doi: 10.1111/j.1365-3121.2012.01067.x
13
A syn-depositional age for EarthÕs deepest d C excursion
required by isotope conglomerate tests
Jon M. Husson, Adam C. Maloof and Blair Schoene
Department of Geosciences, Princeton University, Guyot Hall, Washington Road, Princeton, NJ 08544, USA
ABSTRACT
The most negative carbon isotope excursion in Earth history is
found in carbonate rocks of the Ediacaran Period (635–542 Ma).
Workers have interpreted the event as the oxidation of the
Ediacaran oceans [Rothman et al., Proc. Natl. Acad. Sci. USA 100
(2003) 8124; Fike et al., Nature 444 (2006) 744; McFadden et al.,
Proc. Natl. Acad. Sci. USA 105 (2008) 3197], or as diagenetic
alteration of the d13C of carbonates (d13Ccarb) [Knauth and
Kennedy, Nature 460 (2009) 728; Derry, Earth Planet. Sci. Lett. 294
(2010) 152]. Here, we present chemo-stratigraphic data from the
Ediacaran-aged Wonoka Formation (Fm.) of South Australia that
require a syn-depositional age for the extraordinary range of
d13Ccarb values ()12 to +4&) observed in the formation. In some
Introduction
The Ediacaran Period (Knoll et al.,
2006) is the bridge between the Proterozoic world and the animal-abundant
Phanerozoic. Sponges (Love et al.,
2009; Maloof et al., 2010; Sperling
et al., 2010; Brain et al., 2012) appear
before the 635 Ma (Hoffmann et al.,
2004; Condon et al., 2005) terminal–
Cryogenian ice age, but decimetrescale organisms (animals, giant
protists and macro-algae known as
the Ediacaran Biota (Xiao and Laflamme, 2009)) do not appear until
579 Ma (Bowring et al., 2003).
Broadly synchronous with these first
appearances, the deepest carbon isotope excursion in Earth history is
recorded in carbonates from at least
four continents – most famously from
Oman (Burns and Matter, 1993),
South Australia (Calver, 2000), South
China (McFadden et al., 2008) and
southwestern USA (Corsetti and
Kaufman, 2003). The extreme isotopic
depletion seen in these Ediacaran
basins is colloquially referred to as
the ÔShuramÕ anomaly, although global
synchroneity has not been established
independently (Grotzinger et al.,
2011). Chemostratigraphic (Prave et
al., 2009) and sparse geochronological
Correspondence: Jon M. Husson, Department of Geosciences, Princeton University,
Guyot Hall, Washington Road, Princeton,
NJ 08544, USA. Tel.: (609) 258-0836; fax:
(609) 258-1274; e-mail: jhusson@princeton.
edu
2012 Blackwell Publishing Ltd
locations, the Wonoka Fm. is 700 metres (m) of mixed shelf
limestones and siliclastics that record the full 16& d13Ccarb
excursion. In other places, the Wonoka Fm. is host to deep
(1 km) palaeocanyons, which are partly filled by tabular-clast
carbonate breccias that are sourced from eroded Wonoka
canyon-shoulders. By measuring the isotopic values of 485
carbonate clasts (an isotope conglomerate test), we show that
canyon-shoulder carbonates acquired their d13Ccarb–d18Ocarb
values before brecciation and redeposition in the palaeocanyons.
Terra Nova, 00, 000–000, 2012
(Condon et al., 2005; Bowring et al.,
2007) data suggest that the globally
observed anomalies, synchronous or
not, are hosted in sediments younger
than the 580 Ma Gaskiers glaciation
(Bowring et al., 2003) and older than
551 Ma (Condon et al., 2005).
The dominant paradigm among
chemostratigraphers is that d13Ccarb
in carbonate rocks reflects the d13C of
dissolved inorganic carbon (DIC) in
contemporaneous sea water, which is
virtually uniform around the world
due to the long residence time of DIC
and the short mixing time of the
oceans (Kump and Arthur, 1999).
Similarities in shape and magnitude,
and the broadly Ediacaran age, have
been used to argue that the ÔShuramÕ
anomaly fits this model, and is therefore globally synchronous and can be
used for global stratigraphic correlation (Halverson et al., 2005). A primary DIC origin for the ÔShuramÕ is
also the foundation for models involving stepwise oxidation of the terminal
Neoproterozoic ocean (Rothman
et al., 2003; Fike et al., 2006; McFadden et al., 2008) that resulted in
pulsed inputs of light carbon to the
ocean-atmosphere system (Rothman
et al. (2003), and references therein).
