, 18 – Southern Mali, 19 – Tanzania, 20 – Koonwarra, southeastern Australia, 21 – Eromang, Quiensland, 22 – New Zealand, 23 – West Greenland, 24 – North Slope, Alaska, 25 – Vancouver, 26 – Alberta, 27 – Potomac, Maryland, 28 – Baqueró, Santa Cruz, 29 – Fossil Bluff, Alexander Island.

288 Valentin A. Krassilov. Terrestrial Palaeoecology


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with deciduous short-shoots, and by platanoid angiosperms with peltate leaf blades. Leaf mats are common. The diversity of fossil plant assemblages typically does not exceed 10-15 species per plant bed. The conifer rainforest biome, typified by the Cenomanian–Turonian assemblages of Sikhote Alin, shows the highest diversity of conifer genera, most of them with evergreen extant allies in the Taxodiaceae and Cupressaceae. The largest Mesozoic stem diameters, 2.5 m, occur in the stump horizons of this biome. The mixed evergreen biome (the Late Cretaceous flora of Sakhalin) is likewise dominated by taxodiaceous and cupressaceous conifers but differs in a higher diversity of broadleaved components, in particular laurophyllous, and a higher diversity of about 1520 species per plant bed. Leaf mats are rare. The Mediterranean-type woodland–shrubland, represented in the Cenomanian floras of Bohemia, Lebanon, Dakota, etc., is dominated by angiosperms with narrow serrate or narrowly dissected leaves; conifers are rare. A single-bed diversity does not exceed 15 species. Leaf mats occur occasionally. The scrub-type shrubland (the Aptian to Cenomanian floras of Mongolia, Caspian basin, Crimea, etc.) differs in a smaller angiosperm leaf size and a more prominent scale-leaved component representing a drier vegetation. The plant remains are scattered, scarcely forming leaf-mats. The scerophyllous dwarfed shrubland is poorly represented in the megafossil record by single-species accumulations of scattered scale-leaves suggesting a deciduous shrub Table 7. Biomass of Cretaceous biomes (carbon storage in vegetation and soil) based on their ranges and the unit biomass production in extant analogues. Cretaceous biome Mixed deciduous

Area, 106 km2 6.2

Extant analogue Open conifer-broadleaved woodland

C storage t ha-1 177

Total C Gt 109.7

Coniferous evergreen Mixed deciduous evergreen Mixed evergreen Xeromorphic serrato-scale leaved Xeromorphic scale-leaved

6.5

Conifer rainforest

656

426.4

12.9

Broadleaved deciduous and mixed forest Temperate evergreen forest Dry open forest, woodland Scrub, dry woodland

290

374.1

330

369.6

135

275.4

100

128

11.2 20.4 12.8

Total Corg of Cretaceous terrestrial biomass: 1683.2 GT Total Corg of extant terrestrial biomass: 2039 Gt Loss of Corg with the reduction of terrestrial biomass: 355.8 Gt (20% of the extant value) Increase in atmospheric CO2 adequate to the loss of terrestrial biomass (0.47 ppmv per 103 t of released Corg): 167.2 ppmv (50% of extant value).

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habit. The associated palynological assemblages of plicate-elaterifer-rimulate morphotypes indicate a dominance of ephedroid and hirmerelloid components. Two calculations were performed for the mid-Cretaceous biomes: one with the presentday geography, the other with epicontinental seas (Fig. 114). Total mid-Cretaceous biomass of the present-day continents would have been greater than the present-day biomass (2421 Gt against 2039 Gt). The distribution would have been different, with a single peak of biomass production at the high mid-latitudes (Fig. 115). A reduction of terrestrial biomass with expansion of epicontinental seas and an adequate increase in atmospheric CO2 would have raised the mean global temperatures by no more than 2°C relative to the present (an expected effect of doubling atmospheric CO2 is

Fig. 115. Biomass (plant and soil carbon, Gt) distribution over latitudes. A. Present-day distribution (generalized after UNESCO classification, 1973); (1) Equatorial and tropical evergreen forests (451.7), (2) Tropical woodland, scrub and savanna (346.7), (3) Tropical semi-desert scrub and steppe (46.7), (4) Mediterranean evergreen forest, scrub and woodland (159.9), (5) Temperate broadleaved deciduous and mixed forest (390.6), (6) Temperate grassland, semi-desert (195.8), (7) Conifer rainforest (19.7), (8) Boreal forest, taiga (327.5), (9) Tundra (178.7). B. Mid-Cretaceous distribution (Fig. 114): (1) Xeromorphic scale-leaved biome (128), (2) Xeromorphic serrato-scale leaved biome (275.4), (3) Mixed evergreen biome (lauro-coniferous of Mediterranean type) (369.6), (4) Mixed deciduous biome with evergreen elements (sequoio-platanoid with laurophylls) (374.1), (5) Coniferous biome of temperate rainforest type with Sequoia and Cupressinocladus (426.4), (6) Mixed deciduous (parataxodio-platanoid) biome (109.7).


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Table 8. Biomass of mid-Permian biomes (carbon content in vegetation and soil) based on their ranges and the unit biomass production in extant analogues.

Permian biomes Tethyan Gigantopteridian Catamerian Gigantopteridian East Gndwana Glossopteridian Euromerian coniferoid-callipterididian West Gondwana Glossopteridian Eurangarian peltaspermian Angarian Ruflorian Antarctic Glossopteridian

Area, 106 km2 16.2

Extant analogue Tropical rainforest

C storage t ha-1 430

Total C Gt 698.7

23.1

Humid subtropical evergreen forest Dry subtropical woodland

260

600

135

291.6

Mediterranean scrub and woodland Warm-temperate evergreen forest Temperate deciduous forest

150

444

330

343.2

290

498.8

21.6 29.6 10.4 17.2

estimated as 1.4–5.8°C: VII.2.4). The rest of the approximately 5°C increase inferred on the basis of zonal shifts (VII.4) should have been contributed by other sources. Calculations for the mid-Permian biomes (Table 8, Figs. 52, 116) suggest a distribution of the present-day two-peak type. For the Permian biomass of 2876.9 GT, contain-

Fig. 116. Biomass distribution (plant and soil carbon, Gt) for the mid-Permian biomes mapped in Fig. 52: (SCaAm) Tropical (Southern Catamerian), (NCaAm) Northern humid subtropical (Northern Catamerian), (EG) Southern humid subtropical (East Gondwanan), (EuAm) Northern dry subtropical (Euramerian), (WG) Southern dry subtropical, (EuAn) Warm-temperate (Eurangarian), (AnAt) Temperate (Angarian, Antarctic).

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ing about 30% more carbon than at present, the amount of CO2 circulating in the biota – atmosphere system would have been more than twice the present value.

VIII.3. Diversity Biological diversity is presently a matter of international concern. Yet the significance of diversity for the biosphere in general and for human existence in particular is a matter of opinion. Diversity is vaguely related to sustainability (Busaz, 1972; Vermerj, 1974; Wiens, 1977; Grime, 1979; Grant, 1986; Huston, 1994; Naeem & Li, 1997; Sepkoski, 1997), though diverse biotic communities often prove more vulnerable than the less diverse ones. The differences in diversity over climatic zones are sometimes mechanistically explained as resulting from the species/area relationships (discussed in Rohde, 1997) or as an accident of species survival (Caston & Blackburn, 1996). Diversity is an ecosystem attribute scarcely explained by traditional evolutionary theory. From viewpoint of ecosystem theory, diversity is a measure of structural complexity that provides for the efficiency of biomass production. The cycling rates of organic carbon, including the rates of CO2 uptake, increase with diversity (Naeem et al., 1994; Yemelyanov, 1994). High diversity, as in the tropical rainforest, correlates with a high ratio of standing biomass to dead mass production. Conversely, the low diversity biotic communities, such as pioneer vegetation, tend to produce a disproportionately large dead mass. There is ample evidence of a general increase, with fluctuations, in overall taxonomic diversity over times (Raup & Sepkoski, 1968, 1982; for a revised marine diversity see Peters & Foote, 2001; for plant diversity see Niklas et al., 1985; Knoll, 1986; Collinson & Scott, 1987a). Even more instructive are palaeontological data showing a correlation of taxonomic diversity with ecospace differentiation, e.g. the plant diversity – canopy structures (Collinson & Scott, 1987a), the ichnofossil diversity – infaunal trophic niches (Buatois et al., 1998), the ungulate crown-type diversity – vegetation types (Jernvall et al., 1996). Yet the evolution of biological diversity over time poses problems that are not reducible to quantification of the process.

VIII.3.1. Diversity estimates In general, diversity is a measure of structural complexity. In the case of biotic communities it measures a division of labour invested in utilization of trophic resources and the maintenance of ecological structures. The respective functional roles, or niches, are biotopic (thermophils, xerophils, etc.), trophic (producers, consumers, destructors) and housekeeping, responsible for the maintenance and recovery, e.g. the pioneer, successional, climax forms (Clements, 1916) or the explorers, patients, violents (Ramensky, 1953).


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At the ecosystem level, diversity reflects spatial heterogeneity of biospheric structures in respect to land/sea distribution, relief, climate, etc. It depends not only on physiography but also on sensitivity of biotic communities to environmental “grain” that tends to increase with temporal stability. At the community level, diversity is inversely related to dominance, a quantitative variable reflecting relative ecological significance of co-existing life forms. Graphically, dominant forms are marked off by inflection of a numerical distribution curve, separating them from the rest of life forms. A degree of dominance is conveyed by the sharpness of the kinks so that it is minimal when the curve is nearly smooth (on conversion of dominance indices to the indices of diversity see Smith & Wilson, 1996). While life forms are functional units representing ecological niches, the biological species are essentially reproductive units locally represented by their populations (plant ecologists distinguish “coenotic populations” from “geographic populations” of eidologically orientated research). In practice, however, species are recognized as the units of morphological diversity, with an assumed reproductive background. Insofar as morphology reflects function, species reflect life forms, although a one-to-one correspondence is not obligatory. Species richness is a widely used measure of biological diversity applied to biotic communities as well as to regional biotas. A differential approach would distinguish between the communal (coenotic), intercommunal, and interprovincial components of taxonomic richness, or α, β, and γ diversities (Whittaker, 1975b, 1977; see also Bombach, 1977; Sepkoski, 1988). Their relative contribution to total diversity (S) might change with time, as in the case of the Late Palaeozoic invertebrates, the species richness of which increased by 300% since the Cambrian, with 50% supposedly contributed by α, another 50% by β and the rest by γ diversity (Sepkoski, 1988). Direct estimates by counting species recorded for a stratigraphic interval are insecure because of variations in deposition areas and rates, taphonomic losses, splitter/ lumper taxonomic strategies, etc. Less biased are comparisons of basinal floras/faunas studied by one and the same researcher. In the Russian Far East example (IX.9), the regional Early Paleocene taphofloras with no more than 50 species each are typically 20% less diverse than their preceding Cretaceous taphofloras. There were few extinctions among the dominant taxa, such as Parataxodium (“Cephalotaxopsis”), while most of them (notably Sequoia) lost their dominant status. For time intervals comparable with species longevities, species numbers primarily reflect the turnover rates. Dominance as an inverse correlate of diversity can be more objectively assessed owing to an adequate representation (with exceptions: II.3) of dominant components in the fossil record. Total diversity for a time period (St) is the sum of species entering that period (the initial diversity, Ss) and those originated during it (So). Changes in St over time depend upon origination rates while the standing crop diversity (Sp) at each successive time plane depends on the turnover rates (R). This latter variable is calculated as the difference between origination (O) and extinction (E):

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R=O–E While Sp is an ecologically meaningful variable, St is not. For each time plane, tp, the standing crop diversity, Sp’, is a function of initial diversity (Ss), time passed since the beginning of the period (tp – ts), and of mean turnover rate (Rm) for the period: Sp’ = Ss + Rm/t (tp – ts) A crude estimate of species richness for a geographic realm can be obtained from the species-area relationships. A vast literature on the subject (Arrhenius, 1920; Preston, 1962; MacArthur & Wilson, 1967; Malyshev, 1975; Williamson, 1988; Harte & Kinzig, 1997; Gaston, 2000; Williamson et al., 2001) holds the relationships as conveyed by parabolic equation: y = axz (y = species number per area x, a = species number per unit area, z = an empirical constant of spatial heterogeneity). In the following example, the variables and constants of the species-area equation are provisionally calculated on the basis of palaeogeography and the present-day analogues. Area (x). Palaeovegetation zones can be considered as areas of uniform z values characteristic of the respective biomes. Calculations are made for the area of the northern temperate forested (nemoral) zone north of the evergreen/deciduous boundary that is most distinct over times due to the distinctions in the life forms and taphonomy (IV.3, VII.4, VIII.2.3). The ranges of nemoral biome vary with climate and sea-level, spreading to about 38°-35°N in the mid-Permian and mid-Jurassic, to 43°N in the Early Palaeocene, and constrained at 50°N (60°N in Atlantic sector) in the Late Cretaceous and Eocene, when their extent was additionally reduced by boreal transgressions. Spatial heterogeneity constant (z). A compilation of data for vegetation zones indicates typical z values of 0.10 for boreal conifer forests (taiga), 0.14–0.18 for temperate broadleaved (nemoral) zone, about 0.17 for Mediterranean-type vegetation, and 0.20–0.22 for tropical forests (Malyshev, 1975). Since z depends on ecological convergence rather than taxonomy, these values should be valid for the respective palaeobiomes as well. The upper present-day value of z = 0.18 is provisionally accepted here for the Permian to Palaeocene nemoral zones on assumption of their relatively high heterogeneity. Diversity per unit area (a). In the present-day plant diversity estimates, the unit area varies from 102 to 104 sq. km. For palaeobiomes, the natural sampling unit is sedimentary basin supplying material for the basinal taphoflora. A modal area of hydrological basins is about 103 sq. km, within the range of the present day sampling areas. Hence species numbers for well-known taphofloras represent the contemporaneous (a) values. In the present analysis these values are derived from the following taphofloras: – The mid-Permian of Cisuralia (Naugolnykh, 1998), – The Early Triassic of Tunguska Basin and Taymyr (Mogucheva 1973; Sadovnikov, 1997),


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– The mid-Jurassic of Ust-Baley, Angara River (Heer, 1868-1883, Krassilov & Bugdaeva, 1988 and unpublished), – The Early Cretaceous of Bureya, Russian Far East (Vakhrameev & Doludenko, 1961; Krassilov, 1972c, 1973c, 1978a), – The Late Cretaceous of Sakhalin (Krassilov, 1979), – The Early Palaeocene of Amur Basin (Krassilov, 1979). All these are similarly positioned north of the evergreen/deciduous boundary for respective geological epochs. Species richness (y). Fig. 117 represents the diversity estimates for the northern nemoral zone over a sequence of geological epochs, Permian to Palaeogene, based on the species–area equation, with (x) defined by the sequential positions of the evergreen/ deciduous ecotone, (a) derived from six well-known taphofloras of respective ages and (z) extrapolated from the modern analogue. Fluctuations of standing crop diversity seem to have reflected a combined effect of climate change (temperization in the Permian, mid-Jurassic and Early Palaeocene) and the boreal transgressions (maximal in the Late Cretaceous).

Fig. 117. Species richness estimates for the northern nemoral zone based on the species–area equation y = axz, with the areas (x) shown in Fig. 99, the spatial heterogeneity constant (z) extrapolated from the modern analogues (0.18, a higher value for the present-day nemoral vegetation), and the basinal diversity (a) derived from representative taphofloras: (P2) Mid- Permian of Fore-Urals (Naugolnykh, 1998), (T1-2) Early – Middle Triassic of Tunguska Basin and Taymyr (Mogucheva 1973; Sadovnikov, 1997), (J2) Mid-Jurassic of Ust-Baley, Angara River (Heer, 1868-1883, Krassilov & Bugdaeva, 1988 and unpublished), (K1) Early Cretaceous of Bureya, Far East (Vakhrameev and Doludenko, 1961; Krassilov, 1972c, 1973c, 1978a), (K2) Late Cretaceous of Sakhalin (Krassilov, 1979), (KT) Early Palaeocene of Amur Basin (Krassilov, 1979).

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Extrapolation of (z) is a weak point of the estimates because a palaeobiome might differ substantially from its extant ecological analogue in the spatial heterogeneity of (a). Conceivably, if species richness calculated on the basis of the total diversity and turnover rates (above) is in excess of that derived from the species-area equation, then spatial heterogeneity might have been greater than in the present-day analogue, implying the higher values of intercommunal diversity β, interregional diversity γ or both. The heterogeneity constant (z) is maximal in the case of disjunctive floristic regions. Z=0.27 indicates complete geographic isolation of floristic regions when derived from the canonical equation (Preston, 1962): (A/C)1/z – (B/C)1/z = 1 (A, B = species numbers of two floristic regions, C = the number of their shared species). Palaeogeographic implications of the threshold z value may pertain to “exotic terrains” supposedly transported from remote locations (examples in V.6.3). Z greater than 0.27 would confirm a former separation of continental blocks while a lesser z would suggest land continuity. At least for the northern and southern Cathaysian floras (species lists in Li & Wu, 1996; Li, 1997), z ~10 is far beneath the threshold value, indicating continuity of respective “terrains” rather than their wide separation over the eastern Tethys.

VIII.3.2. Species richness and disparity Diversity as a number of different entities in a sample does not tell us how distinct these entities are. Two assemblages of equal species richness may differ in distinctness of their constituent species. Hence, there are at least two components of species-level diversity: the richness (diversity in a common but inexact usage of the term) and disparity in terms of morphological, etological or biochemical distinctness. Richness might grow by addition of sibling species without an appreciable increase in morphological disparity. On the other hand, an introduction of a single extremely deviant form would increase the disparity with a negligible effect on the richness. Disparity is measured as a mean pairwise difference (distance) between taxa estimated for a selected set of characters (Briggs et al., 1992; Foote, 1995). As such, disparity relates to the initial adaptive radiation, subsequent divergence, and, inversely, to survival of intermediate types (Jernvall et al., 1996). A high initial disparity, as in proangiosperms (Krassilov, 1997a), means either saltational or polyphyletic origins, with convergence at a later stage. In contrast, a conventional mode of gradual divergence implies a low initial disparity that increases with introduction of unique characters over time. In the early land plants, a high disparity/low diversity of growth habits, in particular of the dichotomous, monopodial or H-type branching patterns, stands in sharp contrast with


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the uniformity/minor scale differentiation of their reproductive structures. Conversely, the early evolution of flowering plants produced a high disparity of floral morphotypes, some of them (e.g., the Platanaceae) fairly conservative over times, while their species richness is assessed on the basis of the much more variable leaf morphotypes and the cuticle micromorphology. In these groups the problem of species longevity (Niklas, 1985) would have different solutions depending on the choice of key characters. Disparity can be assessed on a semiquantitative basis, by arbitrary weights assigned to morphological species in respect to their relative distinctness. New morphologies result either from recombination of pre-existing characters or from the appearance of an entirely new character. In the following example from Maslova (2001), higher weights have been assigned to deviating morphologies relative to recombinations of characters that were present in the basal stalk. Thus, the presently monotypic family Platanaceae is first recorded in the mid-Cretaceous as a cluster of genera differing primarily in their androecial morphologies. Most of the early morphotypes disappeared before the Cretaceous/Tertiary boundary, few persisted over the Palaeogene giving rise to a new round of adaptive radiation. However, in the Palaeogene round, it was the female, rather than male, organs that diversified at a higher rate. In the early Palaeocene, the taxonomic richness but modestly increased with the appearance of a new floral genus Oreocarpa. However, the disparity increased more significantly owing to the deviant fruit morphology of this genus producing a single ripe follicle per pistillate head (in a family with typically numerous achenes or nutlets per head). Similar changes in evolutionary strategy are provisionally inferred for other early angiosperm groups as well. In taxonomic studies, a deviant morphology is commonly allotted a higher rank. Therefore, in taxonomically diverse groups, the ratios of higher to lower taxa may serve as a crude measure of disparity. In the higher plant taxa, species richness correlates with cosmopolitanism and the life form diversity, being considerably higher in the families that comprise both tropical and temperate forms, woody and herbaceous habits, wind- and insect-pollination ecologies (Ricklefs & Renner, 1994). In such families, speciation does not necessarily result in a higher-rank diversification. The enormous species richness of tropical rainforests indicates prolific speciation that is scarcely accompanied by an adequate increase in the species disparity. High morphological disparity typically associates with the fine-grain adaptive strategy of pioneer species (Van Valen, 1965) that tends to prevail in the post-crisis biotic communities (e.g., the Early Palaeocene Platano-Trochodendroidetum, IX.9.3). Later on, with stabilization, a coarse-grain strategy gains in importance over a succession of florogenic stages, providing for a steady build-up of species richness. Hence an alteration of disparity/richness trends is bound to cyclic changes of adaptive strategies in response to the contrastive ecosystem situations.

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VIII.3.3. Arogenic communities The areas of high species richness are commonly considered as cradles of biological diversity, although they also give refuge to archaic forms. The species-rich tropical ecosystems are controversially interpreted either as cradle or museum of biological diversity (see Jablonski, 1993; Caston & Blackburn, 1996). Many biogeographers consider tropical biota as a source of innovation for temperate realms. In particular, this theory has been applied to the origins of angiosperms (reviewed in Krassilov, 1997a). On fossil evidence, before the appearance of full-fledged angiosperms, there were transitional morphologies combining typical angiosperm characters with those of gymnosperms. Such proangiospermous plants usually constituted a minor element of the Late Mesozoic floras dominated by cycadophytes, ginkgophytes and coniferoids. However, some Early Cretaceous localities of Transbaikalia, Mongolia and northern China are exceptional in containing many such plants (Fig. 118). No less than 15 proangiosperm species of widely disparate morphologies have been recorded (Krassilov & Bugdaeva, 1982; 1999; 2000; Krassilov, 1986) from the Baisa locality on the Vitim River, Transbaikalia, constituting about half of the flora (Figs. 119 – 121). The contemporaneous Manlay and Gurvan Eren localities in Mongolia are likewise dominated by proangiosperms (Krassilov, 1982). Recently, such localities have been discovered in the Yixian Formation of Liaoning and elsewhere in China (Fig. 122). In distinction from the scattered proangiosperm finds, they suggest angiospermization as a coenotic phenomenon of the type-making – arogenic – communities. The Asiatic communities of this type are confined to the Mongolo-Okhotskian transcontinental rift system and its parallel Kunlun-Ginling fault zone. The fossil assemblages occur in the black-shale facies and associated carbonates of the rift valley lacustrine deposits (II.7.3). In distinction from the more conventional Early Cretaceous floras, they contain only sparse fern remains. This seems to indicate a drastic reduction of fern marshes, a widespread Mesozoic wetland formation. At the same time, certain proangiosperm remains are convergently similar to extant marsh herbs in their general habits and the morphology of pappate disseminules (Problematospermum, Baisia, Typhaera, etc.). In respect to their ancestral bennettitalean, caytonialean and proto-gnetalean forms, the growth habits of the Cretaceous proangiosperms is deminutive while their floral structures are highly condensed (Krassilov, 1997a; Krassilov & Bugdaeva, 1999). They seem to have appeared as substitutes of fern marshes, filling in the vacant ecological niche and acquiring angiospermous characters by way of developmental acceleration related to unstable habitats. The proangiosperm remains are associated with great accumulations of aquatic insects and fish indicating mass mortality. A high frequency of mass mortality events over the Lower Cretaceous sequences attests to unstable lacustrine ecosystems of fluctuant water levels and seasonal turnovers of anoxic deep waters. Since most of aquatic insects are highly sensitive to pH levels, their mass mortality suggests the pH


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a

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Fig. 118. The Late Jurassic to Early Cretaceous assemblages from the lacustrine black shale – marl sequences of trans-Asiatic rift system. Plant remains from the newly appearing proangiosperm wetlands show a convergent similarity to the modern sedge/grass ecological analogues, while frequent association of aquatic and terrestrial remains suggests an impact of terrestrial runoff on aquatic life: (a) Ephemeropsis, mayfly larvae covering a marl bedding plane, a mass mortality case, (b) a coprolite with Samaropsis-type seeds produced by a wetland coniferoid, (c) Typhaera, a pappate disseminule resembling that of a reed-mace, (d) Sparganium-like fruiting head, (d) graminoid leaf, (e) sedge-like inflorescence of three spikes, (d) beddingplane association of Potamogeton-like spike and chironomid remains, (e) Baiera, a ginkgophyte leaf overlapping a fish skeleton, (f) Poblematospermum, a pappate bennettitalean seed. (a, b, h – Baisa, Transbaikalia, c – Bon Tsagan, Mongolia, d-g – Manlay, Mongolia, i – Karatau, Kazakhstan).

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Valentin A. Krassilov. Terrestrial Palaeoecology a

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Fig. 119. Proangiosperm remains from a single plant-bearing bed of the Early Cretaceous (mid-Neocomian) lacustrine deposits at Baisa, Vitim River, Transbaikalia: (a, b) Baisianthus ramosus, inflorescences of whorled cupulate sporangiophores, (c) Preflosella nathania, a preflower with bracteate tepals, stalked sporangiophores and a gynoecium of sessile cupulate ovules, (d) Vitimantha cripta, a preflower with bracteate tepals and a gynoecium of follicle-like cupules.

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e Fig. 120. Proangiosperm remains from a single plant-bearing bed of the Early Cretaceous (mid-Neocomian) lacustrine deposits at Baisa, Vitim River, Transbaikalia (continued): (e) Eoantha zherikhinii, a bracteate preflower with a four-lobed cupule bearing orthotropous ovules, attached to a slender axis, (d) graminoid leaf presumably of the same plant, (g) Prognetella bracteata, a spike of cupulate ovules supported by long bracts, (h) Eoantha ornata, a four-lobed cupule with orthotropous ovules.

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Fig. 120. Continued

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Fig. 121. Proangiosperm remains from a single plant-bearing bed of the Early Cretaceous (mid-Neocomian) lacustrine deposits at Baisa, Vitim River, Transbaikalia (continued): (i, j) Baisia hirsuta, disseminules of cupulate ovules with a persistent gynophore bearing a tuft of long bristles, (k) Loricanthus resiniferus, a cone of peltate scales bearing adaxial sporangia with Eucommiidites-type pollen grains, (l) Dicotylophyllum pusillum, a small petiolate angiosperm-like leaf or cataphyll.

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Fig. 122. Extension of lacustrine rift deposits with Lycoptera– Ephemeropsis–proangiosperm assemblages (of the type shown in Fig. 118) to northeastern China: (a) a fern-marsh palaeosol with Coniopteris in tuffaceous deposits of Didao Formation representing a wetland type that became rare after the transboundary Jurassic/Cretaceous volcanic phase (b) a fish bed of Yixian Formation, a mass mortality case, (c) Gurvanella, a fruit-like proangiosperm cupule common in the Yixian Formation, (d) the same species from contemporaneous deposits of Gurvan Eren, Mongolia (Krassilov, 1982).

catastrophes possibly related to volcanism and acid rains as well as an enhanced chemical weathering of the granite bedrock, a massive source of Al and F in the drainage runoff. The same factors might have been responsible for vegetation turnovers in the adjacent wetlands. Thus, the Jurassic to Cretaceous riftogenesis and basaltic volcanism over the transAsiatic Mongolo-Okhotskian zone (V.5.2) provided an ecological situation for a highrate macrevolutionary development. Remarkably, a comparable proangiosperm development took place over the contemporaneous Brazilian basaltic province (Mohr & Friis, 2000). Rift valleys (Fig. 125) are hot spots of basic diversification not only for higher


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plants (Dettmann & Jarzen, 1990; Krassilov & Dobruskina, 1995), but also for the Late Permian/Triassic tetrapods as well as the Plio-Pleistocene hominids (Leaky et al., 1998; Haile-Selassie, 2001). A causal link between environmental change and arogenesis is also illustrated by innovation of aquatic plant communities over the mid-Cretaceous. A burst of floating ferns and angiosperms in the mid-Cretaceous (IV.2.6) was preceded by the appearance of non-marine diatoms that increased the biotic production of lacustrine ecosystems. New potentials for aquatic plant life were explored by submerged macrophytes, such as the aquatic lycopsids developing a high disparity of heterosporous forms (e.g., the overproductive Limnoniobe and the later appearing Monilispora with giant sporangia producing the bead-like files of megaspore tetrads: Krassilov, 1982; Krassilov & Makulbek-

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Fig. 123. Herbaceous marshland gnetalean proangiosperms from the Early Cretaceous rift lake deposits of (a, b) Manlay, Mongolia and (c, d) Koonwarra, Australia (Krassilov, 1982; Krassilov et al., 1998).

