Deposition and Diagenesis of the Early Permian Lower Parmeener Supergroup Limestones, Tasmania

Becky Rogala

A thesis submitted to the Department of Geological Sciences and Geological Engineering in conformity with the requirements for the degree of Doctor of Philosophy

Queen’s University Kingston, Ontario, Canada April, 2008

The Lower Parmeener Supergroup consists of 500 to 900 metres of marine and terrigenous sedimentary rocks, deposited in the high-latitude

Tasmania Basin during the late Carboniferous to middle Permian, at the end of the late Paleozoic ice age. Two bioclastic carbonate units, the Darlington and

Berriedale limestones, are of particular interest due to their formation in this polar, cold-water environment. Both limestones contain ice-rafted debris scattered throughout, signifying numerous icebergs, and are under- and over-lain by glendonitic siltstone indicating near-freezing seawater. Despite the unusual environment, seawater in the Permian Tasmania Basin was, with the exception of an anomalously high δ13C value, isotopically and chemically similar to modern seawater.

These limestones consist of a high-abundance, low-diversity heterozoan assemblage, dominated by large, robust brachiopods, bryozoans, and

Eurydesma bivalves. Sponge spicules and crinoids are locally important constituents. The carbonates are interpreted to have been deposited in mid-shelf environments during sea-level highstands, where the faunal communities were beyond the depths of grounding icebergs, and sufficiently outboard from terrigenous sediment influx and brackish water. Growth and preservation of biogenic carbonates were promoted by up-welling of nutrient-rich water, which sustained high levels of primary productivity in the water column and phosphate concentrations in the sediment.

ii Lower Parmeener Supergroup carbonates were exposed to a complex series of diagenetic processes, commencing on the seafloor and continuing during rapid burial. Limestone composition was further modified by diagenetic fluids associated with the intrusion of Mesozoic igneous rocks. Alteration in the marine paleoenvironment was both destructive and constructive; although dissolution took place there was also coeval precipitation of fibrous calcite cement, phosphate, and glauconite. These processes are interpreted to have been promoted by mixing of marine waters and enabled by microbial degradation of organic matter. In contrast, meteoric diagenesis was insignificant, being confined to minor dissolution and localized cementation, although mechanical compaction was ubiquitous. Chemical compaction was instigated at burial to depths of approximately 150 m, and promoted extensive precipitation of ferroan calcite. Diagenesis may well have ended here, except for the subsequent intrusion of massive Mesozoic diabases and associated injection of silicifying fluids into the limestones. Finally, fractures associated with Cretaceous uplift were filled with late-stage non-ferroan calcite cement.

iii STATEMENT OF CO-AUTHORSHIP

The following manuscripts are my own work, but owe much to the

intellectual and scientific guidance, and precise editing, of Dr. Noel P. James,

who is co-author of each manuscript and supervisor of this thesis research. Co-

authors Dr. Catherine M. Reid (manuscript 1), Clive R. Calver (manuscript 2), and T. Kurtis Kyser (manuscript 3) provided invaluable scientific advice and

discussion that led to the improvement of this thesis.

iv ACKNOWLEDGEMENTS

I would like to express my gratitude to Dr. Noel James for his guidance

through the pitfalls of graduate studies. His boundless enthusiasm for teaching

and for science is infectious. Under his tutelage I learned not only about the

intricacies of carbonate sedimentology and diagenetic processes, but also about

the qualities that make a great teacher and supervisor. I am also thankful for the

incredible opportunities I have been given: field-trips to Bermuda, New York, and

Quebec, field work in Australia, and of course the chance to work on my

presentation techniques for GSA.

I would also like to thank Dr. Kurt Kyser for his patient explanations while I developed my geochemical intuition, and for advising on Lower Parmeener

Supergroup geochemistry. I am also appreciative of the assistance provided by

Kerry Klassen with preparing samples and running stable isotope analyses, and to Dr. Don Chipley and Bill MacFarlane for directing the preparation of samples and running elemental geochemistry analyses.

Catherine Reid and Clive Calver were instrumental to the completion of this thesis. Thank you to both of them for introducing me to the Lower

Parmeener Supergroup. Catherine was particularly helpful in keeping my stratigraphic nomenclature straight, as well as providing insightful advice regarding the depositional environments and stratigraphic relationships. Clive’s

knowledge of outcrops in the Tasmania Basin and keen observations of grain

and diagenetic relationships were invaluable.

v I would like to acknowledge two of my fellow graduate students who were always available for impromptu thesis discussions. A big thank you to John

Rivers for a willingness to chat about all things carbonate, particularly anything related to geochemistry. It has been interesting sharing an office with John during these last four years between ducking paper missiles, playing the “Price is

Right”, listening to his plans to be a truck driver instead, and enduring seemingly endless seasons of NFL – go Eagles. Thanks for keeping it fun. I’m also grateful to Mike Johnson for always being available for discussions on sequence stratigraphy. I’m sure he explained it to me a half dozen times before it sunk in – thanks for your patience.

Of course my acknowledgements would not be complete without expressing my gratitude to Tom Hamilton, my husband and field assistant. This thesis would not have been possible without your expert fossil extraction skills and core-splitting abilities. Thank you for following me down to southern Ontario and Tasmania, for distracting me just enough to stay sane, and for the continual encouragement along the way.

vi TABLE OF CONTENTS

Abstract …………………………………….………………..………………….…. ii

Statement of Co-Authorship …………………………….….……….………….… iv Acknowledgements ……………………………………..…………….……...…… v

Table of Contents ……………………………………….………………………… vii

List of Tables …………………………………….………………………………… viii

List of Figures ……………………………………….………………..…………… viii

Dedication ……………………………………………………………..…………… x

CHAPTER 1: General Introduction ……………………………………...………. 1

CHAPTER 2: Deposition of polar carbonates during interglacial highstands on an Early Permian shelf, Tasmania (Journal of Sedimentary Research, v. 77, p. 587-606)

Rogala, B., James, N.P., and Reid, C.M. ………………….……. 12

CHAPTER 3: Diagenesis of Early Permian high-latitude limestones, Lower Parmeener Supergroup, Tasmania (Sedimentology, in review)

Rogala, B., James, N.P., and Calver, C.R. …………………...… 66

CHAPTER 4: Geochemistry of Permian brachiopods and eurydesmid bivalves from the Lower Parmeener Supergroup in Tasmania: implications for deducing paleoceanographic conditions from biogenic carbonates (Journal of Sedimentary Research, in review)

Rogala, B., Kyser, T.K., and James, N.P. …………………….… 111

CHAPTER 5: General Summary and Extended Conclusions ……………...… 141

References ……………………………………..…………………………..……… 149

Appendix A: Core Logs ……………………………………………………….…... 182

vii LIST OF TABLES

CHAPTER 1

Table1.1: Summary of Tasmanian, Australian, and Pangean Events …... 6

CHAPTER 2

Table 2.1: Summary of lithofacies in the Lower Parmeener Supergroup 24-25

CHAPTER 3

Table 3.1: Summary of calcite cement properties ………………………..… 80

CHAPTER 4

Table 4.1: Summary of diagenetic stages recorded in the petrography of limestones from the Permian Lower Parmeener Supergroup … 119 Table 4.2: Brachiopod isotopic and elemental geochemistry …………….. 123 Table 4.3: Eurydesmid isotopic and elemental geochemistry …………….. 125

LIST OF FIGURES

CHAPTER 1

Figure 1.1: Reconstruction of Pangea ……………………………………...... 5 Figure 1.2: Reconstruction of Late Paleozoic ice extent ……………………. 8

CHAPTER 2

Figure 2.1: Location map ………………………………………………….....… 14 Figure 2.2: Stratigraphic relationships of units in the Lower Parmeener Supergroup ………………………………………………..…..…… 17 Figure 2.3: Lower Parmeener Supergroup nomenclature used in this study 18 Figure 2.4: Regional geology and distribution of drill holes ………………... 20 Figure 2.5: Extent of glaciation ………………………………………………... 21 Figure 2.6: Lower Parmeener Supergroup cross-section ………………….. 26 Figure 2.7: Siliciclastic-dominated lithofacies ……………………………...… 28 Figure 2.8: Carbonate lithofacies ……………………………………………... 35 Figure 2.9: Spiculitic limestone lithofacies …………………………………… 39 Figure 2.10: Plan and cross-section views of facies distribution — Asselian 41

viii Figure 2.11: Plan and cross-section views of facies distribution — Sakmarian 43 Figure 2.12: Plan and cross-section views of facies distribution — late Sakmarian/early Artinskian …………………………………..…… 45 Figure 2.13: Plan and cross-section views of facies distribution — Artinskian 47 Figure 2.14: Sequence stratigraphic section of the Lower Parmeener Supergroup ………………………………………………………… 49

CHAPTER 3:

Figure 3.1: Reconstruction of eastern Pangea ……………………………… 69 Figure 3.2: Position of the Darlington and Berriedale limestones within the Lower Parmeener Supergroup stratigraphy ……………...… 70 Figure 3.3: Regional geology of Tasmania …………………………………... 73 Figure 3.4: Thin-section photomicrographs of representative facies ……… 79 Figure 3.5: Carbonate cement stratigraphy …………………………..……… 81 Figure 3.6: Thin-section photomicrographs of carbonate cements ………... 82 Figure 3.7: Thin-section photomicrographs of siliceous cements …………. 88 Figure 3.8: Paragenetic sequence of Lower Parmeener Supergroup Cements ……………………………………………….…………… 92 Figure 3.9: Diagenetic burial history curve …………….…………………….. 93

CHAPTER 4:

Figure 4.1: Location of the Tasmania Basin ……………….………………… 115 Figure 4.2: Stratigraphy of the Lower Parmeener Supergroup …………..… 117 Figure 4.3: Typical effects of alteration on Lower Parmeener Supergroup Shells …………………………………….……………………….… 121 Figure 4.4: δ18O and δ13C values for brachiopods and eurydesmids ……... 124 Figure 4.5: Brachiopod and eurydesmid δ18O and δ13C values compared to previous work …………………………………………………….… 125 Figure 4.6: Minor element geochemistry for brachiopods and eurydesmids 127 Figure 4.7: Trace element geochemistry for brachiopods and eurydesmids 128 Figure 4.8: Comparison with modern brachiopods ………………………….. 130

To Tom, my family, and friends:

Thank you all for your support over this last decade or so of higher education,

particularly the last five years of which has culminated in this Ph.D. thesis

x CHAPTER 1:

GENERAL INTRODUCTION

Carbonate sediments serve as sensitive indicators of paleoceanographic

conditions, affording us an invaluable record of climatic variation that can be

used to interpret our presently changing environment. Initially this record was

limited to warm, low-latitude regions of prolific carbonate production since these

areas were easy to access. While the focus was on warm-water carbonates,

there was early recognition that cool ones were present (Chave 1952; Lees and

Buller 1967; Lees 1975). Work began in earnest, however, in the 1970s and early 1980s with the recognition of large cool-water carbonate occurrences off the coasts of New Zealand (Carter 1975; Nelson 1978; Nelson 1988) and

Australia (Connolly and Von der Borch 1967; Wass et al. 1970). It was realized that carbonate sediments were forming abundantly in modern oceans in water temperatures between 10-20°C, albeit with significant biotic differences from their

warm-water counterparts. These variations in skeletal composition were

originally assigned a plethora of descriptive names based on the specific biota

present (e.g., Nelson 1988). James (1997), however, unified the classification of

carbonate sediments into two categories. Photozoan carbonates, frequently

associated with warm-water accumulations, contain photosymbiotic organisms, green algae, and inorganic precipitates, whereas heterozoan carbonates, typically found in cool-water settings, do not.

1 Rao (1981, 1983, 1986) extended the concept of cool-water carbonates

one step further, proposing that Permian Lower Parmeener Supergroup

limestones in Tasmania had been deposited in a polar setting based on the

paleo-latitude derived from paleomagnetic studies, their proximity to glacial facies

and presence of iceberg-rafted debris within the limestone units. Research on these polar carbonates has been relatively slow, primarily due to the paucity of modern analogues, basically restricted to a few Antarctic basins (Taviani et al.

1993; Rao et al. 1998; Taviani and Beu 2003) and the Svalbard Shelf (Andruleit et al. 1996), deep cold-water coral mounds (Freiwald et al. 1997, 2002; Lindberg and Mienert 2005), and limited recognition of ancient polar limestones beyond these Permian examples.

As a consequence, there remains uncertainty regarding oceanographic conditions under which cold-water carbonates form, how the sediments have changed over time, how they behave during diagenesis, and how they are preserved in the rock record. The Permian Lower Parmeener Supergroup of

Tasmania has been revisited to address some of these questions. The cold-

water affinity of these limestones, deposited at high southern latitudes during the

end of the late Paleozoic glaciation is well-established, yet these carbonates

have not been studied in terms of their paleoceanography or their diagenetic

history.

Besides simply understanding the processes that affect polar carbonates

or their detailed diagenetic history using modern techniques and understanding,

investigating the Lower Parmeener Supergroup limestones also provides insight

2 into the dynamics of deglaciation during the Early Permian. This glacial-

interglacial period has been in the spotlight in recent years (Isbell et al. 1997,

1993; Jones and Fielding 2004; Jones et al. 2006). Much of the current research

on late Paleozoic sedimentary rocks is focused on high-latitude siliciclastic

basins and low-latitude carbonate platforms. Paleoceanographic information

from high-latitude carbonates, however, has the potential to provide a more

detailed record of this critical period.

THESIS OBJECTIVES

Previous studies of the Lower Parmeener Supergroup have focused

primarily on biostratigraphic correlations, as well as environmental interpretations

using glacial indicators, faunal assemblages, and geochemistry. The purpose of

this thesis is to examine the petrography and stratigraphic relationships of the

Lower Parmeener Supergroup limestones using modern techniques. The specific objectives are (1) to describe the carbonate facies and their distribution in a sequence stratigraphic framework; (2) to interpret the physical and chemical paleoceanographic conditions that promoted their accumulation; (3) to make inferences about ocean chemistry during the Permian; (4) to examine the deglaciation history of the late Paleozoic ice age; (5) to determine how cold-water carbonates are affected by diagenesis and are preserved in the rock record; and

(6) to compare cold- and cool-water carbonates in terms of sedimentology and diagenesis. This work incorporates observations from outcrop and 18 drill cores from locations across the Tasmania Basin. A total of 245 petrographic sections

3 were examined using standard petrographic techniques, cathodoluminescence, and scanning electron microscopy. Finally, 26 brachiopods and 9 eurydesmids were processed and analysed for carbon and oxygen isotopes and trace element geochemistry.

PANGEAN AND AUSTRALIAN CONTEXT

Tectonics

Deposition in the Tasmania Basin was governed by southern Pangean tectonic events, climate, isostatic rebound, and local controls such as fluvial input and ocean currents. Pangea had formed by 320 Ma (Early Pennsylvanian) and began breaking up at 230 Ma (Early Triassic). There are many versions of paleogeographic reconstructions for the Pangea, based on paleomagnetism and fossil correlations. While the distribution of large continents is well established, the distribution of microcontinents, such as India, and the Tethyan terranes remain controversial (Dickins, 1994; Metcalfe, 1994). This thesis uses the Early

Permian paleogeography devised by Li and Powell (2001), shown in Figure 1.1.

Figure 1.1. Reconstruction of Pangea (after Li and Powell 2001). Black square indicates Tasmania, the study area. Surface current directions are based on models by Winguth et al. (2002).

The earliest stages of Pangean rifting (Stage A), expressed in Eastern

Australia by extension and platform volcanism, began in the Early Permian

(Table 1.1) (Veevers et al., 1994). The tectonic style changed in the Late

Artinskian, however, as a magmatic arc, approaching from the east, collided with

Eastern Australia, first in the north and later in the south (Veevers et al. 1994).

This collision created a series marine sag basins along the east coast of

Australia. The arc, shown in topographic reconstructions by Ziegler et al. (1997),

appears to form a protective barrier around Tasmania during the Sakmarian and

nearly isolates it from the open ocean during the Artinskian. However, many

continental reconstructions using paleomagnetism do not show the arc extending

this far south (Scotese and Langford, 1995). The presence of the magmatic arc

may have additional implications for Tasmania, such as influencing relative sea

level, surface-water circulation patterns and upwelling locations.

Climate

The Pangean Supercontinent provided an ideal configuration for developing an extreme climate regime. The large size effectively insulated the continental interior, leading to severe aridity that was exacerbated by the rain- shadow effect produced by high topographic relief along the Tethyan margin

(Barron and Fawcett, 1995). Conversely, these coastal mountain barriers

6 promoted precipitation and concentration of monsoon and winter storm activity

on the Tethyan Ocean side (Fig. 1.1).

Much of the Pangean land-mass was positioned in southern hemisphere

temperate to polar regions throughout the late Paleozoic (Fig. 1.1), favouring the

expansion of extensive glaciation across the Gondwanan continents. There were two peak glacial periods, one during the Moscovian and one in the

Asselian/earliest Sakmarian of the Permian (Isbell et al. 2003; Jones and Fielding

2004).

While this glaciation is incontrovertible, with evidence occurring from the

South Pole to the Arabian Peninsula, there has been disagreement regarding the nature and timing (Fig. 1.2). Recent evidence suggests that the Pennsylvanian glaciation may not have been in the form of thick extensive cratonic ice caps as originally envisaged by Veevers and Powell (1987). Dickins (1996) points out

that many of the Late Carboniferous rocks in Australia that were formerly referred

to as tillites actually appear to be marine debris flows. Furthermore, Australian

varvites may only indicate the seasonal presence of ice and some of the striated rocks are not of glacial origin. Dickins (1996) pointed out that montane glaciers dominated the Pennsylvanian glacial peak. This view is supported by the work of

Isbell et al. (2003) in the Transantarctic Mountains where it appears that many

Transantarctic basins were not glaciated until the Permian, while others were not glaciated at all.

Figure 1.2. Reconstruction of several interpreted ice-sheet extents for the Carboniferous and Permian (from Isbell et al. 2003). The rectangle marks the position of Tasmania.

The Early Permian glacial peak is now thought to have been the larger of

the two episodes (Dickins, 1996; Isbell, 2003). Dickins (1996) also proposed that

Permian glaciers were wet-based and flowed from areas of high relief rather than continent-wide dry-based ice sheets. This opinion stems from the large amount of till concentrated in valleys. Evidence from some Australian basins suggests that the glacial debris was deposited in rapidly subsiding areas that lacked previous glacial scouring. Lindsay (1997) and Eyles et al. (1997) maintain that there were thick ice sheets, based in east Antarctica, that covered Antarctica,

8 Australia, southern Africa, parts of South America, and India during the Permian glacial peak.

Pangea moved progressively northward from the Late Carboniferous through to the end of the Permian. This change in latitude caused reciprocal variations in paleoclimate. Northern Pangea, particularly Siberia, began as a temperate zone in the Late Carboniferous and ended the Permian as a polar zone with evidence of marine ice (Barron and Fawcett, 1995; Ziegler et al.,

1997). Northern Canada moved from a subtropical zone in the Late

Carboniferous to a temperate zone in the Early Permian and entered a polar climate realm in the Late Permian, as seen by a progressive decrease in the diversity of the biota and a change from from evaporites and photozoan carbonates to heterozoan limestones, spiculites and possible dropstones in the

Late Permian (Beauchamp, 1994; Ziegler et al., 1997).

Southern Pangean climates, in contrast, experienced a progressive warming trend (Dickins, 1996; Ziegler et al. 1997). Globally correlated eustatic sea-level rise in the late Sakmarian tends to be attributed to the melting of glaciers associated with this warming event (Veevers et al., 1994; Dickins, 1996), although Isbell et al. (2003) expressed some doubts that the ice sheets were extensive enough to cause major sea-level fluctuations. By the Late Permian,

Africa and Argentina were covered in lakes and coal swamps with a temperate biota. Antarctica, which was heavily glaciated during the Early Permian, was covered by Glossopteris forests and a variety of ferns and horsetails typical of humid, cool temperate climates by Late Permian (Ziegler et al., 1997).

9 Oceanography

Large-scale ocean circulation patterns of Panthalassa, the one large ocean during the Permian, remain unclear as there have been few attempts to model them, particularly during the Early Permian glaciation. Winguth et al.

(2001), however, modeled the warmer Kazanian and determined that deep water should have formed in northeastern Panthalassa during the northern winter and in the southeast during the southern winter. This would have set up a deep- water circulation pattern such that there was a strong westerly current at 70°N and 70°S latitude moving equatorward along the east coast of Pangea (Figure

1.1). Surface currents would have flowed in the opposite direction, moving warm equatorial water poleward along eastern Pangea. A strong easterly equatorial deep current is interpreted to be disrupted by islands in the eastern Tethys Sea and areas of upwelling to be concentrated along the west coast of Pangea

(Winguth et al., 2001).

Causes of Glaciation

Many climate models for the Pangean Supercontinent cite decreasing atmospheric CO2 levels as the main reason for ice-sheet formation (Gibbs et al.,

2002; Winguth et al., 2002), although there may have been other contributing

factors. Veevers and Powell (1987) suggested that glaciation was initiated by

large-scale uplift. Crowley and Baum (1992) disagree with this idea based on

their models, which indicate that geographic changes, solar luminosity, and

atmospheric CO2 were important. Scheffler et al. (2003) proposed that albedo

10 effects, low atmospheric CO2, and the influence of warm currents bringing

moisture into polar zones were key to the formation of glaciers. Fawcett et al.

(1994) found that adding an open equatorial seaway to their models provided a

better agreement between predictions and geologic data for Australia and India,

although the best-fit scenario kept this seaway open throughout the Late

Permian.

Conversely, increasing atmospheric CO2 is the main reason cited for

deglaciation (Gibbs et al., 2002; Winguth et al., 2002; Scheffler et al., 2003).

Wopfner (1999) observed that deglaciation at the beginning of the Sakmarian was synchronous and rapid along the Tethyan margin. Izart et al. (2003) correlated this event on virtually every continent, noting that deposits associated with deglaciation contain localized high organic content, which is interpreted to indicate a sudden increase in bioproductivity relative to glacial periods. Wopfner

(1999) suggests that this increase was due to a global increase in temperature and atmospheric CO2 content, which built up naturally during the glaciation

period. Increased CO2 was explained by widespread volcanism coupled with

reduced amounts of photosynthesis during glaciation, which served to increase

atmospheric pCO2 while removing the CO2 sink provided by photosynthesizing

organisms. Alternatively, Scheffler et al. (2003) proposed that deglaciation,

initiated or enhanced by the closure of the equatorial seaway, caused a

breakdown in the strong ocean circulation pattern that augmented ice formation.

11 CHAPTER 2:

DEPOSITION OF POLAR CARBONATES DURING INTERGLACIAL

HIGHSTANDS ON AN EARLY PERMIAN SHELF, TASMANIA

Rogala, Becky, James, Noel P., and Reid, Catherine M.

ABSTRACT

The lower Permian Lower Parmeener Supergroup contains two bioclastic limestones that accumulated during the period of prolonged deglaciation of southern Gondwana as the region moved equatorward from 80° S to 70° S. The

Darlington and Berriedale limestones formed in neritic environments, in areas where abundant ice-rafted debris testifies to numerous icebergs and glendonites

indicate near-freezing seawater. These limestones consist of argillaceous and

clean bioclastic floatstone, rudstone, and grainstone that contain a high- abundance, low-diversity heterozoan assemblage of calcareous invertebrates.