Absolute time constraints, although
not available, are necessary to quantify the light carbon fluxes and oxidant budget required, and thus test
the modelÕs viability.
Alternatives to the global-DIC
hypothesis have sought to explain
the ÔShuramÕ anomaly, and include
(1) meteoric and (2) burial diagenesis
models. Under the meteoric model,
the depleted d13Ccarb values and covarying d13Ccarb–d18Ocarb (features of
most, but not all, Shuram-like anomaly profiles) are interpreted as a record
of remineralising fluids charged with
DIC issued from organic matter respiration in soils (Knauth and Kennedy, 2009; Swart and Kennedy,
2012). The burial diagenesis model
invokes post-burial fluid-rock interactions at depth, wherein a high pCO2,
low d13C fluid, developed from buried
organic matter, mixes with an 18O-rich
basinal brine (Derry, 2010). Both
diagenesis models therefore argue for
acquisition of negative d13Ccarb after
the carbonate sediments had been
deposited (although in the meteoric
case, alteration can happen immediately after deposition). If the ÔShuramÕ
anomalies are a record of diagenesis,
one must ask (1) what type of process
would lead to diagenetic alteration of
d13Ccarb, synchronously or not, in
Ediacaran basins around the world,
and (2) why might Ediacaran sediments be uniquely predisposed to such
intense alteration (Grotzinger et al.,
2011).
We present data from six measured
stratigraphic sections (Fig. 1B,C)
from the Ediacaran Wonoka Fm. of
the Adelaide Rift Complex (ARC) of
South Australia that (1) demand syndepositional acquisition of its low
d13Ccarb values (down to )12&) and
the observed covariation between
d13Ccarb and d18Ocarb, (2) rule out the
1
Isotope conglomerate test • J. M. Husson et al.
Terra Nova, Vol 00, No. 0, 1–8
.............................................................................................................................................................
N
139
o
o
140
LITHOFACIES KEY
Debris flow breccia
Intraclast breccia
Microbialite
19
mo
ADELAIDE
RIFT
COMPLEX
m POUND
300
(B)
800
700
600
7
500
6
400
5
300
4
200
200
3
0
2
1
9
10
5, 6
900
C ANYON-SHOULDER SECTION
10
9
8
100
C ANYON-FILL SECTION
19
67
84
–20 –15 –10 δ18O, δ13C 5
10
100
WONOKA FORMATION
400
WONOKA FORMATION
500
WONOKA SECTION LOCATIONS
Measured sections
Canyon-fill section
Canyon-shoulder section
Bunyeroo Gorge section
(from Calver 2000)
(A)
Grainstone/sandstone
1000
Ribbonite/ribbonite with silt interbeds
Siltstone
700
600
o
MAP LEGEND
Undifferentiated sediments
Cambro-Ordovician volcanic rocks
HAWKER – LAKE FROME
GROUPS (542–505 Ma)
WILPENA GROUP (635–542 Ma)
UMBERATANA GROUP (~710–635 Ma)
BURRA GROUP (~777–710 Ma)
CALLANA GROUP (~840–777 Ma)
Undifferentiated PalaeoMesoproterozoic rocks
Major fault
Turbiditic sandstones
m POUND
800
11
31
84
nges
Ra
Flinders
km
Gam
10
67
Stuart
Shelf
nR
9
Map area 80
of Fig. 3
o
N
ang
es
5,6
30
0
Basal carbonate breccias
with clast δ13C range
(see Fig. 2C, 5A)
(C)
–20 –15 δ O, δ C 0
18
13
5
Fig. 1 (A) Simplified geological map (adapted from Preiss and Robertson (2002) and Rose and Maloof (2010)) of the study area
within the Adelaide Rift Complex (ARC). Locations of measured stratigraphic sections are denoted by yellow symbols and
labelled with numbered white squares. Representative physical stratigraphy of (B) a canyon-shoulder section and (C) a canyon-fill
section, paired with d13Ccarb and d18Ocarb (in & notation reported relative to the Vienna Pee Dee Belemnite carbon isotope
standard) data from six localities. (canyon-shoulder sections: 19 – Parachilna, 67 – Saint Ronan, 84 – Black Range Spring; canyonfill sections: 5–6 – Mount Thomas, 9 – Oodnapanicken, 10 – Mount Goddard). Wonoka Fm. carbonates consist mostly of
limestone; dolomite only is found at its base as the thin (0.1)1 m), but persistent, ÔWearing DolomiteÕ (Haines, 1988), and in the
uppermost 150–200 m (units 9 and 11). Haines (1988) subdivided the Wonoka Formation into 11 informal lithological units, which
are broadly recognisable throughout the central Flinders Ranges and labelled here on our canyon-shoulder sections. The chemo–
and lithostratigraphies of canyon-shoulder and canyon-fill sections are remarkably similar from locality to locality, although the
thickness of canyon-fill sections (500–1000 m) is more variable than canyon-shoulder sections (700 ± 100 m), depending upon
the depth of canyon incision.
burial diagenesis model, and (3) constrain the styles of meteoric diagenesis
that could be invoked to explain the
observations from Australia.