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ov, 1995, Fig. 4). These developments contributed to the organic-rich bottom deposits, the seasonal turnover of which fertilized the hypolimnion. The trophic situation was ripe for an entry of floating macrophytes, and these bursted over the Cretaceous greatly enhancing the eutrophication rates. The early aquatic ferns, having their origins with the fern-marsh progenitors, were homosporous, with rosettes of dichotomous leaves convergently similar to the thalomes of Riccia, a floating liverwort (Fig. 123). They gave rise to a highly disparate array of heterosporous forms that included such bizarre ones as Heroleandra with megaspores and microspores jointly dispersed as amphisporions (Krassilov & Golovneva, 1999, 2000; Fig. 50). The simultaneous radiation of aquatic angiosperms generated most of their extant higher order diversity, reshaping the lacustrine ecosystems into a fairly modern aspect toward the end of Cretaceous (Krassilov, 1976a). In arogenic communities of the above examples, high morphological disparity is scarcely matched by the specie-richness, far not the highest among the contemporary plant assemblages. Generally, in a search for cradles of biological diversity, disparity may seem a better guide than species-richness.

VIII.3.4. Diversity and environmental stability Since the pioneering work by Wallace (1870), diversity has been linked to environmental stability, yet the causal relations remained obscure (historical overview in Ricklefs & Shluter, 1993). The widely held view of a direct diversity/stability correlation making the diverse communities superior in terms of sustainability cannot be accepted without qualifications. A short-term stability does not guarantee, or even imply, a longterm sustainability. The notion of stability combines homeostasis, a power of damping the induced fluctuations, with resilience, a power of recovering the equilibrium after disturbance. These aspects of stability play different roles in respect to environmental change. Their correlation is inconsistent (Grime, 1979; Huston, 1994) or at least non-monotonic (Stone et al., 1996). Moreover, the environmental situations must be differentially assessed in respect to their spatial versus temporal components, short-term versus long-term processes and their rates. A detrimental effect of environmental change, such as inflicted by tectonic activity and volcanism, is compensated for by the effect on spatial landscape heterogeneity, notably of the mountain ranges or rift valleys, which are the most productive diversification foci (e.g. the centres of floristic diversity: Krassilov, 1995d). Although all environments are changeable, some are obviously more changeable than others in terms of their spatial heterogeneity and/or temporal continuity. Environmental conditions can be defined as constant (steady to mildly fluctuant), inconstant (conspicuously fluctuant), predictable (constant or regularly fluctuant) or unpredictable (liable to erratic fluctuations). Different adaptive strategies are required for coping with erratic versus pre-


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Fig. 124. A build-up of floating phytomass and diversity of freshwater ecosystems with invasion of ferns and angiosperms in the Cretaceous: (a, b) “Hymenophyllites”, an aquatic fern with dichotomous leaves bearing terminal sporocarpia, the Cenomanian of western Kazakhstan (V.A. Vachrameev’s collection, Geological Institute, Moscow); (c, d) Quereuxia, a floating leaf rosette and submerged leaves from the Late Cretaceous of Sakhalin (Krassilov, 1979); (e-h) nymphaeaceous seeds and a rhizome from the Early Palaeocene of Tsagajan, Amur Province (Krassilov, 1976a).

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Fig. 125. Rift valleys, the foci of terrestrial and lacustrine diversities: (a) Rift Valley landscape near Nairobi, Kenya, containing the early hominid localities, (b) Dead Sea Rift, Israel, with salt marshes of Tamarix tetragina and Arthocnemum glaucum and with freshwater Holocene deposits over the valley, (c) Gusinoye Ozero, Transbaikalia, with plant and dinosaur localities on surrounding slopes, (d) Bighorn Basin, Wyoming, a segment of extensive epirift subsidence over the Palaeozoic rift system, with plant and mammal localities in the redbed deposits.

dictable fluctuations of whatever amplitude. Species-rich multistratal forests provide for environmental stability, but their regeneration rates are lower than in species-poor shrublands. Here biotic complexity increases the homeostasis but decreases the resilience. Typically, an environmental constancy generates a niche-splitting tendency, predictable fluctuations entail the broader tolerances, whereas the high recovery rates are critical for survival under erratic fluctuations. Plant communities respond to seasonal frost, drought or fires by physiological adaptations that expand the limits of their tolerances (as in winter crops or pyrogenic vegetation), whereas erratic disturbances place emphasis on the resilience. Erratic droughts, hurricanes or fire may negatively affect climax components, while their effect on overall diversity can even be positive when, for instance, they let more light in to an understorey (Grace & Pugesek, 1997) or increase the spatial heterogeneity by creating gaps, corridors, etc. Gaps are important in maintaining the diversity of tropical rainforest, hence a correlation with canopy disturbances. A frequently encountered in ecological studies increase in diversity at a mild level of disturbance (Connell & Orias, 1964; Connell, 1978; Malino & Sabatier, 2001) probably results from interplay of the conflicting


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forces. A benign level of disturbance is such that the climax species are not irreparably destroyed while at the same time the pioneer species are promoted. By this reason, the species richness is maximal at the pre-climax stage (Beklemishev, 1951). However, a severity of episodic disturbance need not be confused with long-term stressful conditions that are never promotive of biological diversity (IX.3). Landscape heterogeneity in terms of habitat mosaics or patchiness is commonly considered as a positive correlate of species richness (reviewed in Kassen, 2002) but there is evidence to the contrary (Turner, 1989). Intraspecies competition increases with spatial heterogeneity that has but a minor effect on interspecies competition (Shorrocks & Sevenster, 1995). The spatial and temporal aspects of environmental heterogeneity may differently influence the populational strategies, with dissimilar genetic effects. The genetic polymorphism is found to be higher in the spatially variable environments, whereas the phenotypic plasticity increases with temporal variation (Reboud & Bell, 1997). Conceivably, the correlation depends on a prevailing adaptive strategy (Fig. 126). Eventually, the diversity–stability correlations are consequential to the prevailing populational strategies, which in turn change in response to the level of environmental impacts. Sensitivity to environmental grain may prove a critical factor, with the coarse-grain and fine-grain strategists diversifying at different levels of environmental heterogeneity. In effect, species richness grows with niche splitting in the spatially heterogeneous, temporarily constant environments but it is halted at a relatively low level in the predictably fluctuant environments owing to the necessity of maintaining broad physiological niches. The environmental evaluations are certainly relative because what is stable for one species may prove unstable for another. More specifically, a fine-grain strategist would assess as homogeneous/constant an environmental substrate that is felt as heterogeneous/fluctuant by a coarse-grain strategist. Hence a lesser adaptive effort is re-

Fig. 126. Schematic causal links between biological diversity and environmental stability through the adaptive population strategies: environmental disturbances promote a redundant population growth conferring instability on a biotic community halted in its development at an early successional stage; the species-rich climax stage is reduced and eventually cut off.

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quired in the former strategic domain than in the latter while their contributions to overall diversity are disproportionate. As examples of arogenesis may show (VII.3.3), fluctuant ecosystems enhance disparity rather than species richness. They tend to shed small populations that more often survive in relatively constant ecosystems which serve either as museums of relict species or as cradles of speciation by genetic drift, or both. The present-day species-level endemism correlates with species richness but the peaks do not coincide (Williams et al., 1994), conceivably owing to segregate occurrences of the palaeo- and neoendemics. The extant biomes of a relatively recent origin, such as the tropical rainforest, temperate grasslands and Arctic tundras, contain most of the neoendemics while the conservative ones, like the Arcto-Tertiary type deciduous forests of eastern North America or laurophyllous forests of southeastern Asia, provide refuge to archaic forms (VII.1.3). Hence the cradle/museum dilemma obviously relates to the history of vegetation, each of the newly arising biomes, not only the tropical ones, assuming a cradle role. However, the arogenic and species-rich areas scarcely coincide because of the different environmental incentives for innovative and collateral speciation. Nature conservation is commonly thought of as stabilization, which is questionable for two reasons at least. First, not all species gain from stability. If a pioneer species is targeted, then destabilization may prove more efficient. Furthermore, an attempt at conserving a species-rich climax phase at the expense of a few pioneers, though apparently justified, may eventually create a precarious situation of a non-resilient community virtually incapable of repair. In natural plant communities, the biological diversity is conserved by adaptation to the environmentally imposed disturbance – stabilization cycles. In a persistent forest ecosystem this results in a close canopy – gap pattern optimal for the maintenance of both the climax and the pioneer–seral phases.

VIII.3.5. Niche creating trend Stability promotes to co-adaptation, which is a vehicle of biological diversity. Since interaction is a sensitive stage in the life cycles of the partners, most stringently constrained by environmental conditions, co-adaptation is promoted by constant, rather than fluctuant, environments (e.g. advanced tropical forms of insect pollination in comparison with the temperate ones). Co-adaptaion adds new dimensions to ecospace providing for an increase in biological diversity (VIII.1.2). With new species in turn developing coadaptation ties the process is self-accelerating. The dimension can be taken literally, as in the case of aboveground and underground plant growth (below) or reefal construction. Yet the appearance of warm-blooded animals is also a new ecospace dimension, thermostatic in this case, and the burst of their symbiont diversities testifies to the effect. Even human technologies, although invented for the benefit of a single species only, unintentionally increase biological diversity, of the microbial world at least.


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Quantum steps in the process are owing to the sequential appearances of nichegenerating forms that enhance the diversity of their co-evolving organisms by the multiple co-adaptation ties. As a rule, the chronologically later appearing dominant forms, such as angiosperms, are superior to their predecessors (gymnosperms) in the capability of stabilizing their physical environments (e.g., soil conditions in hardwood forests comparative with those in conifer forests: Lewis, 1986) and creating new ecological niches up the trophic cascades as well as collaterally (for the rhizospheric symbionts, epiphytes, herbivores, pollinators, etc.). With the early land plants advancing from prostrate to erect habits, their co-evolving fungal and animal life was elevated above the ground (as attested by the epiphytic fungi and by arthropod feeding on sporangia: Edwards et al., 1995). More of such niches were added by the later appearing arboreal plants generating a rich above–ground diversity. Eventually, in the multistratal forests, canopy life has become more diverse than ground life. Human ancestry belongs to the canopy life. Hence the origins of humanity are owing to the early land plants that raised their sporangia above the ground and furthermore to the flowering plants that provided for the frugivorous habits. In this example, the diversification drive spreads upward both spatially and in terms of trophic cascades, from primary producers to the climbing vertebrates on top. Simultaneously, the diversification of flowering plants is driven top down by the anthophilic arthropods and vertebrates. In both cases, a build-up of diversity is accelerated by the positive feedbacks in the plant → herbivore → plant circuits. At all evolutionary levels, structural complexity increases with positive feedbacks that increase with evolutionary progress and over developmental sequences (cyanophytes excrete substances that suppress their own growth while mammals sustain their habitats, as in the grazer/grassland example; over an ecological succession, the pioneer species undermine their own existence making their habitats more suitable for subsequent seral stages; even the modern human technologies tend to advance from devastating to sustainable). Co-adaptation pertains to the ability of interaction with a positive feedback, whereas maladaptive forms generate negative feedbacks of an adverse effect on biological diversity. A species entry either decreases the pre-existing diversity by competitive elimination of resident species or inflames diversification by opening niches for further entries. Since ecosystem efficiency correlates with diversity, the progressive evolutionary trend is associated with the niche creating entries alone. In other words, the evolution of ecosystems (as well as human social systems) is owing to those organisms (human individuals) that create new ecological niches rather than those that compete for the established ones.

VIII.3.6. Diversity and niche overlap A build-up of biological diversity associates with diversification of habitats (Benton, 2001), a trend that is taken for granted in evolutionary studies, but is actually a function

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of niche overlap (inverse correlation) and the adaptive strategies in respect to environmental grain (direct correlation with the coarse-grain strategy). Both these vary with ecological situations. A generalist species tends to decrease the diversity by competitive elimination of its rival species. The rates of elimination are directly correlated with the niche overlaps in the competing species. The fine-grain strategist experiences the environment as homogeneous, providing no incentive for differentiation. There is no a priori reason why the entire ecological space might not be taken up by a single generalist species (which is what the human species is presently striving at). Such species should be an extreme fine-grain strategist tolerant of a wide range of environmental conditions and a forceful competitor. These qualities rarely if at all combine, because the fine-grain strategists usually are poor competitors. Moreover, as ecological observation convincingly asserts, broad tolerances can only be sustained when the trophic resources are plentiful (Odum, 1959). The tolerance ranges would inevitably narrow down with a rise of population density and exhaustion of trophic resources, impelling the dominant species to abandon its marginal habitats (as modern humans are impelled, by resource limitations, to cut off certain programs of polar exploration, desert reclamation, etc.), therefore leaving room for other species. Simultaneously, a stronger emphasis is placed on the efficiency of resource utilization. The trophic strategies shift from indiscriminate to selective, setting on a resourcepartitioning trend. While populations are becoming more sensitive to environmental grain, a build-up of species richness is enhanced by the positive feedback: each species entry opens potential niches at the higher and lower trophic levels, thus encouraging more species entries (a spectacular example is a regeneration of biotic community after the Krakatau explosion: Thornton, 1996). In reality, the species entry curves are exponential in their proximal segments alone, showing saturation at the widely discrepant levels (May, 1978). These are primarily defined by two critical variables, the niche overlap and the sustainable population density. To formalize the relationships, a plateau species number can be represented as a function of the community modal values of the pairwise niche overlaps and the sustainable densities: SF ~ TR/NOm + Dm, (SF = finite diversity, TR = trophic resources, NOm = modal niche overlap, and Dm = modal sustainable density). SF increases with a reduction of N0m, Dm, or both. The NOm coming down to zero would mean no competition, an ideal situation not yet achievable either in nature or in human society. For Dm there should be a limit, below which the population is endangered. The formula helps to understand the mechanism of diversity regulation through population strategies. In fossil assemblages, a co-occurrence of morphologically similar species (that are by inference ecologically dissimilar, with a minimal niche overlap) is usually correlated with a high overall diversity. Inversely, a high diversity can be inferred on the basis of such occurrences. A high ratio of species to genera has a


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similar meaning in terms of niche overlap. For example, while the upland diversities tend to be grossly underrepresented in the fossil record, a co-occurrence of several species of Gingkoites, a dominant upland genus, in allochthonous assemblages, contrasted to a prevalence of single-species (though highly polymorphic) Phoenicopsis, Czekanowskia and Pseudotorellia of the contemporaneous riparian vegetation (Krassilov, 1972b), indicates a substantially higher diversity in the uplands irrespective of the species counts. Similarly, a higher species to genera ratio in the Aptian flora in relation to the Valanginian one (Krassilov, 1972b) is evidence of an increase in plant diversity over these Early Cretaceous ages in spite of the nearly equal number of species in the respective localities. As for the population density, the post-crisis Trochodendroidion is an examples of a low-diversity assemblage of few species abundantly represented in the leaf mats (IX.9.3). Such examples are too familiar to dwell upon them, yet their meaning is not fully elucidated.

VIII.3.7. Diversity and redundant density The Darwinian theory was based on the then (and even later) widely held notion that each species, if not constrained by other species, strives to swarm the whole world with its progeny. Hence competition and natural selection of the fittest. It was accepted as an eternal law of nature that each species (each individual within it) is driven by an inherent urge to outcompete other species (individuals) in purpose to secure more room, more trophic/reproductive resources to its progeny. In compliance with this concept, fitness was measured by proliferation. It has long been ignored that the most prolific are, as a rule, the most oppressed, and vice versa (which may be also true for humans; as said in the Exodus, the more they were oppressed, the more they multiplied). Indeed, if being oppressed is a goal, then prolificacy is a criterion of success. However, when Aldous Huxley wrote of “Darwinian specimen”, a competitive human being ever struggling for existence, he obviously had in mind a possibility of an alternative, perhaps nobler, life strategy. Ancient Romans assigned a lower (sixth and below) social grade to the “proletarians”, who serve their state by proliferation alone. Presently we are coming to the understanding of social developments as directed by the imperative of sustainable existence rather than that of unlimited population growth. Prolificacy is only adaptive as a compensation for high mortality inflicted by either natural or artificial (as in pest control) hazards. A dramatic increase in redundancy in pest populations that temporarily escape pesticide control is a typical example of the adaptive strategy acquired in response of a severe extermination (the Niobe strategy: Krassilov, 1992a, 1995a; Fig. 127). The oppressed are compelled to produce a redundant density that can be spared to overkill without loss of the population. The

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Fig. 127. Redundant reproduction (the Niobe strategy) is correlated with high death rates. After an initial success, Niobe strategists give way to the more efficient Leto strategists.

same concerns humans, until quite recently one of the most severely exterminated (self-exterminated) species. In organisms of a high reproductive potential, population growth is regulated by negative environmental feedbacks inflicting large density fluctuations (cyanophytes, Drosophila, lemmings are the familiar examples). Population densities in excess of environmental capacities are cut by an overcompensatory mortality (Clutton-Brock et al., 1997). Since redundant growth is devastating to trophic resources, such populations are maladaptive. In much the same way, organic pollutants instigate a redundant biomass production in aquatic organisms followed by a fall of diversity in hypertrophic reservoirs (Heinonen, 1980; Barinova, 2000). Plant diversity, population density and patchiness are likewise controlled, both directly and indirectly, by compensation mechanisms (Huntly, 1991). Alternatively, in a friendlier environment, sustainability can be achieved at a lower density level. Density regulation falls under internal control (as in the hatch size regulation in birds or estrus periodicity regulation in mammals), thus minimizing a redundant component. Self-regulated populations, with density-dependent fecundity, are more constant than the environmentally regulated ones that develop a redundant density compensating for enormous losses. Henceforth populations become less destructive to their habitats, a successful adaptation strategy (the Leto strategy: Krassilov, 1992a, 1995a).


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Both disaster and pioneer species tend to be highly redundant, the Niobe strategists. Their broad tolerances entail the likewise broad niche overlaps, hence low diversity. A coarse-grain strategy comes later, with stabilization, allowing for a decrease of redundant density and therefore opening way to diversification. At the community level, homeostasis depends on the ratio of self-regulated to environmentally regulated populations and is inversely correlated with redundant densities. Redundant populations are not only themselves fluctuant but they also confer fluctuations over the trophic chains. The biotic communities loaded with a high proportion of Niobe strategists are thereby highly unstable. Insofar as co-adaptations require constant (predictable) densities of populations involved, the communities of Leto strategists are better co-adapted, with elaborate trophic webs, advanced forms of biotic mutualism and a higher overall structural complexity. With redundancy understood as being adversely related to sustainability, we are no longer obliged to identify fitness with population growth, the redundant component of which endangers long-term survival rather than contributing to it. This part of the traditional evolutionary theory must be seriously reconsidered. Fitness as a measure of population efficiency is to be evaluated not so much through overall growth rates as by the constancy of demographic structure. Owing to a great disparity of environmental situations and ecological roles, both Niobe and Leto strategists survive in natural ecosystems. Yet this need not lead us to a false conclusion that they are equal in terms of evolutionary advancement. Redundant reproduction is a source of excessive dead mass. Therefore, at the ecosystem level, a contribution of redundant populations to the biomass/dead mass ratio (a criterion of ecosystem efficiency, VIII.2) is negative. A decrease in this ratio over times signifying progressive developments is due to the Leto strategists alone.

VIII.3.8. Diversity evolution The above considerations seem pertinent to our understanding of the causal relationships between diversity and environmental stability. Schematically, they are as follows. In natural ecosystems and the biosphere as a whole, a diversification tendency is sustained by the imperative of efficiency in converting non-living matter into biomass with a minimal dead mass production. The biomass to dead mass production ratio depends on partitioning of trophic resources – splitting of trophic niches entailing an increase of overall complexity of which the diversity is the measurable equivalent. Environmental factors administer regulation of diversity through population strategies, in particular the niche overlaps and redundant densities. Redundant populations of high reproductive rates controlled by negative feedbacks are a source of erratic fluctuations and an excessive dead mass production. They thus make their residential communities maladaptive. On the other hand, minimization of redundancy gives way to resource

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partitioning and diversification. These processes are enhanced by co-adaptation and the stabilizing feedbacks allowing for low species densities and a variety of interspecies configurations (rainforest situation). Adaptation (co-adaptation) extends as far as the redundancy can be dispensed with, which depends on environmental situation. The more constant the environments, the less imperative is the demand of buffering populations from catastrophic mortality by developing excessive densities. In other words, the carrying capacity of ecosystems (biosphere) grows with stabilization. The trend is reversed with a rise of redundant population densities in response to geological, climatic or human impacts, their destabilizing effects depressing the carrying capacity of global ecosystems, and of the biosphere as a whole. Crisis situations are marked by the low-diversity communities of few abundant species. Such communities comply with the Darwinian principle of survival through proliferation that entails the struggle for existence. With stabilization, the normalizing selection decreases releasing a latent variability as a substrate for a subsequent diversification. Hence the diversity cycles of the fossil record reflect the global environmental (climate) change. So the choice of diversity as a major criterion of evolutionary progress is fully warranted by theoretical considerations binding this variable to population strategies and environmental situations. Even crude estimates of diversities over the geological record show distinct rises and lows. The boundaries of geological eras (the Palaeozoic/Mesozoic/Cenozoic) have been drawn (by Philips in 1840) at the major lows of biological diversity rendering them fairly natural in spite of the repeated claims to the contrary from the orthodox Darwinian camp. The validity of the boundary lows is confirmed by the modern numerical estimates of the total biological diversity over times (Raup & Sepkosky, 1982). A loss of diversity might result from an increase in extinction rates or a decrease in origination rates (McGhee, 1988), or both. Anyway, the evolution of diversity involves several regulation levels. Extinction not compensated for by origination is evidence of a decreasing structural complexity (a cut-off of one or more structural levels: IX.6), inevitably affecting other ecosystem variables. Since diversity is related to efficient utilization of environmental resources by their partitioning and the respective division of functional roles, it is mandatory for a build-up of biomass. Actually, a direct correlation of diversity, standing crop biomass and the trophic effectiveness (with nitrogen as the main limiting nutrient) has been found in a number of case studies, in particular of grasslands, wetlands and macrophyte aquatic communities (Tilman et al., 1997, 2001; Engehhardt & Ritchie, 2001). A co-ordinate rise of diversity and biomass is obvious over the early evolution of terrestrial life, with both these variables dramatically increasing after the advent of higher land plants, arboreal habits and multistratal vegetation. It is commonly believed that a supply of nutrients enhances biological diversity (Vermeij, 1995). Since, according to a well-known ecological rule, only about 10% of the biomass can be consumed at each of the trophic levels, the complexity of trophic cascades eventually depends on the width of the trophic base. The tropical diversities


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are sometimes explained as a result of the exceptionally high primary productivity enhanced by a high level of incoming solar energy (Connell & Orias, 1964; Pielou, 1975), although actually the life under the rainforest canopy primarily depends on shade tolerance. A much higher biological diversity on land (despite the earlier start of marine life, which had already approached the present-day level of diversity in the Palaeozoic (Peters & Foote, 2001) reflects not only, and not so much, the difference in the diversity of habitats (Benton, 2001) but also a proportionately higher primary production of terrestrial ecosystems. However, for some as yet poorly understood reasons, the co-variance of diversity and productivity is non-linear, asymptotic at the levels of half or less of the maximum diversity (Hector et al., 1999; Schwartz et al., 2000), frequently assuming a unimodal hump-back form (Grime, 1997; Stafford et al., 2001; Chase & Leibold, 2002). The threshold of co-variance is due to an abrupt switch from direct to inverse mode (Rosenzweig & Abramsky, 1993; Rosenzweig, 1995). Explanations hitherto provided for such intriguing correlations involve a habitat heterogeneity being, after a threshold, reduced with a further rise of the trophic level (Rosenzweig & Abramsky, 1993) or, alternatively, a changing selection pressure (Wiens, 1977; Grime, 1979; Grant, 1986; Huston, 1994; Stafford et al., 2001), such as favouring the eutrophic adaptations at the expense of the oligotrophic ones (Hallock, 1987). A paleobiological interpretation of diversity fluctuations in Jurassic radiolarians with alternation of the eutrophic/oligotrophic conditions complies with the latter model (Danelian & Johnson, 2001). Of special interest are the co-ordinate changes in biomass and diversity with eutrophication of aquatic ecosystems (Henderson-Sellers & Markland, 1990; Barinova, 2000). An initial build-up of biomass is accompanied by a rise of diversity. Organic pollution inflicts hypertrophic conditions, under which the biomass/diversity correlation fails and is eventually reversed (Heinonen, 1980; Barinova, 2000). This signals a retrograde development, with a redundant growth not accompanied by an adequate increase in structural complexity, resulting in a high rate dead mass production. A succession of ecological dominants is likewise retrogressive, with coccoid green algae and cyanophytes replacing the relatively advanced centric diatoms. Hypertrophic ecosystems are unstable, and are also a source of instability for all basinal ecosystems to which the excessive dead mass is exported. Such phenomena deserve further causal analysis for the processes linking diversity with productivity are of primary importance for ecosystem evolution, involving both the diversity/stability and the diversity/redundancy correlations (VIII.3.4; VIII.3.7). In general terms, a low diversity means an inefficient use of trophic resources, hence a low productivity level. Ecosystem evolution is directed toward optimization of productivity via the build-up of diversity. In stable ecosystems, optimal levels of primary production are sustained by the trophic cascades. However, an excessive influx of nutrients over eutrophication trends disrupts the top-down regulation, generating instability that favours a redundant population growth at the expense of diversity.

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In the final count down, any disturbance is a removal of biomass (Grime, 1979) that is exported or accumulated as dead mass. Since the biomass/dead mass production rates are affected, a change in diversity is eminent. Typically, a biomass removal would inflict a decrease in diversity. However, in view of the parabolic relations (above), a removal of biomass over the ascending or descending limbs of the curve may invoke entirely different diversity responses. Diversity may even increase following a cut-off of the excess productivity. Population strategies are thus a central issue to the diversity/productivity correlations. Collapse of the positive diversity/productivity correlations (inflection of a parabolic curve) is related to a switch from the coarse-grain environmental strategy to the finegrain one and, simultaneously, from the Leto-type demographic strategy to that of the Niobe-type (VIII.3.7), both induced by an excessive supply of nutrients. The richest biotic communities thrive on poor substrates, whereas the marine diversities never reach the level of the terrestrial ones because of instability of their trophic base fluctuating with turnovers of deep water circulation and terrestrial runoff. Owing to the parallelism of ecological succession and evolutionary sequence (VIII.4), a decrease of biological diversity with eutrophication provides an actualistic model of past biotic turnovers as consequential to breakdowns of a positive biomass/diversity correlation, with an incongruous biomass production inhibiting the diversity. Even in the late dinosaur ecosystems, an enormous biomass (B9land & Russel, 1978) appears to have been in excess of a decreasing diversity.