The components are dominated by large, robust brachiopods, bryozoans, and

Eurydesma bivalves. Sponge spicules and crinoids are also common, whereas

coralline algae and conodonts are conspicuously absent.

Carbonates were deposited on the middle shelf during sea-level

highstands, below the iceberg grounding line, where the faunal communities

remained undisturbed. In this setting the organisms were outboard of significant

terrigenous sediment influx and brackish water, which were trapped on the inner

shelf by bathymetry and icebergs. Strong bottom currents also prevented the

deposition of fine-grained siliciclastics by continuous winnowing of the pure

12 limestone facies. Upwelling of nutrient-rich water, inferred from the distribution of phosphate, promoted high primary productivity, which fueled this carbonate factory and inhibited dissolution of the biogenic carbonate.

INTRODUCTION

The greater part of carbonate production occurs in photozoan-dominated, tropical, warm-water settings (Bathurst 1975; Tucker and Wright 1990) and much less so in heterozoan-dominated, temperate, cool-water environments (Nelson

1988; James 1997). Part of the reason for this distribution is the increased solubility of CO2 in colder water, which leads to decreased concentrations of

2 CO3 ¯, which in turn inhibits CaCO3 precipitation and creates sluggish biotic precipitation rates (Mutti and Hallock 2003). These characteristics are even more acute in polar, cold-water settings. Nevertheless, polar limestones do form in modern environments (Taviani et al. 1993; Andruleit et al. 1996), and are inferred to have done so in the past (Beauchamp and Baud 2002). Carbonate strata deposited in these cold environments need to be compared to extant cool- water flora and fauna in order to facilitate their recognition in the rock record. In addition, it is important to study such deposits in terms of the chemical and physical oceanographic mechanisms that allow their precipitation and preservation.

The Lower Parmeener Supergroup (Forsyth et al. 1974) is a 500- to 900- meter-thick marine and terrigenous succession deposited at a high paleolatitude in the Tasmania Basin (Fig. 2.1) during the late Carboniferous to middle Permian.

13 Marine lithologies consist of both siliciclastic and carbonate rocks that have long been interpreted as cold-water in origin (Banks 1957). Previous work has focused on biostratigraphy (e.g., Clarke and Farmer 1976; Reid 2003) and interpretation of lithofacies and paleoenvironments of specific units (Brill 1956;

Clarke and Banks 1975; Brill 1982; Rao and Green 1982). Lithofacies have

recently been reviewed, and the stratigraphic nomenclature revised, by Reid et

al. (in press). There is no integrated sedimentological analysis of these important

rocks.

Figure 2.1. Reconstruction of Pangea, from Li and Powell (2001) with the study area shown in rectangle. Present-day continents are shaded dark gray, and the extent of Permian shelves is shown in light gray. The directions of modeled deep ocean current directions are indicated by dashed lines. (Winguth et al. 2002).

This paper presents the results of a recent detailed sedimentological analysis

of core and outcrop across Tasmania. Four phases of deposition are recognized

and used to demonstrate the effects of changing sea level, ice cover, and

14 circulation patterns in the Tasmania Basin during early Permian deglaciation, that then provide insight into the mechanisms that enabled carbonate formation.

GEOLOGICAL SETTING

Tectonics

The Pangean supercontinent had largely amalgamated by the late

Carboniferous and had gradually migrated northward during the Permian. During

this time Tasmania was positioned near the South Pole (Fig. 2.1), between

Antarctica and mainland Australia, and moved northward from approximately 80°

S in the Pennsylvanian to 70° S in the Permian (Scotese and Langford 1995; Li and Powell 2001).

The tectonic evolution of eastern Australia and Tasmania during this

period was complex, with alternating periods of compression and extension.

Prior to the late Carboniferous, during the latter stages of Pangea formation, this region was accretionary as continental fragments collided with the eastern margins and caused regional tectonic shortening (Veevers et al. 1994). In contrast, the late Carboniferous and earliest Permian were marked by overall extension due to thermal subsidence and rifting (Veevers et al. 1994).

Compression returned with collision between a volcanic arc and eastern Australia during the Sakmarian but did not affect Tasmania until the Artinskian. Shortening continued throughout the region until the Triassic, when Pangea began to break

apart (Veevers et al. 1994).

15 Climate

Pangean glaciation in the Southern Hemisphere began in the middle

Carboniferous (Frakes et al. 1992) and continued with several interglacial-glacial

cycles until the mid-Permian (Frakes et al. 1992; Isbell et al. 2003; Jones and

Fielding 2004). Recent evidence from Australia and Antarctica suggests that late

Paleozoic glaciation was primarily in the form of montane glaciers rather than

large continental ice sheets, and that the most extensive glacial peak occurred

during the latest Carboniferous-Asselian (Dickins 1996, Isbell et al. 2003).

Southern Pangean climates became progressively warmer as Pangea

moved northward throughout the Permian (Ziegler et al. 1997). This climatic

amelioration resulted in the growth of Glossopteris forests, ferns, and horsetails in Antarctica (Ziegler et al. 1997) and cool-temperate swamps in Australia

(Lindsay 1997) by the middle to late Permian. Warming also caused melting of glaciers and concurrent global sea-level rise in the Sakmarian (Dickins 1996).

Weather patterns became increasingly seasonal, dominated by strong, intense monsoons along the Tethyan margin (Crowley 1994; Gibbs et al. 2002).

STRATIGRAPHY

Chronostratigraphic boundaries within the Lower Parmeener Supergroup

are well constrained across Tasmania from both invertebrate and microflora data

(Clarke and Banks 1975). This biostratigraphy, however, is difficult to correlate

with global timescales because of species provinciality.

16 Lithostratigraphic names used in the Tasmania Basin (Fig. 2.2) are

numerous, despite overall similar facies relationships. Reid et al. (in press) describe the broad development of facies throughout the basin and propose

modifications to the lithostratigraphic nomenclature based on previously

established names. This is the format used in this study (Fig. 2.3).

Figure 2.2. Stratigraphic relationships between formations in the Lower Parmeener Supergroup (after Clarke and Forsyth 1989). The locations referred to are shown on Figure 2.4. Formation names used in this study are provided, corresponding with Reid et al. (in press). Wavy lines and hatched areas represent unconformities.

Figure 2.3. Lower Parmeener Supergroup stratigraphy for Hobart region and central Tasmania showing general stratigraphic distribution of lithofacies (modified from Clarke and Forsyth 1989). Stratigraphic position of Tasmanite oil shale is given despite its absence as a discrete bed in the Hobart region.

18 REGIONAL GEOLOGY

The pre-Permian basement geology of Tasmania can be divided into distinctive western and eastern sectors (Fig. 2.4). The boundary between the two is traditionally defined by the southern extension of the Tasman Line, represented in Tasmania by a zone of recurring crustal weakness known as the

Tamar Fracture System, which also marks the easternmost occurrence of

Precambrian rocks (Veevers et al. 1994). New gravity and seismic results, as well as recent mapping, have shown that the Tamar Fracture System more likely denotes a suture between Proterozoic ultramafic rocks in the west and Cambrian mafic and ultramafic oceanic crust in the east (Direen and Crawford 2003).

Siliciclastic and carbonate rocks were deposited across this zone during the

Cambrian, Ordovician, and Devonian, and all were subsequently intruded by late

Devonian granites (Clarke and Forsyth 1989).

Tasmania was glaciated throughout the late Pennsylvanian (Clarke and

Forsyth 1989; Dickins 1996). Ice sheets are interpreted to have flowed from the west (Fig. 2.5) and from topographic highs in the northwest, generally moving eastward in topographic lows and converging along the Tamar Fracture System.

Glaciers scoured preexisting depressions and led to the deposition of thick marine tillite and glaciolacustrine rhythmites of the Wynyard Formation in fjord- like seaways (Clarke and Forsyth 1989; Hand 1993). This sediment was onlapped by the Asselian Woody Island Formation siltstone, which contains few fossils but abundant glendonite. These fine-grained sedimentary rocks are overlain by the fossiliferous Sakmarian Bundella Formation, which includes the

19 Darlington Limestone (Fig. 2.3) in the Hobart and Maria Island region. These

Sakmarian strata are typified by abundant dropstones and the cold-water

Eurydesma fauna (Clarke and Banks 1975; Clarke and Forsyth 1989).

Figure 2.4. Tasmania Geology (from Clarke and Forsyth 1989). Drill holes studied are shown as an “X”, and important cities and outcrop names are shown. Line A-B-C-D shows the position of the cross section in Figure 2.6 (this study).

Figure 2.5. Interpreted glaciation and directions of ice movement during the late Carboniferous (after Domack et al. 1993; Hand 1993; Clarke and Forsyth 1989; Reid et al. in press). Glaciated areas are shaded gray, and unglaciated highlands are unshaded.

These glaciomarine rocks are overlain by late Sakmarian-earliest

Artinskian facies of the Liffey Group, which represent a cold-climate, coal-bearing coastal to fluvial siliciclastic succession with minimal evidence of sea ice (Martini and Banks 1989). The pronounced relief produced during glaciation was filled in by Artinskian time, and subsequent sedimentation of the upper marine unit was relatively uniform across the basin. Deposition of the Cascades Group, including the Nassau Siltstone and the Berriedale Limestone, signals renewed marine flooding. Dropstones are abundant in layers within the limestone, and the

21 Eurydesma fauna, although present, is much reduced (Clarke and Forsyth 1989).

Thin interbedded meta-bentonite layers indicate distant volcanism (Clarke and

Forsyth 1989). Sandstone turbidites, variably glauconitic sandstone, and fossiliferous siltstone characterize Kungurian to Capitanian deposition. These strata contain the last occurrences of dropstones and the cold-water Eurydesma fauna (Clarke and Forsyth 1989).

This whole succession is overlain by a 600-meter-thick upper Permian to

Triassic freshwater succession of the Upper Parmeener Supergroup, consisting of estuarine, fluvial, and alluvial sandstones and coal deposits (Clarke and Banks

1975; Clarke and Forsyth 1989). The Parmeener Supergroup was later intruded by Jurassic dolerites during rifting of Pangea (Clarke and Forsyth 1989).

METHODOLOGY

This research is based on sedimentary rocks logged from 18 drill holes across the Tasman Basin, selected for their distribution and lack of alteration

(Fig. 2.4). Drill holes near Granton, Bicheno, Tunbridge, and Eaglehawk Neck were chosen for detailed sampling because of the completeness of the sections they represent. Surface stratigraphic sections were also studied at Mount

Nassau, near Hobart, and Maria Island. Lithofacies were identified in core and confirmed using 213 thin sections.

22 LITHOFACIES

The Lower Parmeener Supergroup is composed of seven recurring lithofacies associations (attributes summarized in Table 2.1), exhibiting a general north-to-south, onshore-to-offshore polarity (Fig. 2.6). For ease of discussion the shelf system is temporally subdivided into: (1) a “segmented inner shelf”, in which antecedent glacigene topography resulted in an island-and-basin bathymetry on inboard parts of an otherwise homoclinal shelf (Asselian and Sakmarian); and (2) a “continuous inner shelf”, in which previous deposition had largely filled in topography and the surface had a smooth shoreline-to-basin transition

(Artinskian).

Diamictite, Rhythmite, and Glendonitic Siltstone Lithofacies Association

The most geographically extensive outcrop of this facies association is found at Wynyard (Fig. 2.4), where bedded diamictite, consisting of pebble- to block-size clasts in a poorly sorted silt to shale matrix, is overlain by centimeter- scale, locally contorted siltstone and sandstone layers (Fig. 2.7A). Diamictite beds range from decimeter to decameter scale, and accumulations are reported to range from absent to 450 m thick at Cygnet (Farmer 1985) and more than 550 m at Wynyard (Clarke and Farmer 1976). Beds of massive and laminated siltstone and pebbly mudstone both underlie, and are laterally equivalent to, diamictite (Hand 1993). Southwest of Wynyard the diamictite is structureless or horizontally bedded (Domack et al. 1993). Bedding surfaces are striated, there are local roches moutonnées, and some diamictite has imbricated clasts that are

23 Table 2.1. Summary of Lithofacies in the Lower Parmeener Supergroup. Lithofacies Paleo- Stratigraphic Description Fossils Associations environment Unit Diamictite, Diamictite: angular to Rare brachiopods, Glaciomarine Wynyard Fm Rhythmite, and rounded clasts of mixed bryozoans, (Hand 1993; Glendonitic lithology, striated clasts, 2- bivalves Clarke and Siltstone 50 cm diameter; faceted Forsyth 1989) grains, silt and clay matrix; locally bedded Rhythmites: sandy siltstone; Rare plant Glaciomarine and Wynyard Fm interbedded medium and fragments, fossil Glaciolacustrine coarse-grained sandstone; insects, (Hand 1993; Clarke locally contorted; current- arthropod tracks, and Forsyth ripple cross-lamination spinose 1989) acritarchs, Glaciomarine Woody Island burrows (Hand 1993; Clarke Fm, rare in Glendonitic Siltstone: Tasmanites oil and Forsyth Liffey Group unfossiliferous siltstone; shale 1989) microscopic to 3 cm glendonites; variably pyritic, rare dropstones

Cross-bedded Sandy conglomerate, coarse Rare plant Terrestrial – Liffey Group sandstone, sandstone with rare fragments braided and conglomerate, pebbles; trough cross- meandering river, mudstone, and bedding floodplain, coal Sandstone with fining and Plant fragments floodplain ponds thinning upward common (Martini and Banks sequences, cross-bedded 1989) base, well-rounded pebbles Thin sequences of well- Plant fragments, sorted fine sandstone coal alternating with organic- rich siltstone; ripple cross- stratification, climbing ripples, flaser bedding Organic-rich mudstone Plant fragments, interbedded with coal coal, root traces seams

Herringbone Herringbone cross-stratified Skolithos Tidal channels and Liffey Group cross-stratified sandstone, and massive ichnofacies, flats (Martini and sandstone and sandstone ostracodes, rare Banks 1989) bioturbated forams mudstone Bioturbated sandy mudstone Few plant interbedded with fine- fragments grained sandstone; flaser bedding present, rare ice deformation features

Pebbly Fossiliferous pebbly Brachiopods, Shoreface lag Rayner Sand sandstone sandstone to bryozoan and deposit stone conglomerate, bioturbated other skeletal fragments

24 Lithofacies Paleo- Stratigraphic Description Fossils Associations environment Unit Bioturbated Variably bioturbated Rare benthic Inner Shelf Woody Island mudstone and mudstone; ice deformation foraminifera and Fm, lower sparsely features, rare dropstones ostracodes in Bundella Fm, fossiliferous bioturbated Hickman Fm, siltstone mudstone Cascades Gp, Moderate to heavily 1-3 cm thin-shelled Malbina Fm bioturbated siltstone with brachiopods, sparse fossils and skeletal diminutive fragments; dropstones bryozoans, present ostracodes, foraminifera in siltstone Fossiliferous Moderate to heavily Abundant 3-5 cm Mid Shelf Bundella Fm, siltstone and bioturbated siltstone with thick-shelled Darlington limestone densely packed fossils brachiopods and Limestone, and skeletal fragments; small 5-8 cm Cascades Gp, dropstones common; rare Eurydesma Hickman Fm, phosphates alternating with Malbina Fm bryozoan-rich layers; rare small crinoids Rudstones and dense Common to floatstones with abundant 8-10 calcareous siltstone cm thick-shelled matrix; abundant Eurydesma and dropstones; phosphates 3-5 cm thick- common as nodules and shelled phosphatized bryozoans brachiopods, common large robust bryozoans, minor crinoids Rudstones, floatstones, and Abundant 8-10 cm grainstones with cement thick-shelled or minor lime mud matrix; bivalves and 3-5 common dropstones; cm thick-shelled phosphates common to brachiopods, abundant as nodules, common phosphatized bryozoans bryozoans, and resedimented grains crinoids, and Spiculitic limestone, plant fragments phosphatic spiculite; mud Abundant sponge or cement matrix; spicules, small preserved in phosphate foraminifera; nodules minor to common brachiopod, bryozoan, and crinoid fragments Fossiliferous Interbedded siltstone and Common Outer Shelf Deep Bay Fm sandstone and fine-grained sandstone, brachiopods and siltstone, dropstones common but bryozoans sandstone decrease in abundance turbidites offshore; discontinuous thin granule conglomerate layers Table 2.1. Summary of Lithofacies in the Lower Parmeener Supergroup.

25 striated in one or more directions (Clarke and Forsyth 1989). Fossils are rare,

although brachiopod and bryozoan fragments are present in siltstone beds

interbedded with diamictite in drill core from Ross-Quoin, Bicheno, and

Eaglehawk Neck (this study) and along the south coast (Hand 1993).

Figure 2.6. Cross section of Lower Parmeener Supergroup along line A-B-C-D (Fig. 2.3), showing drill holes at A) Bicheno, B) Tunbridge, C) Granton, and D) Eaglehawk Neck.

Figure 2.7. A) Ice-contact features: contorted sandstone beds in siltstone overlying bedded diamictites at Wynyard; camera case for scale (10 cm high). B) Small pyritized glendonite in thin section from the Liffey Group. C) Plant fragments in core from the Liffey Group, central Tasmania; scale in centimeters. D) Fossiliferous siltstone from the Bundella Formation in drill core near Granton. Scale in centimeters. E) Flaser bedding from the Liffey Group in drill core from central Tasmania. Scale in centimeters.

27 Graded rhythmites are reported from northern Tasmania (Clarke and

Forsyth 1989) and from drill core in the south (Hand 1993). The fossil insect

Psychroptilus burrettae, plants such as Botrychiopsis plantiana and Aphlebia sp., arthropod tracks of Tasmanadia twelvetreesi, and lacustrine spinose acritarchs

(Hand 1993) are present in some northern outcrops of the rhythmites.

Conversely, shallow marine trace fossils occur in southern exposures of the rhythmites (Hand 1993).

Glendonitic siltstone makes up the Woody Island Formation and part of the Liffey Group. In the Woody Island Formation the fine-grained sediment is pyritic and variably bioturbated. It contains rare dropstones, and abundant glendonite crystals and crystal arrays that can reach several centimeters in diameter. Rare climbing-ripple cross-stratification, thinly laminated siltstone, and lenticular bedding are reported from the Woody Island Formation in northern

Tasmania (Domack et al. 1993). Oil-shale beds, rich in the green alga

Tasmanites, occur approximately 30 m above the base of the Woody Island

Formation and commonly contain dropstones (Clarke and Forsyth 1989; Domack et al. 1993). Individual beds are up to 2 m thick. Glendonitic siltstone, although present, is rare in the Liffey Group (this study), and the glendonite crystals are typically only a few millimeters across (Fig. 2.7B). In this instance glendonitic siltstone is interbedded with the herringbone cross-stratified sandstone and bioturbated mudstone lithofacies association (see below).

28 Interpretation.---Diamictite is interpreted as the deposits of glaciomarine debris

flows because of their association with marine fossils (Hand 1993). Rhythmite

beds are inferred to be either glaciolacustrine or glaciomarine on the basis of

fossil identification within the beds (Clarke and Forsyth 1989; Hand 1993).

Interpretations of the glendonitic siltstone and Tasmanite oil shale have proven more controversial. Revill et al. (1994), citing the presence of fine-scale lamination and of scouring, proposed that the oil shale was shallow water in origin. Domack et al. (1993) and Domack (1995) disagreed, suggesting that the same features are evidence of distal turbidites and that coarse, pebbly deposits should be associated with nearshore environments in glaciomarine systems, but these are lacking. The geographic distribution of the Tasmanite oil shale supports a deeper-water explanation, although it is proposed here that its accumulation is restricted to isolated sub-basins on the inner shelf. Where the topography was more subdued and sedimentation rates were higher (towards the south), Tasmanites is more dispersed and the rocks are composed mostly of glendonitic siltstone.

Overall, Tasmanian glaciomarine diamictite and rhythmites are interpreted to have been deposited in a fjord-like environment, with glaciomarine siltstone first accumulating in the southeast as glacial retreat continued westward (Hand

1993), similar to marine environments in McMurdo Sound today (Bartek and

Anderson 1991). McMurdo Sound is characterized by fine-grained sediment, derived from both eolian processes off the top of glaciers and subaqueous rain-

29 out, with minor coarse debris deposited by rafting and turbidity currents originating from the glacier grounding line (Bartek and Anderson 1991).

Cross-Bedded Sandstone, Mudstone, and Coal Lithofacies Association

Martini and Banks (1989) have documented this lithofacies association in detail, and their work provides a contextual basis for the limited sections observed in the drill core in this study, as well as providing detail for associated lithofacies that were not studied.

Meter-scale upwards-fining and -thinning sandstone successions are common, and contain thin pebble layers and cross-bedding at the base that grade up into massive or planar-laminated sandstone beds and cross-laminated siltstone. Some examples of this facies association are dominated by 80- centimeter-thick sandy conglomerate beds. The sandstone and sandy conglomerate lithologies are interbedded with flaser-bedded and organic-rich siltstone (Fig. 2.7E) containing thin, discontinuous coal seams with a

Glossopteris flora (Martini and Banks 1989). This lithofacies association is found primarily in the north and northwest parts of the Tasmania Basin, with local occurrences in the northeast.

Interpretation.---These rocks are interpreted as terrestrial debris-flow conglomerate and braided-stream deposits that formed around topographic highs and graded into meandering streams, floodplains, floodplain ponds, and crevasse splays in the valleys (Martini and Banks 1989).

30 Herring-bone Cross-stratified Sandstone and Bioturbated Mudstone Lithofacies

Association

This lithofacies association also occurs in the Liffey Group but is

characterized by massive, graded, cross-stratified and herringbone cross-

stratified fine- to medium-grained sandstone beds, ranging from 5 cm to 1 m

thick. Individual sets of tabular and trough cross-stratification are 1 to 5 cm thick.

Sandstone beds are interbedded with bioturbated mudstone beds 2 to 50 cm

thick. Laminated mudstone, mudstone drapes, paired mudstone drapes, and

flaser bedding are associated with herringbone cross-stratified sandstone that

contains some pebbly laminae. These beds exhibit small flame structures and

rare ball-and-pillow structures. Interbedded sandy mudstone is heavily

bioturbated and commonly contains small Skolithos burrows and local Chondrites

burrows. Plant fragments up to 5 cm long are present on sandstone bedding

surfaces (Fig. 2.7C) and within locally pyritic, carbonaceous mudstone.

Centimeter-scale contorted bedding is also present where thin sandstone layers

are interbedded with mudstone. Convolutions are 3 to 10 cm high and have the appearance of jagged, sharp-crested anticlines with broad synclines that are separated by several decimeters of offset.

Interpretation.---These facies have been interpreted as peritidal deposits that include intertidal channels and bioturbated mudstone tidal flat facies, and deltaic facies (Martini and Banks 1989). Herringbone cross-stratified sandstone and

Skolithos burrows indicate deposition of shifting subaqueous dunes along a

31 coastline with tidal influences (Gerdes et al. 2000). This lithofacies association is

interbedded with bioturbated mudstone and poorly fossiliferous siltstone from the

bioturbated mudstone and fossiliferous siltstone and sandstone lithofacies

association (described below). The presence of Chondrites burrows indicates

high organic content in the sediment. This could be the result of organic input

near river mouths (Löwemark et al. 2004), and combined with the abundance of

load features may suggest a deltaic environment. The contorted bedding and

load structures are similar to sedimentary structures produced by shore-fast ice

on tidal flats of the St. Lawrence (Dionne 1998). These features could also be

attributed to several other mechanisms, including, but not limited to, frost action,

water escape, and pressure changes from passing storm waves.

