lMuuDlNG
ISOTOPE GEOSCIENCE
ELSEVIER
Chemical Geology 125 (1995) 73-99
Sr and Nd isotopes at the Permian/Triassic boundary:
A record of climate change
E.E. Martin 2y*, J.D. Macdougall
Scripps Insrirubon of Oceanography, University of California, San Diego, La Jolla, CA 92093-0220, USA
Received 8 July 1994; accepted 1 March 1995 after revision
Abstract
We present a detailed curve of seawater 87Sr/86Sr for the Middle Permian to Triassic based on analyses of conodonts from
overlapping sections in the U.S.A. and Pakistan, correlated using conodont biostratigraphy. The isotope ratio decreased in the
Middle Permian at an average rate of 0.000062 Ma- I, reached a minimum in the Capitanian (257-258 Ma), and increased in
the Late Permian at an average rate of 0.000097 Ma-‘. The Late Permian rate of increase was roughly two and a half times
greater than the average increase over the past 40 Ma, and approximately equal to the highest Cenozoic rates, which occurred
over much shorter time intervals.
Modeling results suggest that decreasing Middle Permian 87Sr/86Sr ratios were driven by changes in the riverine Sr flux to
the oceans, while increalsing ratios in the Late Permian/Triassic are attributed to both increased riverine “Sr/‘%r and flux. The
reduced Middle Permian riverine flux coincides with extreme continental aridity associated with the formation of Pangea and
recorded by massive evaporite deposits. In addition, mountains in the equatorial region of Pangea may have created a rain
shadow, thereby minimizing precipitation in regions that currently contribute the bulk of chemical weathering products to the
ocean.
Increasing riverine 87Sr/86Srin the Late Permian is suggested by the observation that ‘43Nd/ 144Ndvalues decrease at the same
time; however, the source of radiogenic Sr is not known. Frequently cited mechanisms for increasing 87Sr/86Sr in runoff, such
as glaciations and continent-to-continent collisions, coincide instead with decreasing seawater 87Sr/86Srin the Middle Permian.
One possible source may have been deep erosion into older orogens, associated with a dramatic increase in chemical weathering
in the Late Permian. The cause of enhanced weathering appears to have been increased levels of atmospheric CO2 with associated
global warming and increased humidity. Proposed sources of COZ include dissociation of gas hydrates and oxidation of organic
matter during extreme sea level regression, as well as volcanic emissions from Siberian Traps eruptions. Continental floral and
fauna1 distributions are consistent with this interpretation, as are oceanic 613Cpatterns and variations in shallow-water sediment
lithologies.
1. Introduction
The period of time near the Permian/Triassic (P/
Tr) boundary was marked by a number of anomalous
[=I
* Corresponding authcr.
*Present address: 1112 Turlington Hall, Department of Geology,
University of Florida, Gainesville, FL 3261 l-7340, USA.
0009-2541/95/$09.50 (0 1995 Elsevier Science B.V. All rights reserved
SSDIOOO9-2541(95:100081-X
events in the history of the Earth. The magnitude of
many of these events, including the biotic crisis, sea
level regression and seawater geochemical anomalies,
is unsurpassed during Phanerozoic time. The geochemical anomalies include dramatic variations in seawater
S34S (Claypool et al., 1980), 613C (Baud et al., 1989)
and *‘Sr/*%r (Burke et al., 1982; Denison et al., 1994))
74
E.E. Martin, J.D. Macdougall /Chemical Geology 125 (1995) 73-99
suggesting major chemical changes in the oceans. Specifically, the seawater 87Sr/86Sr curve for this interval
reaches a minimum value for the Phanerozoic (Burke
et al., 1982; Popp et al., 1986b; Holser and Magaritz,
1987; Denison et al., 1994). Changes before and after
this minimum appear to be rapid, comparable to the
well-documented rate of increase over the past 40 Ma.
This paper evaluates the timing, magnitude and rate
of change of seawater *‘Sr/‘%r for the time interval
spanning the P/Tr boundary. Temporal variations in
this isotopic ratio reflect changes in the riverine (continental) and hydrothermal inputs to the ocean. Specifically, the isotopic signature and magnitude of the
continental flux are controlled by the interaction
between tectonics and climate. Tectonic uplift and
landmass distribution determine the age, type and quantity of rock exposed to weathering, while climate controls the intensity and rate of weathering. In contrast,
the hydrothermal flux is presumably controlled by the
rate of sea-floor spreading.
Many models of the seawater Sr isotope system have
focused on the rapid increase in s7Srla6Sr during the
Cenozoic (Palmer and Elderfield, 1985; Hodell et al.,
1989; Capo and DePaolo, 1990; Hodell et al., 1990;
Richter et al., 1992). The general increase has been
attributed to tectonics (Raymo et al., 1988; Edmond,
1992; Raymo and Ruddiman, 1992; Richter et al.,
1992), while climate has been proposed as the cause
of some of the shorter time-scale changes within the
general increase (Hodell et al., 1989, 1990; Capo and
DePaolo, 1990; Miller et al., 1991; Zachos, 1993).
Recently, Edmond ( 1992) proposed that Himalayanstyle collisional tectonics are required to produce dramatic large-scale increases in 87Sr/86Sr because this is
the only mechanism whereby radiogenic Sr can be
redistributed into easily weathered phases.
The rapid rise of seawater s7Srla6Sr in the Late Permian to Triassic is different from the increase over the
past 40 Ma in several respects. First, the Permian is one
of the least radiogenic intervals for seawater Sr during
the Phanerozoic. Second, it is a period during which
the model of Richter et al. ( 1992), which accurately
reproduces many sections of the Phanerozoic 87Sr/86Sr
curve, does not match the measured data. Finally, it is
a major extinction boundary, likely to be characterized
by dramatic environmental changes that could affect
the Sr cycle. We would like to emphasize that the focus
of this paper is to identify the cause(s) of the isotopic
variation, nor the cause(s) of the extinction event.
These may or may not be related.
We evaluate the P/Tr Sr data from both a forward
and inverse perspective. On the one hand, we use the
timing, magnitude and rate of change of seawater s’Sr/
‘?Sr variations to constrain the processes occurring during the Permian. On the other, we apply our knowledge
of geologic and environmental conditions at that time
to evaluate the Sr cycle.
Much of the data in the literature for paleoseawater
87Sr/86Srcome from analyses of carbonate sediments
sampled at widely distributed localities. In contrast, all
of the data presented in this paper are from conodonts
collected from a series of overlapping sections correlated by their conodont biostratigraphy. Conodonts are
teeth-like elements from an extinct soft-bodied chordate group (Briggs et al., 1983; Aldridge, 1987) and
are composed of francolite, a carbonate fluorapatite,
which is believed to be less susceptible to diagenetic
alteration than carbonate (Bemat, 1974; Kovach, 1980;
Staudigel et al., 1985; Bertram et al., 1992; Ruppel et
al., 1993). Another advantage is that carbonate fluorapatite incorporates high concentrations of both Sr and
Nd (Wright et al., 1984; Shaw and Wasserburg, 1985).
Conodonts evolved rapidly, therefore their biostratigraphic zones cover short time intervals and relative
sample ages are well defined.
In addition to the Sr data, we present detailed records
of seawater ‘43Nd/‘44Ndevolution for the Proto-Pacific
(Panthalassa) and Tethys oceans in the Permian. These
data impose added constraints on the sources of seawater a7Srla6Sr variations.
2. Conditions and events during the Permian
Information about the tectonic and climatic conditions for the Permian through Triassic creates the
framework for the evaluation of Sr isotopic variations
in seawater. The timing of many of the major events of
this time period is summarized in Fig. 1. As already
mentioned, the Permian was a unique period in Earth’s
history in many respects. The final consolidation of
Pangea created a supercontinent that probably experienced severe climatic conditions (Crowley et al., 1989;
Crowley and North, 1991; Crowley, 1994), further
intensified by extremely low sea level stands. Global
sea level curves by Hallam ( 1984), Holser and Magar-
E.E. Martin, J.D. Macdougall /Chemical Geology 12.5 (1995) 73-99
stecle
Bee
.u
247.1
3
.”
$
Scythian
-251
.O
E
Changsingian
*g
8
P
253.5 c
.E
3 Wuchiapingian
g!
3
256.0 I-
s*
Capitanian 87Sr/86Sr minimum
S
-2156.5
Kazanianl
Wordian
Kungurianl
Leonardian
N.
I
E
a,
I
I
I
s
._
E
;
0
)I
Z
m
w
287.5
-
--496.0
Rq
-
Carboniferous
r.
I
Fig. 1. Timing of climatic, tectonic and geochemical events during the Permian. Data from: ’Frakes ( 1979), and Caputo and Crowell ( 1985);
’ Epshteyn (1981); 3 Ziegler ( 1989), and Scotese and McKerrow (1990); 4 Nalivkin (1973), Zonenshain et al. (1984), and Scotese and
McKerrow (1990); 5 Baksi and Farrar (1991), Dalrymple et al. (1991). and Renne and Basu (1991); 6 Zharkov (1981); ’ Hallam (1984),
Holser and Magaritz ( 1987), Ross and ROSS( 1987)) and Wignall and Hallam (1993); 8 Popp et al. (1986a), Magaritz et al, (1988). Baud et
al. ( 1989), Gruszczynski et al. ( 1989). and Holser et al. (1989); ’ Claypool et al. (1980); lo Delaney et al. (1989); ‘I Sepkoski (1982) cited
by Holser and Magaritz (1987), and Raup and Sepkoski (1986); ‘* Erwin ( 1993). The time scale used in many of these publications varied
slightly from the one presented here. In these cases, we made our best effort to correlate between the data sets. As a result, the age ranges
presented here do not always match the ranges in the original publications.
itz ( 1987), Ross and Ross ( 1987), and Hallam ( 1992)
outline a regressive phase from the Early Permian to
the P/Tr boundary, with an abrupt drop at, the boundary, or just before the boundary with transgressive conditions at the boundary (Wignall and Hallam, 1993)
(Fig. 1). According to Hallam ( 1984) this lowstand
represented the lowest sea level during the entire Phanerozoic. Estimates of the magnitude of the regression
range up to 210 m (Forney, 1975) or even 280 m
(Holser and Magaritz, 1987). In contrast, the sea level
curve by Vail et al. ( 1977) indicates general regressive
conditions from Mid-Carboniferous
to Jurassic, with
the lowest point occurring about Mid-Permian. This
interpretation is less consistent with geologic evidence
of a worldwide unconformity
at or near the P/Tr
boundary.
Gradual sea level variations are commonly attributed
to changes in the volume of ocean ridges (Hays and
Pitman, 1973). In the Permian this could have occurred
by slower spreading rates or actual loss of ridge seg-
76
E.E. Martin. J.D. Macdougalll Chemical Geology 125 (1995) 73-99
ments. The assemblage of a single huge continent
implies that many of the plate boundaries were continent-to-continent. As Schopf (1974) has suggested,
this might reduce the free movement of plates and consequently decrease spreading rates. Consumption of
ocean basins during continent-to-continent collisions
would also destroy segments of the mid-ocean ridge
(Hallam, 1977).
The Early Permian marked the end of the Late Paleozoic ice age, which began in the Early Carboniferous
when a portion of Gondwana was located at the south
pole. During this glaciation extensive ice sheets covered South America, south-central Africa, India, Antarctica and Australia (Caputo and Crowell, 1985).