2
Geological setting
The ARC (Fig. 1A) was part of a
continental margin formed to the
present-day east of the Stuart Shelf,
and the Wonoka Fm. is part of the
Ediacaran Wilpena Group (Preiss,
2000; Preiss and Robertson, 2002).
2012 Blackwell Publishing Ltd
J. M. Husson et al. • Isotope conglomerate test
Terra Nova, Vol 00, No. 0, 1–8
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The Nuccaleena Formation, the distinctive cap dolostone to the glacial
Elatina Formation (Plummer, 1979;
Williams, 1979; Rose and Maloof,
2010), forms the base of the Wilpena
Gp., and frequently is correlated with
the younger Cryogenian glacial cap
units found around the world (Halverson et al., 2005) (dated to be
635 Ma in Namibia (Hoffmann
et al., 2004) and South China (Condon et al., 2005)). The Wonoka Fm.
varies between 400 and 1500 metres
(m) thick, and coarsens and shallows
upwards into the siliclastic, Ediacara
Biota-bearing Pound Subgroup. At its
base in the central Flinders Ranges
(e.g., 19, 67, and 84 in Figs 1B and
2D), the Wonoka Fm. is a deep-shelf
sequence of turbidites with climbingrippled sands and allodapic (i.e.,
transported down-slope) carbonate
beds. The Wonoka Fm. transitions
upwards into a shallower, storm-dominated mid-shelf sequence containing
(A)
abundant hummocky and swaley
cross-stratified sands (Fig. 2E), and
wavy limestone laminites and grainstones. At the top of these units, the
Wonoka Fm. coarsens abruptly into
thick (50–100 m), trough cross–bedded fine-to-medium grained sandstones (unit 10 under the terminology
of Haines (1988); see Fig. 1 caption).
In contrast, the northern Flinders (5,
6, 9, 10 in Fig. 1C) record deeper
outermost-shelf and slope settings,
with 1000-m deep palaeocanyons
that cut into the underlying Bunyeroo
and Brachina formations (Fig. 2A,B)
(Haines, 1988; Christie-Blick et al.,
1990; DiBona and von der Borch,
1993; Giddings et al., 2010) and are
filled with mixed carbonate and siliciclastic turbidites and tabular-clast
limestone breccias (Fig. 2C). Henceforth, we refer to outer shelf sections
as Ôcanyon-shouldersÕ, and sections
with palaeocanyons as Ôcanyon-fill.Õ
The uppermost Wonoka Fm. is char-
acterised by 80–100 m of microbialite and stromatolite bioherms (unit 11
of Haines (1988)), which blanket the
canyon-shoulders and some canyonfill sections (Fig. 3).
Methods
Carbonates were sampled at 1.0 m
resolution whilst measuring six stratigraphic sections from across the
Adelaide Rift Complex (ARC; Fig. 1
A). Clean limestones and dolostones
with minimal siliciclastic components
were targeted. A total of 1049 samples were slabbed and polished perpendicular to bedding and 5 mg of
powder were micro-drilled from individual laminations for isotopic analysis. Sections 5–6, 9, 10, and 19
(Fig. 1B,C) were measured at the
University of Michigan Stable Isotope Laboratory on a Finnigan MAT
Kiel I preparation device coupled
directly to the inlet of a Finnigan-
Paleaeocurrent (B)
directions
N
2m
Wonoka Fm
Brachina
Wonoka
(C)
5 km
Brachina Fm
(D)
Wonoka Fm
100 m
Bunyeroo Fm
(F)
(E)
10 cm
Fig. 2 (A) Google Earth image (acquired 2 ⁄ 26 ⁄ 2010) of a 1000-m-deep Wonoka palaeocanyon, with palaeocurrent directions
from Eickhoff et al. (1988) – photos B and C are located on the north side of the palaeocanyon near the ÔWÕ in the Wonoka Fm.
label in (A); (B) angular and erosive contact (black line) between Brachina Formation canyon wall and Wonoka Formation
canyon-fill (white dashed lines depict change in bedding attitude) at Mount Thomas (sections 5–6); (C) a 2-m-thick debris flow
breccia, dominated by tabular clasts (1–75-cm long) of micritic marine limestone; (D) canyon-shoulder outcrop near Parachilna
Gorge (19); (E) small-scale hummocky cross-stratified (HCS) carbonate grainstone typical of canyon-shoulder outer shelf facies
from Saint Ronan (67); (F) 20–40-cm thick grainstone (red in colour)-to-microbialite (yellow) couplet typical of upper Wonoka
Fm. stratigraphy at Parachilna Gorge (19).
