Origin, sequence stratigraphy and depositional environment of … · Origin, sequence stratigraphy and depositional environment of an upper Ordovician (Hirnantian) deglacial black

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Palaeogeography, Palaeoclimatology, Pa

Origin, sequence stratigraphy and depositional environment of an

upper Ordovician (Hirnantian) deglacial black shale, Jordan

Howard A. Armstronga,T, Brian R. Turnera, Issa M. Makhloufb, Graham P. Weedonc,

Mark Williamsd, Ahmad Al Smadie, Abdulfattah Abu Salahe

aDepartment of Earth Sciences, University of Durham, Science Laboratories, Durham, U.K.bDepartment of Earth and Environmental Sciences, Hashemite University, Zarqa, Jordan

cDepartment of Geography, University of Wales, Swansea, Singleton Park, Swansea SA2 8PP, U.K.dBritish Antarctic Survey, High Cross, Madingley Road, Cambridge, CB3 0ET, U.K.

eNatural Resources Authority, P.O. Box 7, Amman 11118, Jordan

Received 26 January 2004; received in revised form 30 December 2004; accepted 14 January 2005

Abstract

The upper Ordovician succession of Jordan was located ~608S, less than 100 km from the Hirnantian ice sheet margin. New

graptolite dates indicate glaciation ended in Jordan in the late Hirnantian (persculptus Biozone). The succession records two

glacial advances within the Ammar Formation and the subsequent deglaciations. Organic-rich black shales (Batra Formation)

form part of the final deglacial transgressive succession that in-filled an existing low stand glacial continental shelf topography.

The base of the black shale is coincident with the maximum flooding surface. During transgression, interfluves and sub-basin

margins were breached and black shale deposition expanded rapidly across the region. The top of the black shales coincides

with peak highstand. The bexpanding puddle modelQ (sensu Wignall) for black shale deposition, adapted for the peri-glacial

setting, provides the best explanation for this sequence of events.

We propose a hypothesis in which anoxic conditions were initiated beneath the halocline in a salinity stratified water

column; a fresher surface layer resulted from ice meltwater generated during early deglaciation. During the initial stages of

marine incursion, nutrients in the monimolimnion were isolated from the euphotic zone by the halocline. Increasing total

organic carbon (TOC) and d13Corg up section indicates the organic carbon content of the shales was controlled mainly by

increasing bioproductivity in the mixolimnion (the Strakhov model). Mixolimnion nutrient levels were sustained by a continual

and increasing supply of meltwater-derived nutrients, modulated by obliquity changes in high latitude insolation. Anoxia was

sustained over tens to hundreds of thousands of years. The formation of black shales on the north Gondwana shelf was little

different to those observed in modern black shale environments, suggesting that it was the nature of the Ordovician seas that

pre-disposed them to anoxia.

D 2005 Elsevier B.V. All rights reserved.

Keywords: Hirnantian; Black shale; Sequence stratigraphy; Depositional models; Jordan

T Corresponding author.

E-mail address: [email protected] (H.A. Armstrong).

0031-0182/$ - s

doi:10.1016/j.pa

laeoecology 220 (2005) 273–289

ee front matter D 2005 Elsevier B.V. All rights reserved.

laeo.2005.01.007

H.A. Armstrong et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 220 (2005) 273–289274

1. Introduction

Early post-glacial, transgressive, organic-rich black

shales were a ubiquitous feature of the northern

Gondwana continental shelf during the late Ordovi-

cian to early Silurian. These rocks, which contain up

to 10% total organic carbon, have an elevated gamma

log response (up to 200 API, Andrews, 1991) due to

uranium-enrichment, and constitute one of the world’s

major hydrocarbon source rocks. The origin of these

black shales is uncertain, and resolution of this

problem, the aim of this paper, could provide

considerable palaeoenvironmental, palaeoclimatic

and palaeooceanographic information.

Elucidation of the origin of these rocks is

hampered by the fact that modern open shelf settings

are only subject to seasonal anoxia (Tyson, 1996;

Tyson and Pearson, 1991) and black shales are not

currently forming on modern shelves, despite anthro-

pogenic inputs, making present day shelf seas the

most nutrient-rich in Earth history. Further, authigenic

U-enrichment (Anderson et al., 1989) and the wide-

spread preservation of non-bioturbated, millimetre-

scale laminations in these rocks would be unsustain-

able with seasonal oxia (Wignall, 1991). This led

Wignall (1991) to ask the question: if black shales are

not forming on continental shelves at the present day

then why should they have formed in the past?

2. Existing models for anoxia

Studies of modern oxygen-poor environments have

shown that oxygen deficiency may be due to

decreased oxygen supply or increased oxygen demand

beneath areas of high productivity (Demaison and

Moore, 1980; Wignall, 1991). Present day black shale

formation is occurring in two distinct settings: in

upwelling zones and silled basins, both of which are

appropriate analogues for the peri-glacial setting of

northern Gondwana.

2.1. Upwelling model

Wind-driven coastal upwelling zones are the most

important, and the vast majority of the world’s

dysaerobic sediment forms in these areas. At the

present day, anoxic and dysoxic conditions only occur

in coastal upwelling zones at the intersection of an

intensified oxygen minimum zone (OMZ) and the

upper slope or continental shelf break, in water depths

typically N50 m (Wignall, 1994). Typically, radiolarian

and/or diatoms are abundant and the sediments are

silica-enriched. The upper and lower boundaries of the

OMZ are marked by phosphate enrichment (Demaison

and Moore, 1980; Loughman, 1984). Organic-rich

shale accumulation in upwelling zones is due to

elevated productivity levels in a regime of vigorous

circulation and high organic loading. Oxygen defi-

ciency aids the preservation of organic carbon in the

sediment (Parrish, 1982; Finney and Berry, 1997). For

example, despite the very high productivity around

the margins of Antarctica, the exceptional oxygen-

richness of the water prevents high organic carbon

values in the sediments (Demaison and Moore,

1980). Typically, OMZs associated with upwelling

zones are only stable on a decade time-scale

(Wignall, 1990), whilst oxygenation events from

the influx of turbidites tend to be once-in-a-hundred-

year events (Souter et al., 1981).