Gondwana migrated north during the Late Carboniferous and Early Permian, and most of the ice had melted
by Sakmarian time (Frakes, 1979) (Fig. 1). Final
dropstones are recorded from Australia and Antarctica
in the lowermost Kazanian (throughout this paper we
refer to the stages by their Russian names only for the
sake of brevity) (Caputo and Crowell, 1985). Minor
Kazanian-aged glacial dropstones have also been identified in the northern hemisphere on the Kolyma block
of Siberia (Epshteyn, 198 1). This cooling reflects the
encroachment of Pangea on the north pole. Climate
models predict conditions suitable for extensive glaciation at both poles during the latter part of the Permian
(Crowley et al., 1989; Kutzbach and Gallimore, 1989).
The absence of substantial ice development led Erwin
(1993) to speculate that atmospheric CO, concentrations may have been high at that time, resulting in
global warming.
For much of the Permian extreme continentality governed the climate in the interior of the supercontinent.
Climate models predict daytime summer temperatures
as high as 35~lS”C (Crowley et al., 1989; Crowley,
1994) with seasonal variability of 30°C; a range currently experienced only in Siberia and northern Canada
(Crowley and North, 1991), although the absolute temperatures in these areas are lower. Arid conditions prevailed in the interior regions, with estimates of
precipitation at N 50% of the modem value (Kutzbach
and Gallimore, 1989; Crowley, 1994). In contrast,
strong monsoonal circulation may have delivered seasonal rains along the eastern continental margins (Robinson, 1973; Kutzbach and Gallimore, 1989). Parrish
(1993) suggests that the intensity of monsoonal circulation increased from the Late Carboniferous to a
maximum in the Triassic. As Robinson ( 1973) points
out, the distribution of climate-sensitive rocks, such as
coals, eolian deposits, laterites, red beds and evaporites
is consistent with these climate model results.
Massive evaporite beds were deposited in the Middle
and Late Permian and the Mid-Triassic (Gordon,
1975). In a detailed analysis, Zharkov (1981) estimated the volume of salts and sulfate rocks preserved
from each stage of the Paleozoic (Fig. 1). He concluded that one third of all the salt and sulfate rocks
deposited in the Paleozoic were deposited during the
interval he labels Kungurian. Another 6% of the Paleozoic salts and sulfate rocks are preserved in the Middle
to Late Permian stages of the Kazanian and Tatarian.
Fischer ( 1964) and Stevens ( 1977) suggested that the
Middle to Late Permian ocean may have been stratified
because of this extensive evaporite deposition. They
speculate that dense brines, formed during evaporite
precipitation, could have been stored in the deep ocean,
while surface water salinities were reduced by as much
as 3-3.5%0. Both authors discussed this phenomenon
in terms of the effects of brackish conditions on shallow-water fauna1 extinctions. Holser and Magaritz
(1987) noted that the associated reduction of the Sr
concentration in shallow-water habitats could account
for the very rapid changes in 87Sr/86Srobserved in Late
Permian samples. However, fluid inclusion data from
Horita et al. ( 1991) indicate that salinities in shallowwater environments were similar to modem seawater
values.
Zharkov ( 1981) calculated that 1.48. lo5 km3 of
sulfate rocks are preserved from the entire Permian. At
1000 ppm Sr, 3.62. lOi mol Sr would be stored in
these rocks, equivalent to 3% of the modern ocean’s Sr
reservoir. Evaporites are highly susceptible to weathering, thus the original volume deposited may have
been much greater. Stevens (1977) theorized that the
weathered fraction could easily be 50%, implying that
the Permian evaporites may have removed N 6% of the
Sr from the ocean. This is a relatively minor change
and is consistent with the fluid inclusion data (Horita
et al., 1991).
Isotope curves for Si3C and S34Salso record large
variations in the Permian. In general, Late Carboniferous to Early Permian seawater seems to have been
characterized by high values of S13C, ranging up to
u + 4 to + 6%0(Poppet al., 1986a; Holser and Magaritz, 1987), with intermittent periods of lower Si3C-
E. E. Martin, J.D. Macdougall / Chemical Geology 125 (1995) 73-99
values. A final high spike preserved globally in the
Tatarian (Holser et al., 1986; Baud et al., 1989) contrasts with generally decreasing values from the Kazanian to the P/Tr boundary, with a particularly rapid
decrease at the boundary (Fig. 1) (Baud et al., 1989;
Gruszczynski et al., 1989; Holser et al., 1989). Constant A ‘3C-values ( 6’3C,,b - SL3C,) spanning the P/
Tr boundary suggest that the carbon isotopic variations
record a global, whale ocean signal (Magaritz et al.,
1992). Erwin ( 1993‘) describes three possible sources
of light carbon which are consistent with conditions at
that time: ( 1) an influx of juvenile carbon from volcanic degassing (8°C = -5%0) during the Siberian
Traps eruptions, (2) exposure, weathering and oxidation of buried organic carbon ( SL3C= - 25 to - 20%0)
during the Late Permian sea level regression, or (3)
release of methane ( 613C= -65%0) during gas
hydrate dissociation related again to the sea level
regression. Mass-balance calculations indicate that the
volcanic input would only have a minor effect, but
either of the other two sources would be capable of
producing the observed decrease (Erwin, 1993). In
addition, the relatively high 613C-values preceding the
decrease could be related to enhanced burial of organic
carbon relative to carbonate.
Sulfur isotopic variations recorded in sulfates are
even more dramatic (Claypool et al., 1980). S34S
decreases gradually throughout the entire Permian
(Fig. 1) from a value of - + 14%0 at the PermoCarboniferous boundary to the Phanerozoic minimum
of - + 11%0 in the earliest Triassic, then increases
rapidly to - +26 or + 28%0 in the Scythian. The
decreasing values probably represent a period of net
pyrite oxidation and. erosion, while the rapid increase
indicates sulfide removal through pyrite burial. To
address the very rapid increase in the Early Triassic,
Holser (1977) proposed catastrophic mixing between
the surface ocean and a peripheral basin containing
elevated 634S-values created by pyrite deposition.
Tectonic events may have a significant influence on
seawater 87Sr/86Sr. For example, evidence suggests
that the rapid rise over the last 40 Ma is related to uplift
and unroofing of the Himalaya (Raymo et al., 1988;
Hodell et al., 1989, 1990; Palmer and Edmond, 1989;
Edmond, 1992; Krishnaswami et al., 1992; Raymo and
Ruddiman, 1992; Richter et al., 1992). The two major
continent-to-continent collisions around Permian time
are reflected in the Hercynian and Uralian Orogenies
II
(Fig. 1) . Clockwise rotation of Gondwana led to a NE
to SW progression of continental collisions along the
Hercynian megasuture (Scotese and McKerrow, 1990)
which peaked in the Late Carboniferous (Ziegler,
1989) and ended in the Early Permian. The Uralian
collision between Baltica and Kazakstan also began in
the Carboniferous (Nalivkin, 1973; Zonenshain et al.,
1984). In this case, the main collision event occurred
in the Early Permian (Nalivkin, 1973; Scotese and
McKerrow, 1990). Sediment analysis indicates that the
Urals were high, snow-capped peaks in the Sakmarian
to Artinskian. Thin fanglomerate and sandstone
sequences suggest that compression had essentially
ceased by Kungurian time. And by the Late Permian
the Urals are described as “low and eroded” (Nalivkin, 1973). Thus, most of the effect of these orogenies
on seawater 87Sr/86Srshould have occurred during the
Late Carboniferous and Early Permian.
Massive flood basalt volcanism produced the Siberian Traps around P/Tr time (Fig. 1). This may represent the largest flood basalt province in the
Phanerozoic, with an estimated original volume of
> 1.5 lo6 km3. The eruption interval for this event
appears to have been very short, possibly < 1 Ma
(Renne and Basu, 1991). Recent attempts to date the
Siberian Traps with high-precision techniques yielded
ages roughly equivalent to the P/Tr boundary (Baksi
and Farrar, 1991; Dahymple et al., 1991; Renne and
Basu, 1991). Campbell et al. ( 1992) found that zircon
206Pb/238Uages for rocks comagmatic with the Siberian Traps and for the boundary clay in a Chinese section agreed within analytical uncertainty.
Analysis of the boundary clay in China indicates that
it may be altered ash from a silicic volcanic event that
was unrelated to the Siberian Traps (Zhou and Kyte,
1988). The timing of both the large-scale basaltic and
smaller-scale silicic volcanic eruptions coincided
approximately with the end of a period of stable
reversed magnetic polarity known as the Kiamen Long
Reversed Superchron (Fig. 1). The relationship
between field reversals, massive volcanism and extinction events is unclear; however, it appears that the peak
of the gradual extinction event in the Late Permian and
most of the volcanism occurred several millions of
years after the reversal (Fig. 1) (Raup and Sepkoski,
1986; Holser and Magaritz, 1987).
??
78
3. Stratigraphic
E.E. Martin, J.D. Macdougall/Chemical
background
The rate of isotopic variation in seawater that is calculated from the measured data is highly dependent on
the time scale applied. Unfortunately, there is little consensus on the absolute ages, or even the nomenclature,
for Permian stages. This difficulty can be attributed to
the extremely low sea level and low fauna1 diversity
and abundance at that time. Particularly in the Late
Permian, sediments were frequently deposited in isolated basins and the high proportion of endemic fauna
complicate inter-basin correlations.
We adopted the sequences of stages outlined in Wardlaw ( 1995). Assigning absolute ages to biostratigraphically defined stages is a rather uncertain process.
The time scale adopted here is modified from Harland
et al. ( 1990). Since the completion of that time scale,
ion microprobe dates for zircons from a bentonite
boundary layer at the Meishan section, Changsing,
China, have been published, which give an age of
251.2 f 3.4 Ma for the P/Tr boundary (Claod-Long
et al., 1991) . This is significantly older than the Harland
et al. ( 1990) estimate of 245 Ma. The number of radiometric ages from the Permian is limited, and the implications of the older period-boundary date for the timing
of the stage boundaries is unclear. In general we have
maintained the length of each stage as given in Harland
et al. ( 1990), but shifted their boundaries by 6 Ma;
however, our Kazanian and Ufimian stages are 1 Ma
longer and the Kungurian and Artinskian stages are 1
Mashorter (B.R. Wardlaw, pers. commun., 1993) than
those of Harland et al. ( 1990). With the exception of
the Kazanian/Ufimian boundary, our age assignments
fall within the minimum and maximum values plotted
on the chronograms in Harland et al. (1990). Sweet
( 1992) suggested that the Late Permian may have been
much shorter than is assumed in this paper. In this case,
rates of isotopic change over this interval would be
even more rapid than we have calculated.
No single stratigraphic section provides a complete,
unaltered, Permian through Triassic marine sequence
with abundant conodonts. Therefore, we pieced
together several overlapping sections from the U.S.A.
and Pakistan for this study (Fig. 2). Sample density is
greatest for the Middle and Late Permian, with just a
few samples representing older Permian marine sediments. All of the conodonts for this study were very
generously provided by B.R. Wardlaw (U.S.G.S., Res-
Geology 125 (1995) 73-99
ton, Virginia) who collected, processed and identified
the specimens.