2012 Blackwell Publishing Ltd
3
Isotope conglomerate test • J. M. Husson et al.
Terra Nova, Vol 00, No. 0, 1–8
.............................................................................................................................................................
N
Pound subgroup
Pound subgroup
unit 11
Upper canyon sandstone
Wonoka Fm.
.
Basal breccias with
1 km
Bun
yero
o
80
Unit 11
Unit 10
.
Fm.
Canyon-shoulder
silt interbeds
Bunyeroo Fm
Won
oka
Fm
Units 1 - 7
67
Fig. 3 Bing Maps image (acquired 11 ⁄ 17 ⁄ 2004) of the Saint Ronan canyon-fill and canyon-shoulder complex (map area marked
with black outline in Fig. 1A). Black lines indicate conformable contacts between units or formations, while white lines indicate
sequence boundaries. Isotope conglomerate data of section 80 (Fig. 5B) come from breccias located within the red box in the
bottom–left of the figure. Field mapping shows that the microbialite reef facies of unit 11 caps both the canyon-fill and canyonshoulder sequences, thus indicating that canyon cutting and filling occurred before terminal Wonoka deposition.
(A)
and measured precision is 0.1 (1r)
for both values. For a more thorough
discussion of these methods, see Rose
et al. (2012).
Results
The canyon-shoulder sections of Parachilna Gorge, Saint Ronan and Black
Range Spring (19, 67, and 84 in
Fig. 1B) exhibit very low d13Ccarb
n
values and d13Ccarb–d18Ocarb covariation (see Fig. 4A). Isotopic values are
most negative at the base of unit 3
(down to )12&), and the profile
recovers smoothly and gradually towards positive d13Ccarb values. At the
top of unit 8, the curve crosses 0&,
and d13Ccarb remains positive in units
9–11. The isotopic profiles of canyonshoulder units 1–11 are remarkably
consistent between sections over
(B)
C
Fi
MAT 251 triple collector isotope
ratio mass spectrometer. All other
samples were measured at Princeton
University on a GasBench II preparation device coupled directly to the
inlet of a Thermo DeltaPlus continuous flow isotope ratio mass spectrometer. d13C and d18O data are
reported in the standard delta notation as the difference from the VPDB
standard (Vienna Pee Dee Belemnite),
C
1 mm
(C)
P
S
G
I
M
W
W
Fig. 4 (A) d13Ccarb – d18Ocarb cross plot showing data from all stratigraphic sections, coded by section number, lithofacies, and
lithology (ÔlsÕ stands for limestone; ÔdlÕ for dolomite). All canyon-shoulder sections show d13Ccarb–d18Ocarb correlation, especially
for d13Ccarb values less than )5& (r2 = 0.49, 0.49 and 0.44 for 19, 67 and 84 respectively; for all sections, P<<0.001). Such
correlation is not observed in the fine-grained, allodapic carbonates of the canyon-fill sections (r2 = 0.003 and P = 0.285 for all
three sections combined). In the figure legend, lithofacies are organised in order of increasing permeability (i.e., ranging from finegrained micritic wavy laminites to coarse grainstones and sandy carbonates). No pattern of dependence between isotopic values
and lithofacies is observed. (B) Photomicrograph of a cm-scale fining–upward carbonate turbidite from the base of the Saint
Ronan canyon-shoulder section (67; see Fig. 1B) that shows the lack of coarse recrystallisation observed in Wonoka Fm.
carbonates. Even the fine-sand fraction of carbonate grains, which are most susceptible to recrystallisation are preserved in
primary sedimentary textures (e.g., the dashed white line marks the erosive beginning of another turbiditic sequence). These fabrics
contrast with rocks from unit 11 (C), where the first evidence of recrystallisation and growth of dolomite rhomboids is observed
(one such crystal is outlined (C), with dashed lines depicting orientation of cleavage). Dolomites are only found in the upper 150–
200 m of the Wonoka Fm., where carbon isotopic values are at their most positive (+2 to +8&). The isotopic values of (B) and
(C) are depicted in panel (A) with lettered symbols; both samples are 85% carbonate by weight.