Pebbly Sandstone Lithofacies Association

The Rayner Sandstone is a 1-meter-thick, poorly sorted coarse to pebbly

sandstone that overlies the herringbone cross-stratified sandstone and

bioturbated mudstone lithofacies association and penecontemporaneous marine

units from the lower Hickman Formation. It is variably fossiliferous, mainly

containing brachiopods, bryozoans, and skeletal fragments, and also contains rounded pebbles ranging from 1 to 5 cm in diameter. This bed is interpreted here to be a shoreface lag deposit, similar to the shoreline described by Mack et al.

(2003) from the Permian of New Mexico.

32 Bioturbated Mudstone and Poorly Fossiliferous Siltstone Lithofacies Association

Bioturbated mudstone and poorly fossiliferous siltstone form the base of

the Bundella Formation, are interbedded with sandstone in the Liffey Group and

correlative marine Hickman Formation, and dominate the Nassau Siltstone (Fig.

2.2). These facies are relatively thin across the Tasmania Basin and are most

extensive in the central and northern regions. In the north, buff to light brown,

massive, sandy siltstone up to 15 m thick, containing pebbles and rare small brachiopods and bryozoans, grades southward into finer-grained facies. These southern facies consist of dark gray, massive, heavily bioturbated siltstone and

mudstone ranging from 5 to 25 m thick. The marine biota is dominated by

ostracodes and subhorizontal burrows in the central and north part of the basin, and small, thin-shelled brachiopods and diminutive bryozoans in the south.

Interpretation.---This lithofacies association is interpreted here as a nearshore

marine deposit. The ostracode-dominated bioturbated mudstone in the north

suggests a stressed environment (Curry 1999). The stratigraphic proximity to

fluvial facies is interpreted to indicate a freshwater influx sufficient to create

nearshore hyposaline conditions that decreased in importance offshore. The

bioturbated mudstone grades upward and southward into the poorly fossiliferous

siltstone, the latter of which is clearly marine in origin, although the small size of the fossils is again interpreted to reflect environmental stress, probably limited hyposalinity.

33 Fossiliferous Siltstone and Limestone Lithofacies Association

Medium to dark gray fossiliferous siltstone (Fig. 2.7D, 2.8B,C) passes

eastward and southward into dark gray limestone with a calcareous silt matrix

(Fig. 2.8A, D), medium gray to buff pure limestone (Fig. 2.8E, F, G), and spiculitic

limestone (Fig. 2.9).

Fossiliferous Siltstone.---Fossiliferous siltstone occurs in the Bundella

Formation, the Berriedale Limestone, and parts of the Nassau Siltstone. It

consists of moderately to densely packed fossils in a matrix of massive to heavily

bioturbated siltstone to pebbly or sandy siltstone with a floatstone to rudstone

texture. The siltstone is relatively homogeneous through 20 m-thick sections of

drill core, with variations in fossil abundance and biota on a centimeter to meter-

scale. Fossils are predominantly brachiopods and bryozoans that are larger and

more robust than in the poorly fossiliferous siltstone lithofacies, and Eurydesma

bivalves are also locally present. Brachiopods do appear to be aligned in death

position, but bryozoans (Fig. 2.8B,C) are in all cases horizontally oriented and locally folded. Rare small crinoid ossicles are present in distinct horizons only a few centimeters thick.

Figure 2.8. A) Outcrop at the Fossil Cliffs on Maria Island showing Bundella siltstone (1, 2, 3) and Darlington Limestone (4, 5). Layer 1 is rich in large dropstone boulders crushing underlying layers and bryozoans. Interbedded fossiliferous siltstone and argillaceous limestone (2) are cut by sandy brachiopod bryozoan rudstone channels (3). These beds are interbedded with Eurydesma shoals (4) up to 3 meters thick and are overlain by fossiliferous sandy siltstone beds (5). B) Large foliaceous trepostome bryozoan from the Bundella Formation on Maria Island. Scale in centimeters. C) Large dendroid trepostome bryozoan from the Bundella Formation on Maria Island. Scale in centimeters. D) Close-up of Eurydesma rudstone in Darlington Limestone showing barnacle borings. E) Berriedale Limestone at Rathbone Quarry near Hobart with conspicuous dropstones (arrows). F) Phosphate (dark areas) and brachiopods from the Berriedale Limestone at Rathbone Quarry. G) Phosphatized trepostome bryozoan in Berriedale Limestone from Tolosa Street Quarry near Hobart.

35 Argillaceous Limestone.---Argillaceous limestone, in both the Darlington and

the Berriedale, consists of brachiopod-bivalve rudstone and densely packed

bryozoan floatstone with a calcisiltite matrix, and forms beds a few centimeters to

several meters thick. On Maria Island and a few other isolated localities, outcrops of the Sakmarian Bundella Formation contain Eurydesma bivalves,

typically bored (Fig. 2.8D) by barnacles (Reid et al. in press), that form beds up

to 3 meters thick and have large-scale cross-bedding (Fig. 2.8A). A few brachiopods and pectens are also present in such beds. Paleocurrent measurements of cross-bedding and oriented Eurydesma shells in beds from

Maria Island indicate a unidirectional flow towards the northeast (present), whereas the shells on the tops of beds show a bidirectional northwest-southeast component approximately perpendicular to the Permian shoreline (Brill 1982; this study). Due to the limited geographic and stratigraphic distribution of paleocurrent data it could not be determined whether these current directions are

local or reflect the overall paleocurrent patterns of the Tasmania Basin. Densely

packed and well-cemented brachiopod bryozoan quartzose sandy rudstone (Fig.

2.8A3) occurs as lenses within the Eurydesma beds and becomes progressively

more abundant towards the top of the Bundella Formation. Rao (1981) and Brill

(1982) report conchoidally fractured quartz in the silt matrix of the argillaceous

limestone. Dropstones are numerous and large, up to 1 meter in diameter, and

occur both isolated or in clusters (Fig. 2.8A1). These dropstones both penetrate

underlying beds and crush macrofossils. Phosphate nodules, phosphatized

36 bryozoans, and plant fragments were discovered during this study, but they are

not common.

Pure Limestone.---The argillaceous rudstone and floatstone are interbedded

with clean limestone in both the Sakmarian Bundella Formation (Darlington

Limestone) and the Artinskian Cascades Group (Berriedale Limestone) (Fig.

2.8E). Pure limestone beds are thicker and more numerous towards the east

and southeast. The clean carbonate consists of brachiopod rudstone (Fig. 2.8D),

brachiopod bryozoan rudstone, and crinoid grainstone, all of which are in beds

10 to 50 cm thick. The matrix, when present, is predominantly productid spine grainstone, although micrite exists in isolated patches, particularly associated with intragranular pore spaces. Brachiopods, bivalves, bryozoans, and crinoids are typically large and robust compared to those in the argillaceous limestone;; dropstones are present but are smaller and less abundant (Fig. 2.8E).

Phosphate is present as dark gray isolated nodules, as replaced bryozoans, and

locally as resedimented grains (Fig. 2.8F,G). Plant fragments are common to abundant in the limestone, especially in crinoidal grainstone.

Spiculitic Limestone.---Spiculitic facies have not previously been described from these rocks. They are found in both the Darlington Limestone and

Berriedale Formation, and they are mostly preserved in phosphate nodules where they form phosphatic spiculite to phosphatic foraminifera spiculite.

Spicules in nodules retain their siliceous composition and grade to spicules

37 replaced by calcite at nodule margins. Stained thin sections (Fig. 2.9) show evidence of dissolved spicules within pure limestone, replaced with different generations of calcite cement, in quantities large enough to warrant the name spiculitic limestone. Spicules are much more common than has been originally documented and form a significant part of the deep-water biota in the Tasmania

Basin.

Interpretation.---Overall, this progression from poorly fossiliferous siltstone to pure limestone reflects increasingly open marine conditions as facies graded laterally from inner shelf to outer shelf; crinoids indicate normal marine conditions, and the diversity and size of the fauna is greater than in any other lithofacies association. Shell beds with broken and abraded fossils show evidence of episodic reworking, particularly in the argillaceous limestones.

Eurydesma beds were reworked into shoals in middle-shelf environments

(Hobart and Eaglehawk Neck areas) but also formed barrier bars in intertidal areas along headlands (Maria Island) (Brill 1982). Brachiopods and bryozoans become more abundant in interpreted deeper-water facies to the southeast, similar to the distribution seen in modern cool-water environments (James 1997;

Lukasik et al. 2000). Spiculitic facies are typically associated with the coldest and deepest facies both in the Permian Arctic (Beauchamp and Baud 2002) and in modern Antarctic polar carbonate (Taviani et al. 1993).

Figure 2.9. Photomicrographs of the spiculitic limestone lithofacies, stained with Alizarin red S and potassium ferricyanide; arrows show examples of spicules. Both sections are from the Berriedale limestone at Eaglehawk Neck. A) Siliceous and replaced sponge spicules in phosphatic limestone. B) Sponge spicules replaced with late-stage calcite.

Fossiliferous Sandstone, Siltstone, and Turbiditic Sandstone Lithofacies

Association

This lithofacies association is restricted to the late Artinskian Deep Bay

Formation, and the Malbina Formation and equivalent formations (Reid et al. in

press). Only limited occurrences were encountered in the drill core used in this study. The Deep Bay Formation is confined to the southwest, and consists of light gray siltstone interbedded with buff, fine-grained sandstone, minor

39 mudstone, and thin discontinuous granule conglomerate laminae. In the

northeast, the Deep Bay is coeval with the lower part of the Malbina Formation,

as well as the thinner and glauconitic Marra Formation (Clarke and Forsyth

1989). The facies of the upper Malbina Formation, which overlie the Deep Bay

Formation in the Hobart region, are fossiliferous pebbly sandstone and siltstone

(Farmer 1985).

Interpretation.---The Deep Bay Formation and the lowermost, coeval part of the

Malbina Formation are interpreted to represent deposition by sediment gravity flows in the deepest part of the basin (Reid et al. in press). The overlying

Malbina Formation is interpreted as a shallow marine, nearshore deposit (Farmer

1985).

GEOGRAPHIC AND STRATIGRAPHIC VARIATION

The Lower Parmeener Supergroup can be divided into four time segments that reflect the progressive lateral and stratigraphic change in lithofacies associations along the shelf profile between the late Carboniferous and the middle Permian.

Phase 1: Late Carboniferous-Asselian

Sediment deposited during this time period reflects regional deglaciation

(Fig. 2.10). Local topographic relief was up to 1 kilometer, and the terrain was rugged (Clarke and Forsyth 1989; Hand 1993), particularly in the Tasmanian

Figure 2.10. A) Geographic distribution of Asselian glaciomarine facies (after Reid et al. in press). The line indicates the position of the profile. B) Shelf profile of Wynyard and Woody Island Formations during the late Carboniferous to Asselian (this study).

41 interior, and so deposition took place on a segmented inner shelf. As glaciers

retreated and sea level rose, deep depressions and U-shaped valleys were filled

by glaciomarine diamictite and local glaciolacustrine rhythmites (Wynyard

Formation) (Clarke and Forsyth 1989), with coarse conglomeratic facies shed off

topographic highs. Clasts were largely derived from northern Tasmanian and

Antarctic lithologies (Clarke and Forsyth 1989). Pyritic glaciomarine siltstone

(Woody Island Formation) containing large glendonite rosettes and scattered

dropstones onlapped these glacigene rocks across the basin. Tasmanite oil

shale accumulated in sub-basins on the inner shelf. Rare shelly biota included

Trigonotreta brachiopods and Deltopecten, Etheripecten and Eurydesma

bivalves (Clarke and Forsyth 1989).

Phase 2: Sakmarian

The inner shelf remained segmented throughout the Sakmarian (Fig.

2.11). Bioturbated mudstone and sparsely fossiliferous siltstone floored the inner shelf, grading outboard into mid-shelf fossiliferous siltstone and argillaceous

limestone. Argillaceous limestone was deposited as Eurydesma barrier bars that

were cut by channels, filled with brachiopod rudstone, in intertidal areas near

Maria Island; Eurydesma also accumulated as widespread shell lags and shoals

(Hobart and Eaglehawk Neck regions). Non-argillaceous limestone was

deposited in relative abundance only on the middle shelf near Granton and

Eaglehawk Neck and in isolated beds on Maria Island. Phosphate formed locally

near Granton in a zone transitional between the argillaceous and the pure

Figure 2.11. A) Geographic distribution of Sakmarian Bundella Formation facies (modified from Reid et al. in press). Note the limited distribution of plant fragments and phosphates in the south. The line indicates the position of the profile. B) Shelf profile of Bundella Formation and Darlington Limestone during the Sakmarian (this study).

43 limestone. Brachiopods were dominated by Trigonotreta, although Tomiopsis,

Notospirifer, Pseudosyrinx, Sulciplica, Streptorhynchus, Strophaosia, Pyramus,

Schzodus, and Myonia were also notably present (Clarke and Forsyth 1989).

Bryozoans were abundant, particularly fenestrates and trepostomes (Reid 2003).

Plant fragments accumulated in the transitional zone and in non-argillaceous limestone. Dropstones are common throughout all marine facies, but are particularly abundant within argillaceous limestone in the eastern part of the

Tasmania Basin. Some clasts originated in northern Tasmania (Brill 1982), although provenance information is not complete.

Phase 3: Upper Sakmarian to Lower Artinskian

This was a time of relatively low sea level and widespread terrestrial sedimentation. Alluvial plains prograded across the basin during the late

Sakmarian-early Artinskian, finally filling in paleotopography on the segmented inner shelf (Fig. 2.12). Plant fragments were common in thin coal seams that developed in peat bogs on the plains (Martini and Banks 1989). Herringbone cross-stratified sandstone interbedded with bioturbated mudstone formed in the intertidal to subtidal environment. Contorted bedding developed at zones of sea ice contact (cf. Dionne 1998). Glendonite crystals are small and rare, occurring only in marginal-marine mudstone from central Tasmania. Inner-shelf facies were predominantly bioturbated mudstone with ostracodes and thin-shelled brachiopods together with sparsely fossiliferous bioturbated sandstone.

Brachiopods consisted of both spririferids and productids, predominantly

Figure 2.12. A) Geographic distribution of terrestrial Liffey Group and marginal-marine Hickman Formation (modified from Reid et al. in press). This distribution reflects a marine incursion towards the northeast. Marginal-marine facies are typically restricted to the south. The line indicates the position of the profile. B) Shelf profile of the Liffey Group and Hickman Formation during the latest Sakmarian (this study).

45 Tomiopsis and Cancrinella (Clarke and Forsyth 1989). Middle-shelf facies were not developed within the study area during this phase.

Phase 4: Artinskian

The Artinskian was marked by a relative sea-level rise and a return to widespread marine conditions (Fig. 2.13). The inner shelf was no longer strongly segmented and facies were more uniformly spread across the basin. The nearshore was composed of poorly fossiliferous sand and silt with a few ostracodes and thin-shelled brachiopods. Fossiliferous calcareous silts and argillaceous carbonate graded outboard into pure carbonate that dominated the mid-shelf facies. No Eurydesma shoals or barrier bars were present, but winnowed brachiopod and bryozoan shell beds were common and crinoids became important constituents. During this time the biota was most diverse, dominated by strophalosids and linoproductids with some spiriferids (Clarke and

Forsyth 1989). Bryozoans, predominantly fenestrates and trepostomes with rare cryptostomes and cystoporates, were numerous (Reid 2003). Plant fragments and phosphate are abundant in carbonate deposits across the basin.

Dropstones are also present but are smaller than in the Sakmarian and generally confined to mid-shelf facies. Fine-sandstone and fossiliferous siltstone turbidites characterized the outer shelf.

Figure 2.13. A) Geographic distribution of Cascades Group (Berriedale Limestone and Nassau Siltstone) facies during the Artinskian (modified from Reid et al. in press). Note phosphates and plant fragments now occur across the mid-shelf. The line indicates the position of the profile. B) Shelf profile of Cascades Group and Deep Bay Formation during the late Artinskian (this study).

47 SEQUENCE STRATIGRAPHY

Special consideration is required when assessing the sequence

stratigraphy of mixed carbonate-siliciclastic systems because it is influenced by

the biota as well as by the input of terrigenous sediment (Handford and Loucks

1993; Cunningham and Collins 2002; Caron et al. 2004). In both the cool-water and the cold-water realms there are no ooid shoals and no photozoan assemblages that would produce carbonate buildups. Thus, there is nothing to absorb open ocean wave action, and so deposition is controlled by accommodation space formed by the sea floor and wave base, as in siliciclastic settings. Furthermore, the sequence stratigraphy of glacially influenced basins is additionally complex because of delayed isostatic rebound (Brookfield and

Martini, 1999; Jin et al. 2002). The Lower Parmeener Supergroup can be resolved into two sequences that roughly correspond to the old lower marine and upper marine units of Clarke and Banks (1975).

Sequence 1

The base of the first transgressive systems tract is defined by the bedrock glacial erosion surface, which marks the glacial ravinement surface and sequence boundary (Fig. 2.14). This surface is overlain by glaciomarine diamictite, rhythmites, and siltstone. In general, maximum flooding surfaces can be identified by the presence of condensed sections reflecting distant sediment sources at the height of transgression (Posamentier and Allen 1999). The maximum flooding surface is difficult to pinpoint in the Woody Island siltstone

48 because there is no single condensed bed. Instead, we suggest that a

maximum-flooding zone is represented by the thin series of condensed beds

associated with the Tasmanite oil shale within the Woody Island Formation.

Above the oil shale the siltstone grades into fossiliferous siltstone and limestone

facies, reflecting a gradual deepening in the basin. Highstand limestone

deposition typically lags behind maximum flooding (Handford and Loucks 1993)

and in the Tasmania Basin does not occur until the Sakmarian. The falling-stage systems tract is interpreted to begin near the end of the Sakmarian at the top of the Bundella Formation, where the lithofacies grade from limestone to fossiliferous siltstone to fossiliferous sandstone and indicate shallowing of the basin. Regression continued with the deposition of fluvial facies of the Liffey

Group through the end of the Sakmarian and into the earliest Artinskian.

Figure 2.14. Composite facies profile and sequence stratigraphic interpretation of Lower Parmeener Supergroup shown on an idealized NW to SE cross section. Note that the Rayner sandstone is only 1 m thick and is not portrayed on this diagram.

49 Sequence 2

The Artinskian ravinement surface and sequence boundary is herein defined by an erosional contact between alluvial deposits and marine strata in

the north and northeast of Tasmania near Bicheno, Wynyard, and Tunbridge that

is coincident with the top of the Liffey Group (Martini and Banks 1989). The base

of the Rayner Sandstone, a thin, fossiliferous pebble conglomerate lag bed,

marks the ravinement surface within marine units (cf. Mack et al. 2003). The

maximum flooding surface again is not clearly defined in Sequence 2, but it is

interpreted to lie at the top of the Nassau Siltstone between the siltstone and

limestone beds, similarly to that in Sequence 1. Above this surface there is

evidence for condensed beds, such as the phosphate-rich strata in the Berriedale

Limestone and the glauconitic sandstone in the Deep Bay Formation, that

indicate a stillstand during the highstand systems tract. Above the Deep Bay

Formation a final falling-stage systems tract is denoted by deposition of the

shallow marine Malbina Formation sandstone overlain by Abels Bay Formation

estuarine deposits (Clarke and Forsyth 1989). This sequence culminates in the

lowstand terrestrial facies of the Upper Parmeener Supergroup.

DISCUSSION

Paleoclimate and Paleoceanography

Plant Fragments.---Plant fragments are found throughout the entire deglaciation

sedimentary record of the Tasmania Basin, and their relative abundances are an

important indicator of climate. Hand (1993) reports rare plant fragments

50 associated with Asselian glaciomarine tillite and glaciolacustrine rhythmites. We also found plant fragments within the Sakmarian Bundella Formation and more

commonly in the Artinskian Cascades Group limestone. Thin coal beds are

present in the upper Sakmarian Liffey Group and thick coal beds in the Upper

Parmeener Supergroup (Clarke and Forsyth 1989). Whereas coal was best

developed in interpreted lowstand periods, there is an overall increase in the

abundance of plant material upward in the early Permian strata, reflecting the

change to a more temperate climate. Clarke and Forsyth (1989) summarize the

microflora and note that it became more diverse towards the middle and late

Permian. They also identified the macroflora as Glossopteris, Gangamopteris, and Noeggerathiopsis, which would have formed in cold-climate peat bogs similar to the modern Arctic (Hawke et al. 1999) and indicate that the Permian

Tasmanian climate became more humid and temperate later in the Permian.

Biotic Abundance and Diversity.---The relative abundance and diversity of

microflora, microfauna, and macrofauna also reflect a change in the biologically limiting factors across the shelf through time. Cold temperatures are thought to be the primary influence that caused the overall low diversity of organisms in the

Tasmania Basin during the Early Permian (Clarke and Farmer 1976). Biotic abundance and diversity are at their lowest in the Woody Island Formation and the Liffey Group (and coeval marine correlates of the Hickman Formation).

Whereas periodic sea-ice cover could have contributed to this situation during the Asselian, fresh-water influence is interpreted to have been important during

51 the late Sakmarian. Phosphate is absent from both the Woody Island Formation

and Liffey Group, implying that upwelling was minor and, therefore, the basin overall had normal marine water with relatively low trophic resources.

Water Temperature.---The Tasmanian ocean is considered to have been cold during the Permian (Ziegler et al. 1997). Oxygen isotope studies of bivalve and brachiopod shells by Rao (1981, 1983, 1988) and Rao and Green (1982) from

both the Darlington Limestone and the Berriedale Formation limestone deposits

have been used to estimate paleotemperature. Data show a strong overlap

between the two limestone units along a meteoric alteration trend, with estimated

ocean water temperatures comparable to modern Antarctic water near 0°C (cf.

Taviani et al. 1993).

These cold marine temperatures are confirmed by the occurrence of

glendonite, pseudomorphs of ikaite (Suess et al. 1982), within the siltstone of the

Woody Island Formation and the Liffey Group. Ikaite, CaCO3·6H2O, is

precipitated in water temperatures near 0°C in marine and nonmarine

environments in quiet-water, organic-rich sediment (Suess et al. 1982; Jansen et

al. 1987), in mixing zones (Whiticar and Suess 1998; DeLurio and Frakes 1999;

Buchardt et al. 2001; Omelon et al. 2001), and at methane seeps (Greinert and

Derkachev 2004). The presence of phosphate and alkaline water increases the

stability of ikaite (DeLurio and Frakes 1999; Bischoff et al. 1993) and sulfate

reduction (Greinert and Derkachev 2004). Quiet-water, organic-rich sediment is

52 envisioned for the Tasmanian glendonites, with enhanced precipitation due to

increased alkalinity from decomposing organic matter.

While both of these temperature indicators suggest that the water

remained very cold throughout deposition of the Lower Parmeener Supergroup,

lack of glendonites in the limestone may indicate slightly warmer temperatures

during their deposition or may simply be a result of higher energy. Ikaite forming

in modern Arctic springs shows seasonal alteration zoning and can decompose

into a “crystal mush” at temperatures as cold as 0°C (Omelon et al. 2001). The high energy present during deposition of the Darlington and Berriedale limestones would have physically destroyed any evidence of these delicate crystals.