Analogy with modern systems predicts that ancient

black shales forming in association with coastal

upwelling should be palaeogeographically restricted

to low latitudes. However, given the right climatic

circumstances, upwelling zones could form parallel to

coastlines, in which case they would have a long,

narrow sheet-like geometry. The interval of black

shale formation should coincide with evidence for

increased current activity. Black shales should be

intimately associated with chert (diatom in high

latitudes and radiolarian at low latitudes) and phos-

phate-rich facies. These facies associations are rarely

observed in early Palaeozoic high latitude settings as

diatoms had not appeared by this time and mecha-

nisms exist for rapid phosphorous recycling into the

upper water column where it is utilized by the

biosphere (van Cappellen, 1993).

Black shales, indicative of late Ordovician–early

Silurian anoxia, have been attributed to oceanic

upwelling. At this time, much of the northern margin

of Gondwana (including Jordan) lay in high palae-

olatitudes within a zone of westerly winds, which

according to computer simulations generated intense

and in some areas seasonally restricted upwelling

(Parrish, 1982; Moore et al., 1993). The same

simulations, however, indicate that upwelling contin-

Strakhov model

Low salinity layer

Freshwaterrun off

Organic-richblack shalein anoxic

marine water

Nutrients displaced intomixolimnion (van Cappellen effect)

by rising monimolimnionreduction in the depth of the

mixolimnion = increasing productivity

a

marineinflux

Deuser model

Low salinity layer

Organic-richblack shalein anoxic

marine water

Progressive reduction in the depth of themixolimnion = decreasing productivity

Freshwaterrun off

b

marineinflux

halocline

Fig. 1. Conceptual models for the formation of basin anoxia in a

silled basin setting.

H.A. Armstrong et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 220 (2005) 273–289 275

ued into the Wenlock, long after deposition of black

shale had ceased; hence, more complex processes

must have been operating. However, the black shales

in Jordan fail the low latitude restriction and facies

geometry predicted by the upwelling model.

2.2. Silled basin model

The modern Black Sea, which is the largest anoxic

basin in the world, has been identified as a good

analogue for ancient black shale (Pompeckj, 1909;

Tyson et al., 1979; Wignall, 1994). The basin is

almost landlocked and its link with the Mediterranean

Sea through the Straits of Bosporus is silled, and

locally b30 m deep. The Black Sea receives large

volumes of fresh water from rivers draining its

northern hinterland as well as marine Mediterranean

water. Due to the density contrast between these water

types, there is a well-developed halocline at ~200 m,

which severely restricts advection of oxygen-rich

surface waters to depth (Wignall, 1994). Lateral

advection is also restricted due to the small amount

of water entering the basin, and circulation beneath

the halocline is sluggish, taking ~1000 years for

complete mixing of the water beneath the halocline

(Brumsack, 1989). The combination of poor circu-

lation and moderate oxygen demand from decaying

and descending organic matter results in an anoxic

monimolimnion that contains free hydrogen sulphide.

The sediments accumulating in the deepest parts of

the Black Sea basin consist of fine-grained siliciclastic

turbidites derived from the basin margins (Lyons and

Berner, 1992) and organic-rich, microlaminated,

coccolith oozes and sapropels. Two conflicting

models are available for the formation of the

organic-rich sapropels, both of which rely on the

influx of marine water (Fig. 1). In the Deuser model

(Deuser, 1974), marine water produces a puddle in the

deeper parts of the basin which gradually expands

upwards. Organic carbon accumulation is the result of

enhanced preservation beneath the anoxic water

column. Changes in surface productivity are not

considered important in this model and progressive

restriction of the size of the mixolimnion would

gradually restrict nutrient availability and thus could

lead to a decline in productivity and the resulting

sedimentary organic carbon. In contrast, the Strakhov

model (Calvert, 1990; Calvert and Fontugne, 1987;

Calvert et al., 1987; Strakhov, 1971) invokes elevated

productivity. During initial stages of marine incursion,

nutrients in the monimolimnion are envisaged to have

remained isolated from the euphotic zone by the

halocline. The gradually rising monimolimnion then

releases nutrients to the mixolimnion (bvan Cappellen

effectQ) and triggers a high productivity episode,

which is sustained by organic matter production in

the euphotic zone. Thus, the elevated sedimentary

organic carbon values record a high flux of organic

matter to the sediment at a time when much of the

water column was fully oxygenated. This model does

not take into account the initial concentration of black

shale deposition in the deepest parts of the basin and

must be allied with the presence of weak bottom

currents to concentrate sediment and organic material

(Huc, 1988).

Whilst the Black Sea is not an ideal analogue for a

peri-glacial shelf, it does provide diagnostic criteria

for the recognition of black shales that formed in a


silled basin setting. Most organic-rich shales initially

accumulate in topographic lows, due to the concen-

tration of low density organic particles by weak

currents (Huc, 1988), and would have a restricted

distribution pattern. In the Strakhov model, black

shales formed beneath the halocline and are interbed-

ded with normally oxygenated facies. This results

from climate-driven changes in water advection

(Wignall, 1994) or, with increasing sea level, a breach

of the sill, which allows large influxes of oxygenated

marine waters (Tyson and Pearson, 1991). The Deuser

model predicts declining euphotic zone bioproductiv-

ity, in contrast to the Strakhov model which predicts

sedimentary organic matter will record patterns of

euphotic zone bioproductivity. During deglaciation, in

ice margin settings, melting of the ice sheet and hence

nutrient flux and euphotic zone bioproductivity would

be modulated by changes in high-latitude seasonal

insolation.

High latitude seasonal insolation is controlled to a

large degree by ~21 kyr precession cycles, and low-

latitude seasonal extremes are entirely dominated by

precession (Berger, 1978; Budziak and al, 2000;

Molfino and McIntyre, 1990; Perks et al., 2002;

Rostek et al., 1997). However, over the annual cycle,

precessional insolation anomalies sum to zero at all

latitudes (Liu and Herbert, 2004). In contrast, obliq-

uity variations (~41 kyr periodicity) change the annual

insolation received as a function of latitude (Ruddi-

man and McIntyre, 1984). A high obliquity config-

uration increases insolation at latitudes above 458 in

both hemispheres and decreases the solar energy

received in the tropics and subtropics. Therefore,

obliquity changes the amplitude of the seasonal cycle

and alters the equator-to-pole insolation gradient.