VIII.4. Sere as reiteration of ecosystem evolution In the course of ecological succession (sere), the pioneer populations tend to be finegrained and overproductive using their redundant densities as a buffer against unpredictable environmental impacts. As a consequence, their dead mass production is large relative to their standing crop biomass that increases over seral stages. The sequentially appearing seral species are relatively coarse-grained, therefore less redundant and more effective in their use of trophic resources (Bridham et al., 1995), conferring a stabilizing influence over their habitats. Species richness increases with stabilization, in parallel with the standing crop biomass, to a saturation at the pre-climax stage (Beklemishev, 1951; explanation in VIII.3.4). Thus, the general trends of biomass, dead mass and diversity over a sere are the same as over geological time (VIII.2, VIII.3.8). In both cases, the limits of progressive development are imposed by environmental stability, increasing with a feedback stabilization by biotic communities. As discussed in (III.3.1), the pioneer and seral components persist alongside with the climax phases providing for the resilience of coenotic systems. Yet in constant environments, the pioneer/seral components are competitively eliminat-


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ed, hence a drop of overall diversity before reaching the climax. Conversely, erratic disturbances are destructive for the climax but promotive for the pioneer/seral phases. The parallelism of seral and evolutionary sequences involves both normal and retrograde developments. Thus, an initial build-up of biomass and diversity often starts with phylogenetically archaic forms, such as cyanophytes, followed by the relatively advanced unicells and, eventually, by macrophytes, the successions of which may also parallel their phylogeny, as in the case of quillworts, an ancient group of submerged lycopsids, entering aquatic communities at an oligotrophic stage, followed by water ferns, nymphaeoids and monocots, with duckweeds terminating both the eutrophication sere and the sequence of evolutionary first appearances (IV.2.6). Retrograde developments, on the other hand, are inflicted by environmental stresses of ecological and evolutionary time scales, respectively. Thus, with organic pollution, cyanophytes replace diatoms in freshwater ecosystems, while a rise of lichens and lycopsids (Selaginella), the most ancient terrestrial organisms, to a dominant status in the tundra and tundra-steppe communities is provoked by stressful conditions over the Pleistocene glacial cycles. The consequences of environmental impacts on seral systems to a certain extent depend on the severity of disturbance. Mild disturbances might even increase overall diversity, as predicted by the Hubbard Brook model of shade tolerance as a critical factor of regeneration (in the case of a beech-hemlock forest disturbed by tornado, an increase in species richness was primarily due to a rise of shade-intolerant shrubs and herbs: Peterson, 1995). However, under long-term disturbances, the loss of climax species is no longer compensated for by the rise of stress-tolerant components and the community degrades to a disclimax state. When affecting dominant ecosystems on global scale, the long-term stresses provoke biospheric crises. Hence the ecological and geological time-scales are linked by parallelisms of the biomass/diversity trends, owing to which an ecological succession can be considered as a brief reiteration of ecosystem evolution (Krassilov, 1992a, 1995a). The build-up of biomass and diversity over geological times, as well as their losses at the critical levels of biospheric evolution find their analogy in the seral/climax/disclimax developments. The ecogenetic rule of sere/evolution parallelism is equivalent to the “biogenetic law” of the ontogeny/phylogeny parallelism, both reflecting the systemic nature of multilevel developments.

VIII. 5. Speciation cycles Speciation is currently ascribed to either an adaptive divergence of genetically insulated populations or to the stochastic phenomena of genetic drift, founder effect or bottlenecking (Wright, 1931; Mayr, 1954; Carson, 1975) based on the presumption of selective neutrality. Whatever its role in speciation (critically reviewed in Galiama et al., 1993), the founder effect does not produce a major evolutionary novelty as the original theory held.

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A ubiquity of neutral variation inferred form the assessments of mutation load in natural populations (Kimura, 1983) inflamed a burst of non-selectional speciation models. As a compromise, the maintenance of neutral loci was ascribed to their linkage with the selected ones. This so-called “hitchhiking model” is not universal, however, while a varying selection pressure over a broad range of environmental situations provides a viable alternative (Nevo et al., 1984, 1997). The foregoing discussion implies that the adaptive strategies of biological species are defined primarily by their roles in the development of biotic communities over the pioneer to climax sequences and, in parallel, over the transition from non-coherent to coherent evolution. The contrasting adaptive strategies also pertain to the colonizing versus residential habits (explorers/patients/violents: Ramensky, 1953), niche breadths (generalists/specialists), environmental grain sensitivity (fine/coarse grain strategists: Levins, 1968) and demographic options (Leto/Niobe strategists: Krassilov, 1992a, 1995a). All these can be considered as aspects of the so-called K/r strategies (McArthur & Wilson, 1967), yet an indiscriminate approach does not help understanding the complexity of ecosystem evolution. In the course of ecosystem development, the constituent populations are channelled towards one strategic option or another. Periods of coherent evolution promote the coarsegrain environmental strategy in conjunction with the Leto strategy of parental care/low reproduction rates. The prevailing mode of morphological change is the anaboly (Severtsov, 1939), an adaptive modification at the terminal stages of individual development. Conversely, in disturbed ecosystems, population strategies switch to a fine-grain mode, with prolific reproduction compensating for the high mortality rates (the Niobe strategy). The associated ontogenic mode is an accelerated development resulting in a condensation of developmental stages and paedomorphic transformations. In this model, a neutral polymorphism at an increasing number of loci would appear in extreme situations of either a suspended stabilizing selection in undersaturated communities or a restriction of ecological niches and consequential loss of function in oversaturated communities. Co-adaptation introduces the interaction niches that come to dominate the ecospace differentiation decoupling the partners from certain environmental limitations that are thereby neutralized (VIII.1.3), entailing a loss of adaptive meaning for variation at the respective loci. On the other hand, a neutral variation may assume a selective value with a new dimension added to selective environment. Thus sexual selection imposes a symbolic meaning upon individual distinctions of little if any adaptive significance in respect to environmental variables. Human personality is mainly based on variation of the kind. Since speciation takes too long to be studied experimentally, we have to rely on palaeontological data for verification of theoretical models. The commonest situation in the fossil record is an evolutionary appearance of a species simultaneous over the range of its ancestral species (e.g. of a schizaeaceous fern Anemia dicksoniana replacing A. (Ruffordia) goeppertii over the Northern Hemisphere at the Albian/Cenomania bound-


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ary: Krassilov, 1979 and elsewhere). It will be sufficient to remind, not going into detail, that a long-distance stratigraphic correlation is based on the simultaneous over the range replacements of widespread species. Such speciation events are definitely not in favour of the common believes in geographic disjunction and the founder effect, but rather indicate a diffuse, over the range, speciation processes. A plausibility of duffuse speciation is justified, first, by compelling evidence of adaptive genetic differentiation in spite of a considerable gene flow (Endler, 1973; Nevo et al., 1984, 1997; Galen et al., 1991; McKay & Latta, 2002). A correlation of genetic changes with microenvironmental differentiation has been recently convincingly demonstrated by the ecogenetic studies over heterogenous landscapes – the “evolutionary canyons” (Li et al., 2000a, 2000b; Belyaev et al., 2001). Predictably, the differentiation is related to the coarse-grain adaptive strategy in respect to environmental heterogeneity. However, the grain strategy is a consequence of selective environment to be changed with it. A switch to the fine-grain strategy would sweep out the pattern of the down-to-grain diversity. The heterozygosity would increase by mixing of the ecotypic gene pools, with an appreciable effect on polygenic traits (McKay & Latta, 2002). Genetic introgression would result in a generalist species, homogeneous over its geographical range, with individual variation masking interpopulation differences, as in the Early Palaeocene Trochodendroides arctica (IX.9.3). The underlying genetic processes are as yet not fully understood but their (long denied) correlation with ecosystem events has become more apparent recently. Thus, genetic polymorphism fluctuates with cyclic environmental changes (Scheiner & Goodnight, 1984; Kirzhner et al., 1995; Korol et al., 1995, 1996). Stressfull environments tend to enhance mutation rates (Wills, 1984; Parsons, 1992), partly owing to an effect on the recombination system (Zuchenko & Korol, 1985; Cullis, 1990; Korol, 1999). In plants, the cytoplasmic genomes are transmitted via seed flows that in the colonizing species typically prevail over pollen flows (Ennos, 1994; Oddou-Muratorio et al., 2001). Here the genetic transmission through cytoplasmic genomes directly dependent on the environmental grain strategy. Insulation of gene pools from intrusion of alien genetic material is important for the coarse-grained populations of stable biotic communities. Yet at a non-coherent stage of ecosystem build-up, the pioneer (disaster) populations may benefit from enrichment of their gene pools by genetic introgression as well as by non-sexual (horizontal) transduction of genetic material as a source of additional genetic variability. We arrive at a model of cyclic splitting – fusion dynamics related to the alternative population strategies in respect to environmental fluctuations. Speciation is preceded by adaptive differentiation down to environmental grain resulting in a pattern of distinct ecotypes that increase genetic variability over the species range. Morphological distinctions arise through subtle developmental anabolies inflicted by the adaptive adjustments. This stage involves the coarse-grain strategists that with a change of selective environment give way to the fine-grain strategists bringing with them a mixing and genetic introgression over the

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pre-existing pattern of ecotypic differentiation. In fluctuant ecosystems, the morphological disparity increases through developmental acceleration rendering a macromutaional phenotypic effect (Goldschmidt, 1940). Subsequent splitting of macropolymorphic species may result in superspecies configurations (Krassilov, 1989b). This model is further discussed in relation to genetic assimilation (VIII.6) and biospheric crises (IX.9.3).

VIII.6. Genetic assimilation In the 1890’s-1920’s, many palaeontologists shared orthogenetic views later sunk into oblivion due to the fierce opposition of neo-Darwinians. In essence, the orthogenetic theory stated that closely related species represent successive stages of individual development (Eimer, 1898). Evidence in favour of orthogenesis came from all branches of palaeontological research (e.g., Hyatt, 1986; Cope, 1887; Plate, 1920). Classical examples are furnished by the evolutionary successions of ammonoids, horses, humans, etc. In the early land plants, the progressive webbing of syntelomic branchlets on the way to leaf blades and a concurrent differentiation of vascular system are likewise orthogenetic, fixed at different developmental stages over the arrays of co-existing forms. Morphological modes of such developments were thoroughly investigated (Severtsov, 1939; De Beer, 1940; Rensch, 1954), yet the genetic mechanisms remained obscure until, in the 1970’s, a correlation was revealed between the interpopulation variability in developmental traits (the emergence time, growth rates, metamorphosis, germination, etc.) and the differences in the genomic structures, in particular, the repetitive DNA (Schroeter & Hewitt, 1974, John & King, 1977; Beverstock et al., 1976; Lubs et al., 1977). Fast adaptive shifts were related to a repatternings of repetitive DNA (Bernardi et al., 1976; Mukai & Cockerham, 1977) causing developmental heterochronies. Certain genetic polymorphisms were interpreted as paedomorphic, a retention of juvenile isozymes patterns in the adults (Maters & Holmes, 1975; Nair et al., 1977). Geographic clines of polymorphic traits were shown to conform to their developmental sequences. A correlation of developmental heterochronies at morphological and molecular levels was extended from conspecific populations to closely related species in various groups of plants and animals (Laird & McCarthy, 1969, Maher & Cox, 1973; Narayan & Rees, 1977; Vladychenskaya & Kedrova, 1977; Buongiorno-Nardelli et al., 1977, Macgregor & Andrews, 1977) thus supporting the orthogenetic model. A denial of orthogenesis pertains to the difficulties it poses for the conventional mutation/selection mechanism. The genetic differences between local ecotypes are conventionally ascribed either to drift or selection, or else to recruitment of pre-adapted genotypes from a polymorphous source. However, the differentiation is too rapid for such mechanisms to play a substantial role, suggesting a more direct adaptive response either through an effect on intercellular environment modifying the transcriptional activ-


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ities of the genes involved (Schwartz, 1971) or mediated through a sensitivity of noncoding DNA to environmental induction (e.g., Romanov & Vanyushin, 1977). An increase in the content of repetitive DNA species correlated with the trophically induced heritable modifications in Linum usitatissimum (Durrant, 1962; Callis, 1973 and elsewhere) testifies to the validity of the latter mechanism. These mechanisms give some substance to the slightly ambiguous theory of genetic assimilation (Waddington, 1962), obviously related to the phenomena previously known as the Baldwin effect (Baldwin, 1902) or the Dauermodifikationen (Jolos, 1913). In its original version, the theory postulates a higher probability of fixation for genetic changes channelling the development in the direction of an adaptive phenotypic modification, also interpreted as a higher frequency of advantageous mutations (Hall, 1990). The theory thus admits that mutations are not occasional but are instead induced by environmental effects on development. This is fully justified by the systemic nature of the genome provided with the molecular repair mechanisms that screen most of occasional changes detrimental to the programmed development. Since genome is a system of developmental memory, the environmentally induced changes of development should be memorized in one way or another. In both the brain and genomic memories, the incoming information is processed and stored in a retrievable form. Structural changes, involving nuclear DNA, occur over the memorizing processes at both the neuronal and genomic levels (Wallace, 1974; Kazakhashwili, 1974; Black et al., 1987). New experience is assimilated by learning the genomic form of which is the genetic assimilation. It may seem that genomic structures are far less responsive to direct impact of new experience than the neuronal configurations are. Yet learning plays a restricted role also in respect to subconscious neuronal memory unless it is elevated to the conscious level. So the difference may be of a degree only. In the same way as brain functioning is not explicable at the level of individual neurones, the genetic memory operates at the level of the total genome rather than a particular nucleotide sequence. The holistic nature of genomic memory reveals itself in the retrieval of morphological images over a succession of genetic changes. The phenomenon of retroconvergence is a spectacular example. Retroconvergence is a similarity between composite organs and incorporated basic structures (Krassilov, 1995c). Thus, the composite strobili of conifers, originating from aggregations of simple cones, are convergently similar to the latter in their gross morphology. The same applies to seeds and seed-bearing cupules (Fig. 128). Likewise, in compound leaves, fusion of leaflets results in a composite leaf that looks exactly like the leaflets. There can be several rounds of retroconvergence. The phenomenon has been studied in detail in a Triassic peltasperm, Lepidopteris (Scytophyllum), in which the marginal fusion of leaf segments produces a lobed leaf blade, with the original pinnate venation retained in the lobes (Fig. 129). At an advanced stage, the lobes are no longer discernible but the conservative venation betrays the composite nature of the blade. Eventually, the venation is also modified in the direction of the initial dichotomous pat-

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a

c

d

b Fig. 128. An example of retroconvergence of seed/cupule structures: (a, b) Moresnetia, horned cupules of a Devonian progymnosperm (courtesy of A.R. Ananiev), (c) Lazarospermum, a Late Permian seed (Krassilov, 1999b) and (d) Eoantha, horned cupule of an Early Cretaceous proangiosperm (Krassilov & Bugdaeva, 1999).


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c

a d

b

e

Fig. 129. Fusion and retroconvergent modification of leaf segments in a Triassic peltasperm, Lepidopteris (Scytophyllum) vulgare (Krassilov, 1995c): (a, b, d) incomplete fusion of pinnules forming a lobed lamina, with interstitial pinnules incorporated in the web, betrayed by their persistent venation pattern (arrow); (c, e) further fusion producing an entire lamina, with pinnules, lost in fusion, still occasionally marked by fasciculate proximal veins; the distal veins appear dichotomous as in unmodified pinnules.

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tern. The composite leaf blades now appear exactly as the simple pinnae of the prototype, except for the occasional “bushy” veins still bearing evidence of their secondary origin. Retroconvergence may seem an absurd circularity of morphological evolution when considered in terms of natural selection: an advantage of a derived composite structure, if any, is lost by returning back to the initial form. It is, however, meaningful as evidence of a genomic capability to hold an image (of initial pinnate leaf structure in the above case) and retrieve it after a series of morphological events. As in the neuronal system, there should be a short-term dispensable memory opening the gateways to a long-term memory based on structural changes at the molecular level (Waldrop, 1987). Insofar as the developmental patterns are extremely conservative, most developmental changes involve a short-term memory alone, rapidly resolving into a stereotype mode. Genetic assimilation means that, when repeatedly occurring over generations, the short-memory events might enter the long-term memory. This would result in parallelism of genomic trends during the development and over evolution, making the former a brief reiteration of the latter, as the “biogenetic law” compels. Directional evolution of genomes is invoked by a variety of epigenetic phenomena (Løevtrup, 1974; Jablonka & Lamb, 1995), including a non-random segregation of alleles (the miotic drive), the silencing of an allele through interaction with other alleles (paramutation: Brink, 1973), a selective usage of codons (Gouy & Gaultier, 1982; Duret & Mouchiroud, 1999; Urrutia & Hurst, 2001), etc. There is evidence of methylation being engaged in paramutagenicity (Walker & Panavas, 2001) as well as in a hormone-induced activation (Xu et al., 2001) or silencing of developmental genes (discussed in Quon & Berns, 2001). A significant correlation is found between gene expression and the selective usage of codons (Gouy & Gaultier, 1982; Duret & Mouchiroud, 1999; Urrutia & Hurst, 2001) indicating a possible role of differential gene functioning in a directional change of nucleotide composition. An exchange of nuclear and plasmid DNA (such as a high frequency escape of mitochondrial DNA to the nucleus in Saccharomyces: Thorsness & Fox, 1980) is another potential factor of directional genomic change. It is becoming increasingly clear that genome evolution is a result of multilevel regulation. At the level of post-translational modification, proteins with different functions are produced from a single coding sequence by the splicing mechanism. Recent studies reveal a high number of alternatively spliced genes, about 50% in the human genome (Brett et al., 2002). An environmentally induced change in a gene functioning (or a coordinated change of multiple gene functioning) may result in the loss of one or another splicing alternative (simultaneously in a number of functionally related spliced genes) with a major effect on the direction of genome evolution. At the level of sequential genomic structures, a combination of flexible repeated elements and the relatively constant unique elements not only complies with the two kinds of memory but also provides a mechanism of transforming one kind into the other. Contrary to a widely held view, genomic structures do not remain the same during devel-


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opment (discussed in Krassilov, 1980). The highly repetitive satellite DNA undergoes detectable changes in early embryogenesis (in Cyclops: Beerman, 1977; for reviews see Yunis & Yasmineh, 1971; Tobler et al., 1972; Woodcock & Sibatani, 1975). A quantitative variation in repetitive DNA has also been observed in adult tissues (Perlman & Stickgold, 1977; Storm et al., 1978). Although largely transcriptionally inactive, heterochromatin casts a position effect over the adjacent functional genes modulating their transcription rates (in Saccharomyces: Tauro et al., 1968; Mortimer & Hawthorne, 1973). Heterochromatin influences chromosome pairing through its effect on chiasma frequencies (MacGregor & Andrews, 1977; Jones, 1978; Yamamoto & Miklos, 1978) that change with age (Lathardt et al., 1973; Wallace et al., 1976; King & Hayman, 1978) and are altered by environmental factors, such as trophic conditions or temperature (Speed, 1977; King & Hayman, 1978). Yet a developmental role of heterochromatin mainly pertains to its effect on cell proliferation rates. A shedding of heterochromatin during development, as in Cyclops (Beerman, 1977) may reflect an evolutionary event of developmental acceleration in the phylogeny of the genus. Loss of satellite DNA falls in the category of underreplication that may have farreaching developmental consequences in respect to interspersed repeats. Homeotic mutations shifting cells to a different developmental pathway involve deletions/amplifications of the intermediate repeats (see Quweneel, 1976). Species of repetitive DNA rich in adenine and thymine are slow-replicating as a rule. Most of the intermediate repeats interspersed between unique sequences belong to the slow-replicating class and, as such, are potentially underreplicated. Actually, underreplication of intermediate repeats – and possibly also of the adjacent unique sequences – has been observed in the larval DNA of Drosophila (Woodcock & Sibatani, 1975). An underreplication of intercalary heterochromatin regularly occurs at polytenization (Barr & Ellison, 1972). Even the unique sequences that are transcriptionally inactive are underrepresented in polytene chromosomes (Leibovitch, 1977) showing a functional correlation between the transcription and replication activities. This gives us a clue to the phenomenon of genetic assimilation: those functional modifications that alter transcriptional activities of certain genes also affect their replication eventually inflicting a directional structural change. Since interspersed repeats take part in regulation of sequential gene activities as elements of promoter sequences (Britten & Davidson, 1971; Charlesworth et al., 1994), their deletions/amplifications may alter the development by accelerating/decelerating the activation signals or else by juxtaposing the elements of different functional genes. In particular, the promoter activation in phages is primarily determined by the length, rather than by composition, of the spacer region and is modified by deletions in the latter (Hawley & McClure, 1983; Graña et al., 1985). Such genomic events conceivably generate an abbreviation or overlap of developmental stages, of which the precocious sexuality is a familiar example. Notably, studies of underreplication in various species of repetitive DNA have shown a correlation with developmental rates, in particular, with retention of paedomorphic characters (Yunis & Yasmineh, 1971).

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Insofar as developmental heterochrony is a source of major evolutionary novelty (De Beer, 1940; Rensch, 1954; De Beer & Swinton, 1958; Clark, 1962, Gould, 1977; McNamara, 1983; Bernis, 1984; Miller & Wagner, 1991), any shifts of sequential gene activities are potentially of a considerable evolutionary significance. In plants, heterochrony typically associates with fusion of homologous, as well as non-homologous, structures producing the chimerical fusion meristems of a great morphological potential (on their roles in morphological evolution of angiosperms see Krassilov, 1997a). Whatever the specific mechanism, the multiple developmental effects of fluctuations in various species of repetitive DNA complies with its role of a flexible short-term genetic memory pertaining to transient developmental situations but capable of activating the long-term genetic memory under a persistent environmental induction. Schematically, the causal links over genetic assimilation are: an environmental impact → a change of intercellular environment affecting relative concentrations of certain transcripts → an adequate change of transcription rates in respective genes → an adjustment of the genomic regulatory systems to the altered gene activities through delitions/amplifications of their repetitive elements inflicting a tendency of an earlier activations in the environmentally induced functional units. While in modern organisms a response to environmental induction is mediated by a cascade of genetic events, we can expect retention of a more direct environmental control over development in certain archaic forms. Actually, in amoebae, the transition from growth to development (of multicellular spore-bearing fruiting bodies) takes place when a supply of their forage bacteria is exhausted. Hence the development is triggered off by starvation. However, a switch to developmental program is anticipated by induction of several developmentally regulated genes in response to a decreasing bacterial density (measured by secreting a signalling protein, the “preservation factor”, concentration of which on the acceptors is inversely proportionate to the bacterial concentrations). In these organisms, genetic mutations affect the development either by interfering with the excretion of the preservation factor (Primpke, 2001) or through the responses to concentration levels of the latter. As in many other functional traits, regulation tends to be internalized in the course of evolution, rendering the development less dependent on environmental accidents. Freeliving unicells react to emission of developmental signals (Gurdon & Bourillot, 2001) in much the same way as the cells of a multicellular organism respond to a signal morphogene. Yet environmental induction of a developmental trait, e.g. of an early sexual maturation in Drosophila (Kern et al., 2001), acts upon a system of morphogene gradients rather than targeting a solitary gene product. As discussed above, reproductive rates are defined by population strategies that are determined by environmental situations. At the organismic level, reproductive rates are related to developmental rates. At the genomic level, the developmental rates are accelerated through a decrease of cellular genomes at the expense of non-coding elements. Thus ecology is linked to genetics through a trade-off between the ecologically con-


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trolled population variables, their related developmental variables, the respective genetic change and the ensuing directed evolution of the genome.

VIII.7. Examples of palynomorphological evolution Spores and pollen grains are promising for evolutionary analysis as the relatively simple haploid structures of a single or few cells produced by diploid sporophytes, but functionally controlled by their dispersal environments. The uniformity in antithetic development and the relative ploidy levels in the sexual and asexual phases of the higher plants as compared with a much more variable situation in the higher algae attests to a role in adaptation to terrestrial life. The sexual phase evolves to a parasitic habit, with pollen grains remaining the only part time free-living stage of sexual cycle in the seed plants. The adaptations of pollen grains to parasitic life are mainly biochemical, while the morphology pertains primarily to the free-living stage. The functional roles of surface structures are related to the modes of release and dispersal, volume regulation, storage of signalling substances, catchments by receptive structures and germination. Since pollen grains are endowed with a limited potential for morphological complexity, most of the surface structures are by necessity multifunctional. Studies in evolutionary palynomorphology are focused on the apertural types and sporoderm structures. Thus, Kupriyanova (1969) has recognized three major stages in the evolution of apertural structures: the Palaeozoic of proximal laesures (catasulcate), the Mesozoic of distal laesures (anasulcate) and the Cenozoic of transequatorial colpi or pores. These are the sequentially prevailing types, each of which extends beyond the time of its dominant occurrence. In particular, the catasulcate apertures allegedly defining the “prepollen” of primitive seed plants are found at all stages, occasionally even in the present-day angiosperms. A transition from one apertural type to another remains largely unexplained, although several attempts have been made to relate the development of equatorial apertures to structural changes in tetrad configuration (Wodehouse, 1935) or functional changes in respect to germination (Hughes, 1984), or production of syphonic pollen tubes (in contrast to haustorial tubes in sulcate forms). Notably, the diverse tetrad configurations in angiosperms are correlated with the diversity of apertural types, including the periaperturate ones. However, most monocots retained a solitary polar aperture in spite of their variable tetrad configurations. The pollen/stigma interactions are generally regarded as a crucial factor of pollen morphology in angiosperms, the first appearing dispersed pollen grains of which are recognized by their semitectate reticulate surface and the prevailingly columellate infrastructure, both related to a self-incompatibility syndrome (Zavada, 1984). Columellate infratectum might serve for storage of recognition substances while a periaperturate morphology might provide for germination regardless of a positioning on the stigma.