Dropstones and Ice Cover.---Dropstones are common in the Lower Parmeener

Supergroup, particularly associated with the limestone units. They are, however, rare in Asselian glaciomarine siltstones. Temporally, dropstones are most abundant in Sakmarian and Artinskian deposits, particularly in mid-shelf silty limestone.

These outsized clasts have always been interpreted as dropstones derived from icebergs, but there are several other mechanisms that must be ruled out, such as biological rafting and sea ice (Bennett et al. 1996). While tree- borne clasts up to 3 m have been reported (Bennett et al. 1996), outsized clasts transported by plants are typically 3 to 100 mm in size, even in the Amazon, where trees are quite large (Doublet and Garcia 2004). Glossopterids were the

53 size of small shrubs or small trees (Taylor and Taylor 1993), so it is doubtful that

they were capable of transporting the large dropstones found in the Lower

Parmeener Supergroup. Transportation by sea ice is more difficult to distinguish from iceberg rafting, although the large size, striations, and abundance of clasts within the Lower Parmeener Supergroup are more consistent with icebergs (cf.

Bennett et al. 1996; Dowdeswell et al. 1998).

The Asselian should have been a time of abundant icebergs and possibly

sea ice because it followed a period of glaciation, yet dropstones are rare,

implying that there must have been a mechanism to minimize the occurrence of

icebergs in the Tasmania Basin. Sikussaks, thick accumulations of sea ice,

described from modern Greenland fjords are effective at trapping icebergs

(Syvitski et al. 1996) and such features may have been present in Tasmania.

Sea ice serves to slow the deposition of coarse iceberg debris and promotes the

deposition of the silt and sand fraction. Development of sikussaks is generally

found along grounding lines and in embayments (Syvitski et al. 1996), so

segmentation of the Tasmanian inner shelf may have been important in this

regard.

The occurrence of Tasmanites-rich beds in the upper Carboniferous lower

Woody Island Formation is also interpreted to indicate the presence of sea-ice

cover. Revill et al. (1994) suggest that Tasmanites, a green alga, inhabited sea

ice, similarly to living relatives. The Tasmanite beds, however, are also

associated with dropstones (Domack et al. 1993). The inferred presence of

icebergs suggests melting of sea ice in the Tasmania Basin in order to allow the

54 icebergs to enter the basin. The melting of the sea ice would have released the

Tasmanites and allowed it to accumulate on the seafloor. For this interpretation to be plausible the Tasmania Basin had to be a stratified water mass so that the algae accumulations could be preserved (Palliani et al. 2002). Seawater stratification is supported by the abundance of pyrite in the siltstone, indicating reducing bottom waters, and generally no benthic fauna. Development of a fresh-water lens similar to situations in some Holocene fjords (Tamburni et al.

2002; Dallimore et al. 2005), combined with periodic sea ice and the segmentation of the inner shelf, could have inhibited circulation (Ogorodov et al.

2005) and caused stratification of the water column.

The Sakmarian Bundella Formation environment is interpreted to have been similar to the modern Antarctic and Arctic, dominated by icebergs and possibly seasonal sea-ice cover (cf. Syvitski et al. 1996; Stemmerik 2000). The sikkusak would have disappeared by this time, allowing icebergs to move freely through the basin, although segmentation of the inner shelf probably prevented icebergs from entering some sub-basins.

Dropstones are generally less numerous, smaller, and more widespread across the ancient shelf in the Artinskian (Cascades Group and Deep Bay

Formation), and there is no direct evidence of a nearby glacial source. This situation is interpreted to indicate periodic movement of icebergs along the coast and into the basin, similar to conditions in the Holocene in the modern north

Atlantic (Dowdeswell et al. 1998), where interglacial periods experience only sparse, seasonal iceberg and sea-ice rafting derived from distant glaciers.

55 Salinity.---Salinity stress may have been important during deposition of the

Lower Parmeener Supergroup. Sea ice is interpreted to have covered much of

the Tasmania Basin during the Asselian. Such freezing of ocean water would

have increased water salinity on the inner shelf, possibly promoting glendonite

formation (Ogorodov et al. 2005; Papadimitriou et al. 2003), although this was

probably important in the Tasmania Basin only during the Asselian.

During the sea-ice-free periods of the Sakmarian and Artinskian (Phases 2

and 4), inner-shelf sediment would have been influenced by runoff and fluvial

influx, as well as melting of icebergs. This freshwater is interpreted to have had

the combined effect of creating a lens of somewhat brackish water and a

relatively high suspended-sediment load, such as in the Kangerdlugssuaq Fjord

in Greenland (Syvitski at al. 1996). These two factors likely explain the abundance of ostracodes and the absence of other filter-feeding organisms.

There is nevertheless an overall increase in the faunal diversity during the Early

Permian as Pangea moved north and the environment became more temperate

(Lindsay 1997; Ziegler et al. 1997).

Phosphate and Upwelling.---Phosphate occurrences in sedimentary rocks are often interpreted as an indication of upwelling but may also originate from a terrestrial source. This latter interpretation is possible in the case of the

Tasmanian phosphates described here as microscopic plant fragments are found in the limestones, signifying some offshore transport. If phosphate was derived from a land, however, there should also be enhanced concentrations of other

56 terrestrially-derived nutrients, such as iron (Bruland et al. 2001). Glauconite is

also found in the Tasmania Basin, sometimes in association with phosphates, but not until the uppermost sedimentary units of the Lower Parmeener Supergroup

as the basin was filled in with marine siliciclastics. Thus, phosphate in the

limestones is interpreted here to be a result of upwelling.

Upwelling is further supported by Middle to Late Permian ocean-circulation

models, which hypothesize a strong, deep ocean current that moved northward

along the east coast of Pangea (Fig. 2.1), with upwelling thought to occur at

approximately 70° latitude (Winguth et al. 2002). This model is substantiated by the distribution of brachiopod provinces during the Early and Middle Permian (Shi and Archbold 1996). Cold-water brachiopods of the Early Permian Indoralian

Province extend from Tasmania through New Zealand, the Shan-Thai terrane, and India, and persist as a cold-water to cool-water brachiopod assemblage until the end of the Permian, indicating the existence of cold currents throughout.

Archbold (1997, 2000) postulates that these cold-water currents actually intensified and reached progressively higher latitudes between the Sakmarian and Artinskian, and this interpretation explains the persistence of cold-water marine conditions during terrestrial warming trends recorded elsewhere in

Australia (Jones et al. 2006).

During the Sakmarian Phosphate is scarce in Sequence 1 on the

Sakmarian segmented shelf, occurring only in silty carbonate of the Darlington

Limestone and sandy siltstone of the upper Bundella Formation, and even in

these units it is localized to the southwestern part of the basin near Hobart (Fig.

57 2.12). This distribution is interpreted to be the result of active, but localized,

upwelling. In contrast, the Artinskian strata have a greater amount of phosphate

across the interpreted middle shelf (Fig. 2.13). At that time the shelf was non-

segmented, and thus there was little impediment to circulation. Eurydesma

bivalves were present but their shells were not reworked into large bars as they

were at Maria Island in the Sakmarian. This is possibly because of decreased

bivalve abundance or increased wave activity across the shelf due to reduced

bathymetric complexity. Crinoids are also much more prevalent in the Artinskian

limestones, indicating normal marine conditions and, therefore, a well-mixed

water column. Overall, upwelling is interpreted to have been prevalent, with

open circulation carrying nutrients inboard across the middle shelf.

Circulation Patterns.--- Circulation within the Tasmania Basin supports these hypothesized global patterns. During the Asselian (Phase 1), sea ice is thought to have covered most of the Tasmania Basin, although there were relatively rare periods of open water. Prevalence of silt-size sediment, combined with the presence of glendonite and pyrite, as well as restricted biota, indicate a stratified water body and low-energy conditions. Icebergs were prevented from moving across the shelf, as inferred from the scarcity of dropstones. The most likely explanation for this within a glaciomarine setting would be to have had widespread sea ice.

Open-water conditions prevailed during the Sakmarian (Phase 2),

although the circulation of nutrient-rich currents was not able to penetrate the

58 entire basin. Phosphate deposits exist near Granton and Hobart, but nowhere in the east. Dropstones were also more abundant in the west, implying that islands blocked or redirected upwelling currents. Eurydesma shoals, while irregularly distributed, could also have influenced water movement.

Sea level dropped at the end of the Sakmarian (Phase 3), preventing deep-water currents from permeating onto the shelf. Instead, the interaction of freshwater influx from rivers, tidal currents, and wave activity governed oceanographic conditions on the shallow shelf.

Sea level rose once again during the early Artinskian (Phase 4). Irregular basin bathymetry was filled in during Phase 3 and upwelling currents were free to flow across the Tasmania Basin, as indicated by the widespread distribution of phosphate. Dropstones were also geographically more widespread, although they were smaller. Because the basin did not have as many barriers within it, icebergs were free to float across the shelf wherever it was deep enough.

Furthermore, these icebergs may have originated from outside the Tasmania

Basin, traveling northward from Antarctica.

Mechanisms of CaCO3 Deposition in Cold Water

Deposition of carbonate in cold water requires several criteria that enhance precipitation and inhibit dissolution. Cold, polar seawater is undersaturated with respect to CaCO3 in comparison to warm, tropical seawater, and thus organisms living in cold water are required to expend more energy to extract ions to build and maintain shells (Mutti and Hallock 2003). Furthermore,

59 fresh meltwater in glaciomarine settings enhances undersaturation of CaCO3,

and the corresponding deposition of terrigenous material would dilute any

carbonate that did precipitate. In the case of the Lower Parmeener Supergroup carbonate, upwelling of nutrient-rich water provided the energy resources for building shells by increasing primary productivity. Episodic storm and wave energy would have reduced the sedimentation rate in mid-shelf regions.

Besides the difficulty of initially precipitating carbonate material, shells must also be able to withstand dissolution in cold, corrosive water. The main factors controlling dissolution of biogenic carbonate, other than water chemistry, appear to be total surface area, grainsize, and microstructure (Walter and Morse,

1984; Harper 2000), such that an allochem with a smaller surface area, larger size, and simpler microstructure with large crystals is more resistant to dissolution. This is the case for large, robust brachiopods and Eurydesma bivalves that comprise the best preserved specimens in the Tasmanian limestone.

The presence of phosphate derived from upwelling water may also have been instrumental in preserving the Sakmarian Darlington and Artinskian

Berriedale limestones. Even minor amounts of phosphate can inhibit the dissolution of calcite (Berner and Morse 1974; Morse 1974), and it was clearly abundant during deposition of the Tasmanian carbonate.

60 Polar Limestone

Tasmanian polar carbonate was deposited within mid-shelf facies during

highstand conditions in the Sakmarian and Artinskian. Bryozoans, brachiopods, and crinoids dominated the non-argillaceous grainstone and rudstone deposits, indicating bottom currents that were sufficient to sustain a filter-feeding community, rework the sediment, and prevent the accumulation of mud. This setting is similar to the environments that support the formation of Holocene and modern polar carbonate in the Arctic (Andruleit et al. 1996) and Antarctic (Taviani

et al. 1993).

Sponge spicules are preserved in phosphatized nodules in the Berriedale

Limestone and the Malbina Formation and as molds in spiculitic limestone.

Siliceous sponge body fossils and spicules are frequently found in modern cold-

water environments (Taviani et al. 1993) and are used as indicators for such

environments in the rock record (Beauchamp and Baud 2002). Their relative

scarcity in the Tasmanian deposits can probably be explained by dissolution during diagenesis and reprecipitation as chert in the Darlington and Berriedale

limestones.

Open water and icebergs, as indicated by dropstones, prevailed during

periods of carbonate production. This oceanographic setting would have

supported moderate to high primary productivity, resulting from input of nutrients

from debris shed off icebergs. Several studies have recently documented the

devastation iceberg grounding has on a variety of benthic communities (Gutt and

Starmans 2001; Gerdes et al. 2003), and such grounding could have prevented

61 the establishment of a carbonate factory on the inner shelf. Filter-feeding

assemblages, similar to those recorded in the Tasmanian limestone, are particularly slow in reinhabiting such sites (Gutt and Starmans 2001; Gerdes et al. 2003). Carbonate production, therefore, was concentrated on the middle shelf, probably because the seafloor lay below iceberg grounding depths.

Carbonate production was likely promoted by the upwelling of nutrient-rich water onto the shelf. Upwelling was limited during the Sakmarian, as reflected by the restricted distribution of phosphate near Hobart and Eaglehawk Neck, and the comparatively thin nature of the Darlington Limestone away from the localized intertidal barrier-bar deposits developed at Maria Island. Artinskian upwelling and phosphate production was more extensive, mirrored by the production of abundant pure limestone across the shelf (Berriedale Limestone, in the Cascades Group). This relationship between upwelling and the formation of polar carbonate appears to be important not only in the Permian of Tasmania but also in more recent examples from the Antarctic (Taviani et al. 1993) and the

Arctic (Andruleit et al. 1996).

Comparison with Cool-Water Carbonate

Cool-water carbonate is characterized by a heterozoan biotic assemblage, which backs inorganic precipitates, green algae, and invertebrates with photosymbionts (Nelson 1988; James 1997). Molluscs, benthic foraminifera, coralline algae, and sponges typify inner-shelf facies, whereas outer-shelf facies contain abundant bryozoans, brachiopods, planktic and benthic foraminifera,

62 echinoids, barnacles, and sponges (James 1997). Permian cool-water carbonate

faunas could also contain conodonts but no calcareous plankton, inasmuch as

they had not evolved yet, and barnacles would not have been common

(Anderson 1994).

The Tasmanian cold-water assemblage is superficially similar to that of

cool-water carbonate, but with a few exceptions. In the Tasmania Basin there

are no conodonts or coralline algae in any facies. Recent recognition of sponge

spicules preserved in phosphate nodules, and as molds in limestone, indicates

that siliceous sponges were an important part of the Tasmanian polar biota. The

Eurydesma fauna is a unique addition to the Permian polar assemblage, which is

specifically a Gondwanan fauna associated with cold water (Runnegar 1979).

Besides distinctive faunal components, the Tasmanian cold-water carbonate is

associated with dropstones and glendonite indicating the presence of ice and

cold water.

CONCLUSIONS

1. Lower Parmeener Supergroup strata are interpreted to have been deposited

on a polar cold-water shelf as two depositional sequences. Partitioning of the

basin by antecedent topography during the late Carboniferous to Asselian,

and very cold seawater temperatures, are interpreted to have promoted the

accumulation of shore-fast sea ice, possibly resulting in increased salinity and

reduced mixing of the water column. Post-Asselian inner-shelf environments

are characterized by bioturbated mudstone, siltstone, and sandstone with

63 scarce shelly fossils consisting of ostracodes, benthic foraminifera, small thin-

shelled brachiopods, and diminutive bryozoans. The paucity of organisms at

the time is thought to have been a result of hyposalinity from fluvially derived

fresh water. Middle-shelf facies grade outboard from fossiliferous siltstone to

limestone to spiculitic limestone and are dominated by brachiopods,

bryozoans, crinoids, and lesser bivalves. Plant fragments, phosphate, and

dropstones are concentrated in mid-shelf facies. Outer-shelf facies are

preserved only in the Artinskian strata of the Deep Bay Formation, and are

characterized by siliciclastic turbidites and fossiliferous siltstone and

sandstone.

2. The Early Permian Tasmanian climate was cold during postglacial deposition

but became progressively more humid and temperate towards the Late

Permian. Water temperatures remained cold throughout deposition of the

Lower Parmeener Supergroup, although surficial ice conditions changed.

Two paleoceanographic states are interpreted to have existed during the

early Permian: (1) sea-ice-covered, and (2) open-water with icebergs.

Siltstone deposition beneath sea ice characterized the Asselian, whereas

mixed siliciclastic-carbonate deposition occurred during the Sakmarian and

Artinskian, when open water and iceberg conditions prevailed. Possible

hypersaline conditions developed beneath sea ice during the Asselian,

although most of the early Permian experienced a gradient of hyposaline to

normal marine conditions offshore as a result of fluvial freshwater influx.

Upwelling, as demonstrated by phosphate deposition, was important on mid-

shelf facies during periods of open water during the Sakmarian and

Artinskian. Upwelling was most widespread through the basin in the

Artinskian, when islands on the shelf did not obstruct circulation.

3. Middle-shelf Tasmanian limestone was deposited under open-water

conditions with icebergs during highstands. Carbonate production was

enhanced below the iceberg grounding line by upwelling of nutrient-rich water

and by strong currents.

4. Tasmanian polar carbonate is characterized by a heterozoan assemblage

similar to that of cool-water carbonate. In addition, this polar carbonate

contains the Gondwana Eurydesma fauna, sponge spicules, glendonites, and

dropstones. Conodonts and coralline algae are absent.

ACKNOWLEDGMENTS

This research was supported by the Natural Sciences and Research Council

of Canada through a discovery grant to N.P.J. and post-graduate scholarship to student B. R. The authors wish to especially thank Clive Calver and Max Banks, who introduced us to the region, helped access and interpret cores and outcrop, and advised on Tasmanian geology. Mike Jacobson assisted throughout the core study, and Tom Hamilton provided field and laboratory help.

65 CHAPTER 3:

DIAGENESIS OF EARLY PERMIAN HIGH-LATITUDE LIMESTONES, LOWER

PARMEENER SUPERGROUP, TASMANIA

Rogala, B., James, N.P., and Calver, C.R.

ABSTRACT

The Darlington (Sakmarian) and Berriedale (Artinskian) limestones are neritic deposits that accumulated in high-latitude environments along the southeastern margin of Pangea in what is now Tasmania. These rocks underwent a complicated series of diagenetic processes that began in the marine paleoenvironment, continued during rapid burial, and were profoundly modified by alteration associated with the intrusion of Mesozoic igneous rocks. Marine diagenesis was important but contradictory; although dissolution took place there was also coeval precipitation of fibrous calcite cement, phosphate, and glauconite, as well as calcitization of aragonite shells. These processes are interpreted to have been promoted by mixing of shelf and deep ocean waters and enabled by microbial degradation of organic matter. In contrast to warm-water carbonates where meteoric diagenesis is important, the Darlington and

Berriedale limestones were largely unaffected in this environment. Only minor dissolution and local cementation took place, although mechanical compaction was ubiquitous. Correlation with burial history curves indicates that chemical compaction became important as burial depths exceeded 150 m, promoting precipitation of extensive ferroan calcite. This resulted from burial by rapidly

66 deposited, overlying, thick, late Permian and Triassic terrestrial sediments. This

common diagenetic pathway was, however, complicated by the subsequent intrusion of massive Mesozoic diabases and associated diagenetic fluids.

Finally, fractures associated with Cretaceous uplift were filled with late stage non-

ferroan calcite cement. This diagenetic history confirms that both carbonate

dissolution and precipitation occurred in this high-latitude marine

paleoenvironment. The cold-water diagenetic realm was not simply destructive

in terms of diagenesis. Furthermore, it appears that for the early Permian of

southern Pangea at least, there was no real difference in the diagenetic

pathways taken by cool- and cold-water carbonates.

INTRODUCTION

Early work on cool-water limestone diagenesis found that they were

dominated by burial processes (Nelson et al. 1988; Reeckman 1988), in high

contrast to the prevalence of marine cementation in their warm-water

counterparts (Bathurst 1975; Tucker and Wright 1990; James and Choquette

1990a), leading to the conclusion that cool-water environments were dominated

by dissolution. There has been considerable work since these studies,

supporting the importance of shallow burial diagenesis (Nicolaides 1995; Hood

and Nelson 1996; Caron et al. 2006) but also emphasizing the influence of

marine diagenesis in these environments (Nelson and James 2000; James et al.

2005; Dix and Nelson 2006; Rivers et al. 2007a; Rivers et al. 2007b). Nearly all

of these studies have been done on Cenozoic limestones, while little work has

67 examined diagenesis in older carbonates (Draper 1988; Coniglio and William-

Jones 1992; Philiip and Gari 2005). Further, there have been even fewer studies on high-latitude cold-water carbonates (Rao 1981).

Cold-water carbonates are composed of a heterozoan biota (James 1997) that also may contain ice-rafted debris and glendonites (Taviani et al. 1993;

Andruleit et al. 1996; Rogala et al. 2007). These sediments are typically deposited in water temperatures below ~5°C on high-latitude shelves. Such oceanographic conditions are not generally thought to be conducive to synsedimentary cement formation because the increase in solubility of CO2 at

2 low water temperatures leads to decreased concentrations of CO3 ¯, which in turn inhibits CaCO3 precipitation and promotes dissolution.

The Permian Lower Parmeener Supergroup in Tasmania, a 500-900 m succession of marine and terrigenous sedimentary rocks that was deposited at a high paleolatitude (Fig. 3.1), provides an opportunity to investigate the nature of cementation of cold-water carbonates in the rock record. These marine lithologies consist of siliciclastic and carbonate rocks that have long been interpreted as cold-water in origin (Banks 1957). Of particular interest are accumulations of bioclastic rudstone and grainstone in the Early Permian

Darlington Limestone (Sakmarian) and Berriedale Limestone (Artinskian) (Fig.

3.2).

Figure 3.1. Reconstruction of eastern Pangea from Li and Powell (2001) with the location of the Tasmania Basin shown by the rectangle. Present day continents are shaded dark grey and the extent of Permian shelves is shown in light grey.

The purpose of this study is two-fold. First, cement stratigraphy is used to

interpret the diagenetic environments represented in the Lower Parmeener

Supergroup limestones, and to link these cement stages to specific geologic

events in the Tasmania Basin. Second, the cementation style is compared to

that of Cenozoic cool-water carbonates. It has been suggested that there is a

recognizable difference in the modes of early cementation of warm- and cool-

water limestones (James and Bone 1989; Dodd and Nelson 1998). It is

postulated that dominance of calcitic components in non-tropical carbonates

makes such limestones resistant to early dissolution, delays cementation, and

thus results in considerable compaction (Hood and Nelson 1996; Dodd and

Nelson 1998). The place of cold-water carbonates in this continuum is examined

in light of new cement interpretations in the Lower Parmeener Supergroup.

Figure 3.2. Position of the Darlington and Berriedale limestones within the Lower Parmeener Supergroup stratigraphy. International age relationships are from Clarke and Forsyth (1989).

70 GEOLOGIC SETTING

Tasmania was part of the southern Pangean supercontinent during the late Paleozoic to early Mesozoic and was located between Antarctica and mainland Australia (Fig. 3.1). Between the late Carboniferous and middle

Permian it migrated northward from approximately 80° S to 70° S (Li and Powell

2001). While initially an extensional tectonic setting, the Tasmania Basin underwent compression during the late Sakmarian and Artinskian as a result of collision between Tasmania and a volcanic arc complex (Veevers et al. 1994).

Regional shortening continued until the Triassic when Pangea began to fragment.