Milankovitch theory predicts the Hirnantian ice sheet,

nutrient flux, bioproductivity and preserved sedimen-

tary organic matter (TOC) would primarily record

obliquity cycles.

3. Sequence stratigraphical context for black shale

formation

Wignall (1994) and Wignall and Maynard (1994)

identified two distinct types of black shale within the

transgressive sedimentary record: basal transgressive

(BT) and maximum flooding (MF) black shales (see

also van Wagoner et al., 1990), depending on whether

they rested on the transgressive surface (ts) or the

maximum flooding surface (mfs). Basal transgressive

black shales are further characterised by the fact that:

(1) they only develop in topographic hollows and

basin-centre locations at times of non-deposition on

the basin margin; (2) they have no time equivalent

deeper water facies (Grabowski and Glaser, 1990); (3)

they exhibit no lateral facies variation, but are

interbedded with shallow-water facies, perhaps imply-

ing pelagic depositional processes; and (4) they are

probably shallow water facies as they occur at the

base of the transgressive systems tract (TST). In areas

where sediment influx is low, this surface would be

diachronous towards the basin margin. In such cases,

where the TST is poorly developed, black shale may

closely overlie transgressive surfaces, or the sequence

boundary, transgressive surface and maximum flood-

ing surface may all amalgamate into a single erosive

or non-depositional surface (Wignall, 1994; Wignall

and Maynard, 1994). On the other hand, MF black

shales that rest conformably on the maximum flood-

ing surface can extend a considerable distance on to

the shelf and pass proximally into thicker sections of

shallower, marginal marine facies. They represent a

deep-water, sediment starved condensed section.

Wignall (1991) proposed the bexpanding puddle

modelQ to explain the evolution of black shale basins.

In the puddle, dyoxic/anoxic bottom waters, in basin-

floor lows, are inferred beneath a stratified water

column. Stratification can be vertical gradients of

either temperature (Heckel, 1971; Tyson et al., 1979)

or salinity/density (Baird et al., 1987; Fleet et al.,

1987). For example, Rossignol-Strick et al. (1982)

explained the development of anoxic conditions and

sapropel formation in the Mediterranean Sea by the

presence of a cap of low-density freshwater, which

stopped the sinking of higher density oxygen-bearing

waters. As sea level rises, incised valleys and existing

topography on the basin floor are gradually filled to a

point where the interfluve areas or sub-basin margins

are drowned (Fig. 2). This results in rapid expansion

of accommodation space. If the sedimentation rate is

constant or reduced, the accommodation space cannot

be filled, the basin becomes starved (Wignall, 1991),

and an extensive ravinement surface develops. This

model predicts black shales rapidly expanding out-

wards from the area of greatest basin depth (controlled

a

Continued sea level risebreaches interfluves during the HST

ravinement surfaceonset of deposition

regionally diachronous

Inherited glacial topographyon shelf-faulted graben and

incised valleys

ts TST

HSTmfs

b

Upper Ammar

Lower Batra

Middle Batra

Fig. 2. Sequence stratigraphical model for maximum flooding black

shales in Jordan. (a) Early transgressive filling of the inherited glacial

shelf topography. (b) Expansion of black shale facies, breaching

interfluves, during the HST. Abbreviations: HST, highstand systems

tract; TST, transgressive systems tract; mfs, maximum flooding

surface; ts, transgressive surface (after Wignall, 1991, 1994).


by existing basin floor topography) and the area of

deposition is at its greatest extent during highstand.

4. Stratigraphical and sedimentological context of

the upper Ordovician of Jordan

Ordovician sedimentation occurred on the margins

of the more extensive North African (Gondwana)

shallow-marine shelf, within terrestrial, subtidal,

marginal marine and shelf environments (Amireh et

a

60 So

60 So

30 So

Afri

ca

South America

Southern Europeterranes

Arabia

Jordan

SouthPole

India

b

Syria

Lebanon

Iraq

60 km

Saudi Arabia

Jord

an

Amman

Subcrop of BatraMudstone

Aqaba

33o

32o

31o

30o

35o 36o 37o38o

Southern DesertStudy area

RishaArea

WadiSirhan

Al Jafr

RH-6

WS-3

JF-1JF-2

RH-3

borehole

continentalmargins

ice flow

ig. 3. Locality maps. (a) Outcrop and subcrop of the Batra Formation (in part after Andrews, 1991) and the location of the study area in southern

ordan. (b) Late Ordovician palaeogeographical reconstruction of eastern Gondwana showing the ice sheet (grey shading) and location of Jordan.

F

J

al., 2001; Makhlouf, 1995; Powell et al., 1994), and

by upper Ordovician times the shelf was influenced

by Hirnantian glaciation (Brenchley et al., 1994;

Long, 1993; Marshall et al., 1997; Qing and Veizer,

1994) and the presence of a temperate bLaurentide-scaleQ ice sheet in the continental interior (Turner et

al., 2002) (Fig. 3). The current upper Ordovician

lithostratigraphical framework in Jordan does not

adequately reflect the changing facies associated with

this glaciation or allow for surface and subsurface

correlation of the stratigraphy. We have therefore

adopted an informal scheme prior to submission to the

Geology Directorate of the Natural Resources Author-

ity of Jordan for formal acceptance (Fig. 3).