4
2012 Blackwell Publishing Ltd
J. M. Husson et al. • Isotope conglomerate test
Terra Nova, Vol 00, No. 0, 1–8
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50 kilometres distant (Fig. 1B). The
lower half of most canyon-fill units is
carbonate poor, except for the frequent tabular-clast carbonate breccias
(Fig. 2C). Fine-grained allodapic carbonate beds are common above 300–
400 m of basal canyon-fill, and
d13Ccarb values from Mount Thomas,
Mount Goddard and Oodnapanicken
(5–6, 9, 10 on Fig. 1C) vary between
)8.5 and )5&. Canyon-fill sections
therefore do not record (1) the nadir
seen in the canyon-shoulders, (2)
d13Ccarb–d18Ocarb covariation (r2 =
0.003 and P = 0.285 for all three
datasets combined), or (3) the recovery to positive d13Ccarb values.
In palaeomagnetics, a conglomerate
test is used to determine whether clasts
in a conglomerate were magnetised
prior to transport and deposition
(preserving random magnetic directions) or after deposition (preserving
uniform magnetic directions, despite
random clast orientations). Analogously, we performed an isotope conglomerate test on the breccia units of
Mount Thomas, Oodnapanicken, and
Saint Ronan canyon-fill (5–6, 9, and
80, respectively, on Fig. 5A,B) to
assess the provenance and relative
timing of acquisition of d13Ccarb and
δ13Ccarb (‰)
0
n = 108
r 2= 0.593
P<<0.01
Basin-wide
range of canyon
shoulder
(units 1 – 9)
Limestone
Dolomite
Stratigraphic heights*
(in m)
upper
canyon
574.1
328.2
105.1
4.7
(see 5–6;
9; 10 on Fig. 1C)
–15
Parachilna shoulder
19; see Fig. 1B
9 5,6
–15
Fig. 1B). More distal sourcing cannot
be ruled out, however, because the
canyon-shoulder d13Ccarb profile is
remarkably
reproducible
across
20 000 km2 of basin map-area (Fig. 1
A,B).
Discussion
We interpret the breccia clasts to be
carbonates sourced from units 1–9 of
canyon-shoulder localities (e.g., 19,
67, and 84 on Fig. 1B). In certain
palaeocanyons, 0.05 to 5 m-thick
grey-to-yellow micritic limestone layers are found along the edge of the
canyon wall (Fig. 2A,B). This carbonate is of debated origin (Eickhoff
et al., 1988; Giddings et al., 2010),
and it may represent an additional
source of clasts for the basal breccia
beds of Mount Thomas and Oodnapanicken, although the observed
d13Ccarb range of this in situ carbonate
()6.5 to )9.3&) (Giddings et al.,
2010) covers less than one-fourth of
the full range observed in the carbonate breccias. These limestone beds are
not present at Saint Ronan, where the
lithofacies, colour and isotopic composition of the clasts are identical to
units 1–7 (Fig. 1B) of the canyon-
Saint ronan shoulder
67; see Fig. 1B
n = 377
P<<0.01
r 2= 0.457 Stratigraphic
heights*
(in m)
0
Range of allodapic
–10
5 (B)
18
δ Ocarb (‰)
–5
176.2
98.0
44.3
11.3
0
δ13Ccarb (‰)
5 (A)
d18Ocarb by measuring the isotopic
values of carbonate clasts. The breccia
units range from 0.2 to 11 m in
thickness, are clast-supported in a
matrix of fine sand (Fig. 2C), and are
most common at the base of the
canyon-fill stratigraphy. At Mount
Thomas and Oodnapanicken, the tabular, 1–100-cm long carbonate clasts
consist of two dominant lithologies:
grey, micritic limestone and brown
dolostone. At Saint Ronan, the clasts
are distinctly different from Mount
Thomas and Oodnapanicken breccia
clasts, and identical to the green, wavy
laminites and coarse red limestone
grainstones from the adjacent canyon-shoulder strata (Fig. 4). Individual breccia units record d13Ccarb
variability of 6 and 16& and significant d13Ccarb–d18Ocarb covariation
(Fig. 5A,B). The d13Ccarb range of
canyon-shoulder units 1–9 ()12 to
+4&; 19, 67 and 84; Fig. 1B) is seen
in the breccia units of the more distal
canyon-fill of Mount Thomas and
Oodnapanicken (5–6 and 9 on Fig. 5
A), while the range ()9 to )3&) of the
more proximal canyon-fill of Saint
Ronan (Fig. 5B) matches that of the
locally eroded canyon-shoulder (the
first 550 m (units 1–7) of J67 on