Tasmania was periodically glaciated during the late Carboniferous and early Permian (Clarke and Forsyth 1989; Dickins 1996), as was much of southern Pangea (Frakes et al. 1994; Isbell et al. 2003; Jones and Fielding

2004). Climate warmed as Pangea moved northward (Lindsay 1997; Ziegler et al. 1997). This was recorded in the Tasmania Basin by an increase in bioclastic carbonate production and a decrease in the amount of ice-rafted debris as the glaciomarine environment changed from an Antarctica-style to a north Atlantic- style setting (Rogala et al. 2007). This trend continued throughout the Permian, as the coverage of Glossopteris forests (Ziegler et al. 1997) and cool-temperate swamps (Lindsay 1997) expanded across southern Pangea.

71 REGIONAL GEOLOGY AND STRATIGRAPHY

Precambrian to Devonian Basement

The Precambrian to early Cambrian geology of Tasmania comprises an

eastern and western sector, the boundary between which is a zone of recurring

crustal weakness known as the Tamar Fracture System (Fig. 3.3) (Veevers et al.

1994). This zone is a suture between Proterozoic continental crust and ultramafic rocks in the west and Cambrian mafic and ultramafic oceanic crust in

the east (Direen and Crawford 2003). Subduction was focused along this zone, but was subsequently covered by siliciclastic and carbonate rocks during the late

Cambrian, Ordovician, Silurian and early Devonian (Clarke and Forsyth 1989).

These sediments were then extensively folded in the early to middle Devonian

Tabberabberan Orogeny and intruded by middle to late Devonian granites

(Williams 1989).

Figure 3.3. Regional geology of Tasmania based on Tasmanian Department of Mines 1:500 000 Geological Map (1983).

73 Upper Carboniferous to Upper Permian (Lower Parmeener Supergroup)

This study focuses on the Lower Parmeener Supergroup, and uses the

stratigraphic nomenclature suggested by Reid et al. (in press).

Chronostratigraphy is well constrained by both microfloral and marine

invertebrate data (Clarke and Banks 1975; Clarke and Farmer 1976).

Glacial ice covered Tasmania throughout much of the late Pennsylvanian

(Clarke and Forsyth 1989; Dickins 1996), flowing largely from the west and from

topographic highs to create a glacially scoured landscape. Resulting

depressions were subsequently flooded by the ocean during early Permian

deglaciation. These fjord-like seaways were in turn sites of thick marine till and

glaciolacustrine rhythmite deposition (Clarke and Forsyth 1989; Hand 1993).

During the earliest Permian (Asselian) the Tasmanian Basin was a polar

cold-water shelf (Rogala et al. 2007) floored by basin-wide, glendonite-rich,

poorly fossiliferous silts of the Woody Island Formation. The region of the inner

shelf had been extensively scoured during glaciation. This irregular topography

formed semi-restricted sub-basins during marine transgression where alga-rich mud horizons could be preserved (Domack et al. 1993). Fossiliferous silts

(Bundella Formation) and predominantly argillaceous carbonates (Darlington

Limestone), both with abundant out-sized clasts, were deposited during the

Sakmarian, although the inner shelf was still segmented (Rogala et al. 2007).

This partitioning impeded circulation of up-welled nutrient-rich water, thus limiting the food supply to benthic invertebrates and restricting the distribution of

carbonate sediments. Sakmarian inner shelf sediments consisted of bioturbated

74 muds and poorly fossiliferous silts that graded outboard into mid-shelf fossiliferous silts, and argillaceous and pure bioclastic carbonates. A relative sea-level fall at the end of the Sakmarian resulted in terrestrial sediment progradation across the shelf (Clarke and Forsyth 1989; Martini and Banks

1989).

Renewed sea-level rise at the beginning of the Artinskian led to widespread marine conditions. Inner shelf sediments were poorly fossiliferous sands and silts containing rare ostracodes and thin-shelled brachiopods. These sediments again graded progressively outboard into fossiliferous siltstone, argillaceous limestone, and pure limestone in mid-shelf locations (Rogala et al.

2007). Fine-grained sand and fossiliferous silt turbidites characterized the outer shelf. Relative sea-level once again fell during the Kazanian resulting in shallow marine sand progradation.

Upper Permian to Upper Triassic (Upper Parmeener Supergroup)

Upper Parmeener Supergroup sedimentary rocks, deposited between the mid-Permian and late Triassic, record a change to terrestrial, freshwater deposition as sea level continued to fall. Thin coal beds and carbonaceous silt were deposited in lowland regions across the former shelf, the thickest accumulations of which were in northern Tasmania (Forsyth 1989). Fluvial and deltaic sand dominated highland areas, with volcaniclastic material becoming an important component in the late Triassic (Forsyth 1989).

75 Jurassic to Present

The Jurassic of Tasmania is characterized by an extensional tectonic

regime associated with the break-up of Pangea. Diabases (Fig. 3.3) up to 600 m

thick were intruded into Parmeener Supergroup rocks and spread laterally

parallel to bedding planes (Hergt and Brauns 2001; Leaman 2002). Minor

basalts, geochemically equivalent to the diabases, were also extruded in

southern Tasmania. Seafloor spreading between Tasmania, mainland Australia,

Antarctica, and New Zealand began in the Cretaceous (Williams 1989). This

event was accompanied by faulting, uplift, and syenite intrusion. Prolonged

extension between northern Tasmania and mainland Australia created several

extensional basins, including the upper Cretaceous to Miocene Otway, Sorell,

and Bass basins to the north, and the Derwent and Coal River grabens in the

southeast of Tasmania (Williams 1989). These rocks are overlain by Quaternary glacial sediments.

METHODS

This study is based on information from outcrop and 18 cored drillholes

across the Tasmania Basin, selected for their distribution and lack of alteration.

Drill holes near Granton, Bicheno, Tunbridge, and Eaglehawk Neck were chosen

for detailed sampling because of their completeness. A total of 245 thin sections

were cut from core samples and from surface stratigraphic sections at Mt.

Nassau, near Hobart, and Maria Island (Fig. 3.2). Thirty thin sections were

stained using Alizarin Red S and Potassium Ferricyanide (Dickson 1966), and 15

76 thin sections were viewed under cathode luminescence to examine cements in more detail. A further 6 thin sections were examined under a scanning electron microscope.

MICROFACIES

This study focuses on thin sections of fossiliferous siltstone and limestone lithofacies associations from the Darlington and Berriedale limestones. A detailed description of these associations, and others from the Lower Parmeener

Supergroup, are provided by Rogala et al. (2007).

Fossiliferous siltstone facies contain moderately to densely packed macrofossils in a matrix of massive to heavily bioturbated siltstone. Fossils are predominantly whole robust productid and spiriferid brachiopods and bryozoans, with sporadic Eurydesma bivalves. Small crinoids are present in locally distinct horizons, and gastropods occur sporadically in sandy facies. Dissolution seams, microstylolites and fitted fabrics are ubiquitous (Fig. 3.4A), obscuring any original features on shell margins. Pore-filling calcite cement is rare.

Argillaceous limestone facies consist of brachiopod/bivalve rudstone and densely packed bryozoan floatstone with an organic-rich calcisiltite matrix (Fig.

3.4B). Eurydesma bivalves are common at some localities. More robust shells are frequently intact, but thinner ones are fragmented, fractured, and abraded.

Both brachiopods and bivalves are penetrated by 1-2 mm diameter borings, consistent with the shape of acrothoracid barnacle borings (Reid et al. in press).

Dissolution seams are typical at siliciclastic grain and allochem contacts, and to a

77 lesser degree between carbonate allochems. Sharp-edged pebbles of glacial origin are scattered throughout. Such pebbles tend to be embayed within allochems or otherwise have localized dissolution seams between such pebbles and bioclasts. Phosphate nodules and phosphatized bryozoans, which are both composed of hydroxyl-apatite, and plant fragments are present in this facies, but are not major components.

Clean non-argillaceous carbonates are brachiopod rudstone and grainstone (Fig. 3.4C), brachiopod-bryozoan rudstone, crinoid grainstone, and spiculitic wackestone. Matrix, when present, is predominantly productid spine grainstone, although micrite occurs in isolated patches, usually associated with intragranular pore spaces. Brachiopods, bivalves, bryozoans, and crinoids are larger and more robust than those in argillaceous limestone. These fossils are similarly intact when large, but fragmentation, fracturing, and abrasion are nevertheless common. Brachiopods and bivalves are usually bored by barnacles. Embayed, dissolution seam, or microstylolitic contacts between allochems are typical. Dropstones are less numerous than in the argillaceous limestone facies, but likewise have embayed or dissolution seam contacts with allochems. Phosphate, present as isolated nodules or replaced bryozoans, and plant fragments are also more abundant than in other facies.

Figure 3.4. Thin-section photomicrographs, plane-polarized light. A) Darlington Limestone bryozoan siltstone with microstylolites and dissolution seams. B) Darlington Limestone argillaceous bryozoan brachiopod rudstone. Allochems have microstylolitic and embayed grain contacts. C) Berriedale Limestone brachiopod crinoid grainstone.

79 DIAGENESIS

Calcite Cements

Calcite cements, summarized in Table 3.1 and Fig. 3.5, are best

developed in grainstones and rudstones dominated by brachiopods, bryozoans, and crinoids. These facies occur in the southeast of the Tasmania Basin during the early Sakmarian (Darlington Limestone) and Artinskian (Berriedale

Limestone), with Artinskian carbonates being the thicker and more widespread of the two units (Rogala et al. 2007). Minor cementation also occurs in laterally

equivalent fossiliferous siltstone (Bundella and Berriedale formations).

Table 3.1. Summary of calcite cement properties. Cement Environment Inclusions Iron CL C1 Isopachous Seafloor Yes No Bright yellow fibrous-bladed C2 Small bladed Phreatic Yes No Bright yellow to prismatic red C3 Isopachous Phreatic No Minor Fe-rich Dull to bright red drusy zoning C4 Isopachous Shallow Burial No Fe-rich to Dull to bright equant Ferroan, red, bright yellow zoned C5 Large Shallow Burial No Fe-rich to Dull to bright isopachous Ferroan, red, bright yellow bladed zoned C6 Syntaxial Phreatic to Some in None Dull to bright red Shallow Burial early zones grading to zones Ferroan C7 Blocky Calcite Shallow Burial No Fe-rich to Bright red Ferroan C8 Vein-filling Meteoric No No Red to yellow

Figure 3.5. Carbonate cement stratigraphy. Attributes of carbonate cements are summarized in Table 3.1. The relationship between several carbonate cements (C1, C3, C4, C5, C7) are shown, with various Fe-zones designated by Z1-Z6. Shading is relative to iron content, with increased iron in darker grey tones.

Cement C1.---Isopachous fibrous-bladed calcite spar, the first phase, makes up

only a small proportion of the calcite cement. It is found rimming brachiopod

spines (Fig. 3.6A), brachiopod fragments, and fenestral bryozoan pores. Crystal

rinds are 15-30 µm thick, contain inclusions, luminesce bright yellow, and are non-ferroan. Crystal terminations are broken, corroded and overlain by marine sediment.

Figure 3.6. Thin section photomicrographs under plane polarized light. All thin sections are stained with Alizarin Red S and Potassium Ferricyanide. A) Berriedale Limestone; brachiopod spine rimmed with isopachous fibrous calcite cement (C1) and filled with phosphate (P). White arrows indicate dissolution pits formed on the seafloor. B) Echinoderm with bladed calcite (C2) and syntaxial calcite (C6), Berriedale Limestone. Isopachous equant calcite (C4) rims brachiopod fragments. C) Berriedale Limestone; isopachous drusy calcite (C3) filling a calcitized gastropod, equant (C4), and blocky (C7). D) Bryozoan with isopachous drusy (C3) and equant (C4) calcite spar from the Berriedale Limestone. E) Brachiopod with isopachous drusy (C3) calcite overlain by pyrite (Py) and equant calcite (C4), Berriedale Limestone. Intragranular pore space in fenestrate bryozoans is filled by phosphate (P). Blocky calcite (C7) seals intergranular space, although some areas have been dissolved afterwards and filled with megaquartz (S3). F) Quartzose brachiopod grainstone showing the compactional relationship between quartz grains and allochems, Berriedale Limestone. Voids within productid brachiopod spines are filled with phosphate (P).

82 Cement C2.---Bladed prismatic calcite, the second cement, is present on the undersides of some bioclasts in brachiopod rudstones, and locally on crinoid ossicles preceding syntaxial cement (C6) (Fig. 3.6B). Crystals are non-ferroan, are inclusion-rich, luminesce yellow to bright red, and occur in a layer 160-190

µm thick. This cement is partially dissolved in some locations.

Cement C3.---Drusy calcite forms isopachous rinds around many intraclasts and bioclasts (Fig. 3.6C, D), particularly on bryozoans and within spaces of their fenestral framework. Crystals form rinds 15-40 µm thick, contain inclusions, luminesce dull to bright red, and are typically non-ferroan. In rare instances this cement grades from non-ferroan to ferroan calcite. Drusy spar is abundant in unsilificified limestone facies and is also present in lithofacies dominated by argillaceous limestones and fossiliferous siltstones.

Cement C4.---The fourth cement is isopachous equant calcite, which develops

directly on bioclasts or on top of C3 cement (Fig 3.6C, E). Equant calcite spar is abundant. Crystals form rinds 75-375 µm thick. They are inclusion-free with the

exception of some incorporated isopachous fibrous calcite (C1) and drusy calcite

(C3) crystal fragments, and are composed of zoned ferroan calcite that typically

decreases in iron content outward. These zones have a range of luminescence

from bright yellow (rare) to bright or dull red.

83 Cement C5.---Non-parallel isopachous bladed calcite forms within larger cavities, and is laterally equivalent to, and is overlain by, equant calcite. Crystals are 280-

840 µm long, contain minor inclusions of drusy calcite (C3) fragments, and are zoned (Fig. 3.5). This cement primarily has a bright to dull red luminescence, with some zones having a bright yellow luminescence.

Cement C6.---Syntaxial calcite overgrowths are best developed in micrite-free layers and, when present, occur as the first or second cement on echinoderm ossicles (Fig. 3.6B). Overgrowths are up to 400 µm thick and have as many as three distinct Fe zones, each ranging from 15-300 µm thick. Zones progress outward from non-ferroan calcite with few inclusions to inclusion-free ferroan calcite. Zoning is not always present, and in these cases the syntaxial cement can be ferroan or non-ferroan. Individual zones have a bright or dull red luminescence.

Cement C7.---Blocky mosaic calcite spar is pervasive in grainstones and rudstones, and is the most common type in mud-dominated facies, filling sponge spicule molds and remaining inter- and intra-granular pore-spaces (Fig. 3.6). It locally contains non-ferroan calcite crystal fragments broken off of earlier cements (primarily C1), but is otherwise inclusion-free. This cement is always ferroan, and luminesces bright red.

84 Cement C8.---Fracture-filling cement occurs as inclusion-free, non-ferroan

blocky calcite and cuts across all other carbonate and silica cements. This

cement has a red to yellow luminescence.

Phosphate

Phosphate content increases upwards through the Lower Parmeener

Supergroup. During the Sakmarian there are only a few occurrences in

fossiliferous siltstones of the Bundella Formation. Artinskian phosphate is found

across the Tasmania Basin, but is particularly abundant in the central to

northeastern regions, generally increasing in amount towards the top of the

Berriedale Limestone. Phosphate remains high in the Kungurian Malbina

Formation, again especially in the northeast.

The finely crystalline hydroxyl-apatite and carbonate-apatite cement is concentrated at the bases of brachiopod bryozoan rudstone, floatstone, and grainstone beds, and in spiculitic mudstone of the Bundella and Berriedale

Formations, and in Malbina Formation sandstone. Phosphate tends to occur either as isolated nodules or as a replacement of bryozoans, and as microcrystalline francolite within intraskeletal bryozoan and foraminifera pores.

Sub-spherical nodules, 10-50 mm in diameter, and layers, 10-100 mm thick, typically have diffuse boundaries that show a gradational increase in the amount

of allochem dissolution and ferroan calcite cementation outwards. Locally the

upper margin of a nodule is sharp and overlain by a glauconite crust. In rare

instances phosphate occurs as resedimented, coarse sand-sized grains.

85 Phosphate nodules and layers preserve skeletal components that are absent

from all other microfacies, particularly sponge spicules, and are resistant to compaction, implying an early origin.

Glauconite

Glauconite is a minor component in limestone facies, although it can be a major constituent filling intergranular voids in some marine sandstones of the

Malbina Formation (Kungurian). These glauconitic sandstones, however, are

widespread across the Tasmania Basin, with occurrences in the south on Maria

Island (Clarke and Baillie 1984), southwest of Hobart (Calver et al. 2006), and

throughout the northeast (Calver et al. 1984; Turner and Calver 1987).

Glauconite is geographically more restricted in limestones, mimicking the

distribution of, and often coexisting with, phosphate. It also fills intragranular voids within small foraminifera and bryozoans in both the Darlington and

Berriedale limestones.

Silica Cement

Pervasive silicification is geographically restricted, occurring primarily in the northeast, and in the southeast around Maria Island and at Eaglehawk Neck.

Silicification primarily affects the Berriedale Limestone, although relatively minor amounts of silica cement are present in the underlying Darlington Limestone in the southeast and overlying Deep Bay Formation equivalents in the northeast.

There are three phases of silica precipitation. The first phase is

86 chalcedony (S1), which forms a 40-60 µm thick rim around bioclasts (Fig. 3.7A), and is the first silica cement lining intergranular spaces (Fig. 3.6E), completely filling small voids. The second phase, chert and microcrystalline quartz (S2), fills remaining intergranular pore spaces and molds. In rare instances there is an alternation between bands of S1 and S2. Lastly, there is pervasive megaquartz

(S3) replacement of bioclasts. In this case the entire rock is silicified with the exception of echinoderms where rimmed by syntaxial calcite (C6), isopachous drusy calcite rims (C3) around bryozoans and some brachiopods, and some pore-filling blocky ferroan calcite (C7). All of these remnant calcite cements are dissolved to varying degrees (Fig. 3.7C) prior to or coincident with silica precipitation.

Pyritization

Pyritization is rare and only occurs in geographically isolated locations. It replaces glendonites in the Woody Island Formation, and brachiopod shells in the Berriedale Limestone of the Eaglehawk Neck drill core. Pyrite crystals overlie isopachous drusy cement (C3), and occur concurrently with the most iron-rich zone of the isopachous equant calcite spar (Fig. 3.6E).

Figure 3.7. A) Plant fragment rimmed with chalcedony (S1), with pore spaces filled by chert (S2) and partially dissolved blocky calcite cement (C7). Sample is from the Darlington Limestone. B) Bryozoan in the Berriedale Limestone completely replaced by megaquartz (S3) and rimmed with remnant isopachous drusy calcite (C3). Pore spaces are filled with chert (S2). C) Blocky calcite crystal (C7) in silicified crinoid grainstone from the Berriedale Limestone. Crinoids (Ech) also have dissolved margins and are coated by chalcedony (S1) and chert (S2).

88 Neomorphism

Most components of the limestone facies are neomorphosed. Gastropods and aragonitic bivalves are now composed of non-ferroan calcite that is bright yellow or red under cathode luminescence. Some of these allochems, however, have experienced later, localized dissolution and subsequent precipitation of ferroan calcite along growth lines. Fine textural details are preserved (Fig. 3.6A), yet are crossed by wavy to gently curving intercrystalline boundaries with rare

120° triple junctions. All of these are classic features of calcitization (Bathurst

1975).

Micrite and many of the calcitic allochems, such as thin-shelled brachiopods, bivalves, foraminifera, and bryozoans, have also been neomorphosed. They are presently non-ferroan low-Mg calcite and luminesce bright yellow or red. Some of the calcite crystals in the brachiopods and bivalves have syntaxial overgrowths that extend into the surrounding microspar (cf.

Bathurst 1975). Large, robust brachiopods and Eurydesma bivalves are variably neomorphosed along shell edges and growth lines, and spines in the case of productid brachiopods. In all instances of neomorphism, the calcite is non- ferroan, is bright yellow or red under cathode luminescence, and is consistent with the characteristics of C1 calcite cement.

Dissolution

Dissolution is present at nearly every stage of diagenesis in the Darlington and Berriedale limestones. Small dissolution pits formed on the surfaces of

89 many allochems prior to precipitation of isopachous fibrous calcite cement (Fig.

3.6A). Productid brachiopod pseudo-punctae were preferentially dissolved and filled with non-ferroan drusy calcite (C3). Early non-ferroan calcite cements (C1,

C2, and C3) also show evidence of dissolution and breakage, with fragmented or corroded crystal terminations.

Total dissolution of carbonate allochems is relatively rare prior to silicification; sponge spicules, however, are an exception. These biosiliceous allochems are only preserved in phosphate nodules and some argillaceous layers. Preferential dissolution of sponge spicules took place throughout the period of ferroan calcite precipitation, as indicated by the presence of ferroan calcite from two or three iron zones (Fig. 3.5) inside different spicule molds within a single spiculite-rich layer.

Most allochem and cement dissolution, however, took place immediately prior to silicification (Fig. 3.7). Dissolution fabrics can range from patchy fabric- destructive dissolution of individual grains or cements (Fig. 3.7A), to fabric- retentive dissolution of allochems (Fig. 3.7B), to nearly complete dissolution with only a few remnant, partially dissolved calcite crystals (Fig. 3.7C).

Compaction

Most lithofacies show extensive evidence of mechanical and chemical compaction. Shells in fossiliferous siltstones, argillaceous limestone, and pure limestones have been fractured in situ, prior to the formation of C3 cement. In several instances small crystals have broken off of C1, C2, and C3 calcite.

90 These crystal fragments are randomly oriented, and poikilotopically enclosed by

later C4 and C7 ferroan calcite cements.

Dissolution seams, microstylolites, and embayed grains are pervasive throughout fossiliferous siltstone and argillaceous limestone facies, forming condensed bioclastic layers and fitted fabrics (Fig. 3.6F). Microstylolites are minor to absent in pure limestone facies, although dissolution seam and

embayed grain boundaries are common, particularly when carbonate clasts are

in contact with dropstone pebbles. Chemical compaction features are not as well

developed when extensive early calcite cements are present, indicating that

stylolitization post-dated these events.

Stylolites are also present between silicified and argillaceous layers,

signifying compaction also occurred after silicification. This does not agree with

observations of carbonate cement forming after compaction since the silica

cement was the last precipitate to form. In order to explain this discrepancy there

must have been two compaction events.

Paragenesis

Alteration of Lower Parmeener Supergroup limestones, which occurred in

numerous diagenetic environments from seafloor to burial (Fig. 3.8), can be

divided into five stages. Each stage can be related to discrete syn- and post-

depositional events (Fig. 3.9).

Figure 3.8. Paragenetic relationship of calcite, glauconite, phosphate, silica, and pyrite. Compaction and dissolution events are also shown. Dashed lines indicate intermittent formation of cements and compaction features. The shaded area indicates the timing of iron-rich calcite precipitation.

Stage 1 — Marine.---Carbonate sediments were deposited in high-energy

environments where bioclasts were quickly fragmented, abraded, and bored.

Non-ferroan isopachous fibrous calcite (C1), phosphate, and glauconite are

interpreted to have precipitated in this environment. Isopachous fibrous calcite is

typical of shallow marine environments (James and Choquette 1990a), and has

been documented in mid-Tertiary cool-water limestones (James and Bone 1992;

Nelson and James 2000). Dissolution pits on shell surfaces prior to this synsedimentary cementation indicate that dissolution was occurring on or just

below the seafloor. Marine calcite cement was also partially dissolved, although

it is unclear whether this occurred in the marine environment, in the meteoric

environment or during chemical compaction.