4.1. Dubaydib Formation

The Llanvirn(?)-Ashgill Dubaydib Formation,

which comprises sandstones dominated by hum-

mocky cross-stratification, is divided into three

members (Fig. 4). The lower member, which attains

a thickness of 25 m, is defined by the first occurrence

of the trace fossil Sabellarifex and comprises quartz-

rich sandstones that grade upward into siltstone

(Makhlouf, 1992). The unit is pervasively burrowed

by a variety of vertical, uniserial or U-shaped burrows

with circular cross-sections, attributable to the Skoli-

thos Ichnofacies characteristic of high-energy, shallow

water conditions associated with moving sands. The

middle member comprises 54 m of channelised, fine-


grained, cross-bedded sandstones, with subordinate

very fine, silty sandstone. The upper member is 76 m

thick and contains six coarsening-upward cycles of

greenish siltstone to grey fine-grained sandstone

(Makhlouf, 1992). Trace fossils include Cruziana,

Chondrites and Planolites assigned to the Cruziana

Ichnofacies, which is best developed below normal

fair-weather wave base in well-sorted silts and sands

in relatively quiet water conditions.

4.2. Tubeiliyat Formation

The Ashgillian Tubeiliyat Formation is up to 105

m thick and conformably overlies the Dubaydib

Formation (Fig. 4). It comprises greenish, silty shale

and hummocky cross-stratified sandstone (Makhlouf,

1992), organised into coarsening-upward and less

commonly fining-upward packages, which decrease

in thickness from ~25 m to 10 m upwards. Trace

fossils of the Cruziana Ichnofacies are common in the

greenish silty shales.

4.3. Ammar Formation

Abed et al. (1993) and Amireh et al. (2001)

included all the late Ordovician (Ashgillian) glacial

CHRONOSTRATIGRAPHY

LLANDOVERY

ASHGILL

CARADOC

LLANVIRN

ARENIG

(TREMADOCIAN)

OR

DO

VIC

IAN

SIL

UR

IAN

(par

tim

) RATIYA SANDSMEMBER

BATRA MUDSMEMBER

TUBEILIYAT MB

UPPER

MIDDLE

LOWER

HISWA FORMATI

DU

BA

YD

IBF

OR

MA

TIO

NM

UD

AW

WA

RA

FO

RM

AT

ION

UMM SAHM FORMA

Palaeovalleysst’s

SOUTHERN DESERT JORDAN(OUTCROPS)

Shelf‘Hot Shale’

Shelf

Subtidalchannels

Shoreface

Shoreface

Shoreface

Shoreface

Upper

LowerShoreface

HCS

AMMAR

Braiddeltas

Foreshore

}

200

400

600

800

Met

res

Ammarchanne

Lower

Fig. 4. Currently accepted lithostratigraphy and chronostratigraphy of the

Arabia. The scheme used herein is a modification of this and raises the T

sediments in southern Jordan within the Ammar

Formation (Fig. 5). They divided it into two units,

each underlain by a glacial erosion surface. The upper

unit of the Ammar Formation comprises erosively

based, palaeovalley channel-fill sandstones which

consist of: (1) coarse-grained, proximal, braided

outwash plain sandstones containing glacially faceted

and striated clasts, typically concentrated in the lower

1–3 m, which Abed et al. (1993) and Powell et al.

(1994, Fig. 5) interpreted as tillite; and (2) coarse-

grained, proximal, braided outwash plain sandstones,

similar to those in (1), except that they grade up into

thin marine shoreface sandstones containing trace

fossils and brachiopods (Abed et al., 1993; Amireh et

al., 2001) similar to the Tubeiliyat shoreface sedi-

ments beneath the channel-fill sandstones.

Abed et al. (1993, Fig. 6) considered the glacially

incised lower unit of the Ammar Formation to

comprise a sandy channel lag conglomerate up to 2

m thick, containing glacially faceted and striated

clasts. This is overlain by a 30 m thick, greenish-

grey, massive, well-sorted sandy siltstone, which is

extensively deformed and thought to be completely

void of primary sedimentary structures and fossils

(Fig. 5). A re-examination of the channel-fills in the

Jebel Ammar area shows that they exceed 90 m in

LITHOSTRATIGRAPHY

QA

SIM

FO

RM

AT

ION

CENTRAL AND NORTHERNSAUDI ARABIA

SHARARA MEMBER

QUSAIBA MEMBERTAY

YA

RA

TF

OR

MA

TIO

N

SARAH FMZARQA FM

QUWARAH MB

RA’AN MEMBER

KAHFAH MB

HANADIR MB

SAQ SANDSTONE

FORMATION

Ero

es i o n a

bl as e

opf al a

o

valleys

TONE

TONE

ON

TION

RATIYA FORMATION

BATRA FM

SOUTH NORTH

TUBEILIYAT FORMATION

AMMAR FM

UPPER

MIDDLE

LOWER

HISWA FORMATION

DU

BA

YD

IBF

OR

MA

TIO

N

UMM SAHM FORMATION

SOUTHERN DESERT JORDAN(REVISED)

ls

UpperAmmar

late Ordovician of southern Jordan, and its correlation with Saudi

ubeiliyat and Batra members to formation status.

Tube

iliya

tFm

Am

mar

Fm

Low

erU

pper

Bat

raF

m

Ash

gilli

anLl

ando

very

- Wen

lock

L.S

ILU

RIA

NU

PP

ER

OR

DO

VIC

IAN

250

200

150

100

50

0m

Cla

y

claysi

lt

siltF FM MC C{ {

Marineshelfshales

Marineshelfshales

Fossiliferousmarine shorefacesandstones

Glacio-fluvialsandstones

Glacio-fluvialsandstones

Incision 2 Incision 2

Incision 1

Incision 1

Undeformedshoreface sandstones

Shoreface

Deformedshorefacesandstones

Conglomerate Conglomerate

Transgressivesurface

Maximum floodingsurface

Conglomerate

Sand Sand

Cross-bedding Ripplecross-lamination

Hummockycross-stratification

Fining andcoarsening-upwardsequences

clay silt F M C

Sand

{

Marineshelf

Qusaiba Mb

Glacio-marine

Fluvio-glacial

Incision 2

Incision 1

Tillite

Tillite

Marginal marineshoreface

Sarah Fm

Marineshoreface

Quwarah Mb

NW SAUDI ARABIA

Zarqa Fm

SOUTHERN JORDAN

Fig. 5. Upper Ordovician and lower Silurian stratigraphy in southern Jordan showing the glacial deposits of the Ammar Formation and

equivalent deposits in NW Saudi Arabia (after Vaslet, 1990).


thickness locally and that, despite tectonic and soft-

sediment deformation of the sandstones, they preserve

rare primary ripple cross-stratification, burrows and

brachiopod moulds.