14.1
13.1
12.2
11.3
10.3
9.4
8.9
5.2
4.2
3.3
2.8
1.9
0.5
Range of locally
eroded shoulder
(see 67
on Fig. 1B)
–10
–15
relative to the base of palaeocanyons
80
–15
δ18Ocarb (‰)
–5
0
Fig. 5 Results of isotope conglomerate tests from Mount Thomas and Oodnapanicken canyon-fill [labelled 5–6 ⁄ blue and 9 ⁄ red
respectively, in (A)] and Saint Ronan canyon-fill (labelled 80 ⁄ green in (B)). If the clasts acquired their d13Ccarb values and d13Ccarb –
d18Ocarb correlation in situ on canyon-shoulders, then they should exhibit a random collection of values representing the full
isotopic range present on the canyon-shoulder at the time of canyon filling. In contrast, if the extremely negative d13Ccarb (down
to )12&) and d13Ccarb–d18Ocarb correlation in canyon-shoulders are a result of post-depositional diagenesis, then clasts from
individual breccia beds should either (1) reflect the original pre-diagenetic isotopic values in the shoulder (i.e., not extremely
depleted in d13Ccarb), or (2) a consistent diagenetic value homogenous within breccia units. The clasts of the more distal,
Oodnapanicken and Mount Thomas canyon-fill exhibit the full range of d13Ccarb values observed in units 1–9 in canyon-shoulder
sections throughout the basin (19 and 84; Fig. 1B). The more proximal Saint Ronan canyon-fill appears to have sampled a smaller
canyon-shoulder range; breccias exhibit a smaller d13Ccarb range, but the clasts match the colour, lithologies and isotopic range
observed in units 1–7 of the immediately adjacent canyon-shoulder (67; Fig. 1B; Fig. 3). More distal sourcing, however, is still
possible because the canyon-shoulder d13Ccarb profile is remarkably consistent across the ARC (Fig. 1A,B).
2012 Blackwell Publishing Ltd
5
Isotope conglomerate test • J. M. Husson et al.
Terra Nova, Vol 00, No. 0, 1–8
.............................................................................................................................................................
shoulder section 6.5 km distant. Thus,
these observations require that both
the negative d13Ccarb values and the
d13C–d18O covariation were acquired
in Wonoka Fm. canyon-shoulder carbonates before those carbonates were
brecciated and redeposited in the palaeocanyons.
While the mechanism for the formation of the Wonoka canyons remains
controversial
[subaerial
(Christie-Blick et al., 1995) vs. submarine (Giddings et al., 2010)], there
has been broad agreement that canyon incision occurred early during
Wonoka Fm. deposition (placed variously at a cryptic unconformity within HainesÕ canyon-shoulder units 1–5
(Christie-Blick et al., 1995; Giddings
et al., 2010); see Fig. 1B). Given the
presence of carbonate clasts with
positive d13Ccarb values in canyon-fill
breccia units, our results require a
higher surface (i.e., above unit 8) to be
responsible for at least some of the
canyon cutting at Mount Thomas and
Oodnapanicken. We do not claim,
however, that this proposed surface
represents all canyon cut–fill sequences within the Wonoka Fm., as
workers have documented numerous
intervals of canyon incision throughout Wonoka Fm. deposition (DiBona
and von der Borch, 1993). We therefore propose placing a sequence
700 (A)
Canyon-shoulder
boundary at the abrupt appearance
of unit 10 sandstones, where units 8
and 9 are variably absent (Fig. 3) and
the d13Ccarb profile first recovers to
positive d13Ccarb values. Thus, the
canyon-fill sampled units 1–9 of intact
canyon-shoulder stratigraphy, whose
carbon isotopic range ()12 to +4&)
matches that of the tabular-clast breccias. The most positive d13Ccarb observed from the Saint Ronan breccias
is )3&, corresponding isotopically to
the top of unit 7 of the intact canyonshoulder stratigraphy. Thus, the canyon cutting surface could be placed
lower at Saint Ronan, so long as the
canyon remained open to sediment
infill throughout the deposition of unit
7 on the canyon-shoulder.