The completeness of various P/Tr sections is still
disputed (for details see Sweet et al., 1992), but there
is general agreement that the most complete are from
the Tethyan realm. Uppermost Permian and Lowermost Triassic samples for this study are from the Salt
and Kbisor Ranges in Pakistan (Wardlaw and Pogue,
1995). These specimens were collected from the Amb,
Wargal, Chhidru and Mianwali formations, which
range in age from Kazanian to Smithian (Fig. 2). As
pointed out by Teichert ( 1990) and Sweet ( 1992), the
Uppermost Permian is missing from these locations.
However, data from above and below the missing interval suggest that 87Sr/86Srincreasescontinuously across
the boundary. Detailed descriptions of the formations
are found in the report of the PJRG ( 1985) and Wardlaw and Pogue ( 1995). In general, the sediments are
composed of interbedded limestone, dolomite, marl,
sandstone and shale, representing deposition in intertidal, marginal marine and shallow marine settings. All
of the conodonts were collected from limestone intervals which are believed to have been deposited under
normal marine conditions.
Pakistani sections overlap with U.S.A. sections in
the Kazanian, Capitanian, Smithian and Dienerian.
Most of the U.S. material is from the Glass and Del
Norte Mountains in southwest Texas (Wardlaw and
Grant, 1990; Wardlaw et al., 1990, 1991). The samples
range in age from Artinskian to Capitanian (Fig. 2),
and were collected from the Skinner Ranch, Cathedral
Mountain, Road Canyon, Word, Vidrio and Altuda formations. These sediments were deposited at the southern margin of the Permian Basin in shallow intertidal
or lagoonal to deeper-water shelf edge and slope environments.
Additional U.S.A. samples come from the Baldwin
Creek section at the eastern edge of the Phosphoria
Basin in Wyoming, which includes the Ervay, Franson
and Grandeur members of the Park City Formation and
the Retort Member of the Phosphoria. It has been proposed that all of these sediments were deposited in a
carbonate ramp setting along the Cordilleran margin
(Wardlaw and Collinson, 1986). Three Early Permian
samples are from the Garden Valley Formation in central Nevada, which was deposited along the Cordilleran
margin in shallow shelf to outer shelf settings (Gallegos and Wardlaw, 1992). These specimens are dated
E. E. Martin, J.D. Macdougall / Chemical Geology 125 (1995) 73-99
Pakistan
United States
m
1
B
B
5
AtIe
-
g
-
-
L
;N
Q
247.9
247.1
a
g
249.4
Spathian
Smithian
Dienerian
19
k
Greisbachian
-,251.0
Changsingian
:
5
‘Z
m Wuchiapingian
;i
256.0 k
E
253.5
0t
g
4
-
-
-
-
-
-
-
-
-
-
-
-
-
-
Capitanian
--256.5
Kazanianl
Wordian
g
E
2
sl
=
4
z
262.0
Ufimianl
Roadian
264.1
Kungurianl
Leonardian
--2%
nlne
!scaIO
change
274.6
Arlinskian
3
E
Sakmarian
&
P
*
5
W
-
207.5
-
t
--296.0
Carboniferous
Fig. 2. Proposed Permian time scale and age distribution of conodont samples from the U.S.A. and Pakistan. The time scale is modified from
Harland et al. ( 1990). B.R. Wardlaw defined the conodont biostratigraphy.
as Late Asselian, Early Sakmarian and Early Artinskian. Because there are long, unsampled time intervals
between each specimen, the relative ages are not well
constrained. The lowermost Permian is represented by
two samples from just above the Permo-Carboniferous
boundary in the Mid-Continent Basin, Kansas. They
are from a basal marine regressive limestone in the
classic cyclothem sequences of the Neva Limestone.
Triassic material from the U.S.A. comes from the Crit-
tenden Springs section in northeastern Nevada. These
sediments from the Thaynes and Dinwoody formations
represent deposition in a lagoon to high-energy nearshore setting along the Cordilleran margin (Carr and
Paull, 1983).
Absolute errors for radiometrically dated material
from the Permian range from It 2 to f 10 Ma (Forster
and Warrington, 1985; Harland et al., 1990). This represents a minimum uncertainty because the location of
E.E. Martin, J.D. Macdougall /Chemical Geology I25 (1995) 73-99
80
(4
+ -
Fig. 3a
250
260
270
280
290
i
300
AgO (MaI
(b)
0.7065
??
3i
%
0.7075
-
0.7070
-
,
,
,
r.
h
3
m
Fig. 3b
.
0.7065
’
Triassic
245
’
‘,’ Middle
’ Permian’ ,’EarlvI
1Uf 1 Ku 1 Ar
Ka
Late pennian
Ch 1 Wu 1 Ca 1
255
250
260
265
270
many of the dated samples relative to the stage boundaries is imprecisely known (for further discussion see
Odin, 1985). For correlation between sections the primary concern is relative age estimates. The samples
can be placed quite confidently within biostratigraphic
conodont zones; thus an accurate statement of the relative error in age would be the average length of a
conodont zone, which is remarkably short due to the
rapid evolution of the conodont organism. For material
younger than Sakmarian this value is N 2 Ma, which is
consistent with the spread of data points along the age
axis on the seawater Sr isotope curve (Fig. 3). For
Sakmarian and Asselian material y 5 Ma appears to be
a more accurate estimate.
Thermal maturation of organic material incorporated
in conodont apatite results in progressive color alteration of the specimen. Color Alteration Index (CAI)
values of 1 indicate that the conodont is essentially
unaltered; higher values imply exposure to higher temperatures. Work by Bertram et al. ( 1992) and Cummins and Elderfield ( 1994) suggest that Sr isotopes are
not altered in specimens with CAI-values of < 2.0. All
samples examined in this study had CAI-values (determined by B.R. Wardlaw) of 1, except for the Crittenden
Springs, Nevada, specimens which had values of 2.
Age (Ma)
(c)
. . I
m
0.7065
.
0.7080
-
0.7075
-
0.7070
-
. .
I
I 1 .
I
e@
@.,
: ;;
.
4. Methods
\,j
:.e:’
-
6
2
The conodonts were removed from the host rock by
B.R. Wardlaw using standard procedures. The surrounding matrix was dissolved in 10% acetic acid, the
p
OD
/I\
0.7065
245
250
255
260
265
270
Age VW
Fig. 3. Permian/Triassic seawater s7Sr/*% vs. age.
a. A comparison of Permian/Triassic *‘Sr/% data vs. age from this
study (0 = U.S.A. and ? ?
= Pakistan), as well as A = Veizer and
Compston ( 1974); 0 = Popp et al. ( 1986b); H = Nishioka et al.
(1991); and + = Denison et al. (1994) for the Permian data and
Koepnick et al. ( 1990) for the Triassic data. All ratios were corrected
for interlaboratory bias to an NBS 987 value of 0.710269. Error bars
are smaller than the size of the symbols.
b. Detail of Permian/Triassic *‘S#‘%rSr data vs. age for samples
from this study which define the rapidly decreasing and increasing
segments of the curve. Symbols: U.S.A. -Texas, A = SR, V = ST,
D = TX, U = BM; Wyoming, d = BC; Nevada, 0 = CS; Pakistan
-•O=SA. ? ?
=KL,O=CD, ? ?=KA. @=CW, M =NN.
c. The modified 87Sr/86Sr vs. age data set used to calculate the
smoothed spline represented by the line. Data points enclosed in a
dashed
circle qpear to have been altered under continentalconditions based on a comparisonto our data as well as data from the
literature,thereforethey were excluded from the data set prior to
curve fitting. Outlying data points, delineated with a cross ( X ),
were eliminated using a systematic method described in the text.
Symbols: O-U.S.A.; ? ?
=Paktstan; @ =U.S. data eliminated for
spline fit; ? ?
= Pakistan data eliminatedfor spline fit.
81
E.E. Martin, J.D. Macdougall /Chemical Geology 125 (1995) 73-99
Table 1
Sr isotope and concentration data for Permian/Triassic conodonts
Sample”
Weight
Biozoneb
Stage
AgeC
“SrlwSp
20
(Ma)
(WI
SP
(ppm)
U.S.A.:
Neva Limestone, Kansas:
KN-1
KN-2
37
38
A
B
Assel
Assel
_ 294
-2%
0.708111
0.708132
o.OOoO22
o.oOOO22
1,563
1,614
Assel
Sak
Art
-289
_ 286
-215
0.708878
0.708841
0.707957
o.OOOo22
0.000022
0.000022
2,244
1.832
799
Garden Vallqy Formation, Nevada:
NV-I
NV-2
NV-3
65
12
19
C
D
E
Skinner Ranch, Glass Mountains, Texas:
SR-1
SR-2
SR-3
SR-4
SR-5
SR-6
SR-7
SR-8
SR-9
SR-10
SR-11
SR-12
16
6
5
7
28
68
20
31
57
73
45
36
G
I
J
K
L
L
L
M
M
M
M
M
Art
Art
Kung
Kung
Kung
Kung
Kung
Kung
Kung
Kung
Kung
Kung
267.10
266.90
266.70
266.50
266.30
266.10
265.90
265.70
265.50
265.30
265.10
264.90
0.708011
0.708093
0.708345
0.707605
0.707453
0.707629
0.707643
0.707561
0.707506
0.707513
0.707416
0.707461
0.000022
0.000033
0.000030
0.000026
O.OCOO24
0.000022
o.OOoO22
o.OOoO22
o.OoOO22
0.000022
0.000084
0.000022
1,293
1,517
1,276
1,879