Erosively based channels, which differ from the

upper Ammar glaciofluvial channels, are locally

incised into the top of the Tubeiliyat formation (Fig.

5). These are filled by up to 50 m of massive, pale

grey, fine-grained, fawn-weathering sandstone. The

sandstone is uniformly fine-grained, micaceous,

undeformed and contains small-scale, trough cross-

beds and ripple cross-lamination. The sandstone along

the sides and floor of the channel contains brachio-

pods and numerous trace fossils, including Harlania,

Cruziana and brachiopod resting traces. Thus, two

glacial incision events are recognized in the late

Ordovician of southern Jordan. The first occurs

beneath the disturbed and undisturbed channel sand-

stones locally incised into the top of the Tubeiliyat

Formation, with the former included in the lower part

of the Ammar Formation by Abed et al. (1993). The

second incision occurs at the base of the upper

Ammar Formation glaciofluvial to marine channels.

These are incised into the deformed and undeformed

lower channels, as well as the underlying non-

channelised Tubeiliyat sediments.

At least two episodes of glacial and peri-glacial

deposition are present in Saudi Arabia termed the

Zarqa and Sarah Formations (McClure, 1978;

McClure, 1988; Vaslet, 1990; Young, 1981). Both

formations have glacial unconformities at their bases

and contain glacial, glacio-fluvial and glacio-marine

facies (Vaslet, 1990). We consider the lower and upper

members of the Ammar Formation to correlate with

the Zarqa and Sarah Formations (Fig. 5).

4.4. Batra Formation

The Batra Formation is ~40–120 m thick and, in

the Southern Desert region, it conformably overlies

Fig. 6. Camera lucida drawings of Normalograptus parvulus (H

Lapworth, 1900), �10. a, c, d are from slab 18401-53 and b is from

slab 18401-58. b–d are external moulds. Based on these and

additional specimens from the slabs, the proximal dorsoventra

width is 0.7–0.75 mm, proximal 2TRD (thecal spacing) is 1.25–1.6

mm, distal dorsoventral width (th12+) is ~1.6 mm and distal 2TRD

is ~1.5 mm. These measurements are close to those given by

Loydell et al. (2002) and Zalasiewicz and Tunnicliff (1994) for this

species. N. parvulus has a range of persculptus to early acuminatus

Biozones. It may prove to be conspecific with G. persculptus

Specimens stored in the palaeontological collections of the Geo-

logical Directorate, Natural Resources Authority, Amman.


the Ammar Formation or disconformably overlies the

Tubeiliyat Formation (Fig. 4). In the type area of Wadi

Batn el Ghul, the formation comprises 3 members

(Andrews, 1991). The lower member is found in

shallow wells, apparently restricted to fault-bounded

graben structures. Only the middle member outcrops

and spans the Ordovician–Silurian boundary (Powell,

1989).

The Batra Formation is variably graptolitic. Masri

(1988) identified abundant graptolites, sparse thin-

shelled bivalves and rare trilobites in the mudstone.

Andrews (1991) reviewed the biostratigraphy of this

formation from surface exposures and exploration

wells and concluded that the formation ranged in age

from Ashgill (persculptus Biozone age) to mid-

Wenlock. New collections from black shales from

core in the type area contain abundant Normal-

ograptus parvulus (Fig. 6), which is likely co-specific

with Normaolgraptus persculptus and ranges from the

persculptus to acuminatus biozones. In the Al Jafr

area (well JF-1, 3494 ft depth; Fig. 3), side wall cores

contain Glyptograptus persculptus, whereas in the

Wadi Sirhan area (well WS-6, 1399.8 m depth; Fig. 3)

the lower member contains an acuminatus Biozone

fauna. Monoclimacis flumendosae and Pristograptus

dubius were reported at the top of the formation in the

Risha area (Paleoservices, 1987). Close to the border

with Saudi Arabia, the Batra Formation is overlain by

the Ratiya Formation; the basal 10–15 m of which

consists of mudstone yielding a sedgwickii Biozone

fauna of the uppermost Aeronian (Powell, 1989).

These data indicate that the base and the top of the

Batra Formation young northwards (Fig. 4).

In the type area, the lower member (persculptus

Biozone age) comprises 17.44 m of black shales

(borehole BG14; locality N2930504 E3557410). They

comprise laminated, black siltstone to dark grey

homogeneous claystone couplets. The parallel lami-

nated siltstones are organic-rich, and grade upwards

into mudstone, with irregular patches of siltstone

(?starved ripples) and homogeneous claystones (Fig.

7). The absence of bioturbation indicates anoxic

bottom water during deposition. The siltstone lam-

inites, characteristic of the whole formation, contain

pyrite and marcasite concretions throughout, confirm-

ing an anoxic depositional environment.

The middle member (persculptus–early acumina-

tus Biozone age) similarly consists of grey and iron-

.

l

.

oxide-rich maroon, parallel-laminated siltstones and

grey homogeneous mudstones that become more

sandy upwards; all units are variably graptolitic

(Powell, 1989). These are petrographically identical

to the turbidites of the lower member, except that the

beds thicken upwards and they have been intensively

oxidised following lithification. The couplets are

interpreted as the Td and Te microfacies of distal

turbidites (Fig. 7).

a

b cTe

Td

T6

T4-5

T3

0.5m faci

esm

icro

faci

es

TeT8

T7

T6

T5

T3

T2

T1T0

T4Td

Fig. 7. Photographs illustrating features of the Batra Formation at outcrop, core and thin section scales. (a) Outcrop photograph showing the

centimetre-scale banding picked out by colour variations. (b) Polished, slabbed core of the lower member of the Batra Formation from well

BG14 showing alternating oxic and anoxic layers. (c) Photomicrograph of a sedimentary couplet from the middle member of the Batra

Formation, scale bar is 8 mm and interpretation follows (inset after Stow and Shanmugam, 1980).