Figure 3.9. Diagenetic burial history curve. Diagenetic stages are superimposed on a burial history curve modified from Reid and Burrett (2004), which was based on temperatures calculated from hydrocarbon maturity and assumed a temperature gradient of 35°C/km. A) Constructive diagenetic events. B) Destructive diagenetic events

93 Calcitization and neomorphism can occur in many different diagenetic

environments including seafloor (James et al. 2005), meteoric (Mazzullo and

Bischoff 1992), and burial (Kendall 1975). In the case of the Lower Parmeener

Supergroup limestones, these processes were probably active in the marine

environment as illustrated by petrographic and geochemical similarities between

calcite in the altered allochems and the isopachous fibrous calcite cement.

Luminescence in cements such as the isopachous fibrous calcite can occur

during primary precipitation or during neomorphism in the marine diagenetic

environment, particularly in cool-water settings where slow growth rates allow

increased amounts of Mn2+ to be incorporated into the calcite lattice (Major and

Wilber 1991).

The presence of microcrystalline glauconite and phosphate within intra- and intergranular pore spaces, as nodules, and as crusts, is also characteristic of primary mineralization in condensed marine sediments (Glenn et al. 1994;

Marshall-Neill and Ruffell 2004). Phosphate-glauconite crusts, which were later ripped up in places and redeposited as phosphate clasts, indicate that hardgrounds must have formed on the seafloor during local hiatuses in siliciclastic and carbonate deposition (Carson and Crowley 1993).

Stage 2 — Meteoric Phreatic.---Only minor amounts of bladed calcite (C2),

drusy calcite (C3), and syntaxial calcite (C6) were precipitated during stage 2. All

of these cements are non-ferroan and have pyramidal terminations, indicating a

phreatic environment (James and Choquette 1990b). This cement was

94 precipitated prior to and during mechanical compaction of grains, which can occur in the first few hundred meters of burial (Goldhammer 1997).

Stage 2 diagenesis also occurred in the meteoric environment at depths

amenable to mechanical compaction. In the case of the Darlington Limestone,

these depths may have been reached by the end of the Sakmarian with

deposition of approximately 50 m of overlying Liffey Group terrestrial sediments.

Both the Darlington and Berriedale limestones were likely filled with meteoric

fluids by the late Permian after the deposition of the approximately 300 m of

Malbina and Abels Bay marine sands (Fig. 3.2), and subsequent progradation of

terrestrial sediments. In the absence of cement this overburden would have

been sufficient to fracture the more fragile grains (Goldhammer 1997).

Stage 3 — Shallow Burial.---Mechanical compaction was followed by chemical

compaction as the principal alteration process in stage 3, producing fitted fabrics

and sutured contacts in limestone facies and extensive stylolitization in

argillaceous facies. Pressure solution in argillaceous and micritic facies can

occur as shallow as 150-200m, while grainstones and packstones generally

experience significant porosity reduction due to pressure solution by depths of

300m (Goldhammer et al. 1993).

Stage 3 cements consist of ferroan bladed calcite (C4), equant calcite

(C5), syntaxial calcite (C6), and blocky calcite (C7). All of these cements

precipitated during and after chemical compaction, which is interpreted to

indicate precipitation at depths greater than 300 m. In addition, the calcite spar is

95 ferroan and has textural characteristics consistent with burial cement (cf.

Choquette and James 1990).

Bladed, equant, and blocky calcite are clear, but contain non-ferroan calcite crystal fragments. These crystal fragments have a different origin than crystal silt as defined by Dunham (1969), but may have been formed by similar processes. Crystal silt is derived from low-magnesium calcite crystals that form in intergranular pore spaces in the vadose environment and are then hydraulically fractured by percolating groundwater and redeposited as faintly laminated sediment by the movement of subsurface water in the vadose environment (Dunham 1969; Esteban and Klappa 1983; Reinhold 1998), or as massive crystal silts by gravity in the phreatic environment (Reinhold 1998).

Crystal fragments in burial cements of the Darlington and Berriedale limestones, however, are derived from both earlier marine and phreatic cements.

Some marine cement fragments are only slightly off-set from their original positions. Another important difference between these crystal fragments and crystal silt is the lack of sedimentary features. The crystal fragments float in isolation within the burial cements rather than being deposited as sediment on the floor of the void. It is suggested here that the crystal fragments were fractured, similar to the crystal silt, but remained in suspension within the pore fluid while the burial cements form around them.

Burial cements formed during at least two pulses of iron-rich pore water.

Initially, basinal fluids that precipitated isopachous equant and bladed cements were iron-rich and produced ferroan calcite before being capped by non-ferroan

96 calcite as the dissolved iron was depleted. Local pyritization and Fe-rich zoning

of syntaxial cement around crinoids was simultaneous with the precipitation of

this ferroan calcite. Siliceous sponge spicules were dissolved at this time and,

along with internal voids within foraminifers, were filled with blocky calcite cement

of varying iron concentrations. A second iron-rich pulse produced another

identical Fe-rich-to-Fe-poor calcite. Fluids that formed this cement zone also

fuelled the precipitation of blocky iron-rich calcite in the remaining pore spaces

and skeletal molds.

All of this Stage 3 diagenesis is interpreted to have resulted from loading

by Upper Parmeener Supergroup sediments. Chemical compaction depths

would certainly have been obtained with the addition of approximately 500 m of

coal-bearing terrestrial sediments (Forsyth 1989). Calcite dissolved during

chemical compaction is interpreted to have mixed with iron- and sulfur-rich

basinal brines to precipitate calcite cements C4 and C7, as well as producing

localized pyritization. Iron and sulfur may have been scavenged by groundwater

from overlying coal-beds, as is the case in modern organic-rich environments

(McBeth et al. 2003), or from underlying pyrite-bearing lithologies in the Woody

Island Formation.

Stage 4 — Silicification.---Burial calcite cementation was succeeded by extensive, localized dissolution. Silicification probably occurred simultaneously or followed soon after such dissolution because ghost textures of the original limestone grains are preserved. Late-stage iron-rich calcite cements proved to

97 be more resistant than non-ferroan calcite, preserving partially dissolved

allochems or occurring as inclusions within the silica. The intrusion of thick

diabase sills is interpreted to have led to the chemical compaction that formed

stylolites between the silicified and non-silicified layers.

There were two possible sources for silica in the Tasmania Basin. The

first would be the extensive dissolution of sponge spicules. Spicules, however,

were dissolved over a prolonged time period and were mostly gone prior to silicification, as indicated by the variable Fe concentrations of the calcite filling the spicule voids, mimicking the variable Fe content of the cements. This is most likely a result of evolving basinal brines.

A second source would have been the extensive diabases and basalts that were emplaced during Jurassic break-up of Pangea. Silicification of carbonate rocks during dike emplacement is a relatively common process

(Dutrow et al. 2001). Tasmanian diabases were part of the Ferrar continental flood basalt province and are characterized by high silica and low Fe, Ti, Na, and

P contents compared to their Mg-number (Brauns et al. 2000). Diabase feeders intruded the Lower Parmeener Supergroup across the eastern half of Tasmania and the continental shelf to the southeast (Williams 1989; Leaman 2002). Silica- rich fluids derived from these intrusions could then have been transported along the Tamar Fracture System (Fig. 3.3), which extends beneath the silicification occurrences in the northeast and southeast. Finally, the lithostatic load resulting from the diabases probably caused the late stage compaction event that post- dates silicification.

98 Stage 5 — Late stage calcite.---Stage 5 cement is characterized by large, clear, non-ferroan blocky calcite spar. This fills fractures that cut all previous

cements and allochems. These fractures are interpreted to be a result of erosion

and tectonic uplift that took place in the late Cretaceous to early Paleocene

during continued extension and rifting from mainland Australia, and lithostatic rebound following Quaternary glaciation (Williams 1989). Meteoric water would have flowed along these faults and fractures, dissolving some of the calcite and

reprecipitating it as blocky calcite cement.

Comparison between Darlington and Berriedale limestones

Cement variation between the Darlington and Bundella limestone facies

(Sakmarian) and the Berriedale Limestone (Artinskian) are interpreted to have

resulted from differences in biological components, depositional oceanography,

and burial history. Crinoids are more abundant in the Berriedale Limestone,

producing a higher proportion of syntaxial cement. There is also a change in the type of brachiopods, from predominantly spiriferids in the Sakmarian to

productids in the Artinskian. The increased number of productids resulted in an increase in the quantity brachiopod spines, producing a proliferation of grainstone facies across the basin, and thus more open-framework lithologies where cement could precipitate.

Ocean circulation patterns also changed in the Tasmania Basin between deposition of the Darlington and Berriedale limestones (Rogala et al. 2007).

Upwelling of nutrient-rich water was limited to the southeast during the

99 Sakmarian (Jones et al. 2006), but was well established and widespread across the whole basin by the Artinskian. As a result, phosphate and glauconite are most pervasive in the Berriedale (Artinskian) and Malbina (Kungurian) rocks.

Increased nutrient supply in the Artinskian also promoted the growth of benthic invertebrates, resulting in more grainy facies. The consequent accumulation of rudstones and grainstones, with their concomitant high porosities, provided pore- spaces for cement formation.

Lastly, the Darlington and Berriedale limestones had different early burial histories. Darlington carbonates were overlain by terrestrial facies at the end of the Sakmarian, placing them in the meteoric realm wherein mechanical compaction and phreatic cementation began. Berriedale carbonates, however, did not enter the meteoric environment until the Kungurian. Consequently, there was a higher percentage of meteoric phreatic isopachous drusy calcite precipitated in the Darlington Limestone compared to the Berriedale Limestone.

This early cementation, particularly in the Darlington Limestone, is interpreted to have formed a protective barrier that prevented further dissolution of marine isopachous fibrous-bladed calcite and allochems, and reduced the effect of pressure solution between allochems. The different lengths of time spent in the meteoric realm did not have a large overall effect on the diagenetic pathways, however, because the amount of cementation and dissolution that took place in this environment was minor.

100 DISCUSSION

Comparison with previous work

Rao (1981) examined carbonate cements in the Berriedale Limestone and

suggested that all of these cements were formed on the seafloor or in a mixing

zone developed from glacial melt-water entering the basin. The conclusion of a

marine origin was in part based on the preferential occurrence of length-slow

calcite formed by the accretion of rhombohedral plates and the rarity of

hexagonal prisms, which, based on the work of Kinsman and Holland (1969), were thought to indicate formation temperatures in waters below 3°C. Additional evidence included observations of micrite overlying cement, borings that cut all cement, and crushing of cement by dropstone clasts. Mixing zone spar was described as inclusion-rich calcite plates that incorporated crystals dissolved from earlier phases of cementation.

Results of this study indicate that only a minor part of the cement is marine in origin. The isopachous fibrous calcite cement (C1) is indeed overlain by micrite in places, and so is also interpreted as marine. Macro-borings are present in a variety of allochems, but were not observed in this study to cut any cement. Two features were noticed that could have been misinterpreted as borings. First, linear holes occur along preferentially dissolved pseudo-punctae, which are filled with non-ferroan calcite spar. Second, round holes were observed that initially appeared to be borings, but are in fact sponge spicule molds. This is confirmed by a gradational change from unaltered siliceous spicules outward to dissolved molds filled with various phases of ferroan calcite

101 near diffuse phosphate nodule boundaries. Furthermore, all cements except for

C1 and C2 post-date compaction. Since such compaction requires a minimum burial depth of 150m (Goldhammer et al. 1993), the majority of the cement cannot be seafloor or shallow subsurface. Quartz clasts, interpreted as small dropstones based on core and outcrop observations, are in contact with cement.

These contacts, however, are frequently embayed or stylolitic indicating that the

“crushed” appearance of the cement was a result of compaction (Fig. 3.6F).

It is unlikely that the later carbonate cements were formed in a mixing zone. Neither this study nor that of Rao (1981) observed any evidence of microdolomite crystals or inclusions within the calcite spar, which are typically observed in mixing zone environments (Frank and Lohmann 1996; Kim and Lee

2003). Other diagnostic criteria should include a finely alternating texture of luminescent and dull zones under cathode luminescence with internal truncations

(Kim and Lee 2003), and alternating zones of HMC and LMC (Frank and

Lohmann 1996; Kim and Lee 2003). Since none of these features were observed in the Lower Parmeener Supergroup limestones, a more parsimonious explanation is that the cements formed in a meteoric to burial environment.

Mechanisms for Dissolution and Cementation

Marine.---Polar marine settings are often envisaged as environments of dissolution because the solubility of CO2 at low water temperatures leads to

2 decreased concentrations of CO3 ¯, which in turn inhibits CaCO3 precipitation and promotes dissolution. This is true of the oxygen-poor, nutrient-rich Antarctic

102 Bottom Water (AABW) today (Dittert and Peng 1989), but is not necessarily a

universal trait of all cold-water masses because they can be quite variable in their

oceanographic properties. Most of the Arctic Ocean seafloor lies above the

calcite lysocline (Jutterström and Anderson 2005) and the oxygen-rich, nutrient-

poor North Atlantic Deep Water (NADW) is slightly supersaturated with respect to

calcite (Broecker and Peng 1989). This is generally thought to be due to discharge from glacio-fluvial systems that increases the alkalinity of the bottom water (Zahn et al. 1991; Huber et al. 2000; Jutterström and Anderson 2005), and

low rates of organic-matter degradation (Jutterström and Anderson 2005). Melt-

water discharge must be distal, however, because the addition of fresh water can

create undersaturated conditions (Fairchild et al. 1999).

The dominant water mass in the Permian Tasmania Basin is here interpreted to be similar to the calcite saturated NADW, although the upwelling water mass in thought to have been more akin to the AABW. Alkalinity in

Tasmanian shelf waters could easily have been increased by run-off entering the

basin (Martini and Banks 1989). Carbonate sediment deposition, however, was far enough removed from such interpreted influx that fresh water did not affect

salinity (Rogala et al. 2007), otherwise crinoids would be absent. It is probable,

then, that deposition was also beyond the influence of undersaturated conditions

created by the fresh-water mixing zone.

The idea of saturated Tasmanian seawater is also based on the quality of

the preservation in the Darlington and Berriedale limestones. Admittedly, the preservation of invertebrate skeletons is due to many factors. Dissolution of

103 biogenic carbonates can be controlled by water chemistry, mineralogy, total surface area, grain-size, and microstructure (Walter and Morse 1984; Harper

2000). The limestones are primarily composed of large, robust organisms with simple shell structures (Rogala et al. 2007), making them more resistant to dissolution. In addition, phosphate in upwelling Tasmanian seawater may have helped to inhibit calcite dissolution (cf. Berner and Morse 1974; Morse 1974).

More importantly, aragonitic gastropods and bivalves are calcitized rather than dissolved. It is proposed that seawater must have been saturated with respect to calcite, if not near aragonite saturation, to cause this.

The synsedimentary precipitation of calcite spar, neomorphism, and dissolution of allochems and spar requires a more complex explanation. In cool- water environments significant dissolution can take place in water saturated with respect to CaCO3, leading to a 40-90% taphonomic loss (Green and Aller 2001;

James et al. 2005). Precipitation of LMC cement (James et al. 2005) and

neomorphism (Kyser et al. 1998) have been observed in these same conditions.

Such processes are thought to take place in the marine subsurface wherein

pore-water microenvironments are created by microbial activity. Initially, pore

water becomes undersaturated as a result of bacterial degradation of organic

matter (Ku et al. 1999; James et al. 2005), which causes dissolution. Dissolution

2+ 2- of calcite releases Ca and CO3 into the pore-water. These ions build up until equilibrium is reached, at which point calcite precipitation and neomorphism can take place (Ku et al. 1999).

Similar microbial reactions are common in cold anoxic sediments (Arnosti

104 et al. 2005). Thus, it is suggested that microbially induced dissolution and cementation could explain the observed features in the Lower Parmeener

Supergroup limestones, particularly since organic matter is abundant in all of the

limestone facies and dissolution occurred prior to marine cement precipitation.

Dissolution, however, seems less pronounced than expected in such cold waters.

Carbonate saturation may have been buffered by decreased pCO2 as a result of

depressurization of the upwelling water mass and mixing with the shelf waters,

combined with the removal of P (Plant and House 2002) and Mg2+ (Zhang and

Dawe 2000) from the pore-water by phosphatization and glauconitization, which

would have reduced dissolution and enhanced the precipitation of marine calcite

cement.

Burial.---By the time sediments of the Darlington and Berriedale limestones

entered the burial environment (Stage 2) the vast majority of the components

were geochemically stable because they were originally or had been altered to

LMC in the marine environment. This effectively placed the limestones at a

diagenetic hiatus; there were no metastable carbonates to supply CaCO3 for cementation until they were buried deep enough for chemical compaction to begin. Mechanical compaction, however, was active during this time, fracturing both shells and marine cement crystals. It is thought that these crystal fragments did not dissolve because they were composed of LMC, and so were incorporated into later cements.

Once again there was little dissolution throughout Stage 3, with the

105 exception of most siliceous sponge spicules and minor amounts of calcite

cement. Biogenic silica dissolves readily at temperatures ~150°C in carbonate host rocks (Kastner et al. 1977). In closed systems this silica would readily recrystallize to opal CT or chert, but, as indicated by the influx of iron, the

Darlington and Berriedale limestones were diagenetically altered in an open

system. This, combined with the low amount of magnesium in the system

available to bind and precipitate the silica (Hesse 1990), created a diagenetic

environment wherein the sponge spicules were dissolved and gradually replaced

by blocky calcite cement. Silica dissolution would have created temporary,

locally acidic conditions (Hesse 1990) that could have resulted in dissolution of

previous calcite cement.

These conditions changed during Stage 4 with the intrusion of large

quantities of diabase. The abundant silica derived from these intrusions would

2- have combined with the burial fluids to form H2SiO4 (Hesse 1990), producing an

acidic environment where large amounts of calcite could be dissolved.

Comparison with cool-water carbonate diagenesis

Marine diagenesis in cool- and cold-water environments seems to be

dependent on specific circumstances rather than temperature. Nelson and

James (2000) concluded that marine cementation in mid-Tertiary cool-water

carbonates of Australia and New Zealand was enhanced by low sedimentation

rates, high energy conditions, and upwelling of marine waters. The convergence

of these features served to maintain carbonate saturation within shelf waters,

106 resulting in preservation of bioclastic material. Dissolution and precipitation of marine cement, however, may have been a result of microbially-induced pore- water reactions (cf. Ku et al. 1999). A similar set of conditions is interpreted to have occurred in the Lower Parmeener Supergroup.

Otherwise, early carbonate cements are not common in polar and non- tropical carbonates (Nelson 1988). Both the Lower Parmeener Supergroup cold- water limestones and their cool-water counterparts are composed of a heterozoan biogenic assemblage. Such assemblages are dominated by calcite mineralogies in the rock record (Nicolaides 1995; Nelson et al. 2003), although aragonite may have made up a larger proportion of cool-water carbonates when they were first deposited (Cherns and Wright 2000). Large amounts of aragonite can be present in modern cool-water carbonates, yet are dissolved or neomorphosed in the marine environment (James et al. 2005; Rivers et al. in press). Biogenic aragonite is a minor component in polar limestones, even in the modern (Taviani et al. 1993; Andruleit et al. 1996), but there is still evidence of early calcitization and neomorphism (this study). Regardless, this results in delayed cementation because little aragonite survives into the shallow burial diagenetic environment where it can be reprecipitated as calcite cement (Hood and Nelson 1996; James et al. 2005). Lack of early cement has the additional effect of making the limestones more susceptible to burial compaction

(Goldhammer et al. 1993), rendering chemical pressure dissolution a conspicuous feature in both cool-water (Nicolaides 1995; Hood and Nelson 1996) and polar limestones (this study). Despite enhanced compaction, pore spaces

107 tend to remain open to moderate burial depths where they are filled with late- stage cements (Nicolaides 1995; Hood and Nelson 1996; this study). In summary, this study suggests that cool-water and cold-water carbonates are largely the same from the perspective of diagenesis, except that polar carbonates have less aragonitic components originally.

CONCLUSION

1. Diagenesis of the Lower Parmeener Supergroup was the result of marine,

meteoric phreatic, and burial processes. Intergranular isopachous fibrous

calcite, phosphate and glauconite were precipitated in the marine

environment (Stage 1). Dissolution of allochems was also important.

Meteoric phreatic cements and dissolution (Stage 2) were not common

because the carbonate sediments had already been altered to LMC in the

marine environment and were largely non-reactive. The primary process

in this environment, therefore, is mechanical compaction. Most carbonate

cements were precipitated as ferroan calcite during subsequent shallow

burial (Stage 3). Carbonate allochems were dissolved along grain

boundaries during chemical compaction at shallow burial depths, which

provided ions for precipitation of calcite cement. Macro-scale dissolution

and silicification (Stage 4) occurred after all intergranular carbonate

cements were formed, and were most likely associated with the

emplacement of Jurassic diabases. Later uplift, likely during Cretaceous

tectonism, created fractures that were filled with non-ferroan blocky calcite

108 (Stage 5).

2. Dissolution, precipitation of isopachous fibrous calcite, and neomorphism

occurred in these Permian polar marine environments. Seawater was

most likely saturated with respect to calcite, and precipitation was

enhanced by the phosphatization and glauconitization of the sediments.

Additionally, microbial degradation of organic matter in subsurface pore

waters may have been an important contributor to cementation and

dissolution.

3. The two limestones within the Lower Parmeener Supergroup show minor

variations in their early diagenetic histories. This is, in part, due to

changes in the relative abundance of biotic components. Crinoids are

more abundant in the Berriedale Limestone; therefore, syntaxial calcite is

more common than in the Darlington Limestone. Differences in water

circulation patterns between the Sakmarian and Artinskian also produced

variation. Upwelling of nutrient-rich water was more widespread during

the Artinskian than in the Sakmarian, so early phosphate and glauconite

were abundant in the Berriedale Limestone but scarce in the Darlington

Limestone. The Darlington Limestone was exposed to meteoric

diagenetic environments for a longer period of time than the Berriedale

Limestone; thus, meteoric cements are better developed.

4. Lower Parmeener Supergroup limestones are predominantly composed of

calcitic allochems, akin to cool-water limestones. Because of this,

cementation follows similar diagenetic pathways in these cold-water and in

other cool-water limestones, therefore other criteria, such as

109 sedimentology, must be used to distinguish them from one another.

Surprisingly, dissolution does not appear to be more intense in cold-water

limestones.

ACKNOWLEDGEMENTS

This research is supported by the Natural Sciences and Engineering

Research Council of Canada through a discovery grant to N.P.J. and post- graduate student scholarship to B.R. The authors wish to thank Catherine Reid and Max Banks for consulting on the stratigraphy of the Tasmania Basin, and

Tom Hamilton for providing assistance with field and laboratory work.

110 CHAPTER 4:

GEOCHEMISTRY OF PERMIAN BRACHIOPODS AND EURYDESMID

BIVALVES FROM THE LOWER PARMEENER SUPERGROUP IN TASMANIA:

IMPLICATIONS FOR DEDUCING PALEOCEANOGRAPHIC CONDITIONS

FROM BIOGENIC CARBONATES

Rogala, B., Kyser, T.K., and James, N. P.