1,186
1,750
1,829
1,904
1,375
1,810
1,482
M
N
N
N
P
Kung
Kung
Kung
Kung/Uf
Uf
264.75
264.60
264.40
264.00
262.10
0.707336
0.707341
0.707333
0.707388
0.70733 1
0.000022
0.000022
0.000022
0.000022
o.OOoO22
1,642
1,625
;
KaZ
KaZ
CaP
CaP
261.75
261.00
258.25
257.80
0.707229
0.706914
0.707031
0.706948
o.OoOO22
0.000022
0.000022
0.000032
1,742
1,533
1,415
1,543
Q
R
R
S
R
Cap
GP
CaP
GP
Cap
258.25
257.80
256.30
256.10
256.50
0.707247
0.707406
0.707059
0.707027
0.706913
0.000022
0.000022
o.OOoO22
0.000022
0.000022
1,485
1,790
2,109
1,672
1,443
N
N
N
o/u
U
U
V
V
Kung
Kung
Kung
Knz/Uf
KaZ
KaZ
KaZ
KaZ
264.50
264.30
264.10
262.25
260.75
260.00
259.25
258.60
0.707340
0.707404
0.707373
0.707133
0.707208
0.707100
0.70709 1
0.707063
o.OOOo22
0.000022
0.000022
0.000022
0.000022
0.000022
o.OoOO22
o.OoOO22
1,424
1,469
1,441
1,226
1,730
1,339
1,306
1,148
1,666
Split Tank, Glass Mountains, Texas:
ST-l
ST-2
ST-3
ST-4 + 5
ST-6
20
19
66
17
20
Del Narte Mount&s,
TX-3
TX-2
TX-5
TX-6
Texas:
33
17
24
24
Bird Mine, Del Norte Mount&s,
BM-1
BM-2
BM-3
BM-4
BM-5
1,344
1,522
1,598
12
28
36
30
41
Q
R
Texas:
Baldwin Creek, Wyoming:
BC-1
BC-2
BC-3
BC-4+5
BC-6
BC-7
BC-8
BC-9
55
25
61
63
19
31
41
11
82
E.E. Martin, J.D. Macdougall /Chemical Geology 125 (1995) 73-99
Table 1 (continued)
Sample”
Weight
Biozoneb
Stage
Age”
?w
=sP
20
(Ma)
(pgLg)
Sf
(ppm)
U.S.A. (cont.):
Crittenden Springs, Nellada:
cs-1
cs-2
cs-3
CM
24
68
29
29
FF
II
II
JJ
Gries
Smith
Smith
Smith
250.00
248.50
248.40
248.25
0.707657
0.707885
0.707867
0.707940
o.OOOQ22
0.000022
0.000022
o.OOoO22
1,395
1,631
1,960
1,848
T/U
V
V
KaZ
KaZ
KaZ
260.75
259.30
258.75
0.707211
0.707388
0.707117
0.000022
0.000022
0.000022
1,396
2,091
1,871
V
V
Y
Y
Y
Y
Z
AA
KaZ
Cap
Wuch
Wuch
Wuch
Wuch
Wuch
Wuch
260.75
257.25
255.70
255.25
254.80
254.60
254.10
253.75
0.707190
0.707168
0.70720 1
0.707100
0.707102
0.707117
0.707186
0.707240
0.000022
0.000022
0.000022
0.000022
0.000023
o.OoOO22
0.000022
0.000054
1,400
1,748
1,538
1,326
962
1,436
1,450
1,417
258.60
254.30
253.40
252.70
251.80
249.00
0.707130
0.707399
0.707690
0.707322
0.707427
0.707736
0.000022
0.000034
o.OOoO22
o.OOOo22
0.000022
0.000022
2,035
1,891
1,515
1,404
1,164
252.30
252.20
252.00
251.80
251.70
0.707499
0.707568
0.707458
0.707587
0.707469
0.000062
0.000022
0.000022
0.000049
0.000022
1,466
1,431
1.3%
1,490
1,460
PAKISTAN:
Saidu Wali, Khisor Range:
SA-1 f 2
SA-3+4
SA-5 + 6
34
16
14
Kotia Lodlian, Khisor Range:
KL- 1
KL-2+3
KL-4+5
KL-6
KL-7
KL-8
36
16
20
25
11
20
KL-9
23
3
KL- LO
Chhidru Nab, Salt Range:
CD-l +2
CD-3
CD-4
CD-5
CD-6
CD-7
4
6
7
11
10
31
V
AA
cc
cc
DD
HH
Kaz/Cup
Wuch
Chang
Chang
Chang
Dien
1,400
Kafhwai, Salt Range:
KA-1
KA-2
KA-3
KA-4
KA-5
21
25
6
13
17
cc
cc
Chang
Chang
Chang
DD
DD
DD
Chang
Chung
Chatuwala Nala, Salt Range:
cw-1
cw-2
cw-3
cw-4
7
10
31
54
cc
Chang
DD
EE
FF
Chang
Gries
Gries
251.80
251.60
250.75
249.75
0.707440
0.707503
0.707308
0.707626
0.000024
0.000030
0.000022
0.000022
1,556
1,572
1,511
1,425
FF
GG
GG
HH
HH
JJ
Gries
Dien
Dien
Dien
Dien
Smith
249.75
249.40
249.25
248.75
248.60
248.35
0.707856
0.707895
0.708197
0.708178
0.708377
0.708203
o.OOoO22
0.000022
0.000022
o.OOOo22
0.000022
o.OOw22
1,241
n.a.
Nammal Nala,‘Salt Range:
NN-1
NN-2
NN-3
NN-5
NN-6
NN-7
57
18
62
38
58
25
1,400
1,404
1,611
1,223
E.E. Martin, J.D. Maca’ougall /Chemical Geology 125 (1995) 73-99
83
Table 1 (continued)
n.a. = samplesthat were not analyzed for Sr concentration.
“Samplenumbers assigned for this study. For correlation with original sample designations, please contact first author.
%onodont zonation from Wardlaw (1995). A = S. waubaunsensis4 “barskooi”; B = S. "longissimus‘SW. expansus; C = S. fusus-M.
longifoliosa; D = SW. inornatus; E = M. bisselli-Sw. primus; F = M. bissel&Sw. behnkeni; G = M. bisselli-Sw.whirei; H = N. clarki; I = N.
pequopensis-M. gujioensis; I = N. exsculpfus-M. gujoensis; K = N. foliatus; L = N. foliates-M. idahoensis; M = N. prayi-M. i&hoer&; N = N.
sukopilcatus-M. idahoensis; 0 = N. newelli-M. serrata; P = M. serrata; Q = M. asserrata; R = M. postserrata; S = M. nuchalina-M. n. sp.;
T = M. phosphoriensis; U =:M. bitteri; V = Me. diaergens; W = M. “postbitteri”; X = M. liangshanensis; Y = M. leveni; Z = M. guanyuanensis;
AA = M. orientalis; BB =M. “chanxingensis”; CC = M. subcarinara; DD = M. carinata; EE = H.parvus; FF = I. isarcica; GG = Ne. cristagalli;
HH = Ne. pakistanensis; II = Ne. waageni; JJ = N. milleri.
‘Ages reported for combined samples represent a weighted mean of individual samples.
d s’Sr/s%r values for standards equal 0.710260 for NBS 987, and 0.709175 for seawater samples from the North Atlantic and Central Pacific.
All ratios are fractionation corrected to an %3/ssSr ratio of 0.1194. All samples have been assigned a minimum 20 uncertainty of f 22. 10m6
equivalent to the total range of repeat NBS 987 analyses.
‘Sr concentrations were determined by isotope dilution.
and conodonts were concentrated from the 20 to 180 mesh fraction using heavy
liquids (sodium polytungstate) and hand-picked. At
Scripps Institution of’Oceanography, specimens from
each sample were examined using scanning electron
microscopy (SEM) tloidentify possible sources of contamination such as crystal overgrowths. In a few samples small amounts of matrix material still adhered to
the specimens.
Each sample was leached for - 20 min in a sonic
bath with 1.5 M acetic acid. This process was repeated
until each sample 1o;st - 15% by weight. This procedure was selected based on leaching experiments carried out on four large specimens. These specimens were
dissolved in five sequential steps in which the amounts
dissolved were - lO%, - 20%, - 20% and - 25%.
87Sr/86Sr in the first leachate was significantly more
radiogenic, and the second leachate was slightly more
radiogenic, than succeeding leachates, while values for
the third, forth and final leachates generally agreed
within uncertainly. All specimens were re-examined
using the SEM after leaching. In most cases this technique removed any :remaining matrix material as well
as the surface layer of the conodont, thereby exposing
the crystallites which comprise the lamellae (Pietzner
et al., 1968; Barnes et al., 1973). Bertramet al. ( 1992)
also found that minor surface leaching improved the
interspecific correlations of 87Sr/86Sr in conodonts.
Diagenetic alteration of Sr isotopes is always a
potential problem, particularly for Paleozoic specimens. One of the best tests against this possibility is a
comparison of the i;sotopic ratios from widely distributed sites. Because seawater Sr isotopes are a global
residue was wet-sieved
signal, samples of the same age exposed to a variety of
environmental and diagenetic conditions should yield
the same isotopic ratio if they have not been altered.
With the exception of a few outliers, our data from a
range of sites within the U.S.A. and within Pakistan
agree well (Fig. 3b), but more significantly there is
excellent agreement between U.S. and Pakistan samples (Fig. 3c) over the time interval of overlapping
sections. In addition, our data generally agree with published data from other locations and, most importantly,
from other mineralogies, such as bulk carbonate and
low-Mg calcite brachiopod shells. These mineral
phases have different diagenetic susceptibilities than
apatite.
Following surface leaching, conodonts were dissolved in 1.8 iVHC1,spiked for Sr and Sm/Nd concentration measurements, and dried. A sheath of organic
matter commonly remained. This was oxidized by adding 25 ~1 each of concentrated HN03 and 9 N HCl. At
this point the sample was again dried, then processed
through cation-exchange column chemistry using standard techniques to separate Sr, Nd and Sm. Elemental
concentrations, and 87Sr/86Sr and ‘43Nd/‘*Nd were
analyzed by isotope dilution and thermal ionization
mass spectrometry (TIMS) . Blanks for this procedure
are 15 pg for Sr, 13 pg for Nd and < 1 pg for Sm.
Initially, Rb concentrations were also measured using
isotope dilution in order to correct *‘Sr/%r values for
in situ decay of Rb. However, Rb concentrations were
very low ( < 2 ppm) and the correction, which ranged
from O.OOOOO2
to 0.000007, was insignificant compared to analytical uncertainties. Thus the uncorrected
data may be as much as O.OOOOO7
low.