The Batra Formation is synchronous with black

shale deposits around the Gondwana margin (Sutcliffe

et al., 2000; Sutcliffe et al., 2001). In Saudi Arabia, the

subsurface equivalent of the Batra Formation is the

Qusaiba Shale, dated as Ashgill-Ludlow (McClure,

1988) (Fig. 4). The comparable basal bhotQ black

shale facies contains persculptus Biozone grapto-

lites and the overlying Qusaiba Shale yields Rhud-

danian, gregarious Biozone and convolutus Biozone

graptolites.


5. Tests

A critical assessment of depositional models, and

an explanation of the processes of black shale

formation (the Deuser and Strakhov models) in the

peri-glacial setting of Jordan, depend on three tests:

(1) placing these rocks in their correct sequence

stratigraphical context, (2) demonstrating an increase

in TOC and d13Corg (proxies for bioproductivity) up

section, and (3) detecting obliquity forcing in these

proxy data.

Data are drawn from the upper Ordovician to lower

Silurian siliciclastic succession of the Southern Desert

region of Jordan, as well as published and unpub-

lished well data from the Al Jafr, Wadi Sirhan and

Risha areas (Fig. 3). Detailed sedimentary logs and

biostratigraphical analysis provide the framework for

sequence stratigraphical analysis of the succession,

including Fischer plots, following the methodology of

Sadler et al. (1993).13C/12C analyses were performed on de-carbonated

material by combustion in a Carlo Erba 1500 on-line

to a VG TripleTrap and Optima dual-inlet mass

spectrometer, with d13C values calculated to the

VPDB scale using a within-run laboratory standard

(cellulose, Sigma Chemical prod. no. C-6413) cali-

brated against NBS-19 and NBS-22. Replicate anal-

ysis of well-mixed samples indicated a precision of

Fb0.1x (1 S.D.). The carbon isotopic relationships

associated with the production and burial of organic

carbon have been reviewed by Wang et al. (1997).

high

B

Cin

cinn

atia

nM

o.

Ash

gill

Car

adoc

Llan

.

comp.

Hirnantian

Rawtheyan Richmondian

Maysvillian

Edenian

Gamachian

Chatsfieldian

Cautleyan

Pusgillianlinear.

anceps

extra.

acumin.persc.

Grapto-lites

LaurentianStages

Jordanlithostratigraphy

BritishStages

Batra

Tubeiliyat

Dubaydib

Ammar

M

M

U

U

L

L

Fig. 8. Relative sea level curves from Baltica (after Nielsen, 2004), Laurent

lower Silurian biostratigraphy and chronostratigraphy from (Fortey et al

notation on the Jordan sea level curve identifies relative sea level events d

Abbreviations: L, lower; M, middle; U, upper. Graptolie biozones are:

comp., complanatus; linear., linearis.

Total organic carbon was measured on de-carbo-

nated samples using a Leco CS244 carbon analyser.

Typical repeatabilities (relative standard deviations)

were better than 0.7% of the measured carbon values

of analytical standards (0.816% carbon) and 3.8% of

the measured carbon values of the samples. Spectral

analysis of %TOC and d13Corg was conducted using

an algorithm for irregularly spaced data (Press et al.,

1992).

5.1. Test 1—sequence stratigraphy

A relative sea level curve summarizes our sequence

stratigraphical interpretation, which allows correlation

with the better known sections in Baltica and

Laurentia through matching relative sea level events

(Fig. 8, rsl a–e). It is convenient to divide the

discussion into pre-, glacial and post-glacial periods.

5.1.1. Pre-glacial

The Dubaydib and Tubeiliyat Formations are

characterized by wave and storm-dominated shoreface

environments. Deposition was interrupted in the lower

part of the Dubaydib Formation (middle member) by

the development of a regionally extensive, erosively

based channel system (Figs. 4 and 8, rsl a). The

channels have been variously interpreted as rip-current

channels (Makhlouf, 1995) or storm channels (Powell,

1989), but their precise origin and relationship to base-

level change have yet to be determined. Similar

channels, interpreted as subtidal channels, occur at

low

altica

high low

Laurentia

high low

Gondwana (Jordan)

a

bc

de

ia (after Ross and Ross, 1992, 1996) and Jordan. Upper Ordovician–

., 2000). Jordan lithostratigraphy as proposed in this paper. Letter

escribed that can be used to correlate the curve (see text for details).

acumin., acuminatus; persc., persculptus; extra., extraordinarius;


about the same lithostratigraphical level in Libya where

biostratigraphical evidence suggests they are lower

Ashgill in age (Turner et al., 2002). Newly collected

brachiopods from shell lags at the base of the channel-

fill succession in southern Jordan support this age

assignment (Prof. D.A.T. Harper, 2002, pers. comm.).

The Fischer plot for the upper member of the

Dubaydib and Tubeiliyat Formations (cycle thickness

data from Makhlouf, 1992) (Fig. 9a) exhibits two

orders of cyclicity, the individual cycles and a long

term, higher order of cyclicity. The latter defines two

transgressive–regressive cycles within the upper

member of the Dubaydib Formation and the lower

part of the Tubeiliyat Formation (Fig. 8, rsl b; Fig. 9a,

peak HST at cycle 3; Fig. 8, rsl c; Fig. 9a, peak HST at

cycle 8) followed by a general increase in cycle

thickness to the top of cycle 18, close to the top of the

Tubeiliyat Formation.

0.00

5.00

10.00

15.00

20.00

25.00

30.00

35.00

40.00

45.00

0 1 2 3 4 5 6 7 8 9 1011 1213141516171819

Cycle number

Co

rrec

ted

thic

knes

s(m

)

Dubaydib Tubeiliyat

a

b

-600-400-200

0200400600800

1000

0 50 100 150 200

Cycle number

Co

rrec

ted

Th

ickn

ess

(cm

) Lower Member Middle Member

Ordovician Silurian

HST FSSTmfs

Fig. 9. Fischer plots (based on the methodology of Sadler et al.,

1993). (a) Dubaydib–Tubeiliyat formations (raw thickness data

from Makhlouf, 1992). (b) lower part of the Batra Formation.