As the canyons filled completely,
the fine-grained, allodapic carbonate
beds of the upper canyon-fill (upper
700 m of Fig 1C) represent a recycled
and homogenised canyon-shoulder
section, producing d13Ccarb profiles
(5–6, 9, and 10 in Fig. 1C) that are
stratigraphic averages of a canyonshoulder profile (Fig. 6). To test this
hypothesis, linearly interpolated datasets were created for representative
canyon-shoulder (units 1–9; 19 –
Parachilna; see Fig. 6A) and finegrained canyon-fill d13Ccarb curves
(10 – Mount Goddard; Fig. 6B). Linear interpolation serves to weight
0.6
Fine-grained canyon-fill
(B)
0.5
600
Metres
Probability
500
0.4
Interpolated
data
Measured
data
100
0
–10
–5
13
0
δ Ccarb
5
Interpolated
data
Measured
data
–10
–5
13
0
δ Ccarb
5
(C)
Canyon clasts
Fine−grained
canyon fill
Canyon
shoulder
µ = −7.4
σ = 0.7
n = 472
0.3
0.2
200
d13Ccarb values by the stratigraphic
thickness over which they occur, thus
accounting for the stratigraphic
weighting that occurs during brecciation (i.e., thicker units will contribute
more breccia clasts). The distributions
of the Parachilna and Mount Goddard interpolated datasets overlap,
with the variance of Mount Goddard
being much smaller. This result is
expected, as mixing and homogenisation of a canyon-shoulder profile at a
scale below that sampled for isotopic
measurement should decrease the variance, but produce a similar mean;
where canyon-fill is not well mixed at
the scale of isotopic measurement (i.e.,
breccia clasts), the variance of the
distribution is much larger (Fig. 6C).
These results are thus consistent with
the model that the fine-grained, allodapic carbonates of Mount Thomas,
Oodnapanicken, and Mount Goddard
(5–6, 9, and 10 on Fig. 1C) represent
recycled, redeposited canyon-shoulder
units 1–9.
These stratigraphic and isotopic
observations preclude burial diagenesis models for the negative isotopic
values (Derry, 2010) in South Australia. The low d13Ccarb values and
d13Ccarb–d18Ocarb covariation of the
Wonoka Fm. must either be primary
or a relatively early meteoric diagenetic signal (i.e., before the Wonoka
µ = −6.7
σ = 2.5
n = 573
µ = –6.3
σ = 3.8
n = 108
0.1
0
–15
–10
–5
0
5
13
δ Ccarb
Fig. 6 The distribution of stratigraphically weighted d13Ccarb values for representative (A) canyon-shoulder (units 1–9; 19–
Parachilna; see Fig. 1B) and (B) fine-grained canyon-fill (10 – Mount Goddard; Fig. 1C). The resulting distributions (C) overlap
with a similar mean, and are thus consistent with the fine-grained canyon-fill consisting of recycled, homogenised and redeposited
canyon-shoulder carbonates. Also displayed is the observed distribution of breccia clast d13Ccarb (Fig. 5A), which represents
discrete sampling of a canyon-shoulder isotopic profile. Its histogram distribution looks very similar to the canyon-shoulder profile
for d13Ccarb values less than )7.5&, but the breccia clasts have a long tail into positive values that is not as well developed in the
canyon-shoulder. We interpret this observation as a sampling bias; the rarer dolostone clasts were over-sampled to gain clast
diversity in our sampled population, and it is dolomite that carries the most positive d13Ccarb values. A second contributing factor
may be that unit 9 of the canyon-shoulder is truncated in places by unit 10 sandstones (Fig. 4). Thus, unit 9, and its associated
positive d13Ccarb values, may be under represented in the interpolated canyon-shoulder profile.
6
2012 Blackwell Publishing Ltd
J. M. Husson et al. • Isotope conglomerate test
Terra Nova, Vol 00, No. 0, 1–8
.............................................................................................................................................................
canyons began to fill, and certainly
prior to burial). In classic examples of
meteoric diagenesis, negatively altered
d13Ccarb values are associated with
exposure surfaces, indicative of topdown diffusion of the altering fluid
(e.g., 18O-depleted rainwater charged
with isotopically light DIC originating
from remineralised organic carbon
(Allan and Matthews, 1982; Swart
and Kennedy, 2012)). In the Wonoka
Fm., the only physical evidence for
subaerial exposure is fenestral carbonate fabrics in the microbialite reef facies
of the uppermost unit 11. At Saint
Ronan, however, we see this unit developed at the top of both canyon-shoulder and canyon-fill sections, thus
indicating that the canyons were cut
and filled before unit 11 deposition
(Fig. 3). This observation requires that
canyon formation occurred during
ongoing Wonoka sedimentation, and
therefore demands a syn-depositional
age for the low d13Ccarb values of the
canyon breccia clasts and associated
shoulder stratigraphy. If subaerial
exposure of unit 11 resulted in meteoric
diagenesis, it cannot be invoked to
explain the isotopic values of underlying units 1–9 of the Wonoka Fm.