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However, in early angiosperms, an orientation of small monosulcate or tricolpate pollen grains against the relatively large stigmatic trichomes scarcely affected germination. A volume-regulating (harmomegathy) function of symmetrically disposed apertures may seem a more plausible explanation of their advantage in certain dispersal environments. Surface structures apparently pertained to dispersal environments in early spore plants already. Nikitin (1934) interpreted the excessively developed outgrowths of Devonian spores as anchoring structures. He argued that, since the area suitable for an early pteridophytic plant life was narrow, anchoring might have been more important than vagility. Yet, the pre-Devonian spores were less conspicuously ornate in spite of the even more restricted habitats. Hence we have to consider other possible factors, in particular the role of arthropods as spore dispersal agents in the Devonian thalomic mats and telomic thickets (Fig. 130). Developmental correlations of palynological and epidermal characters are obvious in the case of the elliptical paracytic stomata and the boat-shaped anasulcate pollen grains in bennettites and angiosperms. Both pollen grains and stomata increase in size with the ploidy levels (supposedly also in the bennettites: Krassilov, 1978d). Yet, the sporophyte/ pollen grain correlations pertain to few morphological characters only indicating an insignificant pleiotropic effect of the controlling genes. This complies with the functional independence of pollen grains at the free-living stage. Pollen homeomorphy. Spectacular evidence of independent pollen grain evolution is provided by a convergence in palynological characters simultaneous with a divergence in sporophytic characters. The early peltasperms (callipterids), gigantopterids, modern conifers (as different from the earlier coniferoids), vojnovskyans and glossopterids appeared simultaneously or almost so as the result of the major Permian gymnosperm radiations (Fig. 113). At the same time, the taeniate pollen morphology (with the exine split into parallel bands, or taeniae) has spread over both the Laurasian and Gondwana gymnosperm lineages (e.g., Protohaploxypinus morphotype in glossopterids and northern gymnosperms). A no less prominent pollen homeomorphy accompanied the end-Permian to early Mesozoic radiation of the calipterids, corystosperms, pentoxyls, cycads, bennettites, ginkgoids, czekanowskians and their allies. Their boat-shaped monosulcate unsculptured pollen grains are so uniform that even the major gymnosperm orders are scarcely distinguishable in the contemporaneous palynological assemblages. Noteworthy in this respect is also the parallelism of semitectate columellate pollen structures in different groups of early angiosperms at the time of their explosive midCretaceous radiation. Small reticulate tricolpate/tricolporate morphotypes simultaneously appeared in the Cretaceous ranunculid and hamamelid lineages (Krassilov & Shilin, 1996; Krassilov & Golovneva, 2001). Palynogogical convergence is commonly ascribed to climatic adaptation. Taeniate morphology, for instance, is often related to volume regulation in the supposedly arid Permian to Triassic climates. Such a function is certainly plausible, yet at the time of


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Fig. 130. Devonian spores with adhesive processes for anchoring or sticking to arthropod trichomes. Pavlovsk Quarry, Voronezh Region, European Russia.

their widespread occurrences, the taeniate morphotypes were recorded not only from redbed facies of seasonally dry climates but also from coal-bearing deposits. On evidence of pollen loads preserved in gut compressions (Rasnitsyn & Krassilov, 1996, 2000; Krassilov & Rasnitsyn, 1997, 1998; Krassilov et al., 1999d), Permian insects of the orders Hypoperlida, Palaeomanteida, Psocida and Grylloblattida were regularly

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feeding on taeniate pollen of the Lunatisporites, Protohaploxypinus and Vittatina morphotypes produced by conifers and peltasperms. Mixed pollen loads contained two or more such morphotypes. Hence the insects visited different sources of taeniate pollen over their foraging territories. For them the taeniate structures might conceivably serve as guides to the most rewarding pollen sources. This hypothesis explains a high selective value of taeniate morphology for the insectpollinated Permian gymnosperms. The pollinivorous insects might spread this morphology over the diversity of gymnosperm lineages not only through selection, but also by facilitating a microbial transduction of exogenous genetic material (VIII.1.4). The horizontal gene transfer might have been most efficient in respect to palynomorphological structures controlled by a few genes with insignificant pleiotropic effects. Recurrence. A repeated occurrence of a character over times also falls in the category of parallel, yet heterochronous, developments. There were several periods of predominantly saccate pollen grains (in the Late Paleozoic and Cenozoic gymnosperms) and of sculptured sporoderms (in the Late Cretaceous to Eocene angiosperms) alternating with those marked by the prevalence of asaccate and unsculptured ones. A recurrence of amphisporions (the dispersal units of heterosporous plants consisting of megaspores bearing microspores in their sporoderm appendages) is a striking example. Such structures first appeared in the Devonian Kryshtofovichia, a morphotype of spiny megaspore with an apical “androcamera”, produced by some heterosporous lycopsids (Nikitin, 1934). About 300 m.y. later, much similar structures reappeared in a Cretaceous heterosporous fern Heroleandra (Krassilov & Golovneva, 1999, 2000; Fig.50). The amphisporions provide for a joint long-distance dispersal of male and female gametophytes, advantageous for colonizers. During the Cretaceous, fresh water environments were invaded by floating heterosporous ferns filling a vacant ecological niche as the heterosporous lycopsids did in the Devonian. Certain characters of angiosperm pollen grains, considered as typical of the group, are actually recurrent. These include the lamellate endexine, endoapertures and the tapetal surface deposits, all found already in the Permian gymnosperm pollen grains (a lamellate endexine in Cladatina, a porous endoaperture in Protohaploxypinus, a supratectal layer in the in situ pollen grains of Permotheca: Krassilov et al., 1999c, 1999d). Further studies will show how far the Permian gymnosperms anticipated the angiosperm ultrastructures. Disparity of the sporophyte–gametophyte evolution rates. Still another aspect of sporophyte–pollen grain correlation, or lack of such, is a temporal disparity of morphological innovations, with the sporophytes taking the lead at one time, the pollen grains at another. There are examples of an exceptionally conservative pollen morphology that remained virtually unchanged over great intervals of geological time while the sporophytic organs made a conspicuous progress, and vice versa. Thus, the protosaccate Vesicaspora morphotype first appeared in the Pennsylvanian callistophytes, a small group of seed ferns, the geological range of which did


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not extend over the Carboniferous. Vesicaspora, however, survived up to the Permian as a pollen type of callipterids, the supposed descendents of the callistophytes, differing in the strobilate reproductive structures advanced in the direction of the cycadophyte grade. This pollen type had been inherited by the Early Mesozoic peltasperms and cycadophytes and was still retained in the Cretaceous proangiospermous Preflosella (Krassilov & Bugdaeva, 1999). At least two species of xyelid insects fed on it in the Cretaceous (Figs. 131 – 133). A longevity of Vesicasporatype morphology might have been due to the loyalty of an ecologically conservative pollinivorous insect lineage. The counter-examples of a pollen morphology changing rapidly relative to sporophyte structures can be found in the geological history of peltasperms and their derived cycadophytes. Incidentally, a saltational transition from a protosaccate morphology in Permotheca to an asaccate one in its derived Antevsia was not accompanied by an adequate change in either these pollen-producing organs or the associated vegetative morphology.

a

b

c

Fig. 131. Insects and proangiospems: (a, b) pollen grains form sporangia of a preflower, Preflosella and (c, d) the same morphotype from gut compression of Chaetoxyella, a pollen-eating xyelid insect. The Early Cretaceous of Baisa, Transbaikalia.

d

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a


c

d

b

e

Fig. 132. Insects and proangiosoperms: (a) SEM micrograph of Ceroxyella dolichocera, a pollinivorous xyelid insect with a load of Eucommiidites-type pollen grains; (b) distribution of pollen grains over the intestine, (c) pollen grain morphology with lateral furrows, (d) granulate-columellate infarstructure of a lateral furrow, (e) indigested pollen grains at the hind end of the abdomen. The Early Cretaceous of Baisa, Transbaikalia.


a Fig. 133. Insects and proangiosoperms: (a, b) The pollinivorous species Ceroxyella from the Early Cretaceous of Baisa, Transbaikalia, feeding on pollen grains of two proangiosperm species, perhaps changing its diet over season: (a) an individual feeding on Eucommiidites-type pollen grains, (b) abdomen of the same specimen magnified to show pollen grains, (c) light micrograph of a pollen grain, (d) an individual feeding on saccate Vitimipollis edulis pollen grains like those produced by Preflossella, another proangiosperm of the same assemblage (Fig. 131).

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b c

d

Functional changes. Lack of morphological correlation in the evolutionary development of spores (pollen grains) and their producing sporophytes is related to their different functional environments. In spores, conspicuous morphological variation is often associated with a shift from water to wind dispersal that affects their area to volume ratios as well as the adaptations for flotation or anchoring. The germination environments and the longevity of gametophytes, as well as the rates of their endosporic, aquatic, above-ground or underground developments might have had an appreciable effect on spore size, nutrient storage and the durability of sporoderms. Sporivory might play a role in the Devonian (VIII.1.3), but was rare afterwards. In seed plants, the mode of dispersal is likewise a decisive factor placing specific demands on the modes of discharge, adhesion, catchments, volume regulation, etc.

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Morphological changes in respect to these various functions involve multifunctional structures, such as bladders. Their commonly recognized role in wind dispersal may not be the major one, and certainly is not the only one. Bladders regulate pollen volume and, when distally inclined, protect the aperture. They facilitate pollen catchment by spinning wind-borne grains in the air vortices between the spiral cone scales, also assisting in the ascend of captured pollen grains through the micropylar canal of anatropous ovules. Bladders allow fewer grains in the pollen chamber, a sperm competition factor. Protosaccate bladders of Palaeozoic gymnosperms might have attracted pollinivorous pollinators. Their occasional occurrence in the Mesozoic forms, such as Preflosella (Fig. 131), might have been due to a retention of Palaeozoic feeding habits. At present only the larvae of some Lepidoptera and beetles feed on saccate pollen. Vestigial sacci occasionally develop in pollen grains of the extant ginkgo and cycads indicating the origins from a saccate prototype (Herzfeld, 1927). The ginkgophytes and cycadophytes were simultaneously raised to dominance during the Mesozoic. Their nearly identical pollen morphologies probably resulted from parallel evolution. A reduction of sacci took place also in the ancestry of angiosperms. Noteworthy, the reduction of bladders was accompanied by destrobilation of seedbearing structures (Ginkgo, Cycas, flowering plants), related to endozoochory, but with a side effect on insect feeding and pollination. The destrobilized seed organs have lost their value as shelters for insects (retained by the pollen cones alone) while at the same time the appearance of fleshy seeds opened a new niche of larval seminivory. In time these changes might have disentangled pollination from pollinivory. As discussed above, specialized pollination by pollinivorous insects tends to increase morphological complexity, as in the Permian taeniate morphotypes, Mesozoic Classopollis and the allied cingulate-rimulate forms (Fig. 134). Yet a retention of a simple morphology in specialized pollination systems might have occurred when pollen ceased to be a major attractant for pollinators. Thus, extant cycads are pollinated by weevils that feed mainly on the cone rachis and sporophyll tissues (Terry, 2001). The cycad– weevil association probably goes back to the Jurassic, while thrips, another group of cycad pollinators, might initially have sought shelter in the pollen cones, only secondarily acquiring a pollinivorous habit (both larvae and adults now feed on pollen of Macrozamia: Terry, 2001). Evolutionary grades. A transition from the Palaeozoic to Mesozoic grades of palynomorphological evolution was marked by displacement of sulcate apertures from the proximal to the distal pole. This process, which had already started by the Pennsylvanian (Vesicaspora: Millay & Eggert, 1970), continued over the early Mesozoic. It was accompanied by other changes, such as the loss of taeniae, a decline of monosaccate and, in particular, zonosaccate morphotypes, a transition from protosaccate to eusaccate infrastructures and, eventually, a parallel reduction of bladders in the rising Mesozoic gymnosperm orders.


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Fig. 134. Insects and proangiospems: Classopollis, a rimulate pollen morphotype in the gut content of a katydid insect, Aboilus, from the Late Jurassic of Karatau, Kazakhstan (Krassilov et al., 1997).

Over the transition from proximal to distal apertures, there might have been an intermediate stage of bipolar germination, the vestiges of which still occurred in some Mesozoic pollen morphotypes, such as Classopollis. Persistent tetrads of this widespread pollen genus were bound by sexinal filaments protruding from a proximal Y-mark (Fig. 134). Both the triangular Y-mark and the distal porous leptoma were covered with identical granular membranes suggesting a bipolar germination (Krassilov et al., 1997). In this and allied morphotypes, an annular zonosaccus of a Cordaitina-like Palaeozoic prototype became modified as a volume-regulating rimulate/cingulate structure. A radical change in dominant pollen morphologies at about the Permian/Triassic boundary is commonly related to climatic changes, in particular to aridization that placed emphasis on volume regulation and protection of the apertures. At the onset of Mesozoic, however, the volume-regulating structures, such as taeniae or protosacci, were reduced or lost in the dominant gymnosperm groups in favour of a simpler asaccate anasulcate morphology, with the apertures protected by the involute margins alone. In effect, the

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overall pollen diversity decreased considerably relative to the Permian, mainly at the expense of morphologically advanced pollen types. Both morphological diversity and the complexity of pollen structures are correlated with specialized plant–insect interactions. Remarkably, the most sophisticate pollen morphotypes like Protohaploxypinus, Vittatina, Cladaitina, Classopollis, Eucommiidites, etc. were found in the gut compressions of Palaeozoic and Mesozoic insects (these spectacular pollen morphologies are well-known for they fascinated not only insects but also palynologists). A turnover of terrestrial ecosystem at the Permian/Triassic boundary disrupted the plant-insect interactions resulting in a temporarily reduced morphological complexity as well as diversity. During the transitional period, palynomorphological evolution was guided by the pioneer adaptive strategies, in particular, a mixed pollination strategy of autogamy, anemophily and unspecialised entomophily. Pioneer forms commonly produce many small ovules that impose a diminution tendency on the pollen grains. Elimination of specialized pollinators made such attraction structures as taeniae superfluous while a switch to windpollination demanded an instantaneous discharge of free-flowing pollen, hence the loss

Fig. 135. A similarity of reticulate infrastructure in a Permian pollen morphotype Lunatisporites (a) and a surface reticulum in the in situ pollen grains of mid-Cretaceous angiosperm Freyantha sibirica as evidence of paedomorphic surfacing of exinal structures.


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of sticky surface structures and substances. Finally, the pollen structures of a self-incompatibility syndrome were lost in autogamous forms. Thus, climatic changes, rather than directly affecting pollen morphologies, alter them by destabilizing the co-adapted pollination systems. Reproductive strategies degrade to a redundant Niobe mode (VIII.3.7), with respective changes of pollination ecology. The Niobe strategists do well with primitive ecologies and morphologies. A subsequent stabilization sets forth the opposite tendencies, with a gradual recovery of morphological complexity. Early angiosperms might have appeared as pioneer forms of proangiosperm communities (VIII.3.3) producing various pollen morphologies, none of which might have been typically angiospermous. Yet an increase of growth rates dictated a precocious sexual development affecting all gametophytic structures including the pollen grains. Semitectate reticulate surface structures of early angiosperms resemble a reticulate infratectum of some gymnosperm morphotypes (Fig. 135). Such surfacing of exinal infrastructures is evidence of paedogenic transformation. The vestigial bladders were modified as zonosulcate or transequatorial apertures facilitating exovulate germination. With it, the micropylar pollen drop exudates that attract insects to gymnosperm ovules were lost in angiosperms. a

c

b

d

Fig. 136. Perianth nectaries first recorded in the mid-Cretaceous flowers, signaling a major change in the pollination strategy: (a, b) Freyantha sibirica, a racemose inflorescence of small bracteate flowers from the Cenomanian of Yenisey Basin (Krassilov & Golovneva, 2001), (c, d) the spur-like glandular appendages on calyptrate bracts.

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Before the appearance of nectaries (recently described in Cenomanian ranunculids: Krassilov & Golovneva, 2001; Fig. 136), the inconspicuous flowers of early angiosperms were visited for pollen mainly. This gave an advantage to bisexual forms simultaneously enhancing the self-incompatibility functions conferred to the paedogenic pollen structures, such as the perforate – collumellate sexines. In conclusion, the fossil plant record attests to independence of palynomorphological evolution from sporophyte developments. In pollen grains, the morphological changes over times were controlled by the functional environments of the free-living stage. Convergent pollen grain morphologies appeared in divergent sporophyte lineages. Rapid sporophyte evolution concurred with conservative palynomorphology and vice versa. Critical for palynomorphological evolution were the ecosystem disturbances that disrupted the plant – insect co-adaptations conferring the loss of morphological complexity. Spore-pollen structures typically responded to such impacts through substitution of functions and paedomorphic innovations.

Chapter 9. Crises

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IX. CRISES The geological record is punctuated by biotic turnovers that coincide with the major tectonic, eustatic and climatic events (Krassilov, 1969, 1973a, 1978b, 1995a, etc.; Wethey, 1985; Erwin, 1990; Stanley & Yang, 1994; Retallack, 1995). Evolutionary significance of such biospheric events, discarded by the traditional evolutionary theory, presently attracts more attention in view of the nascent environmental hazards. Etymologically, crisis means a turning point. The major geochronological boundaries, such as the Permian/Triassic (PTB) ca. 290 Ma or the Cretaceous/Tertiary (KTB) ca. 65 Ma, are marked by the conspicuous diversity lows, hence they are turning points in respect to the manistream build-up of biological diversity. In historic sciences, patterns emerge from repetition. To arrive at regularities of biospheric crises we have to compare the sequences of geological and biological events across the critical boundaries. As will be shown later in the chapter, the similarities of transboundary sequences involve the geomagnetic, sea-level and magmatic events, the lithological, geochemical and isotopic anomalies, and the turnovers of biotic communities. The Permian/Triassic and the Cretaceous/Tertiary crises, whatever the causes, are generally held to have been the most severe in the history of the Earth. As such they provide a testing ground for evolutionary models, as well as for stratigraphic practices. It must be reminded here that chronostratigraphic boundaries are currently conceived of as time levels defined by selected points (spikes) in their type sections. Conventionality of the boundary choice is rooted in the Lyellian–Darwinian gradualistic interpretation of the fossil record and, at a deeper philosophical level, in the principle of continuity denying natural boundaries as such. Stratigraphic methodology (e.g. Remane et al., 1996) light-heartedly confuses time with history, disregarding the developmental discontinuity of geobiological systems that confer a distinctness on chronostratigraphic divisions. In recent years, the conventional golden spikes have been substantiated, first, by the iridium spike at the KTB and, later, by the isotopic δ13C spikes at this and other major boundaries. Their interpretation as the signatures of occasional events (e.g., impacts of extraterrestrial bodies) inflicting a radical turn upon the course of the earth’s history implies no regularities of evolutionary developments. It remains to be learned if such indeterministic attitudes agree with what is actually known about the turning points.

IX.1. Transboundary trends In both the PTB and KTB examples, geomagnetic signatures indicate a major event before the boundary, following a long quiescent period (Fig. 137). In the Permian, a period of constant polarity lasted from the Sakmarian, ca. 285 Ma, to mid-Tatarian (Capitanian), ca. 260 Ma (Jin et al., 1997, 1998). Starting from the latter date, constant polar-

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Fig. 137. A comparison of palaeomagnetic polarity sequences over the Permian and Cretaceous, both showing long quiescent intervals, with the reversal frequencies increasing towards the end of the periods (compiled after Channel & Ogg, 1982; Harland et al., 1990; Jin et al., 1997; La Brecque et al., 1982; Van Hinte, 1976). R/T – the concomitant regression – transgression trends.

ity changed to a mixed polarity, with the frequency of magnetic reversals increasing in the latest Permian (Changhsingian) and over the PTB at about 250 Ma. Likewise, there was an outstandingly long period of constant polarity in the Cretaceous, from the lower Aptian, ca.114 Ma, to lower Campanian, ca. 77 Ma. The following interval of a mixed polarity encompassed the rest of the Cretaceous and extended over the KTB, with two distinct reversals in the Campanian and four in the Maastrichtian. The time elapsed from the start of reversals until the boundary is about 10-12 m.y. in both cases. The geomagnetic quiescent periods roughly correspond to the global transgressions of epeiric seas interrupted by a few retreats. Closer to the end of the periods, the ascending sea-level trend switched to a major regression (with minor transgressive episodes at about the midway, in the Severodvinian and the mid-Maastrichtian respectively), culminating at about the critical boundaries (yet the Induan/Griesbachian transgression started in the latest Changhsingian prior to the currently accepted PTB based on a conodont zonation).

Chapter 9. Crises

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The Late Cretaceous climatic curve conforms to the concomitant sea-level curve, the warming trend progressing, with transgression, from mid-Cenomanian to Campanian, interrupted by a brief cooling event in the Coniacian (Krassilov, 1975a). This climatic trend was reversed with temperization and a spread of deciduous vegetation in the late Maastrichtian and over the KTB (Krassilov, 1975a and elsewhere; Golovneva, 1994). In the Late Permian, the Kazanian (Wardian) age was the warmest, with thermophilic life forms prominent both in marine and terrestrial records. A turn to temperization in the Tatarian is suggested by the phytogeographic changes, with the Angarian/ Subangarian floristic elements spreading south to northern Cathaysia. To the end of the Permian, the northern Cathaysia lost its floristic distinctness while the tropical vegetation survived in the southern Cathaysia alone (Krassilov, 2000b; Krassilov & Naugolnykh, 2001; IV.3). A cooling over the latest Permian (Changhsingian) to the basal Triassic (early Griesbachian) has been inferred also from conodont provinciality (Mei & Henderson, 2001). Monstrosity, diminution and prolific reproduction in a few generalist species as well as the loss of smaller populations are the biotic signals of approaching biospheric crisis. The loss of β and γ diversities is commonly ascribed to a floral/faunal mixing over new migration routes, in particular over the emerging land bridges. However, the homogeneous Early Triassic vegetation has spread with a transgression of epeiric seaways. The fine-grain strategy and cosmopolitanism of the survivors might have been of a more general significance, obliterating both ecological differentiation and provinciality. The associated δ13C fluctuations indicate a release of the lighter carbon from the biota attesting to a decrease in the total biomass with a drop of diversity.

IX.2. Boundary events Traditionally, the KTB has been defined by the last appearances of ammonoids in the marine facies and of dinosaurs in the terrestrial realm. In the North American stratigraphic practice, the boundary has traditionally been drawn at the base of the first coal seam above the last dinosaur bone, that is, at some distance above the dinosaur extinction level. The iridium spike in non-marine sequences coincides with this conventional boundary (Orth et al., 1980). A definition based on planktic foraminifers, nannoplankton and palynological zonation places the KTB above the macrofossil turnover. The presently widely held theory of extraterrestrial impact gives priority to the “boundary clay” with shocked quartz, microspherules and a high iridium concentration, allegedly a deposition of the impact ejecta, as a universal KTB marker. The extraterrestrial impact story started at Gubbio, Italy (Alvarez et al., 1981), where the iridium spike hits at a montmorillonitic clay 1 cm thick at the base of the lowermost Palaeocene foraminiferal Globigerina eugubbina zone. The section comprises several

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bentonite horizons and the “boundary clay” appears as one of them. The mass extinctions of the typical Cretaceous foraminifers and nannoplankton are recorded about 0.5 m below the “boundary clay”, whereas the macrofaunal extinction level occurs a few meters downsection. In other well-studied sections, such as Caravaça in Spain, the “boundary clay” is much thicker than at Gubbio, and the Cretaceous foraminifers Hedbergella and Globigerinelloides are found above it. Yet the situation in the type Danian section is even more complicated (compare to the Meishan PTB section, below), with Globigerina eugubbina first appearing below the “boundary clay”, and the iridium anomaly extending up to the anoxic black shale facies enriched with sederophile elements (Häkansson & Hansen, 1979). These discrepant occurrences seem rather in favour of several levels of iridium enrichment within the transitional interval. The KTB is marked by widespread rifting with basaltic volcanism, both intracratonic and over the marginal fold belts (V.7). Volcanic island-arcs emerged over the western Pacific margin (VII.6). The most profuse phase of Deccan traps is correlated with marine deposits containing Hercoglossa danica (or a closely allied form), the radiometric dates fall between 65 and 60 Ma (Mahoney et al., 1982; Sahni, 1988; Subbarao & Sukheswala, 1981; Subbarao, 1999). These volcanic events provided an abundant source of geochemically diverse ash deposits. In non-marine deposits, a well-defined iridium spike is confined to a tonstein, the clayey soil of a lower coal seam in the Palaeocene coal-bearing sequences (Orth et al., 1981; Bohor et al., 1987; Nichols et al., 1990). In an outcrop at Sussex, Powder River Basin, Wyoming (Fig. 138), the “boundary clay” between the dark grey Maastrichtian shales and the lower coal seam is a light brownish-grey argillite with pyroclastic clasts, about 10 cm thick, covered with an ochreous crust containing the ferruginous remains of slender stems and roots. The metallic microspherules with iridium are concentrated in the ochreous crust, thus suggesting a post-sedimentary enrichment, perhaps enhanced by humic acids form the peat soil, rather than a direct fallout that would have brought the spherules to the bottom rather than on top of the clay horizon. Proteacidites, the latest Cretaceous palynological zonal marker, disappears below the “boundary clay” while the trilete fern spores (the “fern spike”) are found in it, yet below the iridium-rich ochreous layer. This is interpreted as a record of a radical vegetation change, allegedly devastation of forests followed by a spread of ferns (Nichols et al., 1990). It must be reminded in respect to these palynological findings that fern spores are as a rule far more resistant to chemical weathering than angiosperm pollen grains (this taphonomic factor does not negate the fern spike, but probably contributed to it). Such features of the terrestrial “boundary clay” as the metallic nodules, semifusenite and fern spores, rather than being evidence of extraterrestrial impact, vegetation burning and colonization of bare grounds by ferns, might owe their origin to waterlogging over an ash layer, scavenging of noble metals from below the redox boundary, and sulphur acid fusenization of plant material as is typical of tonsteins (Crawley et al., 1994).

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b

c

d

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Fig. 138. Iridium spike at the Cretaceous–Tertiary boundary at Sussex, Powder River Basin, Wyoming: (a) view of the outcrop, the transboundary sequence dug up, (b) the boundary of the Maastrichtian marine shales and the Palaeocene coal-bearing strata; red pen-knife points to the “boundary clay” beneath the lower coal seam, (c) light grey “boundary clay” on eroded surface of the dark grey Maastrichtian shale; the palynological boundary occurs at the base of the tonstein, iridium is concentrated in the pedogenic ochric layer on top of it, (d) surface of the ochric layer with plant debris.

Since ferns often spread over fresh ashbeds (Kornaú, 1978), the association of fern spores with tonsteins has an ecological explanation (in addition to the taphonomic one, above). In a thick trans-KTB tuffaceous sequence of Augustovka River, Sakhalin (Krassilov, 1979; VII.6), a fern bed, the greenish-grey pelitic tuffite with fronds of Cladophlebis (Osmunda), Dicksonia and Onoclea on the bedding plane, overlays a coaly fossil soil with horsetail rhizomes, indicating waterlogging in association with the ashfall. Fern spikes have also been recorded at the equivalent stratigraphic levels in other parts of the world allegedly indicating a global deforestation event (Vajda et al., 2001). However, if representing a pioneer phase of temperate rainforests, like the presentday tree-fern groves in the Nothofagus belt of southern Australia (Fig. 139), they might have actually signalled afforestation at the expense of the Cretaceous open shrublands (IX.9.1). The iridium enrichment at the KTB has inspired a search for similar phenomena elsewhere over the Phanerozoic sequences. And the anomalies have been actually found at several critical stratigraphic levels from the Precambrian to Eocene, including the

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Fig. 139. Tree-ferns in the understorey of an upland Nothofagus rainforest above the open lowland vegetation, Victoria, Australia: an altitudinal differentiation that would render an increase in fern spores with an upland–lowland shift of vegetation belts.

PTB. In the well-studied Meishan sequence, South China, the Permian/Triassic transitional interval between the latest Permian Paratirolites and the lowermost Triassic Otoceras ammonoid zones contains an easily distinguishable layer (about 10 cm) of “white (boundary) clay”, a hydrolysed ash with microspherules. According to Yang et al. (1995), the “white clay” is traceable over 12 Chinese provinces covering more than one million sq. kilometres. It is non-fossiliferous except a few latest Permian (Changhsingian) conodonts. The overlying black clay, an organic-rich calcareous laminated claystone with even more abundant metallic and glassy microspherules, is enriched with siderophile elements and bears an iridium spike. Mass extinctions of reefal macrobenthos occur well below the transitional interval, within which the highest extinction rates (up to 94%) in foraminifers, ostracods and cephalopods are recorded at the base of the “white clay” and immediately above it (Jin et al., 2000). The conodonts and bivalves show a more gradual decline while a few brachiopod relicts survive up to the secondary mass extinction level in the Early Triassic.

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Lithological equivalents of the boundary clay are recognized over the marine transboundary sequences of the Elburs, Caucasus, Salt Range and elsewhere. The black clay is assigned to the base of the Induan Otoceras woodwardii zone that is typically represented by black shale facies (Yang et al., 1995). Low in the Otoceras zone, the Triassic newcomers are still mixed with the relatively more abundant Permian survivors. The conodont assemblage of the Hindeodus parvus zone (corresponding to the upper part of Otoceras zone in the Griesbachian of Ellesmere Island and elsewhere) still retains a mixed aspect but with the Triassic newcomers coming forth, hence an option of elevating the PTB to this level. In the non-marine sequences of Volga-Uralian region, an area of historical Permian stratotypes, the PTB has traditionally been considered as unconformable, with a prominent hiatus between the Tatarian and Vetlugian stages (Lozovsky & Esaulova, 1998). However, the chronostratigraphic PTB does not necessarily falls at the hiatus. In the Nedubrovo Section, Vologda Region, a smectitic black clay at the base of the Vetlugian (Fig. 140) yielded an essentially Late Permian (Zechsteinian) assemblage (Fig. 141) of peltasperms (Tatarina) and coniferoids (Ullmannia,

a

b

c

d

Fig. 140. A transboundary Permian–Triassic non-marine redbed sequence at Nedubrovo, Vologda Region, European Russia: (a) view of the outcrop, (b) “boundary clay”, a tuffaceous smectitic layer, dark greenishgrey, light brown when dry, with abundant plant debris, (c, d) leaf compression, xylem fragments, leaf and seed cuticles on a bedding plane of the smectitic clay.

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a

b e

d c

f

g

Fig. 141. A transitional Permian/Triassic plant assemblage from the smectitic “boundary clay” of Nedubrovo, Vologda Region, European Russia (Krassilov et al., 1999a, 2000): (a-d) Tatarina (Raphidopteris) antiqua, a Permian peltasperm survivor (pinna fragment, lower cuticle, stomatal papillae), (e) Otynisporites eotriassicus, a megaspore index species of the basal Buntsandstein in Central Europe, (f) Quadrocladus dvinensis and (g) Ullmannia cf. bronnii, the Permian (Vyatkian–Zechsteinian) conifer survivors.

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Pseudovoltzia, Quadrocladus). The palynological assemblage is of a mixed latest Permian/earliest Triassic composition, with Klausipollenites shaubergeri, a Zechsteinian form, associating with abundant Cycadopites and Ephedripites, the essentially Mesozoic pollen types (Krassilov et al., 1999a). It contains also the megaspores Otynisporites eotriassicus characteristic of the basal Triassic or transitional deposits of Central Europe and China as well as Tympanicysta (Fig. 142), the algal filaments formerly assigned to marine fungi. The arthropod and tetrapod assemblages consist both of Permian and Triassic forms. Lystrosaurus, a tetrapod marker of the basal Triassic, came from a somewhat higher horizon in the Vetlugian redbeds. Mixed complexes of this type, with Quadrocladus and Lepidopteris, were found in the lower tuffaceous horizons of the intracratonic trap series in northern Siberia (Sadovnikov, 1997) and elsewhere. In the Lopingian of southern China, the floristic succession parallels that of the VolgaUralian region. The Cordaites-type leaves, dominant over the Permian, last appeared at the roughly equivalent stratigraphic levels in the Uchiapingian and Late Tatarian (Severodvynian Horizon), sequences, as well as at the base of the Siberian trap series (Sadovnikov, 1997). Diverse conifer assemblages of a Zechsteinian aspect prevail both over the Changhsingian and in the uppermost Tatarian (Vyatkian Horizon). And the transitional peltasperm–conifer assemblage of Nedubrovo is equivalent to the latest Changhsingian (upper Talung) plant-bearing facies just prior to the end-Permian transgression (Leonova et al., 1999). The uppermost Permian is represented by tuffaceous facies in the Talung Formation and elsewhere. The above comparisons suggest a correlation of the PTB across the provincial/facial realms. The smectitic black clay with mixed faunistic/floristic assemblages is traceable from marine to estuarine facies, with Tympanicysta common in both indicating an algal

Fig. 142. Tympanicysta, a filamentous green alga from the transboundary Permian/Triassic deposits of Nedubrovo showing chloroplasts and infolded cell joints. An abundant occurrence of this microfossil over the Permian/Triassic boundary in Nedubrovo and elsewhere indicates eutrophication of ponded estuaries spreading over marine shallows (Krassilov et al., 1999a).