The Dubaydib and Tubeiliyat Formations span the

Ashgill (~447–443 Ma, sensu Fortey, 2000) and

these orders of cyclicity are considered third- and

second-order. The increase in accommodation space

within the later Ashgill most likely marks the onset

of glacial isostatic subsidence and increased sedi-

ment supply, and we suggest that much of the

deposition in the Tubeiliyat Formation is glacially

forced.

5.1.2. Glacial

The top of the shoreface sediments (Tubeiliyat

Formation) is truncated and incised by two episodes

of channel incision (Fig. 8, rsl b and c). The first

glacial incision and deposition of thin (b2 m),

conglomeratic outwash plain sediments was quickly

followed by marine transgression and deposition of

marine shoreface sandstones which occupy the greater

part of the fill of the lower channels, incised into the

top of the Tubeiliyat Formation.

The second incision (Fig. 8, rsl c) formed a set of

much larger, erosively based channels, now inter-

preted as sub-glacial tunnel valleys (Turner et al.,

2002). The channels have been previously interpreted

as marking the onset of the Hirnantian glaciation,

coincident with a major eustatic fall in global sea-level

of between 50 and 100 m (Woodcock and Smallwood,

1988; Ross and Ross, 1988; Eyles, 1993; Heredia and

Beresi, 1995). The data presented here indicates that

two phases of ice advance and retreat occurred during

the glaciation, within the Ammar Formation. This led

to the development of two relatively two short-lived,

third-order glacially induced lowstands. The interpre-

tation of the stratigraphically equivalent glacial low-

stand elsewhere in North Africa as second order

(Luning et al., 2000) is incompatible with the short-

lived nature of the Hirnantian glaciation (Sutcliffe et

al., 2000).

5.1.3. Post-glacial

The onset of post-glacial transgression (Fig. 8, rsl

d) occurs in the persculptus Biozone. The channel-fill

of the upper member of the Ammar Formation

comprises basal conglomerates and fluvial cross-

stratified sandstones containing flood-generated fin-

ing-upwards cycles, deposited within perennial

braided channels characterised by fluctuating dis-

charge (Abed et al., 1993). Fluvial sandstones fine

17

a

c

d

b

Fig. 10. Temporal and cyclic variation in TOC and organic carbon isotopic values for the lower Batra Formation. (a) d13Corg plotted against

height above the base of the Batra Mudstone Formation. (b) TOC plotted against height above the base of the Batra Mudstone Formation. (c, d)

Spectral analysis of %TOC and d13Corg using an algorithm for irregularly spaced data (Press et al., 1992). 95% CL=one sided 95% level

confidence levels (Mann and Lees, 1996), bw=band width. Abbreviations: OM, organic matter; TOC, total organic carbon.


upwards into marine sediments with the transgressive

surface (ts) placed at the base of the marine sediments

in the channel-fill succession (Fig. 5).

The base of the Batra Formation therefore repre-

sents the maximum flooding surface external to the

depocentre. This would be equivalent to the ravine-

ment surface that developed when interfluves within

the existing basin floor topography were breached,

with the resulting rapid, lateral expansion of basinal

anoxic conditions. This is seen in the diachronous

onset of black shales northwards, towards the glacial

fore-bulge, between the Hirnantian and early Rhud-

danian. The black shales of the lower Batra For-

mation, therefore, formed between the maximum

flooding surface and the base of the falling stage

systems tract (FSST), and we interpret them as

maximum flooding (MF) black shales that developed

within an expanding puddle.

The Fischer plot for the lower part of the Batra

Formation records an increase–decrease–increase in

accommodation space (Fig. 9b). The initial increase

reflects the continuing transgression. The decline in

accommodation space (bregressionQ) from the peak of

highstand (cycle 11) through into the organic-poor

shales of the middle member could be interpreted in

two ways in a four system model. This decline could

be the result of relative sea level fall, the FSST, or

alternately, it could represent increasing sediment

starvation as the transgression continued. The main

sedimentological change occurs at the base of the

middle member of the Batra Formation with the

appearance of grey, organic-lean shales. These are still

microlaminated (i.e. no bioturbation) indicating oxi-

dation was not syn-depositional. A second major

sedimentological change occurs within the middle

member where turbidites increase in thickness and

grain-size (Fig. 9b, cycle 115; Fig. 8, rsl e). These

changes are interpreted as recording increased sedi-

ment input to the basin reflecting relative sea level fall

at the start of the FSST. Global eustatic sea-level

curves (Ross and Ross, 1996) indicate that this

second-order sea-level rise continued until the late

Rhuddanian or early Aeronian. However, Melchin et

al. (1998) highlighted a possible regression during the

late acuminatus Biozone and this could correlate with

rsl e (Fig. 8).


5.2. Test 2—temporal variation in TOC and d13Corg

d13Corg values from the lower member of the Batra

Formation (well BG14) show metre-scale oscillations

and a rising trend of �31x to �29.5x (Fig. 10a).

TOC values, taken as a proxy for bioproductivity,

increase cyclicly from 2% to 8% (Fig. 10b). TOC

increases in three stages up section. This corresponds

to 2.7 cycles through the TOC time series.

5.3. Test 3—obliquity forcing of TOC and d13Corg

Spectral analysis indicates regular cyclicity in the

TOC data, with a wavelength of 6.38 m (Fig. 10c), but

shows no significant cyclicity in the d13Corg data (Fig.

10d). However, overall, there is a general trend to

more positive d13Corg values and increasing TOC.

Milankovitch theory predicts the cycles in the TOC

data will reflect orbital obliquity. The Ordovician

orbital-obliquity cycle period was ~36,000 years

(Berger et al., 1989) and the time series has a

minimum duration of ~97,000 years.