ABSTRACT

The stable isotopic and elemental compositions of brachiopods and bivalves from the Darlington (Sakmarian) and Berriedale (Artinskian) formations which accumulated in high-latitude neritic environments in the Tasmania Basin during the Permian deglaciation of southern Pangea would reflect the chemistry of seawater during this unique time in Earth history provided alteration has been minimal. Spiriferids and eurydesmids collected from core and outcrop of these limestones have uneven petrographic preservation, but the oxygen isotopic values are low, variable and not consistent with the depositional setting, and the shells are enriched in Fe and Mn. Therefore, all samples are altered to varying degrees. Three diagenetic trends were recognized based on isotopic data, the first two of which are only recorded by the brachiopods. The first trend is distinguished by low δ13C and relatively high δ18O values, consistent with alteration in a meteoric environment. The second trend, consisting of invariant and high δ13C values and progressively low δ18O values, is interpreted to have resulted from low water/rock interactions in a phreatic to shallow burial

111 environment. The third trend, observed in both brachiopods and eurydesmids,

has a positive correlation between δ13C and δ18O with both becoming lower with

increasing petrographic alteration. Based on cement petrography and isotopic

data, this latter alteration is interpreted as shallow burial diagenesis. Comparison

of geochemical data from least altered brachiopods with data from modern

brachiopods indicates that, except for extreme perturbations to the carbon cycle,

Permian seawater was probably isotopically and chemically similar to modern

seawater.

INTRODUCTION

The character of Paleozoic ocean water is typically inferred from

brachiopod geochemistry (Veizer et al. 1986; Banner and Kaufman 1994; Bojar

et al. 2004; Korte et al. 2005). As a result, extensive research has focused on

proving that brachiopods precipitate their shells in isotopic and chemical

equilibrium with seawater (Gonzalez and Lohmann 1985; Carpenter and

Lohmann 1995; Lee and Wan 2000; Brand et al. 2003; Parkinson et al. 2005),

with primary shell layers and specialized shell structures subject to vital effects,

but secondary layers approaching equilibrium values for δ18O and δ13C

(Carpenter and Lohmann 1995; Parkinson et al. 2005). This seems to be the case for most brachiopod species, but not all (Brand et al. 2003). Trace element distributions are prone to ontogenetic effects, with symmetrical variation amongst the layers, although whole shell averages apparently reflect seawater chemistry

(Lee et al. 2004). Since brachiopods are composed of low-magnesium calcite

112 (LMC), they should be less susceptible to diagenesis than bivalves composed of

aragonite or high-magnesium calcite, and therefore are more likely to preserve

original seawater signatures.

Bivalves that precipitate LMC, such as Cretaceous inoceramids shells,

(Elorza and Garcia-Garmilla 1998; Jimenez-Berrocoso et al. 2006), and modern

and ancient pectenids (Bojar et al. 2004; Chauvaud et al. 2005), have also been

used as paleoceanographic proxies. Inoceramids are assumed to precipitate

their shells in isotopic and chemical equilibrium with seawater (Gómez-Alday and

Elorza 2003), whereas studies of modern pectinid shells indicate that they are

precipitated in equilibrium with seawater for oxygen and carbon isotopic

compositions (Owen et al. 2002; Chauvaud et al. 2005) and trace elements

(Freitas et al. 2006), although there is evidence for the influence of vital effects in

some species (Lorrain et al. 2005). Like brachiopods, bivalves such as pectenids that precipitated LMC also existed during the Permian and thus could potentially provide viable paleoceanographic information. While there are no living eurydesmids, they are a type of pectenid, and distantly related to inoceramids,

(Runnegar 1979). Thus, eurydesmids are likely to have also precipitated their shells in equilibrium with seawater.

The purpose of this paper is to assess the utility of spiriferids and eurydesmids in the high-latitude carbonate rocks or the Lower Parmeener

Supergroup, deposited in the Tasmania Basin (Fig. 4.1), as proxies for early

Permian seawater by analyzing their geochemistry and comparing it against the newly proposed diagenetic history of the region (Rogala et al. 2008). This time

113 period represents a critical stage in Earth history, when all of the continents were

amalgamated into the Pangean supercontinent, and at the end of the late

Palaeozoic glaciation. The coincidence of these two important events could theoretically create significant changes in ocean chemistry, which could be elucidated by these proxies.

GEOLOGIC SETTING

Sediments of the Lower Parmeener Supergroup (Fig. 4.2) record the Early

Permian glaciomarine and deglaciation history of the Tasmania Basin.

Glaciomarine tills of the late Carboniferous Woody Island Formation (Fig. 4.2) accumulated directly on Precambrian granites and Cambrian, Ordovician,

Devonian and Silurian sedimentary and volcanic rocks (Clarke and Forsyth 1989;

Hand 1993). Glendonitic silts of the Woody Island Formation were deposited

across the shelf during the Asselian, with Tasmanites-rich muds forming in sub-

basins on the inner shelf (Rogala et al. 2007). The inner shelf at this time is

interpreted to be ice-covered, and circulation restricted by an island and complex

sub-basin topography.

114

Figure 4.1. Location of the Tasmania Basin. A) Boundaries of the early Permian Tasmania Basin are depicted by bold lines, highlands in grey. Sample locations from drill core (X) and outcrops (J) are shown. B) Eastern hemisphere of Pangea (modified from Li and Powell 2001). Modern continent shapes are shown in dark grey, Permian continental extent in light grey. Tasmania Basin is outlined by a box.

115 A change to iceberg-dominated conditions during the Sakmarian led to the establishment of a benthic biota and created a significant change in the facies distribution in the Tasmania Basin. Poorly fossiliferous sandy silts accumulated on the inner shelf, and graded outboard into fossiliferous siltstone and argillaceous limestone, with isolated layers of pure limestone (Bundella

Formation, which includes the Darlington Limestone; Fig. 4.2). Sea level dropped near the end of the Sakmarian, enabling siliciclastic sand (Liffey Group) to prograde across the shelf and fill in the rugged topography. Parts of the shelf, in the Hobart area (Fig. 4.1), remained marginal marine (Hickman Formation) throughout the low-stand period. Sea level rose once again at the onset of the

Artinskian, bringing a return to fully marine deposition. Widespread circulation of upwelling nutrient-rich water led to the extensive accumulation of carbonate sand on the mid-shelf (Berriedale Formation) during the high-stand. Sea level slowly fell throughout the remainder of the Permian, resulting in the progradation of marine (Malbina and Abels Bay Formations) and terrestrial sands (Upper

Parmeener Supergroup) across the shelf.

Pangea began to rift apart during the Jurassic, expressed in the Tasmania

Basin by the intrusion of thick diabase sills and eruption of minor amounts of basalt (Hergt and Brauns 2001; Leaman, 2002). Rifting, associated with faulting, uplift, and syenite intrusion, began between Tasmania and Antarctica, New

Zealand, and mainland Australia in the Cretaceous and was ongoing through much of the Cenozoic (Williams 1989).

116

Figure 4.2. Stratigraphy of the Lower Parmeener Supergroup. Limestone (shaded in darker grey tones) was the primary source of samples, although a few samples were taken from intervening and overlying fossiliferous units (shaded light grey).

DIAGENESIS

Based on petrography and cathodoluminescence, Rogala et al.

(2008) documented five major stages of diagenesis that occurred in three diagenetic environments for the Lower Parmeener Supergroup limestones (Table

4.1). Alteration of skeletal material began in the marine realm (Stage 1).

Bioclasts show evidence of dissolution prior to the precipitation of isopachous fibrous calcite cement, and many of the allochems were neomorphosed or

117 calcitized. In addition, there was abundant precipitation of authigenic glauconite

and phosphate, primarily within intragranular pore spaces.

Lower Parmeener sediments were subsequently exposed to meteoric

conditions once during the lowstand at the end of the Sakmarian, and again

during the lowstand and terrestrial conditions in the middle to late Permian. This

diagenetic environment was dominated by physical compaction of skeletal

material, but very little dissolution and cementation (Stage 2). Cements

precipitated at this time were non-ferroan drusy and bladed calcite that is interpreted to have formed under phreatic conditions (Rogala et al. 2008).

Most cementation occurred in shallow burial environments (Stage 3, 4, and 5; Table 4.1), at depths less than 1 km. By the Triassic, the Lower

Parmeener Supergroup had been further buried by Upper Parmeener

Supergroup sediments. Chemical compaction began at depths of approximately

300 m, driving precipitation of abundant ferroan calcite cements (Stage 3)

(Rogala et al. 2007). Dissolution remained minimal, with the exception of widespread removal of siliceous sponge spicules. Intrusion of thick diabase sills associated with the break-up of Pangea during the Jurassic changed the

cementation dynamics (Stage 4), by introducing silica-rich fluids that resulted in

pronounced localized dissolution and silicification of the limestones. Later uplift and fracturing during the Cretaceous provided conduits for meteoric ground

water, from which latest stage non-ferroan calcite cement precipitated (Stage 5).

118 Table 4.1. Summary of diagenetic stages recorded in the petrography of limestones from the Permian Lower Parmeener Supergroup (after Rogala et al., 2008). Diagenetic Stage Event Diagenetic Characteristics Environment 1 Marine Deposition • Isopachous fibrous calcite (early Permian) • Phosphate • Glauconite • Dissolution

2 Meteoric Burial by terrestrial • Mechanical compaction of Phreatic sediments (middle to allochems and cement late Permian) • Non-ferroan bladed and drusy calcite

3 Shallow Burial Burial by terrestrial • Chemical compaction sediments (late • Ferroan bladed, equant, blocky Permian to Triassic) calcite

4 Burial Intrusion of diabase • Silicification (Jurassic) • Dissolution

5 Uplift Rifting • Non-ferroan blocky calcite (Cretaceous)

METHODS

Brachiopods and eurydesmids were selected from the Darlington

Limestone, the marginal marine Hickman Formation, the Berriedale Limestone, and the Malbina Formation (Fig. 2). Shells were taken from multiple drill cores across the Tasmania Basin, as well as from outcrops in southern Tasmania. To assess diagenesis, spiriferid brachiopods and eurydesmid bivalves were selected from both locations and from sites in close proximity to diabase sills where carbonate sediments were clearly recrystallized. All of the shells were examined

using cathode luminescence and assigned a qualitative alteration designation of

fabric-retentive non-luminescent, fabric-retentive partially luminescent, fabric-

retentive luminescent, and recrystallized luminescent (Fig. 4.3 A and C). Shells

from each alteration category were sampled, although the least altered portions

of these shells were selected for isotopic and trace element analysis.

119 Representative brachiopod and eurydesmid shells from each cathode-

luminescence class were also examined using scanning electron microscopy

(Fig. 4.3 B and D). Each sample was etched in 2% HCl for 20 seconds and

mounted using carbon paint.

Samples selected for geochemical analysis were drilled, washed in

deionized water three times, dried, and analyzed at the Queen’s University

18 13 Facility for Isotope Research. Values of δ O and δ C were measured from CO2 released from dissolution of 0.5 mg of sample in 100% H3PO4 after 4 hours of

reaction at 72 °C. Replicate analyses using these procedures and comparisons with laboratory standards are within 0.2‰ for both δ18O and

δ13C values. Carbon and oxygen isotopic compositions are reported in standard

notation in units of per mill relative to the Vienna Pee Dee Belemnite standard

(VPDB). Samples were also dissolved in ultrapure 2% HCl, diluted in double

deionized water, and analyzed for trace elements using an Element High-

Resolution ICP-MS as described by Kyser et al. (2002).

120

Figure 4.3. Typical effects of alteration on the petrography of Lower Parmeener Supergroup shells. A) Brachiopod samples under cathodoluminescence. Top three are fabric-retentive, while the bottom brachiopod has had its fabric partially destroyed. B) Brachiopods under SEM, showing changes in the fibrous texture and an increasing amount of cement (arrows). C) Increasing luminescence of eurydesmid bivalves from fabric-retentive non-luminescent to fabric- destructive luminescent. D) Eurydesmids under SEM, showing increased alteration of shell layers.

121 RESULTS

Stable Isotopes

The δ18O values of brachiopods range from −7.9 to +0.5‰ (Table 4.2), with the most positive values associated with non-luminescent samples, and an increasing amount of luminescence with more negative δ18O values (Fig. 4.4a).

Covariance with δ18O, and therefore luminescence, is also observed in relation to

δ13C values, which range from +2.1 (luminescent) to +7.1‰ (non-luminescent).

Eurydesmid δ18O values range from −4.9 to −0.3‰ (Table 4.3), and have a marked increase in luminescence with decreasing δ18O values (Fig. 4.4b). The

δ13C values range from +3.9 to +6.7‰ and co-vary with δ18O values and

luminescence.

The isotopic compositions reported here are similar to those of Rao and

Green (1982) and Rao (1988) for brachiopods and bivalves from the Lower

Parmeener Supergroup (Fig. 4.5). Our brachiopod data, however, record more positive δ18O values and include samples that have high δ13C values despite

having low δ18O values, which are not reported in previous studies. High δ13C values from the Lower Parmeener Supergroup invertebrates are typical of

Permian carbon isotopic compositions (Grossman 1994; Dolenec et al. 2003;

Korte et al. 2005).

122 Table 4.2. Brachiopod isotopic and elemental geochemistry, separated into Formations. Oxygen and carbon isotopic values are provided, depicted in Figure 4.4A and Figure 4.5. Ca and Mg concentrations are given in mol%, and Fe, Mn, and Sr in ppm. Mg, Fe, Mn, and Sr are displayed graphically in Figure 4.6A. Sample Stratigraphic Ca Mg Fe Mn Sr δ18O δ13C Number Unit (mol%) (mol%) (ppm) (ppm) (ppm) EN1-67.6A Malbina -2.2 6.6 98.62 0.62 47 1114 617 EN1-67.6B Malbina -3.1 6.9 98.62 0.59 96 1540 516

B2-F1 Berriedale -3.0 5.3 98.00 0.71 34 1046 534 B2-F2A Berriedale -3.3 4.5 97.71 0.93 8 211 1066 BH10-202.2 Berriedale -0.6 3.4 98.68 0.51 9 1022 843 BH10-212.2 Berriedale 0.5 7.0 98.85 0.48 38 740 501 BT1-590.8 Berriedale -2.0 2.1 99.03 0.41 36 1035 493 BT1-627.4 Berriedale -0.8 6.5 98.99 0.41 17 698 740 EN1-080.0 Berriedale -0.8 6.6 98.90 0.41 5 712 725 EN1-117.8 Berriedale -1.0 6.1 98.43 0.83 22 840 760 GR1-063.25 Berriedale -1.3 6.2 98.83 0.42 16 809 651 GR1-068.7A Berriedale -1.3 6.8 98.88 0.45 7 870 821 GR1-082.0 Berriedale -2.4 5.7 99.09 0.37 7 681 821 GR1-097.3 Berriedale -1.7 5.7 98.03 0.79 175 348 711 GR1-103.4 Berriedale -5.1 7.1 97.74 0.61 0 43 787 GR1-227.3 Berriedale -1.0 5.6 99.07 0.33 9 132 912 GV1-98.0 Berriedale -0.1 5.9 98.26 1.11 8 685 566

BP-462.1 Hickman -1.0 6.3 98.47 0.63 76 1786 477

EN1-198.5A Darlington -4.0 5.6 98.09 0.74 14 932 960 EN1-198.5B Darlington -7.9 4.8 98.89 0.39 88 1196 473 EN1-216.0 Darlington -3.6 6.5 98.81 0.44 16 1034 922 EN1-281.0 Darlington -1.6 6.6 98.73 0.46 20 1300 1329 EN1-282.8 Darlington -2.4 6.6 98.16 1.16 41 848 848

123

Figure 4.4. A) δ18O and δ13C values for brachiopods from the Lower Parmeener Supergroup, data from Table 4.2. B) δ18O and δ13C values for eurydesmids from the Lower Parmeener Supergroup, data from Table 4.3. In both diagrams the shading reflects the degree of luminescence.

124 Table 4.3. Eurydesmid isotopic and elemental geochemistry, separated into Formations. Oxygen and carbon isotopic values are provided, depicted in Figure 4.4B. Ca and Mg concentrations are given in mol%, and Fe, Mn, and Sr in ppm. Mg, Fe, Mn, and Sr are displayed graphically in Figure 4.6B. Sample Stratigraphic Ca Mg Mn Fe Sr δ18O δ13C Number Unit (mol%) (mol%) (ppm) (ppm) (ppm) BH10-153.4 Darlington -0.3 6.7 98.47 0.70 20 174 1005 COL-2 Darlington -4.5 4.8 97.94 1.91 14 299 557 GRI-220.9A Darlington -4.3 4.5 97.82 1.56 51 913 861 GRI-220.9B1 Darlington -0.8 6.3 98.70 1.02 46 199 602 GRI-220.9B2 Darlington -4.9 4.6 99.15 0.43 30 409 844 M2-6B1 Darlington -4.8 3.9 96.76 1.52 58 1329 649 M2-6B2 Darlington -3.1 4.0 99.48 0.37 1 14 1180 M-9A Darlington -1.9 4.7 99.07 0.84 13 41 522 M-9B Darlington -2.1 4.4 99.04 0.69 11 41 550

Figure 4.5. Brachiopod and eurydesmid δ18O and δ13C values (this study) compared to Darlington and Berriedale brachiopods, eurydesmids, and cements of Rao and Green (1982) and Rao (1988).

125 Elemental Geochemistry

The Darlington and Berriedale Limestone brachiopods contain 0.3-1.2

mol% MgCO3 (Table 4.2). There is a slight increase in Mg/Ca with decreasing

δ18O values, but Fe/Ca, Mn/Ca, and Sr/Ca show no correlation with δ18O values

(Fig. 4.6a). Trace elements, such as Ni, Zn, and Mo, tend to correlate with Fe

(Fig. 4.7a) and have a similar scattered distribution with respect to δ18O.

Eurydesmids in these limestones are composed of LMC with 0.4-1.9 mol%

MgCO3 (Table 4.3). Mg, Fe, and Mn contents normalized to Ca all have similar

relationship with δ18O, showing a U-shaped distribution with an initial decrease in

minor element concentrations with decreasing δ18O, and an abrupt increase at

δ18O values of approximately −4‰. Normalized Sr is unaffected (Fig. 4.6b).

Once again, the transition metals correlate with Fe, with trends similar to those of

the brachiopods (Fig. 4.7b).

126

Figure 4.6. A) Mg, Fe, Mn, and Sr data normalized to Ca for all brachiopods and plotted against δ18O. Alteration increases towards the left with more negative δ18O values. B) Mg, Fe, Mn, and Sr data normalized to Ca for all eurydesmids and plotted against δ18O. Increased alteration is indicated by more negative δ18O values.

127

Figure 4.7. A) Select trace element data for brachiopods are plotted against Fe concentrations. Ni, Zn, and Mo are scavenged by Fe. B) Trace element data for eurydesmids are plotted against Fe concentrations, demonstrating covariance with Ni, Zn, and Mo, which are commonly scavenged by Fe.

128 INTERPRETATION

The range in the oxygen and carbon isotopic compositions of the

brachiopods is a reflection of alteration, as shown by covariance with cathode

luminescence. Consequently, the most positive δ18O and δ13C values, found in

the non-luminescent shells, are the most pristine samples (Fig. 4.4). By using

non-luminescent brachiopods and eurydesmids to reflect the least altered δ18O

values, assuming a δ18O for Permian seawater of 0‰ (VSMOW), and correcting

for Mg content (Jimenez-Lopez et al. 2004), seawater temperatures were

calculated to range from 14°C to 22°C (Kim and O’Neil 1997). These

temperatures are nearly 14°C warmer than would be expected from modern

analogues where polar carbonates are forming (Taviani et al. 1993). Permian

seawater would require a δ18O value of nearly −3‰ (VSMOW) to obtain reasonable temperatures from the Tasmanian dataset.

Rao and Green (1982) and Rao (1988) suggested that salinity variations resulting from melt-water or fluvial freshwater influx could account for this discrepancy. Assuming a salinity variation of 2.6‰ for every 1‰ change in δ18O,

as calculated by Buening and Spero (1996), seawater in the Permian Tasmania

Basin would have needed a salinity of 25.4‰ to produce the isotopic values of

the brachiopods and eurydesmid shells. This is unrealistic considering that

brachiopods and eurydesmids are found in the same beds as crinoids (Rogala et

al. 2007), that only live in seawater close to normal marine salinity (Jones and

Desrochers 1992).

129 Evidence for alteration is also shown by comparing trace element

concentrations of Lower Parmeener Supergroup brachiopods to modern

brachiopods (Brand et al. 2003; Lee et al. 2004). Lower Parmeener Supergroup brachiopods are enriched in Fe and Mn (Fig. 4.8), consistent with alteration in a reduced environment. Thus, even the apparently most pristine shells are in fact

altered.

Figure 4.8. Comparison of Mn and Fe concentrations from LPS brachiopods (this study) to modern brachiopods (Brand et al. 2003; Lee et al. 2004). Shaded inner box indicates the most common values for modern brachiopods; outer cross-hatched box indicates maximum range of values.

The relationship between δ13C and δ18O values in brachiopod and eurydesmid shells can be separated into three dominant alteration trends (Fig.

4.4). The first trend is only observed in the brachiopod data, and is typified by low δ13C with consistently high δ18O values. This is interpreted as alteration in an

130 environment where the pore water was not significantly different than seawater,

but where some biological mediation was able to influence the carbon isotopic

composition. As this trend probably reflects mixing between unaltered shell

material and recrystallized shell material with low δ13C values, precipitation of

cement along growth lines or punctae is a likely origin for the secondary cement.

Isotopic shifts of this magnitude are common, as reported in other studies (Allan

and Matthews 1982; Rush and Chafetz 1990).

The second trend is also only observed in the brachiopod data, and is

characterized by a decrease in the δ18O values with no change in δ13C values.

Such a trend is most likely due to alteration of the original carbonate under low

water/rock and low temperature conditions, with carbon derived from the original

calcite, but an oxygen isotopic composition influenced by meteoric fluids (e.g.

Algeo et al. 1992). Timing of this diagenetic phase is uncertain. Many of the

limestones are thin isolated beds surrounded by siliciclastic mudstones, which

have low permeabilities, and thus could have created a closed system where

carbon would be derived locally. Alternatively, this trend may reflect marine shallow burial diagenesis, as demonstrated by the Cretaceous Austin Chalk

(Czerniakowski et al. 1984). This mechanism is supported by the lack of meteoric cementation or dissolution around open pore spaces in the limestone beds where these shells occurred, compared to those from which others were collected. Calcitization of aragonite shells and neomorphism of micrite are thought to have occurred on the seafloor (Rogala et al. 2008), as they do on the

Holocene southern Australian shelf (James et al. 2005). Biogenic LMC, which is

131 not as stable as abiotic LMC (Busenberg and Plummer 1989), is prone to

alteration in this environment. Trace element variations due to this trend are

obscured by the third type of diagenesis.

The third diagenetic trend, with decreasing δ18O and δ13C values, is

prevalent in both brachiopods and eurydesmids. This was initially observed by

Rao and Green (1982), who interpreted it as mixing between the original

carbonate and cement derived from freshwater of the melting glaciers (Fig. 4.5).

While Rao and Green (1982) did not specify precisely which, or how, their cements were sampled, almost all of the visible cements are actually formed in a shallow burial environments (Rogala et al. 2008). Thus, this trend should be interpreted as physical mixing between unaltered shell material and recrystallized shell material derived from shallow burial fluids.