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bloom that provided an abundant source of organic matter conceivably contributing to eutrophication and anoxy (IX.3). The black clay level is bracketed between the mass extinction of macrobenthos below and the turnover of microplankton above, at the base of the H. parvus zone. Within this interval, the progenitorial Triassic forms of marine invertebrates, terrestrial plants and tetrapods, such as Claraia, Eumorphotis, Lepidopteris, Lystrosaurus, etc. appeared amidst the Permian survivors. A major phase of the Siberian traps and the coeval magmatic events (notably, the end-Permian Emeishan igneous province of southwestern China: Thompson et al., 2001) fall in the same time interval providing an explanation of the worldwide bentonite layers, such as the boundary clay (Table 9). Table 9. Interfacial correlation of transboundary events: the KTB.

SL

δ13C

R

D

E

E

G

C

R

R

E

E

S

A

S

S

I

E

O N

Marine carbonate facies Iridium spike Base of Globigerina eugubina zone

In “boundary clay” (Gubbio) and additional levels elsewhere

First appearance of Below the G. eugubbina “boundary clay” in Steven Klint and Gubbio Last appearances of In and above the “boundary clay” Globigerinoides, at Caravaça Hedbegella, etc.

Mass extinction of 50 cm below the “boundary clay” Cretaceous at Gubbio micropankton Top of Abatomphalus (globotruncanids, Archangelskiella) mayaroensis zone Carbonates, Macrofaunal black shales few extinctions meters below the “boundary clay”

Non-marine and marginal facies Palynomorph (Proteacidites) extinction, fern spike below the iridium spike

Tonsteins with microspherules at the base of coal seams few meters above the last dinosaurs in North America

Last appearance of Cretaceous survivors in fossil plant assemblages (Nilssonia)

Basaltic tuffs, island arc volcanics in Western Pacific, main phase of Deccan traps

Last appearance of dinosaurs ca. 3 m below the palynological boundary

Sand/shale cyclothems, fan conglomerates at Tsagajan

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Table 9. Interfacial correlation of transboundary events, continued: the PTB. SL

δ13C

South China (Yang et al., 1995; Jin et al., 2000)

R

I N C R E S E

Hindeodus parvus = upper Otoceras zone

I S E

D R O P

D E P L E T I O N

Volga-Severnaya Dvina (Lozovsky& Esaulova, 1998; Krassilov et al., 1999a) Triassic Triassic vertebrate Ryabian newcomers fauna Horizon of prevail over (Wetlugasaurus, Vetlugian Permian relicts Tupilakosaurus etc.), Series in both benthic first appearance of and planktic Pleuromeia in plant assemblages assemblages

Black Clay = lower First appearing Otoceras Triassic woodwardi zone newcomers (Claraia, Eumorphotis), Changhsingian conodont extinction, anoxic event, iridium spike

White “Boundary Clay” (bentonite) with microspherules

Last Changhsingian conodonts

Upper Talung tuffs of Laibin, Emeishan basalts

Gigantonoclea guizhouensis plant assemblage Shallow water benthic mass extinctions, several cm to 4 m below the “boundary clay”

PseudotirolitesPleurodontoceras zone

Floristic assemblage with the latest Vyatkian – Zechsteinian forms (Tatrina, Ullmannia), first appearance of megaspore species Otynisporites eotriassicus, first appearance of Lystrosaurus slightly above this level Transitional vertebrate fauna (Dicynodontidae, Elginia) Transitional palynological assemblage with Ephedripites Pareiasaurid gorgonopsid fauna Tatarina flora Disappearance of cordaitalean leaf morphotype

Taymyr, Siberia (Sadovnikov, 1997) Marinin Ust’kel’terian Horizon with Lepidopteris

Smectitic black clay, Nedubrov o Member, Vokhmian Horizon of Vetlugian Series

Putoranian Horizon, main trap phase, Permian conifers (Quadrocladus) predominate, Triassic forms (Lepidopteris) common

Molomian member of Vyatkian Horizon of the uppermost Tatarian

Lebedevian Horizon, tuffs with transitional flora

Vyatkian Horizon, variegate marls, redbeds

Tutonchan Horizon, vocanomictic, with fern flora

IX.3. Causal model of the boundary events Features shared by the PTB and KTB transitional sequences are the following: – A long, ca 40 m.y., geomagnetic stability terminated ca. 8-10 m.y. before the boundary, with the frequency of reversals increasing from this level on;

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– Global regression from about the same time-level on; – A conspicuous drop of δ13C, with a concomitant rise of carbonate dissolution levels; – Global spread of black shale facies; – Widespread rifting and intracratonic trap magmatism at or near the boundary; – Increase in terrestrial runoff as a source of eutrophication in estuaries and over the shoaling marine basins, with algal blooms (Tympanicysta at the PTB, Spirogyra at the KTB). – Metallic enrichment, with iridium spikes; – Temperization of the mid-latitudes, followed by a rapid warming; – Mixed faunas/floras with survivors outnumbering newcomers over the transitional interval; – Parallel replacements of ecologically dominant forms in the marine and terrestrial ecosystems; – Mass extinctions of macrobenthos prior to the planktic turnovers that mark the boundaries. Certainly, these features are shared not only by these two but also by other highranking geochronological boundaries, such as the Ordovician/Silurian, Frasnian/Famennian (McLaren, 1982, 1983), Triassic/Jurassic or Eocene/Oligocene (VII.6), globally marked by the drops of sea-level, massive flood basalts, in particular, the Late Eocene Ethiopian basalts over the East-African Rift System (Hofmann et al., 1997) and elsewhere, high iridium levels, δ13C fluctuations as well as temperization of the mid-latitudinal climates and vegetation. Shared features impel common causes that first manifest themselves by geomagnetic events originating deep in the earth’s interior, followed by the tectonomagmatic activation, regression, global cooling, and the biotic turnovers. Geomagnetic events. Geomagnetic field is generated by rotation of the earth’s metallic outer core against the earth’s mantle, with a molten shear zone at the boundary imposing a screening effect (reviewed in Vinnik et al., 1998; Lag et al., 1998). As discussed above (V.2), the interior melting zones are caused by a rotationally generated friction at the density boundaries. The geomagnetic events thus reflect the rotational dynamics of the earth. As a feedback, the rotation is decelerated by electromagnetic torque arising in the molten zone. Long intervals of constant polarity, such as the Sakmarian to early Tatarian in the Permian or the Aptian to Santonian in the Cretaceous, are indicative of geomagnetic stability over a rotational deceleration trend. In both cases, a switch to unstable regime of frequent geomagnetic reversals signalled a major global change literally emerging from the depths of the earth’s interior to the surface. Sea level. A correlation of constant polarity with transgressive sea-level trends is no coincidence since both are related to the earth’s rotation rates. With a decrease in angular velocity, the radial (centrifugal) component of rotational forcing induces a convergence of isostatically compensated continental and oceanic crust, with sea water from the shoaling oceanic depressions inundating the low standing continents (VI.1; Fig. 78).

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Conversely, a switch to a positive rotational acceleration, reflected in the geomagnetic reversal regime, simultaneously sets forth a geocratic tendency, with the oceanic floor subsiding and the epeiric seas retreating from continental interiors. A jigsaw pattern of minor fluctuations is caused by isostatic feedbacks of the water load. The geodetic components of sea-level change are at counterphase in the Boreal and Tethys realms. Yet general tendency transpires toward the critical boundaries. Traps. Since spheric areas increase with the radius, an elevation of an earth crust domain is accompanied by its extension. Fluctuations of centrifugal forcing alternatively confer elevation tendency on the continental and oceanic earth crust domains, with continental rifting and volcanism at counterphase with those of the sea floor (VI.I). Over the acceleration trends of transboundary evolution, the rising continental crust falls under an increasing extensional strain, eventually breaking at the intracratonic fault zones that serve as magmatic channels transferring the high-density material from mantle sources to the upper crust. Trap magmatism over thick crust of intercratonic areas, like those of Putorana in Siberia, Emeishan in southwestern China (Thompson et al., 2001) or Deccan Plateau, corresponds to the maximal elevation/extension of continental crust at culminations of geocratic trends. Large basaltic provinces are formed over the (incipient, as in the case of the Emeishan igneous province) rift zones. Huge masses of basaltic sills and lavas substantially increase the density of continental crust, thereby balancing it against the oceanic crust and, eventually, reversing the sea-level trend. Hence traps mark a turning point of transboundary developments. Trap magmatism is a relatively short-term event (in comparison with the causally related marginal fold-belt magmatism) but it leaves a long-lasting geoidal scar. The Putorana, Tunguska, Emeishan and Deccan traps are confined to a global meridional fault zone extending from the Yenisey Gulf in the north to the Gulf of Khambhat (Cambay) in the south. This presently inactive zone is still evident as a broad meridional depression of the present-day geoid (V.3). The role of traps as culmination markers of the geobiological crises is confirmed by their chronological correlation with other markers, such as the “boundary clays”. Eutrophication. The boundary marker beds, known as “boundary clays”, are peculiar products of eutrophication, anoxic sedimentary environments, shallow carbonate lysoclines, extremely low intrinsic sedimentation rates, and metallic pollution. In marine sequences they appear over regressive trends, with terrestrial runoff, no longer trapped by eperic seas, reaching directly to oceanic basins. An increased export of terrestrial dead mass explains such transboundary phenomena as a negative excursion of δ13C and conceivably even the iridium spike (if produced by chemical weathering of ultramafic substrates, see below). A spread of black clay at the PTB is accompanied by a mass appearance of Tympanicysta and allied genera of filamentous green algae (previously described as marine fungi, hence the PTB “fungal spike”). Notably, a rise of Spirogyra, a related genus of zygnematalean green algae, is recorded over the KTB (Douglas Nichols, pers. com.).

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Eutrophication of marine ecosystems is related to export of organic material from estuaries, the coastal sites of profuse biotic production (Mantoura, 1991; McAnally & Mehta, 2001; Pye & Allen, 2000; Woodruff et al., 2001). Presently the yearly primary production of estuarine ecosystems is about 1500 gram per sq. m, three times greater than in the tropical upwelling zones (Alimov, 1989). The coastal eutrophication is assessed as global phenomenon (Nixon, 1995), with an overwhelming effect on oceanic productivity and oxidation (Smith & Hollibaugh, 1993). Generally, at the time of global regression, the ratio of terrestrial inorganic to organic nutrients influxed to estuarine ecosystems tends to increase resulting in their autotrophic functioning, greatly enhanced by algal blooms (Oviatt et al., 1986; Smith & Hollibaugh, 1993; Kemp et al., 1997; Lin et al., 2001; II.1). The organic-rich estuarine facies with Tympanicysta, as at Nedubrovo (above), bear evidence of algal blooms inflamed by terretsrial runoff and contributing to the coastal eutrophication (II.7.2, VII.1.1). A spread of hypertrophic waters over carbonate platforms, as suggested by the presence of Tympanicysta in carbonate facies of the PTB stratotype at Meishan, South China (Jin et al., 1998; 2000) would have increased pH of marine waters, hence a rise of carbonate lysoclines and non-deposition/dissolution of carbonates at the “boundary clay” level. Likewise, the benthic foraminiferal assemblages indicate a rise of oceanic trophic levels in the terminal Maastrichtian followed by an abrupt drop after the KTB (Alegret et al., 2001). A rapid replacement of planktonic foraminiferal assemblages over the Danian (Gerstel et al., 1987) might have been related to the same environmental factor. Carbon isotope ratios. The δ13C fluctuations over both the KTB and PTB are variously explained by a decrease in biological productivity or in the ratio of terrestrial to marine production dropping with sea level, a release of methane through dissociation of oceanic gasohydrates, an emissions of CO2 from volcanic sources, etc. (Jasper & Gagosian, 1989; Kump, 1991; Dickens et al., 1995; Krull & Retallack, 2000, a review of the hypotheses in Pálfy et al., 2001). In the final count down, the 12C to 13C ratio depends on the terrestrial biomass production and its input to the oceans, which involves a number of interrelated variables. A tectonic, volcanic and epeiric activation over the critical boundaries could have caused a disruption of climax ecosystems with a release of light carbon and a respective depletion of δ13C in the shallow-water carbonate reservoir isotopically equilibrated with the atmosphere. Yet the isotopic anomaly need not be a long-lasting one, because soon after the crisis a spread of pioneer vegetation after the retreating epeiric seas would have provided a major sink for the light carbon. Global regressions at the KTB and PTB (or just prior to the current version of the latter: Jin et al., 1998; 2000) must have caused an increase in the terrestrial input and eutrophication of estuarine ecosystems (above), with an appreciable effect on marine productivity and the carbonate dissolution depths. An input of iron as a by-product of voluminous mafic volcanism (traps, island-arc basalts) would have increased the mobility of phosphorus over the sediment–water in-

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terface, thus acting as an amplifier of biotic production. This explains, among other things, the association of δ13C excursions with metallic pollution from the same sources. Hence, the direct effects of tectonomagmatic activation on isotopic ratios through volcanic emissions or tapping of hydrocarbon reservoirs by marginal faulting of the rising continental blocks, etc. were added, and conceivably greatly enhanced, through their impacts on biotic production. Detailed stratigraphic studies reveal a time lag of δ13C fluctuations in respect to the potentially involved geologicalal events (Wilson & Norris, 2001) perhaps reflecting a delayed response of the biosphere. While the opposite trends of δ13C and δ18O fluctuations suggest a climatic component to their regulating system, affecting both the biomass production and carbonate dissolution depths, their positive co-variance, such as a simultaneous decrease over the PTB (Hegdari et al., 2001), seems to have been due to an exposure and diagenetic alteration of carbonates during the non-deposition intervals, the frequency of which tends to increase over the critical boundaries owing to instability of sedimentation systems worldwide. Temperization. Both at the northern and southern mid-latitudes, temperate vegetation expanded over the Maastrichtian (a spread of Nothofagidites and of serrate leaf morphotypes, VII.5-6) indicating that temperization had commenced before the KTB, simultaneous with the major phase of ophiolite emplacement in the eastern branch of the Tethys belt (IV.V.6.5). A temperization prior to the PTB is reflected by penetration of Angarian and Gondwanic elements in the Cathaysian province (IV.3.3), also related to the closure of the Tethys seaways at the end of the Palaeozoic cycle. The mechanism of teleconnection between the Tethys orogenies and the polar ice caps is described in (VII.2.2). Over both critical boundaries, the cooling trends, implied by the vegetation changes, must have been enhanced as a net effect of elevated continents (an increased albedo), deeper oceans with a longer residence time for dissolved CO2, and an additional sink of atmospheric CO2 to fast-growing vegetation colonizing the emergent land. Insofar as sea surface cooling in turn increases CO2 dissolution rates, the process is self-accelerating. The oceanic climate-generating events are bound to the climatically driven developments on land through terrestrial runoff, the amount and composition of which, in particular the ratios of organic to inorganic nutrients, affects autotrophic functioning of aquatic ecosystems, with consequences for the redox stratification, carbon burial and carbonate dissolution rates (II.7.4). The feedback effects of temperization on ecosystem evolution are far-reaching, mediated by population strategies driven to the fine-grain mode by an increase in seasonality of climatic, as well as trophic, resources. The population densities tend to increase, with consequences for genetic polymorphism and speciation rates. The biotic communities are forwarded toward the oligo- to monodominant composition. The net impact on biological diversity is devastative, as predicted by Stanley (1986) in respect to the PTB mass extinctions.

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Dinosaurs, etc. Temperization leads to afforestation of mid-latitudes and a spread of deciduousness (VII.5). Since deciduous forests are inferior to open evergreen vegetation in resilience and in nutritional quality of the leaf mass, afforestation means a drastic reduction of trophic resources for the larger herbivores. Dinosaur extinction is commonly perceived as enigmatic, implying an exotic explanation. Yet dinosaurs were the larger herbivores dependent on phytomass production, and predators dependent on the herbivore biomass. The early diverging insectivorous line supposedly gave rise to the birds. The later mammalivorous branch might perish with multituberculates and other archaic mammals. In the history of dinosaurs there were at least two episodes of semiextinction at the Triassic/Jurassic boundary and in the mid-Cretaceous (the ornithopod/sauropod replacement), both related to vegetational changes. Afforestation over the KTB affected the habitats of the richest dinosaur faunas in North America, Amurland, Mongolia, northern China and elsewhere at the mid-latitudes. The end-Permian afforestation involved the same areas, with the Eurangarian vegetation extending to northern China. On taphonomic evidence, most of the dominant arboreal forms were shoot-droppers. This kind of deciduousness primarily affects the larger browsers that prune leafy shoots at high browsing levels (VIII.I.3). The residual dinosaur communities of forested areas might have foraged on the residual patches of cycadophyte thickets (as suggested by association of dinosaur remains with Encephalartites in the Maastrichtian of Koryak Highlands: Krassilov et al., 1990). Judged by the aquatic morphological adaptations and the gut contents, the Maastrichtian duck-bills consumed a low nutritional aquatic phytomass or gyttja (Krassilov, 1981a; Krassilov & Makulbekov, 1995). The trophic factor might have inflicted a turnover of tetrapod communities at the PTB as well. Most of the Permian terrestrial tetrapods became extinct towards the end of this period. The end-Permian to Early Triassic lystrosaurids were ecologically equivalent to the Maastrichtian hadrosaurids, both switching from browsing to aquatic grazing in response to afforestation and the growing aquatic production (eutrophication, above). Later in the history, afforestation over the shrub-tundra and savannoid “tundra-steppe” landscapes drove to extinction the larger browsers of the Late Pleistocene mammalian faunas (Ukraintseva, 1986). Iridium. The cosmic impact model of the KTB extinction (Alvarez et al., 1980) implies a giant crater, for which the 250 km wide Chicxulub is considered as the most plausible candidate. It is a multiring basin with a central peak ring and a prominent Moho rise beneath (Sharpton et al., 1992; Christeson et al., 2001). The latter feature seems to better agree with magmatic origins. The rock melts and brecciation by allogenic impacts (reviewed in Dressler & Reimold, 2001) are scarcely different from those caused by explosive volcanic eruptions. Whatever the origins of Chicxulub, its correlation with the KTB is unlikely because no evidence of an impact-generated (or a volcanic eruption generated, as it might be) tsunami has been found in the transboundary sequences of the

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area (Keller et al., 1994). In addition, the volume of Chicxulub ejecta has been drastically exaggerated (Morgan et al., 1997), as were also the estimates of thin dust residing in the stratosphere (Rampino & Self, 1982). In the shallow-water marine sequences of North America, the disturbed sediments ascribed to a giant impact occur below the last appearance of scaphitid ammonites while Wodehouseia spinata, a palynological zone marker of the terminal Cretaceous, was found both below and above the presumed impact level. This situation has been interpreted as evidence of ammonite survival into the Palaeocene (Terry et al., 2001). It may instead attest to a but vague correlation between the sedimentary disturbances of whatever origin and biotic turnovers. A single giant impact would scarcely account for the multiple iridium spikes detected over the trans-KTB sections (Crocket et al., 1988). At least some iridium-rich “boundary clays” undoubtedly contain volcanic material. Thus, at Steven Klint, Denmark, the boundary clay is a diagenetically altered volcanomictic (Mg-smectitic) marl bed, with noble metals concentrated on pyrite particles (Elliott et al., 1989). The microlayering of the “boundary clays” indicates metallic precipitation rather than a fallout (Schmitz, 1988). In the non-marine example (IX.3), the boundary clay is a tonstein, with iridium concentrated in the ochreous crust, a product of chemical weathering in a waterlogged rhizospheric environment. The anomaly is thus owing to a post-depositional concentration rather than primary deposition. At the average pelagic deposition rates ca. 1 cm/k.y. an intervention less than 1 k.y. has a slight chance of being recorded (Sadler & Dingus, 1982). No less than 70% of millennial-scale events are obscured by bioturbation (Anderson, 2001). Hence a sedimentary record of allogenic ejecta, with a resident time in the atmosphere of several years maximum, is of exceedingly low probability. Two such records are twice less probable and the probability decreases with each additional record. Since there are more than 80 marine records of iridium anomalies over the KTB, the probability that they reflect a single short-time event is negligible indeed. The same relates to the iridium anomalies at other critical boundaries. A multiple impact hypothesis developed by McGhee (1996) for the Devonian appears more plausible than the single impact model. Yet McGhee ascribed the extinctions to a stepwise cooling that might happen without any cosmic interventions as well. At the same time, the input of iridium to estuaries and marine basins by rivers globally many times exceeds that of an extraterrestrial origin (Anbar et al., 1997). Terrestrial sources of iridium must have been greatly enlarged by the transboundary tectonomagmatic and eustatic events. A major phase of ophiolite emplacements over the circumPacific belt, eastern Tethys and elsewhere falls at the Campanian–Maastrichtian, with the ultramafic bodies, rich in iridium and other platinoids (Garuti et al., 1997), exhumed and exposed to weathering over the KTB. Activation of marginal volcanic belts and island arcs at or close to the KTB is recorded over both the circum-Pacific (Fig. 143) and Tethys belts. High concentrations of platinoids are confined to the mantle xenoliths

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a

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Fig. 143. The Lesser Kuril Islands, a segment of the island-arc system emerging at about the Cretaceous/ Tertiary boundary: (a) view of Tyatya Volcano of the present-day island arc, (b) view of the Cretaceous/ Palaeocene island-arc basaltic sequence, (c) a giant pillow structure in the transboundary basaltic complex, (d) fossiliferous shales with foraminifers, marine mollusks and terrestrial plant remains overlain by lavobreccias.

of island-arc volcanics (Kepezhinskas & Defant, 2001) while the Alaskan-type concentric ultramafic-mafic intrusions contain economic platinoid resources 20 times exceeding the Alpine ophiolite estimates (Garuti et al., 1997; Kepezhinskas & Defant, 2001). Thus, in both the PTB and KTB situations, the iridium enrichment might have been owing to: – Global regressions inflicting erosion cycles, with an increased terrestrial runoff as a source of iridium in marine deposits; – Weathering of ultramafic bodies (in particular the harzburgites, an economic source of iridium) tectonically exhumed and uplifted during the Maastrichtian (V.7); – Volcanic activity tapping the mantle sources over the continental margin/island arc belts, with voluminous ignimbrite eruptions carrying ultramafic xenoliths rich in platinoids (the Maastrichtian of Sikhote Alin and elsewhere in the circum-Pacific belt :V.6.1);

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– Elevation of carbonate lysoclines inflicting a widespread carbonate dissolution leaving the clayey residue impregnated with metallic precipitates (the “boundary clays”); – Scavenging of iridium by Fe-Mn hydroxides in shallow water environments under rapid switches of redox boundary, resulting in iridium concentrations in pyritic spherules; – Similar developments in the waterlogged rhizospheres over ash beds producing the iridium-rich ferruginous crusts on top of smectitic “boundary clays”. Remarkably, an increase in rare elements in dinosaur bones has been recorded well before the KTB, in the middle Maastrichtian of Gobi (Samoilov et al., 1996). Qualitatively at least, the terrestrial sources have an advantage of placing iridium in the context of the transboundary developments involving the well-known tectonic, volcanic and sedimentary events, neither of which are accounted for by the extraterrestrial impact model. Not only iridium but also other metallic and non-metallic anomalies (e.g., the endPermian sulphur anomaly: Kaiho et al., 2001) have been ascribed to extraterrestrial impacts. Geochemical anomalies are the constant feature of the critical boundaries. Yet a much better documented source of anomalous geochemistry is an upheaval of mantle material through the marginal and intracratonic fractures and fissures of the uplifted landmasses. A scheme. Various developments over the critical boundaries make a coherent story when considered in the light of rotational geodynamics and its biospheric effects (Fig. 144). An early signal of approaching crisis comes from geomagnetic perturbations generated at the major density boundaries deep in the earth’s interior. Later on, the hypsometric levels of oceanic and continental domains diverge as a result of their differential, density-dependent, centrifugal acceleration. This activates the marginal shear zones, with the island-arc magmatism as a source of geochemical anomalies. Sea level drops, while the rising continental areas are expanded, rifted and impregnated with mantle material over the trans-cratonic fissures generating the trap fields. With a transfer of angular momentum from mantle to lithosphere the trend is reversed, hence the continental traps mark culmination of the crisis. The geological developments affect biota through their destabilizing effect on ecosystems and the respective changes of population strategies. Yet the direct effects acquire a biospheric significance when amplified by fluctuations of biomass production. Thus volcanic eruptions are source of iron that increases the mobility of phosphorus contributing to eutrophication conferring instability on freshwater and coastal ecosystems and through these on marine ecosystems. An increase of terrestrial runoff to the oceans with a retreat of epeiric seas results in redistribution of biotic production and reallocation of dead mass reflected in the anomalous carbon isotope ratios. Eutrophication of coastal ecosystems provides a prolific source of exported dead mass conferring instability on marine ecosystems.

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Fig. 144. An interpretation of biospheric crises initiated by rotation forcing of interrelated geomagnetic, magmatic, sea-level, climatic and biotic events.

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A long-term climate change is brought about by elevation of the continents and by closure of the Tethys seaways casting a cooling effect on both hemispheres. Temperization of mid-latitudes reduces the trophic resources of larger herbivores triggering a decrease of biomass up the trophic cascades. A shift to the fine-grain adaptive strategy with seasonality of climatic and trophic resources adds to the devastative net effect on the biological diversity.