This result can be corroborated using stratigraph-

ical thickness data from the international boundary

stratotype at Dob’s Linn. Here the ratio of the

stratigraphical thickness of the first post-glacial high-

stand (negative shift in d13Corg during the persculptus

Biozone) to strata representing the glacial maximum

(~extraordinarius Biozone) is 0.72 m (thickness data

from Underwood et al., 1997). Assuming sedimenta-

Low salinity layer

High deglacialrun-off of nutrients

and freshwater

Organic detritus

HST -organic-richblack shale

(Batra)TST - Fluvial to marine sandstone

(Upper Ammar channel fill)

A combination of the van Cappellan effectwith the rising monimolimnion plus addition

of organic detritus from mixolimnion

a

mfs

b

halocline

Fig. 11. Conceptual model for the upper Ordovician organic-rich bhotQ shapuddle during the HST. Abbreviations as in previous figures.

tion rate was constant and the extraordinarius

Biozone represents two eccentricity cycles, as envis-

aged by Sutcliffe et al. (2000), then the duration of the

HST is 144 kyr.

6. Model for the formation of black shales during

deglaciation

Luning et al. (2000) reviewed the occurrence of

anoxic, black shale facies throughout North Africa

and the Arabian Peninsula offering a three-system,

sequence stratigraphical model characterised by the

succession in Libya. They placed the transgressive

surface at the base of the black shales. However, the

interpretation of the early post-glacial black shale as

BT shales (Luning et al., 2000) is inconsistent with the

field evidence and stacking patterns seen in the

equivalent Jordan succession, where the black shale

lies on the mfs. Initially organic-rich black shale was

deposited in glacially incised shelf sub-basins. As the

deglacial transgression progressed, sub-basin margins

and interfluves were eventually breached producing a

rapid increase in accommodation space. Low sedi-

mentation rates meant that this space could not be

filled and the basin became starved. The spread of

black shale away from the depocentres within

glacially incised shelf depressions or sub-basins was

diachronous. We therefore contend that black shales

are MF shales (sensu Wignall, 1991) and that they

Low salinity layer

Deglacialrun-off of nutrients

and freshwater

HST -organic-richblack shale

(Batra)TST - Fluvial to marine sandstone

(Upper Ammar channel fill)

Situation at peak highstand, the"expanded puddle," Strakhov andvan Cappellen effects continue

mfs

Organic detritus

halocline

le in Jordan. (a) Initiation of the anoxic puddle. (b) Expansion of the


fulfil the diagnostic criteria of the expanding puddle

model (Fig. 11).

Anoxic conditions could have been initiated in two

ways: (1) through the upward expansion of a warm

salinity dense oceanic monimolimnion that had

remained anoxic throughout the glaciation or (2)

formation beneath a stratified water column within a

sub-basin on the shelf. In the second scenario, it is

likely that, during initial stages of marine incursion,

nutrients in the monimolimnion remained isolated

from the euphotic zone by the halocline, the gradual

rising monimolimnion released nutrients to the mix-

olimnion (van Cappellen effect), and triggered a high

productivity episode which initiated anoxia. By

analogy with the present Antarctic margin and the

Black Sea, it would seem more likely that stratifica-

tion probably resulted from a salinity gradient with

more saline waters at depth. The fresher surface

mixolimnion would have resulted from ice meltwater

generated during the earliest stages of deglaciation,

coincident with a reduction in density driven thermo-

haline circulation as the ice margin retreated.

Increasing TOC and d13Corg up section indicate

that once euxinic conditions were established, photic

zone bioproductivity was the primary cause of

sustained anoxia supporting the Strakhov hypothesis.

The supply of meltwater-derived nutrients was

directly related to the seasonal melting dynamics of

the ice sheet and was forced by the predicted changes

in orbital obliquity. Anoxia was sustained in this way

over tens–hundreds of thousands of years.

7. Conclusions

The onset of Hirnatian glaciation in Jordan is

marked by the Ammar Formation that lies in the

proximal part of the glacial depositional system. In

Jordan, the second of two glacial incision events

resulted in steep, U-shaped tunnel valleys up to 4 km

wide and 70 km long, trending NNE–SSW, parallel to

palaeo-iceflow. These were to become the initial

depocentres for black shale deposition. The outcrops

of the Ammar Formation delimit the proximity of the

ice margin, at or near the grounding line.

Sequence stratigraphical analysis indicates that the

succeeding organic-rich black shale of the Batra

Formation are MF shales that formed beneath a

salinity stratified water column. Initially anoxia was

developed beneath the halocline as a consequence of a

cap of low density glacial meltwater, declining

thermohaline circulation and the ingress of normal

marine waters at depth. Increasing values of TOC and

d13Corg upwards indicates that once anoxia was

established it was sustained by mixolimnion biopro-

ductivity which itself was sustained by continued and

increasing levels of meltwater-derived nutrients. Ice

sheet melting was modulated by orbital obliquity. As

the deglacial transgression continued, sub-basin mar-

gins were eventually breached resulting in the rapid

lateral expansion of the black shale facies. This

scenario is best encapsulated in the expanding puddle

model for black shale deposition.

Modern day processes, operating in silled basins,

can be used to explain the development of organic-

rich black shales on the peri-glacial Gondwana shelf

in Jordan. Peri-glacial settings are fully oxygenated at

the present day. What distinguishes the Ordovician

black shales in Jordan is the initiation anoxia of early

in the deglacial transgression. Either, the shelfal

basins in which these sediments formed were highly

restricted silled basins, comparable to the Black Sea at

the present day and low levels of oxygen in the

atmosphere–ocean pre-disposed Ordovician shelf seas

to anoxia or, the Hirnantian glaciation was insufficient

to develop an intense thermohaline circulation that

could ventilate the world’s oceans.

Acknowledgements

This research was supported by the U.K. Natural

Environmental Research Council (NERC Grant

NER/B/2000/000068), the University of Durham

and the Hashemite University. We are also grateful

to the Natural Resources Authority of Jordan for

logistical support and providing core samples. TOC

analyses were conducted by Dr. M. Jones, NRG,

University of Newcastle upon Tyne. Organic carbon

isotopes analyses were conducted by Dr. M. Leng,

NIGL, British Geological Survey. Prof. M. E. Tucker

read and improved an earlier version of the manu-

script. We are indebted to journal reviewer J.

Ghienne and editor Finn Surlyk for their helpful

comments and suggestions. This paper is a contri-

bution to IGCP Project 503.


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