This last trend is not typical of burial diagenesis in limestones, which

would have less δ13C variability as does the second trend discussed here. More

negative δ13C values are due to the influence of biological processes operating in

the meteoric realm on pore-water isotopic composition, which are then reflected

in altered shell material. This signature is usually quickly swamped as the

carbonates are exposed to diagenetic fluids that promote the dissolution of

aragonite and HMC components, thus contributing a δ13C signature that reflects

the original marine composition (Allan and Matthews 1982; Budd and Land 1990;

Algeo et al. 1992; McClain et al. 1992). However, the meteoric-like carbon isotopic signatures from the Tasmanian samples are inferred to represent mixing with calcite cements that would have formed at depths of approximately 300 m

132 to, at most, 1 km (Rogala et al. 2008). This abnormality is attributed to the original composition and marine alteration of these cold-water carbonates, which are interpreted to have been composed almost exclusively of chemically-stable

LMC prior to entering the meteoric diagenetic environment. Because of this stability there was only limited early dissolution of carbonate sediments, allowing the biologically-mediated carbon isotopic signature to persist to a greater depth.

Delayed dissolution has also been reported from cool-water carbonates (Nelson

1988; Nicolaides 1995; Hood and Nelson 1996; Caron et al. 2004, 2006) but this isotopic trend has not (Nelson and Smith 1996).

Trace metal and REE data provide few additional details. In both

eurydesmids and brachiopods many of the trace metals, such as V, Ni, Mo, Cu,

and REE concentrations vary with Fe (see Fig. 4.7 for examples) and mimic its

distribution. These elements may be adsorbed from seawater onto Fe-Mn

oxyhydroxide particles or organic matter and then incorporated into the sediment

(e.g., Tribovillard et al. 2006) or onto the organic-rich layers in shells. They are

then released along with the Fe and Mn under reducing conditions and may be

remobilized during early diagenesis. Thus, it is expected that these elements should correlate. Variations in the sensitivity of individual elements to oxidizing

fluids may explain the scatter. Cement petrography indicates fluctuations in Fe

content (Rogala et al. 2008), which is consistent with this hypothesis.

Diagenetic overprinting during three separate alteration events has likely

obscured any one trace element trend (Fig. 4.6a). The trace element trends in

eurydesmids (Fig. 4.5b) should, however, be clearer since they appear to have

133 only been affected by one of the diagenetic events. The most altered eurydesmids show the highest amounts of Fe and Mn (Fig. 4.4b), although not in a linear relationship. This could mean that eurydesmid shells were actually exposed to several stages of diagenesis that are not reflected by the isotopes or, more likely, it is a reflection of the compositional variability of the shallow burial fluids (Rogala et al. 2008).

IMPLICATIONS FOR PALEOENVIRONMENTAL STUDIES

Three diagenetic events contributed to the alteration of the brachiopods and eurydesmids shells from the Lower Parmeener Supergroup, although they affected each shell type differentially. Brachiopods were altered by several stages of diagenesis (Fig. 4.4a), whereas eurydesmids were more resistant to early alteration and record only the burial stage (Fig. 4.4b). These differences most likely reflect shell structure. Tasmanian sprifierid brachiopod shells have a thin fibrous texture that would have been coated in organic molecules. These molecules degrade within 130,000 years (e.g., Endo et al. 1995), thus leaving voids that could act as conduits for early diagenetic fluids. Once these fluids penetrated the shell they would quickly alter the thin fibers. Eurydesmids, however, have a thick prismatic internal structure (Runnegar 1979). While there are still organic-rich layers along growth lines, the interiors of individual shell layers are here interpreted to have been more isolated from potential diagenetic fluid interactions than in the brachiopod shells. Rush and Chafetz (1990) reported a similar difference between fibrous and prismatic brachiopod shells,

134 and suggested that the large prismatic shell structure takes longer to alter because of its low surface area to volume ratio. This would make these shells more resistant to early diagenesis and, therefore, more likely to record only later diagenesis, as is observed here for the eurydesmid shells.

Cathodoluminescence and SEM are the most common techniques employed to determine the preservation quality of fossil shells. The results of this study indicate that fabric-retentive, non-luminescent shell material is no guarantee of having unaltered shell carbonate. Rush and Chafetz (1990) analyzed a series of Upper Silurian/Lower Devonian brachiopod shells sampled across a subaerial exposure surface, using both techniques to test shell preservation. They concluded that apparently unaltered shells had the same isotopic compositions as obviously recrystallized ones. While there is a general correlation between δ18O values and luminescence in the Lower Parmeener

Supergroup samples, all shells have been altered to some degree, even those with no luminescence. Thus, although cathodoluminescence and SEM can be used to assess alteration, they are not infallible techniques.

Diagenetic effects on the primary chemistry of the Lower Parmeener

Supergroup brachiopods and eurydesmids make inferences about the paleochemistry of Permian seawater difficult. The δ13C values, the most pristine of which are near +7‰ (Fig. 4.4), are significantly higher than modern values near +1.5‰ (Veizer et al. 1999), and at the high end of the range of global values for Permian bivalves from previous studies (e.g., Veizer et al. 1999; Korte et al.

2005).

135 The global positive δ13C excursion began in the Carboniferous. Increased terrestrial photosynthesis associated with the proliferation of the gymnosperms, and to a lesser degree marine productivity, is postulated to have preferentially removed 12C from the atmosphere (Beerling et al. 2002). This 12C-rich organic

matter would then have been either buried in extensive Carboniferous coal

swamps or sequestered in the deep ocean (Kump and Arthur 1999; Dolenec et

al. 2003). Additional coal- and oil shale-rich beds were deposited and buried in

many basins directly following glaciation (Berner and Raiswell 1983; Korte et al.

2005). Carbon isotope fractionation by C3 plants may have been further

enhanced by high O2 levels in the atmosphere (Beerling et al. 2002). Oxygen

levels gradually rose throughout the early Paleozoic, reaching a maximum of

35% at the beginning of the Permian (Beerling and Berner 2000). Experiments

conducted under similar elevated oxygen levels show that C3 plants change their

rate of leaf gas exchange, which increases fractionation, thus removing even

more 12C from the atmosphere (Beerling et al. 2002).

This removal of 12C from the atmosphere apparently resulted in a

positive δ13C shift in marine carbonates, averaging approximately +4‰ in the

Tethys Ocean (Korte et al. 2005). Brachiopods from coastal Tethyan basins,

however, tend to record more positive values, similar to the Tasmania Basin.

These basins are restricted and contain coal measures or, in the case of the

Tasmania Basin, oil shales that were deposited directly following glaciation

(Clarke and Forsyth 1989; Revill et al. 1994; Simoneit et al. 1993). A similar

trend is observed in Miocene brachiopods from the Murray Basin (Pufahl et al.

136 2006), and is attributed to increased productivity. Localized carbon sequestration due to elevated primary productivity combined with restricted circulation in the

Tasmania Basin is likewise interpreted to have resulted in δ13C values higher

than the Permian open-ocean signature.

Trace elements in the Permian bivalve shells have been remobilized

during diagenesis, but average original concentrations in brachiopods can be

inferred using regression based on their relationship with Fe. Fe was selected

because concentrations in modern brachiopods are well constrained, primarily to

100 ppm or lower (Brand et al. 2003; Lee et al. 2004). Thus, Fe enrichment

beyond this is an indication of alteration, and trace element trends toward the

lowest Fe values should reflect original concentrations. These results were also

compared to extrapolated values using δ18O as a representation of alteration,

which resulted in a similar range for the original compositions, which were

calculated to be approximately 750 ppm for Mg, 800 ppm for Sr, and 25 ppm for

Mn. All of these values are within the range of modern brachiopods, although Mg and Sr are in the low end of the range. The low Mg values might be expected in

the Tasmania Basin, however, as estimated cold water temperatures should

discriminate against Mg in the calcite structure (cf. Carpenter and Lohmann

1992; Vander Putten et al. 2000).

Based on this comparison between Permian and modern brachiopods, the

paleochemistry of the Permian ocean was not significantly different from modern

ocean chemistry except for carbon isotopes. This interpretation agrees with that

of studies of Permian geochemistry from fluid inclusions in evaporites (Hardie

137 1996; Horita et al. 1991; Lowenstein et al. 2005). It has been suggested that the

chemical composition of modern seawater can be explained by mixing of river

water and mid-ocean ridge hydrothermal brines (Spencer and Hardie 1990).

Despite the differences in continent and mid-ocean ridge distribution, estimated variables such as spreading rates and climate, which would effect river influx rates, are thought to be similar in the Permian and Modern (Hardie 1996).

Therefore, it is reasonable for ocean chemistry to be the same despite the radical difference in climate and in the carbon cycle.

CONCLUSIONS

Stable isotopic compositions and elemental concentrations in brachiopods and bivalves from the Permian Darlington and Berriedale formations in the

Tasmania Basin indicate that all spiriferid and eurydesmid shells have been affected by diagenesis, despite appearing petrographically unaltered.

Temperature calculations based on the most pristine δ18O values are unrealistic,

and cannot be explained by variations in salinity. Mn and Fe concentrations are

elevated, confirming alteration.

Three diagenetic events are recorded in the isotopic and chemical data.

The first event, which is recorded only by the brachiopod samples, is

distinguished by low δ13C and high δ18O values, and is interpreted as occurring in

a meteoric environment that is influenced by cold, seasonally frozen,

groundwater which maintains elevated δ18O values. Alternatively, evaporation

could increase the δ18O values of the fluid, although climatic evidence provided

138 by expansion of Glossopteris swamps indicates a progressively more humid

climate throughout the Permian (Clarke and Forsyth 1989; Ziegler et al. 1997).

The second event, also recorded only by the brachiopods, is characterized by

constant and high δ13C values and low δ18O values. This is interpreted to be the

result of low water/rock processes in a phreatic to shallow burial environment.

The third alteration event, observed in both brachiopods and eurydesmids,

exhibits a positive correlation between δ13C and δ18O values, with both values becoming lower with increasing alteration. Based on cement petrography and

isotopic data, this change is interpreted to be driven by shallow burial diagenesis.

A characteristic of cold-water carbonate diagenesis appears to be

mineralogical stabilization of virtually all allochems to LMC on or just below the

seafloor (Rogala et al. 2008). This means that there is little dissolution in

meteoric environments, and therefore little influence on pore-water chemistry by

the surrounding rock. Low δ13C values produced during biological processes in

the meteoric environment can, therefore, be preserved despite alteration caused

in deeper diagenetic environments.

On the basis of this study Permian seawater was probably not significantly

different from modern seawater, with the exception of having high δ13C values.

Concentrations of Mg, Sr, and Mn from Permian brachiopods in the Tasmania

Basin fall within the range of modern brachiopod data, indicating similar ocean

chemistry despite differing continental and mid-ocean ridge configurations.

139 ACKNOWLEDGEMENTS

This research is supported by the Natural Sciences and Research Council

of Canada through a discovery grant to N.P.J. and post-graduate scholarship to student B. R., as well as NSERC Discovery and MFA grants, CFI, and OIT grants awarded to T.K.K. and N.J. The authors wish to especially thank Don Chipley,

Kerry Klassen, and Bill MacFarlane for their assistance with running and processing sample geochemistry.

140 CHAPTER 5:

GENERAL CONCLUSIONS AND SYNTHESIS OF RESEARCH

INTRODUCTION

The Lower Parmeener Supergroup is a ~500 m-thick succession comprising two sequences of siliciclastic and carbonate rocks deposited in the

Tasmania Basin during the early to mid-Permian. These rocks record deposition on a southern-hemisphere polar shelf during a critical time in Earth history: the end of the late Paleozoic ice age. By contrasting the two sequences it is possible to make inferences about the deglaciation and sedimentation history of the

Tasmania Basin, and associated changes in oceanographic circulation patterns.

Two limestone units, the Darlington Limestone and Berriedale Formation, are of particular interest because they represent the accumulation of carbonate sediments in a cold-water environment. Few studies have been conducted on high-latitude cold-water carbonates, and those that have concern Pleistocene

(Taviani et al. 1993), Holocene (Orpin et al. 1998; Freiwald et al. 2002) and modern (Andruleit et al. 1996; Rao et al. 1996; De Mol et al. 2002) examples.

There are still many uncertainties regarding these unusual carbonates, such as how their composition changes over time, what oceanographic conditions promote their formation, and how they change during diagenesis. Lower

Parmeener Supergroup rocks provide an opportunity to answer some of these questions surrounding cold-water carbonate deposition, with a unique perspective from the rock record.

141 CONCLUSIONS FROM THE RESEARCH

Lithofacies Distribution and Paleoceanography

The Tasmanian climate was cold during the late Carboniferous and early

Permian but became progressively more humid and temperate towards the Late

Permian. Water temperatures remained cold throughout deposition of the Lower

Parmeener Supergroup, although surficial ice conditions changed. Two paleoceanographic states are interpreted to have existed during the early

Permian: (1) seasonal sea-ice-covered and (2) open-water with icebergs.

Seasonal ice-covered conditions prevailed during the late Carboniferous and earliest Permian (Asselian). Partitioning of the basin by glacially-scoured topography during the late Carboniferous to Asselian, along with very cold seawater temperatures, are interpreted to have promoted the accumulation of shore-fast sea ice. This would have resulted in increased salinity and stagnation of the water column, promoting the formation of glendonites. Extensive ice- derived silt covered much of the shelf, with local development of organic-rich mud in isolated sub-basins.

Open water and iceberg conditions prevailed on the post-Asselian shelf, although circulation patterns differed between the Sakmarian and Artinskian.

Bathymetry of the inner shelf remained irregular during the Sakmarian and continued to impede circulation of upwelling nutrient-rich water, as indicated by localized phosphate deposition. By the Artinskian the shelf topography had been reduced by sedimentation resulting in a smooth seafloor, allowing circulation of upwelling water across the entire basin. This interpreted increased nutrient

142 delivery promoted thicker and more laterally extensive carbonate sedimentation

and phosphatization during the Artinskian.

The distribution of Tasmanian cold-water shelf lithofacies was similar

during the Sakmarian and Artinskian, and reflected a gradient of hyposaline to normal marine salinity offshore due mostly to fluvial freshwater influx. Inner-shelf facies are characterized by bioturbated mudstone, siltstone, and sandstone with scarce shelly fossils consisting of ostracodes, benthic foraminifera, small thin- shelled brachiopods, and diminutive bryozoans. Middle-shelf facies grade outboard from fossiliferous siltstone to limestone to spiculitic limestone and are dominated by brachiopods, bryozoans, crinoids, and lesser bivalves. Plant fragments, phosphate, and dropstones are concentrated in mid-shelf facies.

Outer-shelf facies are preserved only in the Artinskian strata, and are characterized by siliciclastic turbidites and fossiliferous siltstone and sandstone.

Carbonate sediments are interpreted to have accumulated only during the highstand systems tract. During lowstand conditions there would have been a decrease in primary productivity corresponding to reduced shelf-width, as well as an influx of siliciclastic sediment that would have overwhelmed the predominantly filter-feeding fauna. As sea level rose, the outer shelf was once again isolated from terrigenous input, although carbonate sedimentation was most likely delayed until the faunal communities re-established themselves.

143 Tectonics

Comparisons between the Tasmania Basin and global or even eastern

mainland Australian events are difficult due to faunal provinciality which makes correlations tricky (Clarke and Farmer 1976; Clarke 1990). Parts of Eastern

Australia underwent a phase of extension and rifting during the early Permian, creating back-arc basins containing siliciclastics and volcanics, while other areas of the coast formed foreland basins as they underwent compression or sag basins during thermal subsidence (Fielding et al. 2001; Holcombe et al. 1997a;

Holcombe et al. 1997b; Veevers 2006). Since there are no volcanic rocks nor any evidence for sediments derived from an offshore arc in the Lower Parmeener

Supergroup it is doubtful that these rocks were deposited in a back-arc or foreland basin setting. The early Permian Tasmania Basin must then be placed in a sag basin or passive margin setting with local isostatic rebound. The regression at the end of the Sakmarian may be related to variations in global sea level caused by a period of glacial resurgence in Australia and Antarctica (Isbell et al. 2003; Jones and Fielding 2004), although constraining the timing of this event with Tasmanian stratigraphy is difficult. A few ash layers occur in the

Tasmania Basin during the middle and late Permian, possibly indicating collision with an arc (Veevers et al. 1994; Reid et al. in press). If this is the case, the arc must have collided diachronously with eastern Australia, impacting northeastern

Australia first.

144 Diagenetic History

Petrography.---Alteration of Lower Parmeener Supergroup limestones was the result of combined marine, meteoric phreatic, and burial processes. The marine

environment was diagenetically complex, with precipitation of intergranular

isopachous fibrous calcite, neomorphism of micrite and aragonite components,

and dissolution of aragonite and calcite allochems occurring. Seawater was

most likely saturated with respect to calcite, whereas precipitation was enhanced

by the phosphatization and glauconitization of the sediments. Microbial

mediation of subsurface pore waters is thought to have contributed to

cementation and dissolution processes.

Since Lower Parmeener Supergroup limestones had already been

stabilized to low-magnesium calcite on the seafloor, there was little impetus for

dissolution during meteoric diagenesis, and thus little cementation. The primary process in this environment was mechanical compaction. Most of the profound diagenesis subsequently took place in the burial realm. The majority of the

carbonate cements were precipitated as ferroan calcite during shallow burial to

depths of 300-1000 m. Carbonates were dissolved along grain boundaries

during chemical compaction, resulting in precipitation of calcite cement in the surrounding pore spaces. Localized pervasive and fabric-retentive dissolution and silicification occurred after all intergranular carbonate cements had precipitated. Silicification was most likely associated with the intrusion of

Jurassic diabases. Later uplift, likely during Cretaceous tectonism, created fractures that were filled with non-ferroan blocky calcite.

145 Geochemistry.---Concentrations of Mn and Fe in both brachiopod and eurydesmid shells are higher than the expected values of similar modern species, indicating alteration under reducing conditions. Oxygen isotopes vary with the degree of luminescence, and obvious petrographic alteration. The least altered of the brachiopods and eurydesmids were used in temperature calculations, but gave unrealistic values that could not be explained by variations in salinity. These results confirm that all brachiopod and eurydesmid shells have been altered by diagenesis, despite some of them having preserved fabrics and being non-luminescent.

Three diagenetic trends are apparent on the basis of isotopic data. The first trend occurs only in the brachiopod samples. It is distinguished by a covariant trend of light δ13C and heavy δ18O values and is interpreted to reflect progressive alteration in a meteoric environment. The second trend, also seen only in the brachiopods, is characterized by constant heavy δ13C values and progressively lighter δ18O values. This has been interpreted to result from the relatively low water/rock interaction in a phreatic diagenetic environment. The third trend, observed in both brachiopods and eurydesmids, has a positive correlation between δ13C and δ18O with both becoming lighter with increasing alteration. Based on cement petrography and isotopic data, this trend is interpreted as shallow burial diagenesis.

146 Comparison of Cool- and Cold-water Carbonate Biota

Sedimentology.---Both cool- and cold-water carbonates are characterized by a heterozoan assemblage that lacks inorganic precipitates, green algae, and photosymbiotic organisms (James 1997), although there are differences between the two. Generally, open, inner shelf cool-water carbonate sediments consist of coarse-grained red algae-, benthic foraminera-, and mollusc-dominated facies that form in water depths primarily <50 m, but up to 100 m (Hayton et al. 1995;

Rivers et al. 2007). These facies grade outboard into bryozoan- and molluscan- rich muds.

In contrast, cold-water carbonate shelves are characterized by impoverished inner shelf facies, consisting solely of benthic foraminera, ostracodes, and minor amounts of molluscs (Taviani et al. 1993; Andruleit et al.

1996; this study). Coralline red-algal facies are absent from all cold-water carbonates. This lack of shallow-water carbonates is most likely due to increased sediment influx associated with post-glacial terrigenous environments combined with substrate disturbance from grounding icebergs (Taviani et al.

1993; Andruleit et al. 1996; this study). Temperatures below 5°C may also contribute by preventing the colonization of these areas by red algae.

Most of the carbonate deposition on these polar shelves occurs at depths between 50 and 200 m, although cold-water coral mounds have been reported as deep as 1200 m (Kenyon et al. 2003). The composition of carbonate sediments below 50 m changes significantly depending on their age. Permian

Lower Parmeener Supergroup carbonates grade outboard from facies dominated

147 by eurydesmids, a specifically Gondwanan fauna, and bryozoans to brachiopod- bryozoan-crinoid-sponge-rich sediment. Pleistocene and Holocene cold-water carbonates additionally contain barnacles in mid-shelf to outer-shelf regions, and planktonic foraminera in deep outer-shelf to slope areas (Taviani et al. 1993;

Andruleit et al. 1996). In addition to this biota, cold-water carbonates may contain dropstones, and are underlain by diamictites. Lower Parmeener

Supergroup limestones are also distinctively associated with glendonite occurrences in under- and over-lying beds.

Diagenesis.---Lower Parmeener Supergroup limestones are predominantly composed of calcitic allochems, as are cool-water limestones. Because of this, cementation follows similar diagenetic pathways in both. Both cool- and cold- water carbonates experience dissolution, neomorphism and cementation on the seafloor. These processes may be more pronounced in cold-water carbonates where even some LMC components have dissolution pits and there are fewer initially aragonite components. Extensive seafloor alteration to LMC means early mineralogical stabilization, and therefore little dissolution in subsequent meteoric environments. Since there is low water/rock interaction, light carbon isotopic signatures produced during biological processes in the meteoric environment can, therefore, be maintained and reflected in alteration caused in deeper diagenetic environments. This has not been reported for cool-water carbonates, and thus could potentially be diagnostic of a cold-water environment where virtually no aragonite enters the meteoric diagenetic environment.

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APPENDIX:

CORE LOGS

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Map of Tasmania with the locations of the drill core used in this study. Pictoral logs of each drill hole are included within Appendix A. The position of the capital of Tasmania, Hobart, and Maria Island are included for reference.

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Drill Hole Name Abbreviation ID Number Easting Northing Bicheno 10 BH10 5977 604500 5373000 Bonney’s Plain 1 BP1 13595 550174 5375100 Bothwell-Thorpe 1 BT1 6149 503563 5306011 Eaglehawk Neck 1 EN1 5258 576000 5238200 Fingal Range 4 FR4 3525 588032 5389999 Fingal Range 85 FR85 3610 591218 5390466 Golden Valley GV1 2292 474911 5391332 Granton 1 GR1 5348 515614 5266492 Hunterston 1 HUNT1 20613 495500 5326400 Margate Hart’s Hill 1 MHH1 5360 520500 5234220 Margate Hart’s Hill 2 MHH2 5361 520340 5233760 Nicholas Range 9 NR9 6147 588701 5401002 Oyster Cove 1 OC1 5358 518110 5228150 Pelham 1a & 1b PEL1 20576 495240 5242000 Porter Hill 1 PH1 6169 528877 5247467 Ross 1 ROS1 13583 536282 5347165 Ross-Quoin 1 RQ1 5349 554738 5331162 Tunbridge 2 TUN2 13978 524510 5334870

All drill hole locations used datum AMG66. ID Number refers to the drill hole records of the Tasmania Department of Mines. All core are stored at the Mineral Resources Tasmania Mornington facility.

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