IX. 4. Mass mortality and mass extinction Public interest in the impact story is due primarily to the alleged dramatic killing of dinosaurs and other charismatic animals. Actually, the transboundary mass extinctions, with a single major and two to several secondary fits, is correlated with a number of physical phenomena, of which the iridium spikes are among the most spectacular. Extinctions result in anomalous biological diversities that are related to biospheric crises in much the same way as the geochemical and isotopic anomalies, hence causally related. However, in the case of dinosaurs, a direct link with the iridium anomaly is unlikely because at the time of the supposed extraterrestrial impact they were just not there. In both the PTB and KTB examples, the dominant elements of terrestrial as well as marine biotas, such as dinosaurs, rudists, cephalopods, etc., were eliminated no less than 100 k.y. before the iridium spikes (Bugdaeva et al., 2001; Pearson et al., 2001; Kennedy et al., 2001). A closer, though not universally consistent, correlation of lithological markers (the “boundary clays”) with microplankton extinctions and palynological events, such as the fern spikes, seems to have been owing to a combination of depositional and postdepositional phenomena (IX.2), including an impact of ash deposition on permeability of substrates and on water chemistry. Taphonomic smearing (the Señor-Lipps effect) should have been of the same scale for microfossils and metallic spherules. Granted the reality of giant impacts, their immediate effect on biota would have been mass mortality that need not be confused with mass extinction. The instability of geological and biospheric systems over the critical boundaries provides multiple causes for mass mortality. By analogy with the presently declining spruce forests of northern Europe, soil acidity and metallic pollution related to volcanic eruptions can be inferred as a critical factor of mass mortality in the forest dominants, such as glossopterids, cordaites, Phoenicopsis, Parataxodium, etc. (IV.3). Dilution by fresh-water injections (Fischer, 1965) inflicting anoxic conditions down the water column is often considered as a critical factor for marine communities. An overturn of anoxic deep waters brings to the surface SO2 and CO2 which are poisonous for certain planktic organisms. Selective CO2 poisoning has been suggested for the Permian mass mortalities (Knoll et al., 1997) and is equally plausible for those at the KTB. For comparison, a prominent oceanic anoxic event at the Cenomanian/Turonian bound-

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ary (Bonarelli Level), about 96.4 Ma, is thought to have been responsible for about 24% of extinct benthic genera. This episode was marked by a short-term (the Milankovitch scale) increase in δ13C (Prokoph et al., 2001) tentatively ascribed to a rise of methanogenic Archaea (discussed in VII.1.2). Pelagic extinctions typically culminate at or slightly before the global black shale markers, such as the PTB black clay of South China, Salt Range, the Griesbachian of Ellesmere Island and elsewhere, as well as the KTB anoxic fish clay of the Danian stratotype, the Cangilia Formation of Greenland, etc. Foraminiferal turnovers occur at about these levels (Reddy & Yadagiri, 2001). Yet a reduction rather than elimination of trace fossils over the black shale intervals suggests but a mild near-bottom anoxy. A not so abrupt extinction of deposit-feeders is considered as nutrient-driven rather than directly related to the anoxy (Smith & Jeffrey, 1998; Jeffrey, 2001). Even more hazardous for pelagic ecosystems might have been a widespread hypertrophy of surface waters. In marine sections, such as at Meishan in South China, an increase of Tympanicysta is recorded before the extinction spike (Jin et al., 2000). Its abundance in the coeval paralic black shales (Krassilov et al., 1999b) suggests an algal bloom in the ponded estuaries, with eutrophic waters splayed over shelves and further, to pelagic environments. Eutrophic ecosystems are unstable, with large-scale fluctuations of primary production inflicting a drop of biological diversity up the trophic cascades (examples in Heinonen, 1980). Mass accumulations of cephalopod shells in shallow water facies are often accompanied by land plant debris indicating the role of terrestrial runoff. To take a few examples, the Cenomanian ammonites of Sel Bukhra locality, northern Crimea, occur in hard pyritic marls with drifted fern pinnules and coniferous scale leaves (Krassilov, 1984; II.7.4). The Lower Triassic (Olenekian) ammonoids of northern Siberia are crowded in pyritic nodules also containing the floating sporangia of Pleuromeia, a coastal wetland dominant of the time (Krassilov & Zakharov, 1975). In the Late Permian of South China, enormous accumulations of ammonoid shell impressions co-occur with cordaite-type leaves in bituminous shales at several levels over the Uchiapingian regressive sequence. Upsection, in the black shales of the Talung Formation, occasional shoots of ullmannioid conifers were found among the abundant ceratitid impressions covering the lavishly exposed bedding plane (Leonova et al., 2000). Such occurrences are evidence, among other things, of high population densities prior to extinction, the “Exodus effect” of redundant proliferation in the oppressed (VIII.3.7). To resume, the chronology of mass extinctions, diachronous as they were within the narrow transboundary intervals, is scarcely consistent with a single-cause model. Although each of the alleged causes might have been responsible for occasional mass mortality, mass extinction may or may not follow. A bedding-plane accumulation of perished organisms attests, in the first place, to a redundant density buffering population from extinction in the face of occasional environmental hazards (VIII.3.7). Yet a spread of such strategies may confer instability on biotic communities potentially causing losses

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of biological diversity. Mass extinction is a rough regulation of biological diversity via population strategies adjusted to long-lasting environmental disturbances.

IX.5. Replacement of ecological dominants The idea of biotic turnovers over the critical boundaries has originated from the replacements of charismatic groups, in particular, of dinosaurs by mammals. This latter replacement signifies the end of the Mesozoic and the beginning of the new era. Opponents of the idea have argued that, since dinosaurs constituted but a minor part of biological diversity, their elimination (typically considered in this line of thought as the result of a competitive replacement by mammals) does not signify a global biotic turnover of whatever, if any, scientific meaning. If we instead focus on the replacement of gymnosperms by angiosperms, then the boundary of geological eras (the Mesophytic/Cenophytic) must be placed in the middle of the Cretaceous. A choice of one or another option for the boundary is therefore a matter of convention. Such considerations have led many evolutionists to believe that stratigraphic levels of accelerated turnover rates can be objectively defined for particular lineages rather than for biota as a whole. The eras of life need not be taken literally as distinct stages of organic evolution. Some lineages do show high turnover rates over the selected boundaries, others do not. Plants have evolved under a different beat than animals while the dinosaur extinctions on land at approximately the same time level as the ammonite extinctions in the sea is a mere coincidence. However, with the rise of biospheric ecology, this logic no longer seems faultless. Ecological niches of the larger dinosaur herbivores seem to have remained vacant for at least 5 m.y. after their extinction, hence competitive elimination is unlikely. Some 150 m.y. long coexistence of dinosaurs and mammals only means that the dinosaurian communities and their contemporaneous mammalian communities were fairly well-equilibrated, with minimal niche overlaps, while the replacement means that this long-standing ecological equilibrium was eventually disrupted. Adaptive radiation of modern mammals postdates, rather than predates, the decline of dinosaurs. Archaic mammals perished simultaneously with dinosaurs, although over a somewhat longer time interval. The buildup of mammalian diversity is due to new types of vegetation, such as grassland, as well as new trophic niches, such as frugivory. The gymnosperm/angiosperm story is fairly consistent with the above version of the dinosaur/mammal story (Krassilov, 1997a and elsewhere). Both only make sense in the context of biospheric evolution. Angiosperms first appeared in the mid-Neocomian, concomitantly with the therian mammals (Aegialodon dawsoni and subsequent records of the metatherian–eutherian grade: Lillegraven, 1969; Lillegraven et al., 1979; Fox, 1980) and the true birds, jointly signifying a major innovation of terrestrial plant communities.

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This coincided with a major innovation of marine biota. Through the Neocomian, the planktic calpeonellid communities, dominant in the oligotrophic ecosystems of a Jurassic aspect were invaded by dinoflagellates, diatoms and primitive hedbergellid foraminifers. The concomitant replacement of green algae by red algae (Poignant, 1983) indicates an involvement of trophic factors related to an input of estuarine production to the oceans at the initial stage of Cretaceous transgressions. Eutrophication of surface waters was accompanied by a spread of black shale facies. In the benthic communities, the hitherto dominant buchian bivalves were replaced, from the Hauterivian on (Kemper, 1982), by the convergently similar, though perhaps more desoxy-tolerant, aucellinids. The bulldozer feeding habits (Thayer, 1979) and durofagy appearing with organic enrichment of marine sediments and the build up of trophic cascades, entailed a turnover of both softground and hardground epifaunal communities (Palmer, 1982; Jablonski et al., 1983). Lacustrine black shale facies have appeared at the equivalent stratigraphic levels over the trans-Asiatic rift system of Transbaikalia, Mongolia and northern China, attesting to highly productive aquatic ecosystems (their significance as diversification centres is discussed in VII.3.3). The next mid-Cretaceous turnover phase was likewise marked by prominent blackshale events in the oceans as well as over the lacustrine domain (VII.2.1). A popular idea of the mid-Cretaceous vegetation change as a rise of angiosperms – decline of gymnosperms is an oversimplification of what is actually recorded, i.e. the simultaneous innovations in all major plant groups correlated with the appearance of modernized plant communities. Thus the conifer-broadleaved forests first appeared during the mid-Cretaceous, with Pseudolarix, Sequoia and Parataxodium replacing the archaic Schizolepis, Athrotaxopsis and Elatides, and with angiosperms entering the understorey. While the archaic cycadophytes and ginkgophytes declined, the Cretaceous records of extant genera Cycas and Ginkgo testify to their involvement in the modernization process (Krassilov, 1978c). Angiosperms first became prominent in marginal (seral) communities such as the riparian platanoid woodlands, xeromorphic shrublands and the aquatic macrophyte vegetation (Krassilov, 1997a). Faunistic innovations were for the most part related to these latter types of plant communities. The appearance of new ecological niches of anthophiles, frugivores and folivores stirred diversification in the preadapted groups of insects and vertebrates, with the pollinivores switching to nectar feeding (see Krassilov & Golovneva, 2001on nectaries in Cenomanian angiosperms), the dinosaurs acquiring aquatic grazing habits, and the mammals exploring arboreal habitats (Lillegraven et al., 1979). The larger browsers consuming great amounts of low-nutritional leaf mass should have been most dramatically affected by a change in the leaf mass to xylem ratio with the microphyllous angiosperms replacing macrophyllous cycadophytes (Krassilov, 1981; VIII.1.3). Giant suropods declined, while an innovation of dental morphologies in herbivorous ornithopods point to a leading role of trophic factor, related to the vegetation change. As discussed above, the aquatic habits in hadrosaurids might have developed with the

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appearance, in the mid-Cretaceous, of a floating leaf mass and the ensuing significant increase in aquatic biomass production. The incipient divergence of marsupials and placentals had occurred in the Albian (Slaughter, 1981), while the earliest records of eutherian-dominated faunistic assemlages came from the Turonian–Santonian of Dzharakuduk in Uzbekistan (Archibald et al., 2001). These events are correlated with adaptive radiation of early angiosperms and a spread of angiosperm-dominated communities (in the Turonian of Central Asia and the Middle East). Yet angiosperms came to prevail, both in terms of dominance and species richness, toward the end of Cretaceous, simultaneously with the replacement of the metatherian–eutherian grade by modern mammals (Lillegraven et al., 1979). High species richness both of angiosperms and mammals was attained not earlier than the Late Palaeocene, when their dominant status had been fully established. When we think of dinosaurs and other charismatic groups in ecological terms, as of those exerting the most powerful influence on their environments, like elephants and hippopotamuses do today, or the ceratopsids and hadrosaurids did in the Late Cretaceous, we come to understanding extinctions and replacements as not occasional happenings, but as an outcome of concerted developments on the way from dilapidated biospheric structures to the new ones. Dominant forms of the fossil record are climax dominants in terms of ecological succession, and their replacements signify the cut-off/ recovery of climax stages.

IX.6. Climax cut-off As discussed in (VIII.4), ecological succession is a brief reiteration of ecosystem evolution through times. Both over evolution and ecological succession, the mainstream development is guided by the imperative of biomass production sustained through buildup of biological diversity. For both crisis means a reversion of the trend. Escalate biomass production not matched by structural complexity is not sustainable. This explains crises as disruptions of the biomass/diverisy correlations. Due to the parallelisms of evolution and seral developments the former can be explained as a time projection of the latter. With evolution, new forms enter certain stages of ecological succession, preferentially the early ones that are less stable and more open. The climax may not be affected due to equifinality of seral developments. Yet a persistence of seral stages indicates that the ecological succession is halted, lately if ever reaching to the potential climax (Fig. 145). When stressful conditions spread over global ecosystems, the truncation of seral sequences would eliminate climax stages. Since these stages are the most species-rich, their cut-off would mean a drastic drop of diversity – the mass extinctions of the fossil record. Evidence for the climax cut-off model (Krassilov, 1992a, 1995a) is furnished by vegetation developments over the KTB, with the mixed conifer-broadleaved Metasequoia–

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Fig. 145. Climax cut-off model: (1) a seral sequence develops over a stable landscape, (2) the late successional stages are eliminated under a long-term environmental impact (symbolized by volcanic eruption), (3) the pioneer phase spreads over the entire ecospace, (4) with stabilization, survivors of the pioneer phase give rise to a new seral system, (5) new forms, appearing in the process, spread over the sere.

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Trochodendroides forests replacing the coniferous Sequoia–Parataxodium forests. Taxonomic diversity of the latter increased considerably over the Maastrichtian owing to the invasion of new angiosperm species at the pioneer–early successional stages. Parataxodium, a highly polymorphous genus of taxodiaceous conifers, has since been lost as a result of its splitting into the progenitorial Taxodium-like and Metasequia-like morphotypes, while Seqouia has receded from its dominant status, with a single relict polyploid species surviving to this day. Of the replacing dominants, Metasequoia is distinguished by the decussate arrangement of leaves and cone scales, a paedomorphic character in conifers. Trochodendroides is an aspen-like leaf morphotype of a deciduous woody angiosperm with exceptionally large panicles of follicular fruits producing great numbers of small winged seeds, a pioneer dispersal habit. Both seem to have first appeared as deciduous components of riparian vegetation with sporadic fossil records over the Late Cretaceous. Their ubiquitous records above the KTB show broad ranges of morphological variation. At the PTB, the vojnovskyalean (cordaitalean) climax forests, persistent over the Permian, disappeared at the onset of trap volcanism while the species-rich pecopterid marshes with helophytic lycopsids and arthrophytes were replaced by the monodominant lycopsid marshes with Pleuromeia. Only the hydroseral conifer–peltasperm communities survived into the Early Triassic as a source of Mesozoic gymnosperm radiations (IV.3). Palaeobotanical examples may include the Dinantian vegetation of Scotland, inhibited from reaching the edaphic climax by seismic debris flows and frequent fires over the volcanic landscapes (Bateman & Scott, 1990), and similar developments in the Miocene (Cross & Taggart, 1983; Collinson & Scott, 1987a, 1987b). In the Cretaceous epeiric seas, a climax of ecosystem developments seems to have been reached with the spectacular morphological complexity of the larger carinate and bicarinate foraminifers (Caron & Homewood, 1983), the exoskeletons in the epifaunal serpulids and bryozoans, the excessive shell ornamentation in ammonites, the durophagous adaptations in echinoderms and vertebrates (Surlyk & Birklund, 1977), and with the hesperornithiform birds and mosasaurs appearing at the top of the trophic cascades. Reefal successions typically started with bryozoans and advanced toward the rudist climaxes (Kauffman, 1974). The KTB extinctions selectively affected the advanced forms of the climax stages, while the pioneer stages were enhanced. Extinction of rudists and a concomitant rise of bryozoans is a typical example. A short-term rise of small foraminifers at the expense of the morphologically advanced globotruncanids definitely falls into the same category. Marine records also show a selective mass extinction of poor dispersers in spite of their competitive superiority (Tilman et al., 1997). At an early stage of the KTB extinction, local species disappeared before cosmopolitan forms (Keller et al., 1994). This confirms that the loss of the γ component owing to cosmopolitanism of the early successional fine-grain strategists could have contributed to the total decrease in biological

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diversity. There are examples of two-stage extinctions, with the second stage involving the pioneer forms that appeared in post-crisis radiative bursts (Newman & Roberts, 1995; Petuch, 1995). Relaxation of environmental stress rendered such disaster species maladaptive.

IX.7. Post-crisis disparity Over the critical boundaries, such as the PTB and KTB, a decrease in biological diversity with a decline of climax species is scarcely compensated for by originations in the surviving pioneer groups that are typically far inferior in their species richness. Yet some progenitors of post-crisis biotas had appeared already at the pre-crisis phase. Examples include the archosaurian reptiles first appearing in the latest Tatarian, the condylarths of the terminal Cretaceous mammalian assemblages, and, among plants, the incipient betuloid, fagoid and ulmoid forms first recorded in the Maastrichtian. Their entries might have been owing to more open biotic communities, the pioneer stages of which expanded in anticipation of the approaching crisis. The taxonomic enrichment at the beginning of the crisis developments complies with an ecological observation that preclimax communities are the richest (Beklemishev, 1951). The same factor could have been responsible for the pre-crisis monstrosities observed in a number of lineages (in particular, ammonoids) on the brink of extinction. Although often ascribed to a mysterious senility of phylogenetic lines, monstrosity, as an extreme form of phenotypic plasticity transgressing the normal range of morphological variation, is evidence of a relaxed normalizing selection in disturbed ecosystems. Disaster populations are buffered from environmental hazards by their redundant densities (VIII.3.7) implying high reproduction rates. These are typically achieved through developmental acceleration impelling an abbreviation/condensation of individual development (VII.6). Hence paedomorphic transformations more frequently occur in crisis situations then over the coherent phases of ecosystem evolution. Paedomorphic features still occur in such post-crisis dominants as Metasequoia with decussate arrangement of leaves and cone scales. A spread of “simple morphologies”, such as smooth asaccate Cycadopites-type pollen grains at the PTB (Krassilov et al., 1999a) or small tricolpates in the Early Palaeocene, are paedomorphic events of gametophyte evolution (Krassilov, 2000a; VIII.7). Dwarfing is a related phenomenon commonly ascribed to a direct stunting effect of environmental stress (Antunes & Sigogneau-Sussel, 1996), but more consistently explained by a stress-driven acceleration of reproduction rates that is usually associated with a decrease in body size. Characteristic of crisis developments, and likewise related to a relaxed competitive environment after the collapse of long-standing climaxes, are the so-called Lazarus species reappearing after a long period of no records, such as Lazarospermum, an essentially Late Devonian–Early Carboniferous preovule morphotype with long integumental

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lobes, making its second entry in the latest Permian (Krassilov, 1999b; Fig. 128). A Cretaceous example is Czekanowskia, a widespread pioneer of Mesozoic peatlands, exceedingly rare over the Late Cretaceous in general, but prominent in the Maastrichtian dinosaur beds of Amur Basin (Bugdaeva, 2001). Encephalartites, a semiextinct Mesozoic cycadophyte, is likewise common in the contemporaneous deposits of Koryak Highlands (Krassilov et al., 1988) while the relatively poor KTB assemblage of Augustovka, western Sakhalin, contains two such cycadophytic genera, Pterophyllum and Cycadites, giving it an archaic mid-Cretaceous aspect (Krassilov 1979; Fig. 146). An association of macromutational newcomers with Lazarus forms is a distinctive feature of crisis communities that combine a low diversity with a high morphological disparity and morphological polymorphism of dominant species (IX.9.3).

IX.8. Recovery In plant communities, recoveries after local crises tend to restore the pre-existing structure. The models of recovery based on recent examples (Carr-Kitchell, 1980; Barry et al., 1991; Rosenzweig & McCord, 1991; Frederiksen, 1994) emphasize the significance of an initial diversity and of speciation rates that are fostered by vacant ecological niches as well as by immigration. The climax cut-off model (Krassilov, 1992a; IX.6) maintains that, in a stressful environment, ecological succession is halted at an early stage never reaching the potential climax. In effect, the recovery of climax species is hampered and, in the long sustained disclimax situations, might never occur. Major ecological crises truncate ecological successions over a wide range of terrestrial and marine habitats, thus driving the climax species to extinction. This explains the great extinctions that are recorded as nearly simultaneous disappearances of dominant forms over a wide range of habitats. The post-crisis recoveries eventually bring forth new dominants that are recruited from among pioneer species of pre-existing, or recently arising, taxa. Thus, dinosaurs recovered, with new dominant forms, after their near-extinction over the mid-Cretaceous crisis but were replaced by mammals over the KTB crisis (IX.5). The recovery or non-recovery of a formerly dominant group basically depends on its ecological status, with severity of disturbance and competition as additional factors. Recovery usually involves pioneer species or, in a broader sense, taxonomic groups having their representatives in the pioneer/early seral stages. It is problematic for the taxa represented by climax forms only. In biotic communities, the pioneer and seral species may belong to the same higher taxon (evolutionary grade) as the climax species. For instance, in broadleaved forests, the climax (lime, oak) and seral (poplar, maple) dominants are of the same evolutionary grade. The recovery of the lime–oak climax is impeded in heavily polluted environments but the broadleaved forest may recover through its persistent seral stages.

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a b

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Fig. 146. Transitional plant assemblages of the tuffaceous transboundary Cretaceous/Tertiary sequence at Augustovka, western Sakhalin (Krassilov, 1979): (a) Pterophyllum, a Lazarus bennettite, reappearing after a long interval of no records in the Late Cretaceous, and (b) Glyptostrobus (a twig with male cones), a Palaeogene newcomer, from the boundary tuff assemblage; (c) fern bed, a tuffaceous paleosol above the boundary, (d) Nilssonia bed few metres upsection, and (e) Corylites bed with fragmentary Metasequoia, representing edaphic variants of an ashfall regrowth.

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In contrast, a coniferous old-growth forest with seral broadleaved forms is “graded” in the sense that its consecutive developmental stages belong to different evolutionary grades of gymnosperms and angiosperms respectively. In such situations, the recovery of a climax grade is improbable, the more so as the pioneer forms represent an evolutionary higher grade. In other words, a graded community with the pioneer – seral forms evolutionary advanced relative to the dominant forms, may persist for a long time, as the Parataxodium forests with early angiosperms survived over the Cretaceous. Yet, with a spread of persistent disclimaxes, a point of no return can be reached for the climax dominants, unless they survive as local edaphic climaxes of a kind, as the bold cypress at present. Major biotic turnovers seem consistent with this pattern. Palaeontological examples include the Mesozoic cycadophyte shrublands (IV.4), with pioneer proangiospermous undershrubs, evolving into the xeromorphic angiosperm scrublands following the cut-off of their dominant cycadophyte species over the mid-Cretaceous crisis. The Late Cretaceous taxodiaceous conifer forests were likewise graded, with the trochdendrioid – platanoid angiosperms as a seral stage and undergrowth. The latter gave rise to the broadleaved riparian forests shedding their conifer component over the KTB crisis. Later on, the survival of residual taxodiaceous forests was due to the persistent hydroseral stages of their Cretaceous forerunners. Likewise, the dinosaur recoveries after the Jurassic/Cretaceous and mid-Cretaceous near-extinctions might have been owing to successional species, such as the smaller iguanodontids or protoceratopsians. Prior to the KTB non-recovery, the vertebrate communities apparently acquired a graded structure, with dinosaurs as the climax dominants and mammals as the pioneer to early successional forms (some perhaps also as the climax co-dominants that perished with dinosaurs). The non-recovery cases are thus related to graded structures, in which the phylogenetically advanced taxa enter ecological succession as pioneer forms then rising to dominance with a new climax. The opposite situation arises when archaic taxa, surviving as pioneer forms of biotic communities dominated by phylogenetically advanced taxa, are brought forth after a cut-off of the latter. Among the most spectacular examples are tundral lichens and Selaginella, the recent triumphs of archaic forms.

IX.9. Example of plant community evolution at the KTB The following analysis is based on monographic studies of the Cretaceous and Palaeogene taphofloras of Amur River Basin, Primorye, Mongolia, Sakhalin, Kuril Islands, with additional data from West Siberia, Kazakhstan and Middle East (Krassilov, 1967a, 1976a; 1979, 1982, 1984, 1988; 1989c; Krassilov & Makulbekov, 1995; Krassilov & Dobruskina, 1995, 1998; Krassilov & Bacchia, 2000; Krassilov & Golovneva, 1999, 2000, 2001; Krassilov & Shorokhova, 1989; Krassilov et al., 1983, 1988; VII.5-6). Taxonomic diversity of Late Cretaceous and Palaeocene assemblages is often exaggerated by neglecting

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leaf polymorphism. Widely disparate species numbers are given for one and the same taphoflora by different researchers (compare Kryshtofovich & Baikovskaya, 1966 with Krassilov, 1976a for the Tsagajan taphoflora). In effect, the boundary is much more distinct at syntaxonomic level.

IX.9.1. Late Cretaceous assemblages The following Late Cretaceous plant communities are inferred on the basis of recurrent fossil plant assemblages and the relevant taphonomic information (III.2): (1) Lauro-Sequoietum, a mixed redwood–laurophyllous lowland (dryland to wetland) assemblage (Fig. 147) dominated by Sequoia (Geinitzia) in association with Cupressinocladus (Chamaecyparis), Protophyllocladus, Araucarites as co-dominant conifers, and with Laurophyllum, Araliaephyllum, Magnoliaephyllum, Liriophyllum and other laurophylls, occurring as both autochthonous and allochthonous remains over a wide range of terrestrial and paralic sedimenatry facies. A xeromorphic variant, the Debeyo–Sequoietum, with narrow thickly cutinized Debeya, Sapindopsis and Proteophyllum morphotypes, as well as a mixed broadleaved variant, the Platano–Sequoietum, with prominent platanophyll component, might have come from topographically/edaphically marginal habitats. (2) Platano–Parataxodietum, a redwood–platanophyllous assemblage dominated by Parataxodium (“Cephalotaxopsis”), a deciduous taxodiaceous conifer morphologically intermediate between Taxodium and Metasequoia, and by broadleaved angiosperms prevailingly of platanoid leaf morphologies. The betuloid morphotypes enter in the Maastrichtian. A typical taphonomic occurrence is a leaf-mat of deciduous shoots and leaves. In northern Asia, this community had gradually replaced the Lauro–Sequoietum north of 50°N, with many platanophyllous species transgressing the phytogeographic boundary. In the terminal Maastrichtian, the Platano–Parataxodion spread south over the ecotonal zones of Sakhalin and Primorye. (3) Cyatho–Platanietum, a fern-platanophyll assemblage represented by abundant platanoid leaves forming leaf-mats in the levee facies in association with cyatheoid (gleichenioid) fronds and with scattered shoots of Sequoia, Parataxodium and other taxodiaceous conifers. With an increase in the latter, this community merged with the Platano–Sequoietum (Parataxodietum). (4) Filici–Nilssonietum, a fern or fern-cycadophyte marsh assemblage including the Osmundo–Anemietum dicksoniae, Osmundo–Cyathetum and the Nilssonio–Cladophlebidetum (Osmundetum) arcicae, merging with (3) but with fewer platanophylls and tending to occur in the vicinity of coal seams as well as in the tidal flat facies. In the fern marshes the formerly dominant Ruffordia goeppertii, Osmunda diamensis and Coniopteris spp. were substituted by Anemia (“Asplenium”) dicsoniana, Osmunda (Cladophlebis) frigida and Cyathea sachalinensis in various combinations with Nilssonia serotina and N. gibbsii.

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b

a e

c

d

f g

h

i

Fig. 147. A mixed conifer–broadleaved assemblage with laurophyll elements from the Late Cretaceous of western Sakhalin: (a) Sequoia, (b) Cupressinocladus, (c, g) Trochodendroides, (d-f) Araliaephyllum (a Sassafras-like leaf morphotype), (h) Debeya, (i) Liriophyllum (Krassilov, 1979).

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(5) Quereuxion, a union of aquatic communities, with submerged isoetalean lycopsids, water ferns and nymphaealean angiosperms. Aquatic monocots entered with a widespread eutrophication of lacustrine reservoirs over the Maastrichtian (Krassilov & Makulbekov, 1995).