84
E. E. Martin, J. D.
Macdougall
/ Chemical Geology 125 (1995) 73-99
Table 2
Nd isotope and concentration data for Permian/Triassic conodonts
Sample
Weight
Age”
Sm
Nd
(CLg)
(Ma)
(ppm)
(ppm)
‘47Sm/‘44Nd
“=Nd/‘“Nd” (0)
‘“Nd?NdC
74.01
62.20
94.99
77.69
41.50
14.56
0.147
0.137
0.139
0.132
0.117
0.126
0.5 12060
0.512136
0.512206
0.512149
0.512190
0.512191
0.511819
0.511912
0.511979
0.511933
0.511998
0.511985
27
38
14
14
18
22
-9.7
-1.9
-6.6
- 7.5
-6.2
-6.5
8.07
7.71
47.67
38.40
34.56
128.80
0.127
0.135
0.224
0.512109
0.512139
0.5 12262
0.511901
0.511918
0.511895
68
18
14
-8.1
-7.8
- 8.2
11.40
19.50
44.70
86.30
0.154
0.137
0.512211
0.512216
0.511959
0.511992
14
14
-7.0
-6.3
18.90
10.84
26.99
29.75
95.74
48.10
109.00
116.90
0.119
0.136
0.150
0.154
0.512171
0.512148
0.512134
0.512116
0.511976
0.511925
0.511889
0.511864
21
27
18
27
-6.6
-7.6
-8.3
-8.8
6.27
2.47
2.23
4.81
39.60
14.41
12.83
29.95
0.096
0.104
0.105
0.097
0.512301
0.5 12343
0.5 12208
0.5 12255
0.512145
0.512173
0.512036
0.512096
14
19
24
21
-3.3
-2.8
-5.5
-4.3
1.26
3.04
4.86
7.60
17.66
27.90
0.103
0.100
0.105
0.512071
0.512104
0.5 12092
0.511902
0.511940
0.511920
62
68
59
-8.1
-1.3
-7.1
44.19
182.00
0.147
0.5 12023
0.511783
14
- 10.4
37.90
95.94
84.63
43.42
141.40
117.00
82.80
164.00
373.00
366.90
194.70
618.00
733.00
372.10
0.139
0.155
0.139
0.135
0.138
0.146
0.134
0.5 12050
0.5 12277
0.512212
0.512280
0.5 12273
0.5 12242
0.512196
0.511819
0.5 12023
0.511984
0.512059
0.5 12047
0.512003
0.511976
17
14
14
18
15
14
84
-9.7
-5.7
-6.5
-5.0
-5.2
-6.1
-6.6
(0
2C+
zNd(0
U.S.A.:
Skinner Ranch, Glass Mountains, Texas:
SR-1
SR-6
SR-7+8
SR-9
SR-10
SRI]+12
16
68
51
57
73
81
267.10
266.10
265.80
265.50
265.30
265.00
18.04
14.00
21.85
16.94
8.06
3.04
Split Tank, Glass Mountains, Texas:
ST-2
ST 3-5
ST-6
19
83
20
264.60
264.30
262.10
Del Norte Mountains, Texas:
TX l-3
TX 4-6
78
70
261.60
258.40
Bird Mine, Del Norte Mountains, Texas:
BM-1
BM-2
BM-3
BM-4
12
28
36
30
258.25
257.80
256.30
256.10
Baldwin Creek, Wyoming:
BC1+2
BC3
BC4+5
BC-6,7,9
80
61
63
67
264.45
264.10
262.25
259.75
Crittenden Springs, Neoada:
cs-1
cs-2
cs-4
24
68
29
250.00
248.75
248.20
PAKISTAN:
Saidu Wali, Khisor Range:
SA l-6
64
260.00
Kota Ladlian, Khisor Range:
KL-1
KL 2-5
KL-6
KL-7
KL-8
KL-9
KL-10
36
36
25
11
20
23
3
260.75
256.50
255.25
254.80
254.60
254.10
253.75
E.E. Martin, J.D. Macdougall / Chemical Geology 125 (1995) 73-99
85
Table 2 (continued)
Sample
Weight
Age”
Sm
Nd
(CLg)
(Ma)
(ppm)
(ppm)
2154.30
2!53.40
:!52.70
:!5 1.80
:!49.00
149.30
94.01
44.82
94.42
6.99
252.25
252.00
251.80
1251.70
‘47Sm/‘*ONd
‘.“Nd/‘“Ndb (0)
‘43Nd/‘UNdc(0
28
G.UCN
631.00
385.70
153.00
434.00
30.56
0.143
0.147
0.177
0.131
0.145
0.512174
0.512122
0.5 12038
0.5 12080
0.512133
0.511940
0.511881
0.511748
0.511865
0.511910
25
30
67
25
34
-1.3
- 8.5
-11.1
-8.8
-7.9
56.12
45.80
40.73
22.93
264.00
168.30
139.30
95.05
0.128
0.164
0.177
0.146
0.512134
0.5 12085
0.512140
0.512095
0.511925
0.511816
0.511851
0.511856
14
50
14
18
- 7.6
-9.8
-9.1
-9.0
251.80
251.60
250.75
240.75
82.20
110.00
30.25
10.32
347.20
405.00
116.90
43.60
0.143
0.164
0.156
0.143
0.512133
0.512164
0.512010
0.512169
0.511899
0.511895
0.511754
0.511935
21
22
19
26
-8.1
-8.2
-11.0
- 7.4
249.75
249.40
249.25
248.75
248.55
248.35
22.95
17.60
6.49
17.25
22.71
4.98
101.60
82.34
25.42
62.06
83.80
24.29
0.136
0.129
0.154
0.168
0.164
0.124
0.512174
0.512107
0.5 12053
0.5 12036
0.5 12022
0.512117
0.511951
0.5118%
0.511800
0.511761
0.511754
0.511915
19
32
23
21
14
25
-7.1
-8.2
- 10.1
- 10.8
-11.0
-7.8
PAKISTAN (cont.):
Chiddru Nala, Salt Range:
6
I
11
10
31
CD-3
CD-4
CD-5
CD-6
CD-7
Kathwai, Salt Range:
46
6
13
17
KAl+2
KA-3
KA-4
KA-5
Chatuwala Nala, Salt Range:
7
10
31
54
cw- 1
cw-2
cw-3
cw-4
Nammal Nab
57
18
58
44
58
47
NN-1
NN-2
NN-3
NN-5
NN-6
NN-7
“Ages reported for combined samples represent a weighted mean of individual samples.
“Nd isotopic value measured from conodonts. The ‘43Nd/‘“Nd value for the La Jolla standardis 0.511859. All ratios are fractionation corrected
to ‘“NdO/ ‘“NdO = 0.242436.
‘Nd isotopic values from the Permian and Triassic corrected for radiogenic production of ‘“Nd. ‘43Nd/‘UNdc,, = ‘43Nd/‘UNdC,,,- 14’Sm/
‘“Nd[exp(ht) - I].
“UncertaintyX 106. All samples have been assigned a minimum uncertaintyof f 14. 10e6, equivalent to the total variability of repeat standard
analyses.
ebNd(r) =
[‘43Nd/‘44Nd~,)I’143Nd/‘~NdlCHUR)
- 1] X l@. ‘43Nd/‘UNdo,,,(,,
All samples were analyzed on a W-54@ single-collector TIMS at Scripps Institution of Oceanography.
Final weights for samples composed of l-5 conodont
elements, or pieces of elements, ranged from 4 to 70
pg. Sr concentrations were 1600 ppm *40%
(Table 1). For isotope ratio measurements, sample
amounts of - 25 ng Sr were loaded on tantalum oxide
on a single tungsten filament and 300 ratios were collected at 1.5 V. The measured value and total variability
for *‘Sr/*‘%r in 7 1 analyses of NBS 987 Sr loaded with
this technique and analyzed over the interval of this
~0.512638. ‘43Nd/‘C(Nd(CHUR)(2MM.)
=0.512316.
project was 0.7 10260 f 0.000022. This external precision represents the minimum uncertainty assigned to
any sample. For smaller samples we used a variation
of this loading technique described by Birck ( 1986).
For lo-ng standards the absolute value and precision
using this small sample method is identical to that of
the tantalum oxide technique; however, for very small
samples the within-run uncertainty was often higher
(Table 1).
Nd concentrations were highly variable, ranging
from 8 to 600 ppm (Table 2). By analogy to fish teeth,
86
E.E. Martin, J.D. Macdougall /Chemical Geology I25 (1995) 73-99
rare-earth element (REE) uptake probably occurred
after deposition of the conodont on the sea floor. Variations in concentration, therefore, may relate to exposure time and redox conditions in this environment. As
might be expected, Sm/Nd ratios are far more consistent. At the beginning of the project the REE fraction
of several samples were combined prior to the Sm-Nd
separation to insure sufficient Nd was present for isotope ratio analysis. This practice was eventually discontinued as knowledge of the expected concentration
from each sample and the chemical yield improved. Nd
samples were loaded on a single Re filament and analyzed as NdO+ at 0.7 V for 300 ratios. Lower intensities
were used and fewer ratios measured for smaller samples. For 19 analyses of the La Jolla Nd standard analyzed during this project the measured e,,-value was
- 15.2 with a total range of k 0.27. Again, the withinrun uncertainty for samples containing very small
amounts of Nd (OS-2 ng) was generally higher
(Table 2).
5. Discussion
5.1. Sr isotope results
The Pakistani and U.S.A. data outline a very rapid
decrease followed by an even more rapid increase in
seawater 87Sr/86Sr in the Late Permian (Fig. 3;
Table 1) . Prior to the late Early Permian the distribution of data points is too sparse to define precisely the
shape of the isotope curve (Fig. 3a). A greater sample
density from the Artinskian/Kungurian boundary to
the Capitanian details s’Sr/%r values which decrease
at an average rate of 0.000062 Ma- ’ to a minimum
value of 0.70706 (Fig. 3b and c) . This rate of change
exceeds the average rate of increase for the past 40 Ma
(DePaolo and Ingram, 1985; Hess et al., 1986; Hodell
et al., 1989, 1990) and the minimum represents one of
the lowest points during the Phanerozoic with the possible exception of the Jurassic (Burke et al., 1982).
From the Capitanian through the P/Tr boundary seawater ratios increase at the very rapid rate of O.oooo97
Ma-‘, approximately two and a half times the average
rate of increase for the past 40 Ma. Based on data that
continue beyond the earliest Triassic, this increase persists into the Anisian stage of the Triassic (Burke et al.,
1982; Koepnick et al., 1990).
Veizer and Compston ( 1974)) Popp et al. ( 1986b),
Koepnick et al. ( 1990), Kramm and Wedepohl
( 1991), Nishioka et al. ( 1991), and Denison et al.
( 1994) have all published 87Sr/86Sr data for this time
interval. Because of differences in the time scales used
for the individual data sets, interlaboratory correlations
are exceedingly difficult. In addition, the method of
stratigraphic correlation is not always defined in other
studies, and it is likely it differs from the conodont
biostratigraphy applied to our data. Veizer and Compston ( 1974) and Popp et al. ( 1986b) assigned sample
ages based on sample location in the upper or lower
portions of a stage. In these cases, we assumed that
their stages are consistent with those utilized in this
study. Nishioka et al. ( 1991) used a very similar time
scale, thus we have used their age estimates shifted by
6 Ma to account for the revised P/Tr boundary age,
Denison et al. ( 1994) used the time scale of Harland
et al. ( 1990)) therefore their data have also been shifted
by 6 Ma and those from the Ufimian through Artinskian
adjusted slightly ( < 1 Ma). Kramm and Wedepohl
( 1991) only place their samples within the Zechstein
stratigraphy, therefore we were unable to correlate their
data exactly. By comparison to our 87Sr/86Sr values
their data appear to be from the Late Kazanian to
Changsingian, which is roughly equivalent to most estimates for Zechstein ages.
As Fig. 3a illustrates, some of the data points from
Veizer and Compston ( 1974) and Poppet al. ( 1986b)
and Nishioka et al. ( 1991) agree well with our values,
while others are much lower. The data recently published by Denison et al. (1994) are the only other set
with sufficient coverage to delineate the curve. The
correlation between their data and ours is impressive,
although their values tend to be lower in the latest
Permian, possibly due to differences in correlation
techniques or the time scales applied. As a consequence, the 87Sr/86Sr minimum appears to occur at a
slightly younger age in their data set than in ours.
As mentioned earlier, the calculated rate of change
of seawater 87Sr/86Sr is highly dependent on the time
scale employed. The numeric ages from Odin (1982)
or Forster and Wanington (1985) would produce a
more gradual change for the decreasing portion of the
curve than that in Fig. 3; however, the rate of increase
in the latest Permian would be _ 1.5 times more rapid.
Little change would result if the time scale of Palmer
( 1983) were used. Middle and Late Permian stages are
E. E. Martin, J.D. Macdougall /Chemical Geology 125 (I 995) 73-99
much longer in the time scale proposed by Menning
( 1989) than in Harland et al. ( 1990), thus both the
decrease and increase :inisotopic ratios would be more
gradual.
5.2. Sr geochemical model
Using equations that describe the Sr geochemical
cycle it is possible to estimate the magnitude and possible causes of change.s in the Sr budget that would be
required to produce the observed seawater 87Sr/86Sr
fluctuations during the Middle Permian to Early Triassic. There are three major sources of Sr to the oceans:
river water, which transports weathered continental
material to the oceans; hydrothermal fluids, which have
interacted with mid-ocean ridge basalts at high temperatures; and sediment pore fluids, composed of seawater that has been altered by carbonate dissolution
and recrystallization at low temperatures. According to
Palmer and Edmond ( 1989), the modem 87Sr/86Sr
ratios for these sources are estimated to be 0.7119,
0.7035 and 0.7084, respectively (the pore fluid value
was taken from Elderfield and Gieskes ( 1982) ) , while
the relative proportions of each are N 65%, * 30% and
* 5%.