Although our proposed sequence
boundary at the base of unit 10 could
be submarine, with canyon cutting and
filling being a subaqueous process of
mass wasting (Giddings et al. (2010),
and references therein), some have
argued that the Wonoka canyons
formed subaerially, with the base level
fall accomplished by basin isolation
and Messinian-style evaporitic drawdown (Christie-Blick et al., 1990). This
scenario would leave the canyonshoulders as much as 1.5 km above
local sea level and susceptible to meteoric diagenesis. However, based on the
isotope conglomerate test (Fig. 5A,B),
in this scenario, diagenesis would need
to occur after the basin was exposed,
but before any substantial canyon
cutting and redeposition of shoulder
rocks into canyon-fill had occurred.
Although our data do not rule out this
possibility, there is no visible evidence
of such alteration in the form of
widespread recrystallisation (primary
sedimentary fabrics (Haines, 1988;
Giddings et al., 2010) are exceptionally well preserved in the Wonoka Fm.;
see Figs 2E,F and 4B). Recrystallisation is only pervasively observed in
unit 11 dolomites, where d13Ccarb val 2012 Blackwell Publishing Ltd
ues are at their most positive (+2 to
+8&; Fig. 4C). Also, as the meteoric
model hinges upon fluid–rock interactions, the lightest isotopic values
should be found in horizons most
amenable to fluid flow. Fluid flux
would be controlled by primary porosity and permeability, which is a function of grain size, shape and packing,
and thus directly related to lithofacies.
Contrary to these predictions, we observe no pattern of permeability-dependent isotope modification (Fig. 4A).
Furthermore, the d13Ccarb profile of
the canyon-shoulders (most negative
d13Ccarb at the base, and increasing in
value towards unit 9 and the top) is
opposite of that expected from topdown buffering with meteoric water
charged with DIC issued from organic
matter remineralisation (Allan and
Matthews, 1982). Finally, whether
Shuram-style d13Ccarb anomalies are
globally synchronous or not, Messinian-style diagenesis would be a local
event, restricted to the Adelaide Rift
Complex and the Wonoka Fm., and
would not explain the negative carbon
isotope signatures of other Ediacaran
basins.
Conclusions
Based on the isotope conglomerate
tests, acquisition of d13Ccarb values
()12 to +4&) in South Australian
carbonate sediments was synchronous
with deposition of Wonoka Fm. canyon-shoulder sediments. The filling of
palaeocanyons occurred during the
latter stages of Wonoka Fm. deposition, and the canyons were cut and
filled before development of upper
Wonoka Fm. microbialite reefs (unit
11 on Fig. 1B). These findings are
inconsistent with a burial diagenesis
origin, and the expected first-order
stratigraphic and microtextural patterns predicted by meteoric diagenesis
are not observed. Therefore, the balance of evidence supports a syn-sedimentary origin for the extraordinary
range of d13Ccarb values seen in the
Wonoka Formation of South Australia.
Acknowledgements
J.M.H was supported by the National
Science Foundation Graduate Research
Fellowship
Program
(NSF–GRFP);
A.C.M. was supported by the Alfred Sloan
Fellowship; A.C.M. and B.S. were supported by NSF grant EAR–1121034 to
Maloof and Schoene, and Princeton University supported field and laboratory
work. Blake Dyer, Laura Poppick and
Brenhin Keller provided assistance in the
field. We thank Jim Gehling for fruitful
discussions on Wonoka Formation stratigraphy. Reviews from Dan Rothman, David
Fike, Frank Corsetti and three anonymous
reviewers greatly improved this work. We
thank the landowners and pastoralists of
South Australia for land access. Rebecca
Marks, Elizabeth Lundstrom, Jake Reznick, Steve Shonts, Jenny Piela and Mary
Van Dyke helped with sample preparation.
Stable isotope measurements were performed at Princeton University by Laura
Poppick and at the University of Michigan
by Lora Wingate and Kacey Lohmann.
Author contributions
Field work was conducted by J.M.H.
and A.C.M. (2 field seasons) and B.S. (1
field season), following initial project
planning by J.M.H. and A.C.M; J.M.H.
analysed the data, wrote the manuscript
and drafted figures, all with input from
A.C.M. and B.S.
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Received 16 December 2011; revised version
accepted 23 February 2012
2012 Blackwell Publishing Ltd