IX.9.2. Early Palaeocene assemblages. In the early Palaeocene of eastern Asia, fossil plant localities occur in the sedimentary basins of intracratonic depressions, intermontane depressions of the marginal volcanic belt, their foredeeps and over the island arc belt (VII.6; stratigraphic correlation in Fig. 148). They represent distinct types of terrestrial and aquatic assemblages. In the intracratonic Amur and Zea-Bureya basins, a relatively complete transboundary section encompasses the Tsagajan Formation of three conglomerate/sand/clay cyclothems, with a rich dinosaur locality in the middle coarse-grained member. A diverse Maastrichtian

Fig. 148. A correlation of transboundary Cretaceous/Palaeocene sequences over the cratonic – marginal zones of Russian Far East

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spore-pollen assemblage and a macrofossil assemblage with small-leaved Trochodendroides, Liriophyllum and occasional Czekanowskia came from the uppermost dinosaur beds (Markevich, 1994; Bugdaeva, 2001). The Upper Tsagajan member consists of thick conglomerates and gravels of a major erosion cycle thinning upward to sandstones/siltstones of a meandering river system. There are 23 fossil plant beds of various riparian facies. In the coarse-grained member with lenticular clays, the poorly sorted fan deposits repeatedly prograded over the flood-plain. The oxbow clays imbedded in the debris-flow gravels contain abundant (hypo)autochthonous remains of few aquatic and wetland species, with both their foliar and reproductive organs preserved on the same bedding plane, as well as allochthonous remains of occasional leaves and samaras. The coarsely laminated crevasse-splay and the thin-bedded levee siltstones of avulsion cycles contain leaf mats of a single or few broadleaved species. The following assemblages are recognized: (1) Trochodendroidion, represented in all facies types by leaf–fruit assemblages, in which Trochodendroides arctica is a single or a single dominant species reconstructed as a small to medium-sized deciduous tree, heteroblastic, with aspen-like polymorphous leaves and with large panicles of follicular fruits producing a great number of small winged seeds. The Trochodendridion includes two distinct variants: (1a) Trochodendroidetum arcticae, represented by the single-species leaf mats on bedding planes of channel sandstones, with Trochodendrocarpus panicles and Trochodendrospermum seeds belonging to the Trochodendroides plant. (1b) Platano-Trochodendroidetum arcticae, represented by the two- to severalspecies leaf mats in the levee and oxbow facies (Fig. 149), with Trochodendroides and platanoids in nearly equal proportions. The relatively diverse assemblages of oxbow clays also contain Taxodium while the biserrate morphotypes Viburniphyllum and Tiliaephyllum occur in occasional leaf-mat patches. A single locality contains Protophyllum and Gleichenites, a residual Cretaceous fern–platanophyll community (above) conceivably survived in few patches of the riparian mosaic. (2) Tiliaephylletum tsagajanicae, represented by the single-species leaf mats of levee siltstones, with overlapping Tiliaephyllum, a cordate biserrate broadleaved morphotype (Fig. 36) representing a patch of shrubby vegetation, perhaps seral, reflecting gaps in the forest canopy. (3) Mixed conifer–broadleaved upland vegetation represented in the oxbow and levee facies by occasional samaras, cone-scales, fragmentary leaves and dispersed leaf cuticles that are preservationally different from the bulk of plant fossils. The coarser levee facies contain abundant plant debris but determinable remains are infrequent. Notable among the latter are occasional leaf fragments and dispersed leaf cuticles of Ginkgo as well as scale-leaves of Araucarites pojarkovae. Since fossiliferous lenses are imbedded in the debris-flow piedmont deposits, the allochthonous plant remains may conceivably represent an upslope vegetation of a single or several altitudinal belts. The conifers are fairly diverse, comprising both “northern” (Metasequoia, Sequoia, Cupressinocladus, Pseudolarix, Pinus) and “southern” (Araucarites, Podocarpus) genera. Alnites,

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Fig. 149. The Early Palaeocene post-crisis plant assemblages of Tsagajan, Amur Province (Krassilov, 1976a): (a) leaf mat of Trochodendroides with a paniculate infructescence of the same plant, (b) a slab association of Taxodium and Nyssa, the dominants of a wetland community first appearing at this level, widespread later on; (c, d) serrate leaves of Tiliaephyllum, supposedly representing a pioneer (post-fire) shrub growth.

a

c

d

Cyclocarya, Papilionaceophyllum and Celtis among the angiosperms testify to a relatively high rate modernization in the uplands. (4) Nysso–Taxodietum with Myrica and Nordenskioldia, dominating the clayey oxbow facies. This is the earliest occurrence of a bold cypress – black gum – wax myrtle community surviving over the Tertiary to the present. Nordenskioldia, supposedly a climber with long drooping racemes, was shed from this type arboreal wetlands in the Miocene (Manchester et al., 1991). (5) Gramino–Carexetum, a primeval herbaceous wetland replacing the Mesozoic fern marshes, with abundant, though taxonomically poorly understood “Arundo” and “Phragmites” morphotypes as well as the modern-looking Carex-type achenes. This appears to be an early representative of hydroseral grasslands persistent over the Palaeogene and supposedly progenitorial both to mesic and xeric grassland types that appeared later in the Tertiary (IV.2.5). (6) Limno–Nupharetum, a diverse aquatic community of rooted and floating angiosperms, with Nuphar, Nymphaeites, Nelumbo, Limnobiophyllum, Potamogeton,

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Hydrocharis, and with a few Quereuxia angulata indicating an affinity with the Late Cretaceous Quereuxion. The simultaneous appearance of nymphaeoids and herbaceous monocots, with extant genera prevailing over the extinct ones (Limnobiophyllum, the Araceae) suggests a high rate modernization of freshwater vegetation in step or even ahead of terrestrial communities. By analogy with the present-day aquatic communities, multiple introductions of floating herbs indicate eutrophication of lacustrine ecosystems. In the coastal volcanic ranges of Sikhote Alin and Chokotsko-Okhotskian belts, fossil plant localities are confined to intermontane depressions cutting across the fold belt lineaments. The Cretaceous sequence of felsitic/andesitic volcanics is terminated by the exceptionally thick ignimbrites assigned to the latest Maastrichtian Wodehouseia fimbriata – Ulmoideipites krempii palynozone (Markevich 1994). Slumped liparitic tuffs of this age contain a diverse, though fragmentary, conifer assemblage with Sequoia, Glyptostrobus, Cupressinocladus, Androvettia, Pseudotsuga and Pseudolarix as well as angiosperm debris, prevailingly of betuloid leaf fragments and dispersed bracts. The conifer assemblage represents a precursory montane taiga while the allochthonous birch remains came from an upslope belt corresponding to the present-day stone birch woodland. The montane communities appear fairly advanced in comparison with the contemporaneous Tsagajan-type lowland vegetation (above). The palynologically defined KTB coincides with a switch from felsitic to basaltic volcanism indicating a transition from compressional to extensional geodynamic regime. The lowermost Palaeocene basalts alternate with the rift-lake deposits of dark brown to black tuffaceous shales intervened by the marginal debris-flow accumulations and by thin coal beds. Such sequences are confined to a series of pull-apart basins over the transcurrent shear zones. The lacustrine shales contain rich plant localities with Metasequoia, pines, and the diverse ulmoid–betuloid angiosperm morphotypes (Krassilov 1989c). The dominant species, Betula pomoera, is abundantly represented by leaves and bracts, the latter occasionally preserved in dichasial clusters (Fig. 150). Obviously, their source community at that time grew closer to the site of deposition than in the Late Maastrichtian before, indicating a downslope shift of the birch belt (Fig. 151). In the foredeep over the Tatar Strait – western Sakhalin, plant remains occur in the coal-bearing deltaic cyclothems and estuarine black shales intercalating with shallow marine deposits. The Cretaceous catenic system ascended from the delta fringe fern– nilssonia marshes over the delta-plain tree fern (Cyathea sachalinensis) – platanoid zone to the lauro–coniferous rainforests (Krassilov 1979). After the Maastrichtian transgression, the advancing seral vegetation, with the typical Cretaceous fern–nilssonia marshes, was invaded by the riparian Platano–Prataxodietum with the first appearing Trochodendroides arctica, a Palaeocene newcomer. In the relatively complete KTB sequence at Augustovka, western Sakhalin, the paralic coal-bearing sequence is terminated by a coarse-grained tuff with a peculiar association of Palaeocene species Glyptostrobus nordensldoldii and Trochodendroides arctica accompanied not only by Nilssonia gibbsii, but also by two Lazarus cycadophytes,

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d

c

Fig. 150. An increase in betuloid remains (leaves, partial inflorescence, a, b) relative to conifer remains (c, d) in the Early Palaeocene of Zerkalnaya depression, Sikhote Alin suggests a downslope migration of the birch woodland belt relative to its altitudinal position in the Maastrichtian (Krassilov, 1999c).

Pterophyllum and Cycadites. The plant remains are coarsely preserved, oblique to the bedding plane, apparently the result of an instantaneous burial by ash flow. The overlying tuffaceous sequence contains nine sequential fossil plant assemblages. In the basal part, a root-bed with horsetail remains is followed by a 50 cm thick fern-bed, a tuffaceous shale with Woodwardia and Cladophlebis columbiana. Several leaf mats upsection are dominated by Metasequoia occidentalis and Corylites protoinsignis (leaves and husks), with Ginkgo, Trochodendroides arctica, Alnites protoschmalhausenii, Liriophyllum sachalinense and Macclintockia kanei as the common species.

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Fig. 151. An interpretation of an upland–lowland migration of altitudinal vegetation belts over the Cretaceous–Tertiary boundary on slopes of Sikhote Alin Ranges: a climate-driven expansion of birch woodland (B) into the conifer belt downslope (G – Glyptostrobus, M – Metasequoia, P – Pinus).

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Thus, over the transboundary sequence, the Cretaceous Platano–Parataxodion was replaced by the newly appearing Glypto(Metasequio)–Trochodendroidetum arcticae and Alno–Coryletum protoinsigni, with a few Cretaceous survivors, Nilssonia, Liriophyllum, and the “Lazarus” cycadophytes still prominent in occasional patches of piedmont vegetation giving the local assemblages an archaic Cretaceous aspect. Rare Potamogetonites attests to an entry of monocots into the aquatic vegetation. An increase in ferns and Corylites conceivably reflects a spread of pioneer vegetation over ash fallouts and burned soil. Fern propagation after ash falls is presently observed in Zambia (Kornaú, 1978) and elsewhere, while hazel shrubs are common in post-fire successions. A comparable transitional KTB assemblage came from clastic interbeds of the islandarc/fore-arc basaltic sequences forming the Lesser Kuril Islands (Krassilov et al., 1988). Resting on the back-arc Maastrichtian turbidites with inoceramid remains, these deposits also contain benthic foraminifers and marine bivalves indicating the uppermost Maastrichtian – lowermost Palaeocene age. The terrestrial plant remains are evidence of a volcanic ridge emerging above sea level after a major thrusting event at about the KTB. As is known from the present-day examples, turbidite deposition on slopes of active island arcs typically post-dates the emergence by several thousand years. The relatively well-preserved conifer shoots represent a residual Late Cretaceous association of Sequoia reichenbachii and Cupressinocladus cretacea. The phyllogenetically younger representatives of the Pinaceae (Picea, Pseudolarix), Cupressaceae (Androvettia) and Taxaceae (Amentotaxus) occur as small debris of fragmentary shoots, needle-leaves and dispersed samaras, apparently a slope downwash. The silicified fossil wood from the overlying lavabreccias belongs to several conifer species of Palaeogene affinities. Among angiosperms, Corylites protoinsignis is the only frequent leaf type supposedly representing a post-fire shrub growth, as in the Augustovkian assemblages of western Sakhalin (above). Among the rare angiosperm morphotypes, Debeya cf. pachyderma is a Cretaceous survivor, while Trochodendroides arctica, Viburniphyllum asperum and Menispermites katiae (similar to M. favosus from the basal Palaeocene of Augustovka) are new arrivals.

VIII.9.3. Vegetation change The regional sequences of fossil plant assemblages over a series of tectonic settings thus attest to a major vegetation change at about the palynologically and microfaunistically defined KTB. The Cretaceous evergreen/semideciduous redwood forests with laurophyllous and platanophyllous broadleaves declined over the KTB and were replaced by the deciduous broadleaved and mixed communities with Metasequoia, Trochodendroides and biserrate broadleaves. The deciduousness of the post-KTB assemblages is sometimes perceived as anomalous related to selective elimination of evergreen species after a catastrophic extrater-

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restrial impact (Wing, 1998). Alternatively, it can be explained by a spread of riparian broadleaves with a greater proportion of deciduous elements than in the dryland vegetation (as in the present-day deciduous–evergreen catenic systems of Florida, Colchis, South China, etc.). Wetland communities were likewise involved in the general restructuring, with the Mesozoic type fern–nilssonia marshes giving way to the first appearing grass–sedge marshes and the Nysso–Taxodietum swamp forests. The diverse aquatic plant communities, with rooted and floating monocots, might have developed on the basis of the monodominant Late Cretaceous Quereuxion. The vegetation change had already commenced in the Maastrichtian, with the invasion of the numerically subordinate Palaeocene newcomers, such as the small-leaved Trochodendroides, biserrate Corylites, Potamogeton-like aquatic monocots, etc., supposedly entering the pioneer to early successional stages of a disturbed riparian vegetation. These invasions coincide with an increase in tectonic/volcanic activity and with concomitant climate change inflicting a southward spread of temperate vegetation (Krassilov, 1979; Golovneva, 1994; Herman, 1999). A culmination of the temperization trend over the KTB is indicated by the prevalence of leaf-mat taphonomy and serrate leaf morphology, as well as by the downslope shifts of altitudinal vegetation belts. Thick conglomeratic sequences, as in the Tsagajan Formation (above), reflect a major erosion cycle in cratonic areas driven by a drop of sea level globally recorded at the KTB. In the bordering fold belts, a geodynamic switch from compressional to extensional regime is attested by the petrochemistry (a transition from felsitic to basaltic volcanism) and sedimentology of the pull-apart transcurrent basins. On fringes of the marginal shear belt, these events correlate with the emergence of volcanic island arcs, e.g., the Lesser Kuril Islands. The concerted tectonomagmatic, eustatic and climatic events generated a stressful environment for the plant communities surmounting the KTB. The Early Palaeocene lowland plant formations (Trochodendridion) apparently evolved from a pioneer stage of Cretaceous riparian forests (Fig. 152). Its dominant species Trochodendroides arctica might have arisen from a small-leaved Cretaceous T. sachalinensis or an allied species (or through a genetic introgression of such species as the polymorphism may suggest), constituting a minor deciduous element of the lauro–taxodiaceous forests. The leaf mats of Corylites and Tiliaephyllum represent an early successional shrubby vegetation of disturbed habitats (a spread of Corylletum over the KTB finds its analogy in the early Holocene abundances of Corylus avelana, a pioneer species declining with the development of dense broadleaved canopies: Tinner & Lotter, 2001). The turnover of plant communities thus corresponds to the climax cut-off model of vegetation change (Krassilov, 1992a, 1995a and above), yet affecting all seral stage. The average species richness is about 65 species for the Late Cretaceous basinal taphofloras against 50 species for the Early Palaeocene taphofloras of the same palaeofloristic realm, a nearly 20% drop off, primarily at the expense of fern–cycadophyte marshland species, archaic conifers (Parataxodium, Protophyllocladus), lobed lauro-

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Valentin A. Krassilov. Terrestrial Palaeoecology VEGETATIONAL CHANGE AT THE KTB PALEOCENE

Climax: Lauro-Sequietum

CERTACEOUS

Riparian/Seral: Trocho-Platanetum

Fig. 152. A schematic interpretation of plant community evolution over the Cretaceous–Tertiary boundary through the climax cut-off.

phylls (Araliaephyllum, Liriophyllum), glandular protophylls of Protophyllum ignatianum type, etc. A comparable drop is recorded over the KTB succession in North America (Wing, 1998; Gemmill & Johnson, 1997). Extinctions and semi-extinctions occur within a transitional interval that encompasses the basal Palaeocene still retaining relict species of Liriophyllum and Nilssonia. The Early Palaeocene plant assemblages are typically mono- to oligodominant, often repre-

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sented by the single-species leaf mats of a considerable extent. A decrease in diversity was apparently accompanied by an increase in population densities. At the same time, the survival of Cretaceous relicts and “Lazarus” species as well as evolutionary innovations (e.g., in hamamelid angiosperms: Maslova, 2001) both contributed to a high morphological disparity of the Palaeocene communities. Occasional patches of the riparian mosaics still bore a residual Cretaceous aspect. The β diversity remained relatively high owing to persistent seral communities of the Coryletum or Teileiaephylletum types. Taxonomic heterogeneity of the Early Palaeocene regional taphofloras of Amur Basin, Sikhote Alin and Sakhalin – Lesser Kuril Islands systematically exceeded that of the respective Cretaceous taphofloras. The subsequent Arcto-Tertiary flora might have evolved from several Early Palaeocene florogenic centres (Krassilov, 1976a). Allochthonous plant material suggests the relatively high innovation rates over the highlands of the end-Cretaceous orogeny, with an incipient altititudinal differentiation of the conifer forest – birch woodland belts. The climate-driven downslope migrations in the Early Palaeocene might have introduced the advanced elements to the lowland vegetation. Altogether, the regional vegetation developments over the KTB reflect a turnover that would have been more likely inflicted by long-term environmental instability rather than by an instantaneous impact. An increase in taxonomic diversity before the KTB, and a sharp drop off after it comply with actualistic ecological evidence of the highest pre-climax diversity as well as with the model of climax cut-off through a truncation of seral sequences in unstable environments. The low diversity – high disparity plant communities prevailing at the beginning of the Early Palaeocene testify to a change in the population adaptive strategies driven toward the fine-grain mode. The deciduousness of the mid-latitudinal Palaeocene biomes (perhaps enhanced by preferential survival of seral stages) is the result of a climatic temperization that started in the Maastrichtian already. The subsequent build-up of species richness over the Late Palaeocene reflects a gradual transition from the non-coherent post-crisis developments to the coherent evolution.

VIII.9.3. Macropolymorphism and morphological innovation The taxonomic diversity of Early Palaeocene floras has been greatly exaggerated by assigning the remains of conspecific polymorphous leaves to different species or even genera. In the Tsagajan flora of Amur Province, two species, Trochodendroides arctica and Platanus raynoldsii, account for about 40 binominals published by Kryshtofovich & Baikovskaya (1966). For comparison, in the contemporaneous flora of Wyoming, four species encompass about 95% of all leaf morphotypes (Gemmill & Johnson, 1997). Such leaf species as Cercidiphyllum crenatum, Tetracentron amurensis, Trochodendroides speciosa, T. smilacifolia, T. richardsonii, T. genetrix, etc. (see synonymy in Krassilov, 1976a) are the more or less distinct leaf variants of T. arctica. They are found in each of the adequately sampled Early Palaeocene leaf populations (Figs. 153, 154).

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Fig. 153. Leaf variants of Trochodendroides arctica, a polymorphous species from the Early Palaeocene of Tsagajan, Amur Province.

Such repeated co-occurrences suggest a stable system of leaf polymorphism including at least ten distinct leaf allelomorphs. For comparisons at the population level, the epithets of synonymized binominals like “speciosa,” “smilacifolia,” “richardsonii,” “genetrix,” etc. can be used as non-taxonomic descriptive designations (as in the population genetics of Drosophila). For comparison, the precursory Cretaceous species, Trochodendroides sachalinensis, is represented by no more than three stable leaf allelomorphs. Another dominant species, “Platanus” raynoldsii, is likewise polymorphous, with different leaf (leaflet) morphotypes previously described as Alnites, Betulites, Quercus, Pterospermites, Ficus, Grewiopsis, etc., allegedly belonging to different angiosperm families (Kryshtofovich & Baikovskaya, 1966; synonymized in Krassilov, 1976a). The cuticular characters unequivocally indicate polymorphous leaf population of a single species.

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Fig. 154. Leaf variants of Trochodendroides arctica, a polymorphous species from the Early Palaeocene of Tsagajan, Amur Province (continued).

The associated reproductive structures combine characters of two or more modern families. Thus, Trochodendrocarpus, an infructescence morphotype of Trochodendroides plant, comprises the large panicles of folliculate many-seeded fruits, so far the largest among the Cretaceous and Palaeocene angiosperms, attesting to an enormous seasonal reproductive effort (Fig. 155). Although usually compared with Cercidiphyllum, it differs in the general structure and in pairing of the follicles anticipating a pairwise syncarpy in the Hamamelidaceae. “Platanus” raynoldsii produced small staminate heads of conventional platanaceous morphology (Krassilov, 1976a). However, the pistillate heads, scarcely larger than the staminate ones, consisted of persistent carpellodia surrounding a single folliculate carpel of a magnolioid, rather than platanoid, morphology (Maslova & Krassilov, 2002). Nor-

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Fig. 155. Examples of morphological experimentation in post-crisis biotic communities: unusual reproductive structures in the Early Palaeocene angiosperms, combining characters of two or more extant angiosperm families: (a) Trochodendrocarpus, a large loose panicle of paired many-seeded fruits, and (b) Nordenskioldia, a long loose catkin with flower buds and persistent calyptras. Tsagajan locality, Amur Province, Russian Far East (Krassilov, 1979, 1997a).

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denskioldia of obscure taxonomic affinities, combining menispermoid, illicioid and trochodendroid characters (Krassilov, 1997a), is a further example of morphological experimentation. With normalizing selection suspended after a collapse of long-standing climaxes, the surviving species had revealed full ranges of their potential variation transgressing the conventional morphological distinctions between angiosperm genera and families even. In both platanoids and trochodendroids, an increase in morphological polymorphism might have been due to genetic introgression over speciation cycles discussed in (VII.5). In particular, certain regularly co-occurring morphotypes of polymorphous Trochodendroides arctica correspond to local Maastrichtian species of small rounded (T. sachalinensis) and elongate-elliptical (T. elliptica from the lower Tsagajan dinosaur beds) leaves. These latter might represent an ecotypic differentiation phase of coarse-grain populations followed by an introgression phase with a switch to the fine-grain strategy induced by the stressful environments. As predicted by Vavilov’s law of homological variation series (Vavilov, 1922, 1931), the polymorphism incorporated in the genome show up with a release of latent morphogenic potentials in the undersaturated post-crisis communities.

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Chapter 10. Conclusions

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X. CONCLUSIONS The global change of our lifetime experience is partly man-made, partly natural. How much of it is natural can be inferred from pre-historic records. If, as the records seem to suggest, the changes are cyclic, it will be useful to know at what stage of a cycle we are at the moment. One guess is as good as another until we arrive at a general if but crude model of complex interactions that drive evolution of biosphere as a geobiological system. The model would encompass the multilevel developments from the geoid to the genome, but as interrelated aspects of a singular story rather than separate stories. The book presents an attempt at such a model from a viewpoint of a palaeoecologist. The role of biosphere in the global change is a stabilizing one, and it will be potentially even more so with an addition of the rational human factor. However, the biosphere, for all its significance, is a thin film stretched and occasionally ruptured by the deep breath of the earth. The breath is rhythmic, paced by the earth’s rotation and revolution, with autocyclic and induced fluctuations creating a hierarchical cyclicity. Periodic gravitational disturbance of the time-scale of the Galactic Year are such as should be inflicted by the huge concentrations of interstellar matter that are encountered by the solar system on its eccentric orbit round the centre of the Galaxy (V.2). The shorter-term rotational perturbations are relatively well known and are the most obvious driving force of the global change. The atmospheric and hydrospheric circulations are driven by the rotational forcing that acts upon the solid earth as well. That the latter effect is far from being negligible is immediately evident form a glimpse at the transcurrent fault system cutting across the mid-ocean ridges and their parallel lineaments: their brocken-stick outlines show maximal displacement at the equator. A rotational origin of the global fault network is beyond any reasonable doubt as is the physical necessity of a discrepant rotational acceleration over the density heterogeneities, such as the continental and oceanic crust domains or the interior divisions of the earth’s core, mantle and lithosphere. The boundaries of the continents with the denser oceanic crust are the shear zones of a vigorous tectonic and igneous activity creating the global fold belts. Friction melting at the interior density boundaries generates magmatic sources as well as perturbations of the geomagnetic field. However the rotational acceleration inflicts not only the lateral but also the radial shear. The centrifugal component of rotational acceleration of plus or minus sign drives, respectively, a divergence or convergence of hypsometric levels for the isostatically compensated oceanic and continental crust resulting in an alternation of geocratic (high standing continents – deep oceans) and talassocratic (low continents – shallow oceans) situations. Insofar as a segment of lithosphere expands while it is raised, the geocratic tendencies are correlated with a continental rifting and trap magmatism, whereas the sea-floor magmatism increases during the talassocratic epochs.

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In this way a rotational forcing triggers off both the earth’s interior and surface processes (V-VII). The latter receive energy from the solar radiation, yet how much of the incoming energy resides and how it is distributed over the globe depends on the earth rotation and its derived geographic variables, such as the relative elevation of the highlatitude and low-latitude areas, land/sea distribution, sea gates, land bridges, mountain ranges, etc. Climate is a system of gradients the steepness of which depends on the roughness of the geography. In particular, the orogenies due to an expansion of the low-latitude rotation bulge affect the temperature/precipitation gradients by hindering heat transfer from the tropical zone. Not only the latitudinal but also the altitudinal gradients become steeper in effect, and a glaciation may start with the low-latitude mountain ice caps (VII.2.2). In fact, the first-order climatic cycles are chronologically correlated with the tectonomagmatic events in the global fold belts, the Tethys belt primarily. The epochs of equable global climate correspond to the Tethys seaway geography, the continental glaciations to the Tethys mountain range geography consequential to the major pulses of the Taconian, Hercynian and Alpine orogenies. There is admittedly a long way from these generalities to detailed regional interpretations, yet they at least provide a meaningful framework for such. A theoretical model is meaningful when it illuminates relationships rendering various observations coherent rather than presenting them as totally irrelevant to each other, a collection of happenings. In the framework of the rotational model, the geomagnetic, tectonic, eustatic and climatic developments make a coherent story that offers the long-sought explanations of cyclic fold belt activities, their correlation with the epeirogenic upheaval/subsidence of landmasses antithetic to those of the sea floor, and the related magmatic events. The geological records of past climates make sense in the context of continental and oceanic developments that affect heat transfer over the globe. The tectonic activities and sea-level fluctuations also indirectly affect climate through their multiple impacts on the carbon cycle. The machine of the carbon cycle is driven through the biosphere by the solar energy and is thus intimately related to climate. The atmospheric concentrations of carbon compounds render a feedback effect on climate through their contribution to the albedo (the greenhouse effect). They are regulated by the exchanges with the much larger oceanic and biotic concentrations. Since their balance is strongly affected by temperature, the temperature → CO2 regulation greatly prevails over the feedback CO2 → temperature regulation. An understanding of greenhouse effect as a feedback of the mainstream climatic regulation may help in evaluating its role as scarcely a leading one, although conceivably essential as an amplifier of climatic fluctuations. The carbon disposed with sedimentation is partly returned to the cycle by such processes as thermal outgassing, anaerobic metabolism of the Archaea or the use of fossil fuels by technological humans, yet the amounts emitted to the atmosphere are negligible in comparisons with those circulating in the system, and the effect is short-


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leaved, overshadowed by the larger fluctuations imposed by the changing oceanic capacities and the related capacities of terrestrial biota and soil, the major variables of the carbon cycle (VII). Climatic regulation thus involves a great number of interrelated variables engaged in the multiple loops of the positive and negative feedbacks. An attempt to reduce an explanation of global climate change to a single variable, such as the atmospheric CO2 concentration, is a short-sightedness bordering on light-heartedness. If the money spent on struggle against CO2 emissions were invested in flood prevention, as the present author has repeatedly suggested (Krassilov, 1992a and elsewhere), the catastrophe of the last (2002) summer might have been evaded. What makes us believe in a rapid melting of polar ice caps is a long-standing misunderstanding of the recent climatic situation as extraordinary, a residual glaciation returning to the non-glacial norm of the global climate. Yet, judging by a persistent temperizing effect of ice caps over the mid-latitudes, polar glaciation is the norm whereas a nonglacial climate is extraordinary. As inferred from the four-phase climate cyclicity (VII.9), we are at the transition form cold dry to warm dry phase of a millennial cycle. Transitional situations are unstable which is fairly evident from the ongoing climatic developments. In the animal world, narrow adaptations are advantageous in stable environments but scarcely so in the unstable ones, and the same with human technologies. In a chase for the benefits of high technologies it will be unwise to dispense with primitive technologies. In respect to the cyclic environmental change, a technological robustness and resilience must be given priority over technological sophistication. Dealing with such immensely complex systems as the global climate we have to distinguish between the effects that we can alter and those to which we have to adapt. On evidence of high frequency severe earthquakes in southern Asia, the transition from the Tethys mountain range geography to the Tethys seaway geography is on the way already. Since the latter is correlated with a shift of terrestrial biomass production from the tropical zone to the higher mid-latitudes (VIII.2.2), we have to expect a great wave of human migrations in the same directions. We need the attitudes and political structures that will facilitate such migrations rather than hindering them. Recent developments in natural sciences attest to our willingness to pay for a sublimation of our subconscious fairs, be it a movie thriller or a scientific theory of the kind. The killer bolid story is a striking example. For more than hundred years we firmly believed in the Darwinian theory of gradual evolution driven by the natural selection. The Cuvierian revolutions were scarcely remembered, let alone Maupertuis with his theory of discontinuities caused by extraterrestrial impacts. And this was shaken over night by an assertion that the history of life has been radically changed, with the Mesozoic era truncated and the new era commenced, by an occasional bolid. Our rush to protect all living things, the biological diversity, whatever the meaning, is certainly commendable, especially in the face of our extremely poor understanding of the threats (there is no protection from bolids by the way).