The rate of change of seawater 87Sr/86Sris described
by the following equation:
&w
~=fwW,(R”-~sw)l
(1)
where N is the number of moles of Sr in the oceans; J,
is the flux of Sr (in mol a- ’) into the ocean from source
n; R, is the “Sr18?Sr ratio of source n; and Rsw is the
seawater 87Sr/86Sr ratio (Hodell et al., 1989; Kump,
1989; Capo and DeP,aolo, 1990; Richter et al., 1992).
As Kump ( 1989) pointed out, the pore-water flux is
small and its isotop:ic ratio approaches the seawater
value; thus, this source can be neglected without significantly affecting the model. Richter et al. ( 1992)
estimated that ignoring this flux introduced an uncertainty of N f 5% in their calculations of other fluxes.
We also ignore this term in the calculations. Thus Eq.
1 becomes:
d&w
N-==J,(R,-Rsw)
dt
+Jt,(&-Rsw)
(2)
where subscripts r and h symbolize the riverine and
hydrothermal inputs, respectively.
87
We test the sensitivity of the model to changes in the
various unknowns in order to determine which values
are most appropriate for the Permian, and to evaluate
the potential of each parameter to control seawater
87Sr/86Sr variations. Of the seven variables incorporated in this equation, the seawater 87Sr/S6Sr(R,,, and
the rate of isotopic change (d&,ldt)
are based on
analytical measurements, leaving five unknowns. As in
previous models (Kump, 1989; Capo and DePaolo,
1990; Richter et al., 1992) the 87Sr/86Sr ratio of the
hydrothermal component (R,,) is considered to be similar to the mid-ocean ridge basalt value and constant
through time. We have chosen a value of 0.703, which
is slightly less radiogenic than the ratio given by Palmer
and Edmond ( 1989), but it is probably a more accurate
reflection of the end-member hydrothermal contribution (Richter et al., 1992). This assumption further
reduces the number of unknowns to four. For the Cenozoic portion of the curve Richter et al. ( 1992) eliminated an additional variable by assuming that the
hydrothermal flux was proportional to sea-floor spreading rates; however, these are unconstrained for the Permian. For relatively recent times it may be safe to
assume that the modem estimates of R, and J,, are reasonable approximations of the true value, but this is
much less certain for the Permian. As already discussed, the combination of continental configuration,
sea level, and resulting climate conditions were probably quite unique for this period. These factors would
likely influence weathering and drainage patterns as
well as hydrothermal circulation, which in turn could
generate large changes in the geochemical cycle of Sr.
However, as discussed below, it may be possible to
constrain changes in the Sr isotopic ratio of the riverine
flux (R,) by examination of seawater Nd isotopes.
Input values for Rsw and dR,,ldt were calculated
from a smoothed spline fit to a modified data set
(Fig. 3c), with outliers eliminated from the data prior
to curve fitting. First the three oldest samples from
Texas were removed because of their anomalously
radiogenic ratios compared to all other data (Fig. 3a).
The increasing and decreasing segments of the seawater
curve were each then fit with a second-order polynomial. Any isotopic ratio which deviated more than
0.00025 ( -0.035%, or ten times analytical uncertainty) from these curves was removed, and new curves
were fit to the remaining data. This procedure was
repeated a second time, eliminating values which devi-
88
E.E. Martin, J.D. Macdougall /Chemical Geology 125 (1995) 73-99
ated by more than 0.00015 (0.02%). This iterative
smoothing process eliminated a total of four out of
thirty-four data points ( 12%) from the decreasing segment, and seven out of twenty-eight (22%) from the
increasing segment; however, four of those were from
a single Triassic locality (Nammal Nala) .
Earlier works (e.g., Burke et al., 1982) fit curves to
the minimum value measured for any given time interval, based on the assumption that diagenetic alteration
in a continental setting would tend to increase the isotopic ratio. However, in Fig. 3a it is clear that some of
the ratios reported in this study, as well as some from
earlier studies, are anomalously low relative to other
samples of the same age that exhibit a variety of mineralogies and come from widely distributed locations.
These anomalous samples probably do not represent
the best-estimate of seawater in the past. However, our
spline fitting technique did eliminate twelve points
from above the curve but only two from below
(Fig. 3c), suggesting that most anomalous samples
were indeed ones for which 87Sr/86Sr has increased.
We note also that a curve fit through the lowest values
would have slightly different absolute values, but
essentially the same shape as our spline-fit curve, and
therefore would not significantly alter our model calculations.
5.3. Nd isotope results
Like Sr, the Nd isotopic composition of seawater
potentially records changes in continental weathering;
however, there are two important differences. First,
there is no balancing input from sea-floor hydrothermal
systems (Michard et al., 1983; Piepgras and Wasserburg, 1985; Bertram and Elderfield, 1993). Instead,
input to the ocean comes entirely from Nd dissolved in
river water and adsorbed on river-borne and eolian
particulates (Goldstein and Jacobsen, 1987). Second,
the residence time for Nd in seawater is less than the
mixing time of the oceans (Elderlield and Greaves,
1982; Piepgras and Wasserburg, 1987; Jeandel and
Peng, 1989). As a result the ‘“Nd/‘#Nd ratios of
individual water masses are distinct, reflecting the
lithology and age of material weathered from the surrounding continents.
Nd isotopic compositions were determined for a
selected group of the samples analyzed for Sr. These
results are presented in Table 2 and Fig. 4, and con-
-12
““““““““1
Triassic
245
1
250
s
r”
w
-4
-
-6
-
-0
-
.‘I””
Middle Permian
Ka
Uf
KU
1
260
1
1 Early
1 A,
265
270
(Ma)
.I ‘1 ‘. 1’.
-2~
)
255
Age
(b)
1
LatePenian
WU ) Ca
Ch
‘. 1 “,
1, ”
I
u
0
-10
-
’
-12
Triassic
245
250
. ’
Late Permian
Ch 1 WU 1 Ca
255
Age
’
1 Middle
1
Ka
260
’ 1
Permian
Early
1 Uf 1 KU 1 Ar
265
270
(Ma)
Fig. 4. Permian/TriassiceNddata vs. age for samples from: (a) the
US; and (b) Pakistan.The arrows highlight decreasingNd isotope
trendsindicativeof increasedcontinentalinput.Fordatapoints without errorbars the erroris less than the size of the symbol. Symbols
as in Fig. 3.
stitute one of the first detailed records of short-term
‘43Nd/‘44Ndvariations in seawater. Changes in oceanic
Nd isotopic composition are considerably less systematic than those of Sr. This probably reflects the short
residence time, and therefore more rapid response, of
seawater Nd to changes in the inputs, as well as variability due to local rather than global phenomena. This
point is illustrated by the offset between Wyoming and
Texas samples (Fig. 4a). Both of these locations were
situated along the PaleoPacific margin during the Permian, but their distinct Nd isotopic ratios imply that
mixing between these sites was incomplete. Another
possible explanation for the scatter could be diagenetic
alteration, although, as we have argued earlier, the low
CAI for these samples suggests that this problem should
be minimal
E. E. Martin, J.D. Macdougall / Chemical Geology 125 (1995) 73-99
As is the case for modern phosphatic fish teeth, conodonts probably contained only ppb levels of Nd when
the animal was alive, and the current high levels of Nd
were incorporated following deposition on the sea floor
(Wright et al., 1984; Shaw and Wasserburg, 1985). To
dispute the claim that this “excess’ ’Nd representspore
fluids altered by surmunding sediment Staudigel et al.
( 1985) pointed out that fish teeth yield the same Nd
isotopic values as Mn nodules which are known to form
at the sediment-water interface in direct contact with
seawater. Several authors have demonstrated that ratios
for Mn nodules closely track seawater values and are
distinct from surrounding pelagic clays (Piepgras et
al., 1979; Elderfield et al., 1981; Goldstein and
O’Nions, 1981). In fact, Piepgras et al. ( 1979) demonstrated that the Nd isotopic ratios from the top of a
Mn nodule, which was in contact with seawater, and
the bottom of the nodule, which was in contact with
sediments, were identical, although the concentration
of Nd on the top was twice that of the bottom. This
suggests that seawater is the ultimate source of Nd to
the fish teeth, and paesumably therefore to conodonts
as well. There is also concern that phosphates continue
to take up and exchange Nd during burial. To address
this problem Bernat ( 1974), Staudigel et al. ( 1985),
Elderfield and Pagett ( 1986) and Wright et al. ( 1987)
all showed that recent fish teeth acquire very high concentrations of Nd wil;hin the top few mm’s of the sediment-water interface, and that there is no systematic
increase in this concentration with burial depth. It is
possible that much of this early uptake coincides with
alteration of the hydroxyapatite from the living fish to
the carbonate apatite of the fossil specimen; although
this association has yet to be studied in detail.
For our samples the most intriguing feature in the
Nd data is the dramatic decrease to more continental
values that occurred in the Late Permian in both the
western U.S.A. (Paleo-Pacific) and Pakistan (Tethyan) samples. This gross trend is apparent at every site
evaluated from 260 to 250 Ma regardless of the associated Nd concentration. Concentrations in the U.S.A.
samples are generally < 100 ppm Nd, while Pakistan
samples range from 100 to > 700 ppm Nd. It has been
suggested the Nd content of phosphates may be related
to their alteration history. In a study of REE patterns
preserved in fish teeth, Elderfield and Pagett ( 1986)
concluded that specimens from oxic, slow sedimentation rate, deep-sea sediments that had high Nd concen-
89
trations were the most likely to represent seawater
chemistry accurately. In contrast, Bertramet al. ( 1992)
found a correlation between Nd concentration and more
continental Sr and Nd isotopic signatures in conodonts,
such that specimens with higher concentrations were
more likely to be altered. We found no correlation
between Nd concentration and the deviation in 87Sr/
‘?jr from the spline fit for our seawater curve, and
conclude that for those samples there is no clear alteration signal for Nd that is related to concentration.
Assuming that the decreasing Nd values do reflect
changes in seawater chemistry, these data help constrain the possible causes of change in the Sr geochemical cycle. The Nd data indicate that the types of rocks
weathered from the continent were increasingly “continental”, which would also affect the 87Sr/86Srvalues.
In both locations the decrease in ‘43Nd/‘44Nd appears
to begin close to the minimum in the Sr curve, although
there may be small differences in timing between the
two groups of samples.
5.4. Sr model sensitivity and results
5.4.1. Sr content of seawater
We first examined the influence of oceanic Sr content
on other model parameters. Currently the number of
moles of Sr in the ocean (N) is 1.19~10”. As mentioned earlier, Holser and Magaritz (1987) have suggested that this value may have been lower in the
Permian. Our calculations indicate that evaporite formation would have decreased seawater Sr concentrations less than w 6%.
In Fig. 5a we show the influence of N on the riverine
flux required to produce the observed *‘Sr/?$r curve.
Calculations are shown for three values of N, ranging
from 30% below to 15% above the modem Sr content
of the ocean. Although somewhat arbitrary, these values encompass a total variation of almost 45% in N,
and are based on the sum of the inputs required to satisfy
the observed Permian seawater 87Sr/86Sr at the beginning of the model (265 Ma), divided by two estimates
of the residence time for Sr in the oceans: 2.5 and 4.0
Ma. Although a modem residence time of 4 Ma is
frequently cited (summarized in Elderfield, 1986),
improved data on the total Sr fluxes to the oceans (Palmer and Edmond, 1989) yield a value closer to 2.5 Ma
(Hodell et al., 1990). The model results were calculated for each value of N using: (1) a modem hydro-
90
E. E. Martin, J.D. MacdougaN/
Chemical Geology 12.5 (199s) 73-99
2.50E+iO
4
E
3j 2.OOE+lO
e
‘C
u
g
1.50E+iO
Fig.
.
l.OOE+lO
5a
. ’ . . ’ ’ I MiddlePermian
’ . ’ . . ’ I. Early
Triassic ChL&Permian
1Wu 1 Ca 1 Ka ] Uf ] Ku 1 A,
245
250
255
260
Age
265
’
O.OOE+OO
Triassic
270
245
’
Late Permian
Ch
Wu 1 Ca
[
250
Age
,
4.50E+lO
3.50E+lO
1
255
(Ma)
(d)
‘,’ Middle
’
Ka
’ ‘,’
Permian
EN
1 LH 1 Ku 1 Ar
260
&
266
270
(Ma)
. . , . .
,
. .
, .
.
-
Modern
Value
2.50E+lO
-
1.50E+lO
0.706
Triassic
245
’ ‘~t~p:rr;lian ’ Lddle~,,i’,’
Ch ( Wu ) Cs 1 Ka
1Uf 1Ku
250
255
Age
260
265
t’eri
1
050E+lO
t
Fig.
5d
....‘..,,‘....‘..,.‘,.,,
TriZSSiC
Ar
270
(MaI
245
250
1
LatePermian
Ch 1 WU 1 Ca 1
255
Age
1
Middle Permian
Earfy
Ka
Uf
Ku ( Ar
260
1
I
265
270
(Ma)
Fig. 5. Permian/Triassic“‘!W8%rmodel results for: (a) the calculatedriverineflux (Jr) given a range of values for the numberof moles of Sr
in the ocean (N), and assuming J,, = 1.47-10” mol a-’ Sr, R, =0.703 and R, =0.7110; (b) the calculatedhydrothermalflux (J,,) given a
rangeof values for the riverineflux (Jr). and assumingN= 1.19-10” mol Sr, R,, =0.703 and R, =0.7110; (c) the calculatedriverineisotopic
ratio (R,) given a range of values for the riverineflux (J,). and assuming N= 1.19-10” mol Sr, J,, = 1.47-10” and R,,=O.703; and (d) the
calculatedriver-meflux (Jr) given a range of values for the river-meisotopic ratio (R,), and assumingN= 1.19* 10” mol Sr, J,, = 1.47- 10” and
R,, =0.703.
thermal flux of 1.47 10” mol a- ’Sr, which is slightly
less than the value given by Palmer and Edmond
( 1989) to compensate for the fact that we ignore the
pore-water flux; (2) an 87Sr/86Sr of 0.703 for the
hydrothermal flux; (3) a riverine 87Sr/86Sr value of
0.7 110 which approximates the modern ratio (0.7 119)
minus the highly radiogenic values for rivers draining
the Himalaya; and (4) an assumption of steady-state
conditions.
The results indicate that the model is not particularly
sensitive to changes in N. A 45% change in N would
be compensated by a maximum change of only 7% in
the riverine flux (Fig. 5a) ; thus a 6% decrease in seawater Sr concentration due to evaporite formation
??
would have a negligible effect on other model results.
Therefore, in further calculations, we assume the modem value of 1.19*10” mol Sr for N.
5.4.2. Hydrothermalflux
In Fig. 5b we plot the hydrothermal flux that would
be required to match the observed seawater curve if
riverine flux values were held constant. Three different
rivet-me fluxes are used: the modem flux of 3.330 10”
mol a- ’ (Palmer and Edmond, 1989), as well as two
lower values ranging down to approximately half the
modern flux. The lower two values were calculated to
yield the observed Rsw at the start of the model (265
Ma) and at the 87Sr/86Sr minimum (257.25 Ma)
E.E. Martin, J.D. Macdougall /Chemical Geology 125 (1995) 73-99
assuming the modern value for J,,, 0.703 for R,,, and
0.7 110 for R,.
As mentioned previously, the hydrothermal flux is
probably directly related to mid-ocean ridge volume.
Complete assembly of Pangea at the end of the Early
Permian (Scotese and McKerrow, 1990) suggests
ridge volume was probably at a minimum at this time.
Based on the extremely low sea level this volume probably remained low into the Late Permian. Thus the
hydrothermal flux was likely to be less in the Permian
than today, requiring a smaller balancing Permian riverine flux than the modern value (Fig. Sb) .
It can also be inferred from Fig. 5b that changes in
Jh alone are not likely to have generated the observed
seawater 87Sr/86Sr variations. Even using the lower
riverine flux values, the hydrothermal flux would have
to have increased by N 15% during the Middle Permian, then decreased by N 50% from the Late Permian
into the Triassic. However, based on the distribution of
ridges, in particular the initial rifting of South China in
the Late Permian, thle hydrothermal flux should have
increased throughout the time interval investigated.
Even if the sense of change were correct, the Late
Permian-Triassic change is unrealistically large,
requiring a 50% decrease in ridge volume over only 10
Ma if the hydrothermal flux is directly proportional to
ridge volume. This compares to a predicted maximum
change of 10% per 10 Ma calculated for a single ridge
system in the Cenozoic (Kominz, 1984).
X4.3. Riverine flux
The riverine Sr flux required to explain the observed
seawater isotopic ratios is strongly anticorrelated with
its *‘Sr/*?Sr ratio. Fig. 5c shows the riverine isotopic
ratios required for thlethree riverine flux values used in
Fig. 5b, again assuming modern values for N, J,, and
Rh. The range of isotopic compositions fall easily
within the range reported for individual modern river
systems (Goldstein and Jacobsen, 1987; Palmer and
Edmond, 1989). However, the model results are for
global average inpms. Note that for the two lower J,values ( -45% and - 55% of the modem flux) high
input *‘Sr/s”Sr values are required for much of the Late
Permian and Early Triassic, We have argued above that
J, in the Permian was probably lower than today; thus
possible reasons for high riverine 87Sr/86Sr need to be
assessed.
91
As mentioned earlier, decreasing seawater Nd isotopes suggest the type of rocks weathered from the
continents were becoming more “continental” during
the Late Permian. Using our Nd data and the relationship between Sr and Nd isotopic ratios of dissolved
river runoff published by Goldstein and Jacobsen
( 1987) yields a shift in 87Sr/86Sr values of 0.7091 to
0.7104 for 259 to 255 Ma from the U.S. data, and
0.7082 to 0.7104 for 255 to 250 Ma from the Pakistan
data. Comparison of the U.S.A. results and Fig. 5c
illustrates that given the modem riverine flux, the calculated change in the riverine isotopic ratio could
account for the entire increase in seawater *‘Sr/*?Sr.
However, the relationship proposed by Goldstein and
Jacobsen ( 1987) is based on data from rivers, not seawater. Applied to the present-day ocean, it would predict *‘Sr/*‘?Sr values that are low compared to those
observed for seawater or global runoff. This may in
part reflect the different behavior of these two elements
in seawater, especially the difference in residence time,
and also the fact that the particulate Nd flux is not
accounted for. The geochemical cycle for Nd is still
poorly understood; for example, estuarine processes
apparently strongly impact the relationship between
riverine and seawater Nd isotopic compositions, but
the effect has not been quantified. For our Permian
samples, the Nd data provide information on only two
regions, but seawater *‘Sr/*?Sr depends on knowledge
of global inputs. Therefore, although the Nd data
strongly suggest that riverine *‘Sr/?Sr increased in the
Late Permian, it provides little information on the magnitude of that change. For the decreasing portion of the
seawater *‘Sr/*?Sr curve, in the Middle Permian, Nd
data are too few and too scattered to help constrain the
Sr input.
The two most likely reasons for rapidly increasing
riverine *‘Sr/@Sr are continental glaciations and Himalayan-style collisional events. Glacial activity can
remove surface sediment and produce rock flour from
old shield material, thereby greatly increasing the surface area of material with high 87Sr/86Sr exposed to
weathering (Armstrong, 197 1; Palmer and Elderiield,
1985; Miller et al., 1991; Zachos, 1993). However, the
only known glacial episode during the interval we
investigated is in the Early Permian and coincides with
decreasing Sr isotope ratios. There is no evidence of
glacial activity during the Late Permian increase in
87Sr/86Sr.
92
E.E. Martin, J.D. Macdougall/Chemical
Edmond ( 1992) argued that it is difficult to alter the
global riverine isotopic value substantially because of
the inverse relationship between Sr concentration and
87Sr/86Sr in most drainage basins (Palmer and
Edmond, 1989; Palmer and Edmond, 1992). However,
modem rivers draining the Himalaya pose an exception
to this relationship (Palmer and Edmond, 1989; Krishnaswami et al., 1992), which Edmond ( 1992) attributes to the formation of high-grade metamorphic rocks
during the continental collision, and the associated
redistribution of radiogenic Sr into phases more susceptible to chemical weathering. But available evidence concerning the sequence of tectonic events in the
Paleozoic leading to the formation of Pangea indicates
that the major continental collisions, and the erosion
from these events, preceded the observed seawater
87Sr/86Sr increase by tens of millions of years. The
youngest of these collisions resulted in the formation
of the Urals; however, this mountain range was not
located in the tropics where it would be most susceptible to weathering. Small Asian microcontinents and
arcs may have collided in the Late Permian (Nie et al.,
1990), but these would have lacked the radiogenic
material necessary to alter the global isotopic ratio of
river runoff. Thus, the source for increasing riverine
“Sr/‘“Sr suggested by decreasing eNd-values is difficult to identify. For this reason we believe it is unlikely
that the increase was dramatic enough to account for
the entire increase in the seawater isotopic ratio.
In the section above, we have shown that if other
parameters are held constant, it is unlikely that changes
in R,, N or J,, by themselves can account for the
observed large variations in seawater 87Sr/86Srduring
the time period between N 257 to u 250 Ma, although
an increase in R, was probably a contributing factor.