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Short-term disturbances may even enhance the diversity by introducing an additional heterogeneity of habitats (such as gaps in the forest canopies). A serious threat comes not so much from an occasional mass mortality (not to be confused with mass extinction) or overkill as from long-lasting disturbances that affect population strategies inflicting a switch from the sustainable demography of the prosperous to the redundant proliferation of the oppressed. These latter bring with them erratic fluctuations of population densities that destabilize biotic communities. A major loss of biological diversity inevitably follows. Presently, with the human technological impact on natural ecosystems rapidly approaching the red line, the theory of evolution must provide us a guidance, which the Darwinian theory failed to do. We cannot afford to forever meddle with theories that resort to chance and occasion. There is nothing occasional in the ecosystem evolution, which is directed toward a sustainable production of living matter and the efficiency of which, measured as a biomass to dead mass ratio, increases through times. This trend is facilitated by diversification reflecting the build-up of structural and functional (spatial, seral, trophic, mutualistic) complexity. The ecosystem evolution is briefly reiterated by an ecological succession of the pioneer – climax stages (in much the same way as morphological evolution is reiterated through individual development). In both, as in any other goal-orientated process, the structures are anticipated by their potential functions. The protective functions gain in importance with development requiring an adequate structural accomplishment. The human intelligence is such an accomplishment. In biological communities, as in human societies, an increase in diversity is owing to creation of new (ecological, social) niches rather then to competition for the established ones. Far from being a vehicle of evolutionary progress, competition decreases biological diversity by eliminations among the species whose niches happen to overlap. Through the evolution, as well as over an ecological succession, the significance of competition decreases with the niche overlaps and redundant densities. The biological progress (as well as the sociological one for that matter) is related to the niche-generating events (e.g. the appearance of land plants creating an additional aboveground and underground ecospace of the multistratal canopies and the rhizosphere, respectively). The technological humans are vigorous competitors intruding ecological niches of many natural species. They are potentially a powerful niche-generating agent also. Palaeoecological evidence attests to a holistic nature of biotic communities maintaining their climax structures over the geologically appreciable time periods. Yet, under long-term environmental stresses, ecological succession can be truncated never reaching the pre-existing climax. In effect, the climax species find themselves at a disadvantage, and some or all of them may perish. Since such species constitute a substantial part of biological diversity, the latter drops off with a cut-off of the climax stage. Stressful environments favour redundant population growth as a buffer against catastrophic mass mortality (the Niobe strategy). Escalation of biomass production rates not matched by an adequate increase of diversification rates disrupts the biomass/diversity correlations. The consequences are devastating, as in the case of hypertrophic ecosystems.


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When the mainstream evolutionary tendencies are reversed, we are justified in recognizing a biospheric crisis, such as over the Permian/Triassic or the Cretaceous/Tertiary boundaries. Both were preceded by an increase in geomagnetic instability and culminated with a rise/expansion of continental crust and the voluminous intracratonic magmatism. At both an upheaval of mantle material caused prominent geochemical anomalies (including the iridium spikes) correlated, through the effects on biological production, with an anomalous isotopic composition of organic carbon and the anomalously low biological diversity. Geological developments affect biota through their impact on biotic production rates mainly (e.g., through an input of iron from volcanic sources enhancing the mobility of phosphorus in freshwater ecosystems: IX.1). Eutrophication spreads with terrestrial runoff over estuaries to marine ecosystems (as recorded by the algal blooms over the Permian/Triassic boundary). The crisis thus starts deep in the earth’s interior and surfaces with the biotic turnovers. With normalizing selection suspended after the collapse of long-standing climaxes, the surviving “disaster” populations reveal the full ranges of their genetic potentials, with the extremes that are otherwise exterminated. Such extremes are qualified as macromutations. In particular, a demand on reproductive rates induces an acceleration of developmental rates resulting in a paedomorphic variation. The same conditions favour an openness of genetic systems to introduction of exogenous genetic material. Horizontal gene transfer may gain in significance in post-crisis communities (on evidence of morphological characters spreading at certain time-levels across the diverging phylogenetic lineages: VIII.1.4, Fig. 113). With stabilization, the macropolymorphous populations might have served as the foci of subsequent adaptive radiations. Thus evolutionary developments at the population and genomic levels are top-down regulated by ecosystem evolution. Speciation is closely related to the biospheric cycles through their impact on population strategies. A coarse-grain strategy highly sensitive to environmental heterogeneities (“grains”) prevails under stable conditions giving way to an indiscriminate fine-grain strategy with erratic disturbances (VIII.3.4). The down to grain adaptation is accompanied by a genetic divergence of local ecotypes. With a switch to the fine-grain strategy, the pattern of the coarse-grain diversity is swept out. Genetic mixing through introgression of the down-to-grain adapted genotypes generates a new system of genetic variation, a speciation event. The genome responds to environmental demands by a change of differential gene transcription rates that are memorized (assimilated) through the respective changes in these gene replication rates. Schematically, the more transcriptionally active genes fall in the fast replicating category, whereas an inactivated part of the genome is slow replicating and can be potentially lost through underreplication. The respective regulatory systems are adjusted to the altered gene activities through a deletion/amplification of their repetitive elements providing for an earlier switch on of the environmentally induced genes, inflicting a directional evolution of the genome.

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The general regularities of ecosystem evolution seem relevant to the human evolution as well. Darwin left us with a choice of either to be a “Darwinian specimen” ever struggling for existence (VII.3.7) or not to be such by suppressing our natural drives, which is no less detrimental. His theory did not accounted for evolution of evolution itself. The self-destructing strategies of pioneer species need not govern the mature stages of evolutionary development. The legacy of early humans, a pioneer species, is still steering us to ecological expansion, unrestrained population growth, overdominance, devastative use of trophic resources and the ensuing intense intraspecies competition. Population growth not matched by diversification is risky, as biological examples show. There are, however, multiple signs of a turning point toward sustainability as a goal of social and technological restructuring, non-redundant demographic strategy and a decrease in the personal niche overlaps opening the way to a build-up of individual diversity. These developments bring humans to the mainstream of biospheric evolution.

Subject index

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SUBJECT INDEX Albedo 187, 208, 209, 210, 212, 214, 219, 221, 226, 234, 238, 241, 243, 353, 388 Alga-like (thalloid) cutinized plants 96, 98, 105, 245 Allochthonous peat (coal) 7, 22, 24, 25 Allotillites 188, 239 Alpha diversity 188, 245, 293, 296, 341 Alpine-type ophiolites 144 Altitudinal temperute gradient 211, 212 anaboly 318, 319 Angarian realm 108, 172, 173 Annular shear zones 141 Aquilapollenites province 211 Arboreal habit 236, 280, 314, 362 Archaea (isotopic signature) 359, 388 Archaeopterids 84, 105, 245, 283 Archaeoptrid zone 105, 245, 375 Arcto-Mesozoic realm 60, 80, 82, 115, 117, 119, 246 Arcto-Permian realm 107, 291 Arcto-Tertiary 79, 82, 84, 90, 121, 196, 213, 246, 247, 273, 308, 381 Arogenesis 303, 308 Arogenic communities 298, 304 Arthropod herbivory 236, 274, 275, 276 Ashfalls 17, 261, 342, 343 Aspect (of vegetation) 59, 79, 80 Autochthonous peat (coal) 24 Autocycles 22, 23, 29, 209, 243 Autotrophic systems 3, 274 Avulsion 22, 28, 69, 259, 373 Back-arc basins 124, 138, 141, 142, 151, 154, 378 Back-swamp vegetation 24, 70, 99 Baisia 15, 31, 35, 75, 298, 301 Basinal biota 62 Basinal vegetation system 11, 61, 62 Benioff zones 140, 150, 157 Beta (intercommunal) diversiy 245, 293, 296 Biomass estimates 286-292 Biomass redistribution 285 Biomass to dead mass ratio 283, 390 Biospheric crises 232, 317, 320, 339, 358, 359 Biotic turnovers 194, 234, 316, 339, 350, 355, 361, 369, 391 Black shales 3, 8, 18, 19, 20, 21, 23, 28, 29, 31, 32, 35, 153, 182, 188, 189, 257, 258, 285, 360, 375 Bond cycle 216 Boreal transgressions 215, 294, 295 Boundary clay 35, 189, 234, 235, 341, 342, 343, 345, 346, 348, 351, 352, 355, 357, 359

Boundary layer (air) 39, 197 Brachyphyllous vegetation 44, 68, 73, 75, 76, 77, 83, 91, 92, 115, 118, 239 Brachyphyllous conifers (coniferoids) 44, 61, 167, 245 Brachyphylls 47, 87, 92 Browsers 80, 94, 277, 278, 354, 362 Canopy structures 197, 292 Carbon isotopes 193, 250, 352 Carbonate cycles 32, 35 Carbonate dissulution (compensation) 32, 181, 182, 353, 358 Carbonate lysoclines 35, 226, 351, 352, 357 Carbonate platforms 5, 352 Catena 55, 62, 65, 68, 72, 73, 74, 76, 77 Catenic systems 67, 73, 74, 75, 76, 379 Cathamerian realm 111, 291 Cathaysian realm 106, 111, 114, 245, 280 Cauliflory 44 Chemical weathering 29, 210, 220, 224, 226, 230, 236, 285, 302, 342, 351, 355 Chronogeographic parallelism 280, 282 Classes (in syntaxonomy) 60, 197, 250, 278 Climate proxies 187-205 Climate tolerances 194 Climax (equilibrium community) 63, 363, 391 Climax cut-off 363, 365, 367, 379, 381 Clorofluorocarbons (CFCs) 218, 233 CO2 phytogeographic effect 238 Coal measures 108, 111, 112, 188, 239, 261, 283 Coal-bearing deposits 20, 63, 70, 83, 192, 329 Coarse-grain strategy 15, 297, 310, 313, 391 Coherent/non-coherent evolution 268, 318, 366 Competition 236, 272, 273, 282, 307, 310, 311, 334, 367, 390, 392 Concurrent basins 18 Constant environments 307, 316 Continental runoff 4, 7, 8, 28, 165, 189, 193, 243, 285 Continental volcanism 180, 351 Convergence 47, 50, 80, 92, 127, 165, 171, 179, 183, 184, 196, 216, 241, 279, 294, 296, 328, 350, 387 Cordaits 5, 39, 67, 84, 91, 105, 106, 107, 108, 110, 112, 245, 280, 347, 359 Coriolis force 134, 173 Cosmopolitanism 297, 341, 365 Cradles/museums 308 Cretaceous/Tertiary boundary (KTB) 249, 258, 339, 340, 341, 342, 343, 349, 350, 351, 352, 353, 354,

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355, 356, 357, 359, 360, 363, 365, 366, 367, 369, 375, 378, 379, 380, 381 Crisis situation 339 Critical latitudes 134, 164 Cuticular coal 105 Cycadeoidea 44, 79, 115, 116, 118, 245, 246 Cycadeoidea zone 245 Cyclic speciation 297, 317-320, 385, 391 Cyclothems 15, 18, 22, 23, 25, 26, 27, 28, 29, 30, 32, 69, 73, 165, 372, 375 Dansgaard-Oeschger cycle 253 “Darwinian specimen” 311, 392 Dawny plants (convergence) 38 Dead mass export 2, 61, 103, 177, 193, 285 Dead mass production rates 12, 274, 283, 284, 316 Deltas 29, 68-72 Density heterogeneities 127, 129, 387 Density regulation 312 Density-dependent acceleration 126, 137, 179 Density-dependent fecundity 312 Density/diversity correlation 313 Detritivorous food web 237, 270, 273 Detritivory 4, 25, 271, 273, 274 Developmental acceleration 108, 199, 268, 298, 320, 325, 366 Developmental homeostasis 236, 268 Diagonal faults 132, 135 Diffuse co-evolution 273 Diffuse speciation (model) 319 Dinosaurs 18, 44, 175, 278, 341, 354, 359, 361, 362, 363, 367, 369 Disaster populations 319, 366 Disparity 129, 131, 296, 297, 303, 304, 308, 313, 320, 330, 366, 367, 381 Dispersers 11, 51, 194, 203, 271, 365 Disseminules 4, 37, 38, 42, 201, 203, 204, 298 Diversity estimates 292, 294, 295 Diversity/ solar input correlation 315 Diversity/ productivity correlation 316 Dominance 11, 58, 60, 63, 69, 70, 76, 77, 80, 82, 96, 102, 119, 233, 245, 290, 293, 334, 363, 369 Dominant forms 60, 78, 96, 293, 309, 350, 363, 367, 369 Drake Passage 149, 214 Dwarfing 366 Ecogenic law 317, 390 Ecospace differentiation 292, 318 Ecosystem collapse 16 Ecosystem efficiency 283, 292, 309, 313, 390 Ecosystem goals 268, 283, 390

Ecosystem turnover 357, 361, 362, 363 Emboly 39, 190 Endemism 308 Endosporangial germination 274 Endozoochory 276, 279, 334 Entropy/negentropy production rates 268 Environmental grain strategy (strategists) 11, 297, 307, 310, 316, 318, 319, 359, 391 Environmental induction (of genomic changes) 321, 326 Eoantha 75, 275 Ëotvös force 164 Epirift basins 17, 18, 21, 135 Equatorial glaciation 211, 212 Equatorial maximal deformation zone 134, 173 Estuaries 5, 17, 27, 35, 189, 274, 285, 350, 352, 355, 360, 391 Estuarine circulation 177 Estuarine production 7, 362 Euramerian realm 106, 108, 110, 111, 291 Eurangarian realm 108, 291 Eutrophication 3, 5, 27, 32, 35, 61, 105, 189, 220, 284, 285, 304, 315, 316, 317, 348, 350, 351, 352, 354, 357, 362, 372, 375, 391 Evapotranspiration 38, 39, 211, 221 Exocycles 22, 23, 209 Exodus effect 311, 360 Exozoochory 203, 328, 329 Extinction 15, 16, 37, 177, 194, 234, 267, 278, 279, 284, 293, 314, 341, 342, 344, 348, 350, 353, 354, 355, 359, 360, 361, 363, 365, 366, 367, 369, 380, 390 Facies realms 22, 81 Fault system (network) 19, 129, 130, 132, 134, 135, 136, 138, 151, 154, 157, 163, 167, 170, 387 Fern marsh 24, 61, 68, 72, 74, 75, 94, 102, 167, 261, 273, 298, 370, 374 Fern-covered mire 24 Fine-grained strategy 11, 307, 310, 316, 359, 391 Fluvial cycles 22, 23, 69, 373 Folivory 276, 277 Forearc basins 166 Founder effect 16, 317, 319 Fracture zones 134, 139, 163, 178 Frugivory 279 Fusenite 25, 342 Galactic cycles 165 Galactic Year 165, 265, 387 Gamma (interprovincial) diversity 293, 296, 341 Gene transduction 282, 279, 319, 330

Subject index Genetic assimilation 320, 321, 324, 325, 326 Genetic drift 267, 308, 317 Genetic introgression 319, 379, 385 Geocratic epochs 179, 232, 240 Geodethic sea-level fluctuations 178 Geofloras 80, 105, 106 Geographic parallelism 80, 280, 282 Geoid configurations 128, 137 Geomagnetic periods 339, 340, 350 Geomagnetic perturbations 357 Geomagnetic quiescent 340 Geomagnetic reversions 339, 350 Giant herbivores 284, 362 Gigantopterids 70, 71, 82, 110-112 Glossopterids 54, 84, 106, 111, 112, 113, 114, 172, 281, 328, 359 Gnetophytes 90, 94, 115, 275 Gondwana realm 105, 172 Gondwanaland 112, 169, 170, 171, 172, 173, 240 Graded/non-graded communities 367, 369 Grass biomes 94, 95, 121 Gravitational forcing (center of the Galaxy) 165 Growth increments 129, 244 Gulf Stream 214, 216 Gut contents (of fossil insects) 37, 54, 55, 277, 331334, 354 Hadley cells 234 Hadrosaurids (feeding habits) 354, 362, 363 Harmomegathy 201, 328 Heinrich event 193, 216 Helophytes 76, 81, 91 Herbivory 4, 87, 222, 236, 273, 274, 275, 276, 277, 279, 282 Heroleandra 77, 103, 304, 330 Heterochromatin (developmental effects) 325 Heterochrony 326 Heterogeneity index 294 Heterospory 77, 96, 274 Heterotrophic systems 2, 3, 274 Hirmerellids 290 Holistic concept of biotic communitis 50 Homeotic mutations 325 Horizontal gene transfer 282, 330, 391 Hypsometric curve 180, 181, 183 Greenhouse (climate) 177, 187, 207, 210, 218-234, 241, 243, 249, 250, 253 Importance value 60 Individualistic species concept 50, 272 Insect pollination (entomophily) 11, 50, 87, 308 Interaction niche 271, 272, 318

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Intracratonal basins 17, 21 Iridium anomaly (spike) 342, 359 Island arcs 17, 18, 150, 154, 165, 166, 249, 355, 378, 379 Island chains 139 Isostatic compensation 124, 127, 179 Isostatic compensation (levels) 165, 184 Isostatic sea-level fluctuations 216 Isotope records 194, 251 Jarman-Bell rule 277 Kerogens 7, 21, 177, 285 K/r strategy 318 Lagoonal circulation 177 Land bridges 168, 214, 341, 388 Late Palaeocene Thermal Maximum 250, 265 Lazarus species 366, 368 Leaf cutting 275, 276 Leaf margin 47, 80, 174, 194, 197, 198, 199, 200, 247, 273 Leaf mats 1, 42, 84, 112, 115, 245, 246, 289, 311, 373, 376, 379, 381 Leaf micromorphology 39 Leaf mimicry 47, 49, 197 Leaf polymorphism 370, 382 Leaf size classes 197 Length of day 129 Lepidophytes 39, 103, 167, 222 Leptocauly 244 Leto (demographic) strategy 312, 318 Life forms 1, 37, 39, 42, 50, 60, 79, 80, 84, 95, 96, 112, 195, 236, 268, 269, 284, 293, 294, 341 Limnobiophyllum 103, 167, 374, 375 Limnoniobe 103, 303 Linear magnetic anomalies 124, 133 Little Ice Age 217, 218 Localities Augustovka 258, 343, 367, 375, 378 Baisa 55, 201, 298 Bureya 51, 69, 141, 143, 198, 295, 372 Gerofit 35, 189 Gurvan-Eren 298 Kamenka 66, 68, 69, 72, 73, 91, 102, 198 Karatau 91, 102, 228, 230 Kavalerovo 255 Khanka 261 Khoindjo Point 255 Kiwda 259 Koonwarra 303 Krynka 261 Kudrintsy 270

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Laibin 70, 106 Manlay 298 Nammoura 35, 189 Nedubrovo 27, 66, 73, 87, 91, 110, 189, 234, 345, 347, 352 Nemegt 4, 53 Pavlovsk (Voronezh) 105 Peski 66, 72 Razdolnaya 262 Sel-Bukhra 93 Sussex 342 Tastakh 253 Tchekarda 55 Tsagajan 13, 196, 249, 258, 370, 372, 373, 375, 379, 381, 385 Umalta 51, 69, 73 Ust-Baley 44, 59, 295 Yixian 298 Yuri Island 258 Lystrosaurids 354 Macromutation 367, 391 Macropolymorphism 381 Madro-Tertiary flora (geoflora) 79 Magnetostratigraphy 256, 260, 261 Mangal ecosystems 250, 274 Mangrove peat 24, 25, 72 Mangroves 24, 28, 44, 51, 83, 96, 102, 112, 120, 208, 250, 252, 253, 286 Marginal basins 21, 23, 151, 165 Marshgas condensates 225 Mass mortality 16, 30, 31, 298, 359, 360, 390 Medieval optimum 263 Meso-Mediterranean realm 44, 77 Metabolic aerosols 233 Methane hydrates (outgassing of) 219, 250 Microbial films 8 Microbial mats 8, 15, 269, 283 Mid-ocean rises (ridges) 124, 132, 133, 164 Milankovitch cycles 23, 208, 216 Mixed (geomagnetic) polarity 261, 340 Molassoids 143, 154, 165, 191, 210, 255 Monilitheca 4, 103 Monsoon intensity (index) 29, 210, 217, 253 Monsoonal circulation 252 Monstrosity 341, 366 Morphological characters 37, 39, 196, 282, 328, 391 Mosaic structure (of fold belts) 141, 164 Mountain torque 127 Multifunctional characters 37-42 Mutagenic effect of ozone depletion 234, 235

Mycorrhiza 39, 194, 224, 236, 271, 272, 279, 280 NADW 193, 209, 214, 216, 217, 220 Nemoral zone 79, 108, 205, 248, 249, 294, 295 Nepheloid layer 17 Neutrality (neutralism, neutralists) 272, 317 Niche creating trend 308 Niche overlap 268, 309, 310, 311, 313, 361, 390, 392 Niche splitting 307 Nilssonia (nilssonia) 43, 44, 49, 50, 68, 69, 70, 73, 84, 198, 199, 258, 370, 375, 378, 379, 380 Niobe (demographic) strategy 311, 318, 390 Non-linear reponse to orbital forcing 210 Normalizing selection 314, 366, 385, 391 Normapolles province 175 North Atlantic land bridge 169, 214, 253 Oceanic plateaux 181, 183 Oceanic ridges 135, 139, 148, 154, 164, 168, 183 Ombrotrophic peat (coal) 25, 108, 165, 188 Ophiolites 134, 138-141, 143, 144, 148, 153, 158, 159, 162, 173-175, 182 Orbital cycles (also Milankovitch cycles) 22, 208 Orestovia 96, 105, 236, 245, 284 Orographic (climatic) effect 105, 123 Orthogonal faults 132, 135 Overbrowsing 77 Overrepresentation 15, 58, 81, 105 Ozone hole 233 Pachycauly 244 Pacific-type ophiolites 144, 153 Paedomorphism 318, 320, 325, 338, 366, Palaeocene/Eocene boundary 250, 251 Palaeosol (fossil soil) 1, 25, 27, 51, 53, 187, 190, 251, 270 Palaeosyntaxa 59-61 Palynological assemblages (spectra) 13, 118, 258, 290, 328 Palynomorphology 327, 338 Paralic coal 24, 193, 286, 375 Paramutation 324 Pecopterids 70, 110, 114 Pedogenic carbonates 110, 190, 227 Pedogenic weathering 177, 223, 227, 242 Permian/Triassic boundary, PTB 339, 340, 342, 344, 345, 347, 349, 350, 351, 352, 353, 354, 356, 359, 360, 365, 366 Phoenicopsis zone 245, 246 Picnocline (model of black shale deposition) 189 Plant - animal interactions 82, 203, 245, 279 Plant - arthropod systems 274-276 Plant - fungal systems 271

Subject index Plant - insect interaction 47, 50, 194, 197, 198, 201, 336 Plant communities 11, 13, 50, 51, 63, 68, 75, 76, 77, 84, 114, 166, 236, 269, 271, 273, 277, 283, 303, 306, 308, 361, 362, 367, 370, 379, 381 Plant growth rates 222, 274 Plant productivity 222, 293 Pleuromeia 42, 43, 72, 96, 249, 273, 360, 365 Polar vortex 233 Polar wander 123, 129, 139, 164, 246 Pollen bladders (sacci) 38, 334 Pollen load 55, 201, 272, 282, 329, 330 Pollinivory 274, 282, 334 Post-crisis communities 385, 391 Post-crisis populations 319, 366, 381 Predictable environments 304 Preflosella 55, 75, 201, 331, 334 Proangiosperms 18, 75, 102, 103, 204, 296, 298 Progression (migration) of volcanic activity 150 Psychrospheric circulation 169, 216, 252 Pteridosperms 54, 76, 87, 106, 276 Pull-apart basins 138, 154, 167, 375 Quereuxia 103, 167, 375 Recovery 63, 284, 292, 306, 337, 363, 367, 369 Recurrence 58, 253, 330 Redbeds 9, 19, 22, 26, 81, 111, 115, 167, 174, 187, 190, 191, 192, 193, 211, 238, 246, 347 Redox boundary 26, 342, 357 Redudant density (redundancy) 311, 313, 314, 315 Relaxation of selection 318, 366, 385 Relicts 196, 258, 344, 381 Repetitive DNA (developmental effects) 320, 321, 325, 326 Replication rates 391 Reproductive strategies 201, 311, 312, 318, 337, 392 Resinite 25, 82, 167 Retroconvergence 321, 324 Rhacophyton zone 105 Rheotrophic peat (coal) 8, 25, 165 Rhizocretions 22, 27, 81, 251 Rhizosphere 8, 22, 39, 191, 223, 236, 271, 272, 284, 357, 390 Rift basins 17, 21, 32 Rift valleys 18, 123, 302, 304 Rotation forcing 123, 126, 179, 184, 241 Rotation rates 127, 129, 137, 178, 179, 210, 218, 238, 241, 350 Rotation stress 127, 173 Samara 37, 87, 203, 204, 261, 373, 378 Savannoid vegetation 95, 121 Scleromorphy 39, 42

397

Sclerophylly 200 S-cracks 132, 164 Señor-Lipps effect 359 Sea gateways 167-169, 214 Sea-floor volcanism 163, 164 Sea-level fluctuation models 177-180 Sedimentary basins 16, 67, 69, 166, 210, 254, 372 Sedimentary environments 1, 26, 214, 221, 351 Seed germination 47, 196 Seed/fruit size 201 Seedlings 44, 87, 96, 112 Seed-plant phylogeny 281 Selective codon usage 324 Seral sclerophylls 42 Sere (ecological autosuccession) 15, 63, 76, 96, 316, 317 Sere/evolution parallelism (law of) 317 Shear faulting 125, 131, 132-134, 137-141, 147-150, 153, 155, 164 Short-term/long-term genetic memory 326 Shrub biome 90-94, 119, 289 Shuguria 96, 105, 236, 245, 284 Smectitic clay 27, 345 Snow line 211, 212, 213, 243 Soil arthropods 52, 224, 271 Solar constant 208, 238 Spatial heterogeneity 293, 294, 296, 304, 306, 307 Speciation 16, 178, 297, 308, 317, 318, 319, 353, 367, 391 Speciation cycles (also cyclic speciafion) 317, 385 Speciation model 318 Species longevity 178, 297 Species richness/disparity correlation 58, 197, 272, 287, 293, 294, 295, 296, 297, 298, 307, 308, 310, 316, 317, 363, 366, 379, 381 Sporangia 51, 69, 84, 201, 271, 274, 276, 280, 303, 309, 360 Sporivory 274, 333 SST 209, 210, 216, 217, 218, 220, 221, 224, 225, 230, 233, 243, 263, 285, 286 Stomatal index 227 Stomatal length index 228 Stratospheric ozone 123, 209, 226, 232, 233, 234, 236, 243 Strike-slip faulting (also shrear fauting) 132, 138 Subangarian realm 108 Subcructal shear 140 Suess effect 237 Superspecies 320 Sustainability 196, 269, 292, 304, 312, 313, 392

398


Symbiotic integration 279 Syntaxonomy 59 Talassocratic epochs 179, 183, 387 Tanaitis 84, 105 Taphoflora 59, 79, 84, 110, 114, 166, 213, 228, 258, 293, 294, 295, 369, 370, 379, 381 Taphonomic bias 11, 15 Taphonomic environments 17, 21, 165, 166 Taphrogenic basins 16 Tasmantis 178, 181 Tectonic cycles 138 Teleconnection (climatic) 214, 216, 217, 253, 353 Temperature - CO2 system 224 Temporal heterogeneity 304, 305 Tethyan transgressions 178, 215 Tethys (belt) 157-162, 164, 213, 265 Tetrapod herbivory 276-278 Thermocline 26, 30, 32, 209, 214, 216, 224, 263 Thrust belts 141, 144, 146, 148, 153, 154, 157, 158, 163, 174, 184 Tidal deceleration 126, 127, 179 Tillites 106, 174, 188, 211, 238, 239, 240 Time lag (of CO2 flactuations) 225, 226 Time lag (of biosphere responses) 194 Tolerance 13, 59, 83, 84, 194, 195, 196, 197, 230, 271, 272, 277, 306, 310, 313, 315, 317 Trade wind belts (shift of) 217 Transcription rates 325, 326, 391 Transcurrent faults 19, 20, 124, 133, 154, 161, 261 Trauscurrent basins 19, 20

Tree ferns 77, 80, 222, 274 Tree line 58, 194, 212, 213, 272, 284 Trichomes 38, 39, 68, 197, 328 Trochodendroides 13, 47, 84, 167, 246, 249, 252, 256, 257, 258, 259, 260, 319, 365, 373, 375, 376, 378, 379, 381, 382, 383, 385 Tundra-steppe 90, 95, 317, 354 Turgay Straight 216 Two-stage extinctions 366 Tympanicysta 5, 27, 35, 347, 351, 352, 360 Typhaera 203, 298 Underground organs 39, 91 Underreplication (of genes) 325, 391 Unions (in syntaxonomy) 60, 61 Variegate deposits (cycles) 110 Varvites 23, 245 Vavilov’s law 385 Vojnovskyaleans 106 Volcanic aerosols 218, 234 Volcanism (climatic effect) 16, 17, 18, 26, 124, 125, 137, 138, 139, 143, 163, 164, 180, 184, 211, 226, 241, 302, 304, 342, 351, 352, 365, 375, 379 Wallace Line 172 Water stress 39 Wind pollination (anemophily) 38, 201 Wind roughness (of leaf morphology) 80, 197, 245 Xeromorphy 39, 42, 90 Xerothermic zone 73, 110, 113, 119, 286 Zonal rotation 134 Zoocecids 275

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