This leads us to the conclusion that there must have
been significant changes in the riverine Sr flux. Fig. 5d
illustrates the required variations in this parameter for
three different values for RP Again R,,, N and J,, were
assigned their modem values. The R, estimates are
based on present-day values: Goldstein and Jacobsen
(1987) report 0.7101 as the weighted average of 13
large rivers today, and this value is similar to the “Sri
‘?jr values calculated from the Nd isotope data; Palmer
and Edmond ( 1989) estimated that they surveyed 47%
of the total river runoff to obtain a global average of
0.7 119; and Capo and DePaolo ( 1990) and Richter et
al. (1992) chose the intermediate value of 0.7110,
Geology 125 (1995) 73-99
which also approximates the isotopic ratio of modem
runoff excluding rivers draining the Himalaya. It is
obvious from Fig. 5d that the riverine Sr flux required
to match the seawater Sr isotope data is quite sensitive
to even relatively small changes in the “Sr/@?Sr of the
river input. The total range in the riverine flux calculated using the three isotopic estimates is l.l-lO1o4.1 10” mol a- ‘, similar to model results for the last
100 Ma (Richter et al., 1992). The highest riverine
87Sr/86Sr 0.7119, requires the least variation in the
riverine Aux; however, this value is derived from the
present, a time when seawater 87Sr/86Sris at its highest
value since the beginning of the Phanerozoic. As discussed earlier, there is no apparent source for such
radiogenic runoff during the Late Permian. The lowest
value investigated, 0.7101, requires the greatest change
in the riverine flux, reaching values that exceed the
modem input. As we have argued earlier, based on the
probable low hydrothermal flux in the Permian, riverine
fluxes were likely to have been lower than the present
day. Thus, a value for R, near0.7 110 seems reasonable.
Assuming that the isotopic ratio of the flux also
increased over this interval, the magnitude of increase
in the riverine flux would be less than that illustrated
in Fig. 5d.
To summarize, we have investigated the sensitivity
of the various parameters affecting seawater *‘Sr/*‘?jr
to change in each of the other parameters. We conclude
that the major cause of decreasing seawater *‘Srjg6Sr
in the Middle Permian was a decrease in the riverine
flux, while the major factors leading to the large
changes in the seawater ratio in the Late Permian-Early
Triassic were probably an increase in the riverine isotopic ratio and the riverine Sr flux. The intermediate
curve in Fig. 5d illustrates a doubling of the riverine
flux over 12 Ma from 1.45*10” mol a-’ Sr in the
Middle Permian to 2.90 10” mol a-’ Sr in the Early
Triassic. This represents the maximum potential
change required in the riverine flux. The magnitude
would be significantly reduced by an increase in R,, the
probably smaller hydrothermal flux, and a lower oceanic Sr content due to evaporite deposition.
??
5.5. Implications of the model results
The model results (Fig. 5d) thus suggest that the
riverine Sr flux decreased by _ 10-152 during the
Middle Permian. This does not define the full extent of
E. E. Martin, J.D. Macdougall / Chemical Geology 125 (I 995) 73-99
the decrease because the seawater “Srls’?jr ratio actually began decreasing earlier in the Early Permian
(Fig. 3a). Geologic evidence suggests that the Hercynian and Uralian orogens were still topographic
highs subject to extensive erosion in the Early to Middle
Permian. Thus the decrease in seawater 87Sr/86Sr
implies that either the rocks exposed to weathering had
low 87Sr/86Sr ratios, or chemical weathering products
were not transported to the ocean.
The decreasing seawater isotopic ratios could be
attributed to extensive continentality and cool conditions during this interval. The vast continental interiors
would have been very dry (Kutzbach and Gallimore,
1989; Crowley and N’orth, 1991) and the proportion of
internal runoff compared to that reaching the oceans
may have been very high. In addition, evidence of glaciations at high southlern latitudes at this time suggests
that the global climate was probably cool and dry,
resulting in less chemical erosion than at warmer times
(Brady and Carroll, 1994), Another factor contributing
to cooler conditions and less intense chemical weathering may have been a drawdown in atmospheric CO*
caused by the initial fstagesof intense weathering associated with the orogenies. As Caldeira (1992) suggested, it is difficult to sustain high rates of chemical
weathering unless there are renewed sources of COZ in
the atmosphere, such as volcanism or metamorphism.
Evidence of arid conditions includes the deposition
of massive evaporites which characterize much of the
period, but are particularly prominent in the Kungurian
and Kazanian (Fig. 1). Additional evidence for
decreased continental runoff comes from a study of
nonluminescent Carboniferous to Kungurian age
brachiopods. Delaney et al. (1989) report that Li/Ca
ratios in these fossils are only -50% of the value
observed for modlern and Devonian specimens
(Fig. 1) , and speculate that this may be the result of
lower Li concentrations in seawater due to a lower flux
from the continents.
The transition from decreasing to increasing seawater 87Sr/8”Sr (and thus, in our model, to an increasing riverine isotopic ratio and flux) occurred in the
earliest part of the Late Permian. This is puzzling
because there are several factors which might be
expected to force a continued decrease in seawater
87Sr/86Sr at this time. These include: reduced rates of
erosion of Pangean oaogenies, which by this time would
have subdued topographic expressions; a lack of evi-
93
dence for glaciation; and climate models which suggest
persistent arid conditions into the Late Permian due to
the continental configuration.
As discussed, it is difficult to understand the source
of increasing riverine 87Sr/86Sr* however Erwin’s
(1993) analysis of conditions in the Late Permian
poses a possible solution to the increase in the riverine
flux. He suggested that atmospheric CO* levels may
have been high at this time, resulting in global warming,
which could lead to enhanced chemical weathering.
More intense weathering might also result in weathering of deeper erogenic roots, particularly in the Hercynian mountains located in the tropics, and therefore
might also explain the increase in riverine 87Sr/86Sr.
Although there is no direct record of paleo-CO2 concentrations, Erwin ( 1993) has compiled a convincing
list of possible Late Permian CO:! sources, including
release of methane from gas hydrates, oxidation of
organic material, and emissions from volcanism. As
sea level fell, large areas that previously lay within the
hydrate window (300-500 m below sea floor) along
the continental shelves and within intracratonic basins
would have been exposed, creating unstable conditions
for any hydrates. The methane released through this
process is an effective greenhouse gas, and it converts
to CO, over short time scales in the atmosphere.
Carbon isotope data provide evidence of another
possible source of atmospheric CO*. As mentioned earlier, Si3C-values were quite high ( N 3-4%0) during the
Early and Middle Permian, then decreased to values
less than zero during the Late Permian (Fig. 1) (Popp
et al., 1986a; Holser and Magaritz, 1987; Magaritz et
al., 1988; Baud et al., 1989; Gruszczynski et al., 1989;
Holser et al., 1989). This decrease may in part reflect
input of isotopically light carbon from release of gas
hydrates. Another viable source of light carbon would
be oxidation of organic carbon associated with falling
sea level and exposure of shallow shelves (Holser et
al., 1989; Erwin, 1993; Grossman, 1994). The net
result of these processes would be a decrease in the
organic carbon reservoir.
Large amounts of CO2 would also have been released
during the eruption of the Siberian Traps, which have
an estimated volume of > 1.5 - lo6 km3 of basalt. The
amount of CO1 released during this eruption could be
as much as lOI mol, or lo’* mol a-’ if spread evenly
over 1 Ma (Erwin, 1993 -extrapolated from McLean,
1985). For comparison, based on Bemer’s ( 1990,
94
E.E. Martin, J.D. Macdougall /Chemical Geology I25 (1995) 73-99
1991) model, the Permian atmospheric CO* reservoir
was two to four times greater than the modern reservoir,
on the order of 10L6-10” mol. Although the Siberian
Traps appear to have erupted after the minimum in
s7Sr/‘%r and the total output of CO* was relatively
small, they may have provided a source of CO* to
sustain enhanced weathering rates as CO* liberated
from other sources was consumed.
Berner ( 1990, , 1991) has modeled atmospheric
CO* levels in the Permian based on a set of criteria
distinct from those considered by Erwin ( 1993). Berner’s model is based on a combination of estimates of
volcanic and metamorphic degassing derived from seafloor spreading rates (ultimately from sea level), sediment burial rates estimated from isotopic data,
weathering rates calculated from the distribution, area
and elevation of the continents, land plant evolution,
and estimates of river runoff. He found low levels of
CO* in the Early Permian, similar to modern values,
but increasing concentrations into the Triassic, which
supports Erwin’s ( 1993) proposal.
A warmer, more humid climate in the Late Permian,
initiated by increasing atmospheric CO*, is also consistent with mid-latitude lithologic, fauna1 and floral
distributions. Yemane (1993) studied a region in
Africa located between 45’ and 60”s during the Late
Permian, for which climate models predict extreme
seasonal temperatures and arid conditions. Yet the geologic record is one of extensive lacustrine deposits,
therapsid fossils and Glossopteris vegetation, suggesting humid, temperate conditions. Taylor et al. ( 1992)
also describe a Late Permian fossil forest from a paleolatitude of W-85%. The fossils represent a deciduous
forest in which widely spaced growth rings and the
absence of frost rings indicate warm conditions at high
latitudes.
Finally, decreasing oxygen isotopes in the latest Permian of the Carnic Alps, Austria (Holser et al., 1989,
1991) support the idea of greenhouse warming.
Although this decrease has not been identified globally
and it could possibly be attributed to diagenesis or
salinity changes, Holser et al. (1991) argue that it represents a - 6°C rise in the temperature of seawater.
Currently most of the chemical weathering of the
continents occurs in a band surrounding the equator. If
this situation were also true for the Permian, orographic
effects of the low-latitude mountains generated by the
Hercynian megasuture may have had a significant
impact on chemical weathering at that time. During the
Late Carboniferous/Early Permian these mountains
may have cast a rain shadow over Pangea (Scotese and
McKerrow, 1990). As they eroded, the rain-laden
equatorial easterlies would have penetrated farther onto
the continent. Parrish (1993) has also suggested that
the intensity of monsoonal circulation (Robinson,
1973; Kutzbach and Gallimore, 1989), and therefore
the hydrologic cycle, increased throughout the Permian
to a maximum in the Triassic. Thus erosion of core
regions of equatorial mountains coupled with the development of monsoonal circulation may have accentuated the increase in chemical weathering, and thus the
riverine Sr isotopic ratio and flux, associated with
increased levels of atmospheric COZ.
Deep-sea sediment accumulation rates and the depth
of the carbonate compensation depth (CCD) are strong
indicators of the amount of chemical erosion from the
continents. Unfortunately, these deep-sea records are
generally unavailable for the Paleozoic. Ronov et al.
(1980) have analyzed sedimentation patterns from
shallow-water environments throughout the Phanerozoic, but their two-part division of Early and Late
Permian is too broad to isolate the increasing and
decreasing portions of the curve. They found almost no
change in average sediment accumulation rates
between these divisions of the Permian; however, they
did find a substantial decrease in the percent carbonate
deposited (by volume), accompanied by a large
increase in the percent marine elastics, in the Late Permian. For shallow-water environments this change in
lithologies is consistent with increased weathering and
erosion from the continents.
Thus we believe that the evidence for changes in
climate, and hence weathering rate, during the Permian
supports the conclusion that changes in the riverine Sr
flux as well as the isotopic ratio of that flux are responsible for changes in oceanic *‘Sr/%r during this time.
There is no need to invoke major changes in ocean
circulation as has been suggested by other authors. For
example, Gruszczynski et al. ( 1992) argued that under
stagnant, stratified conditions the surface waters of the
ocean would become increasingly enriched in radiogenie Sr from river runoff, while the deep waters
received less radiogenic hydrothermal fluids. Sudden
mixing of these water masses would rapidly reduce the
87Sr/86Srvalue of the surface waters. In this scenario,
the isotopic record preserved from shallow,water envi-
E. E. Martin, J.D. Macdougall / Chemical Geology 125 (1995) 73-99
ronments, such as those sampled for this study, would
initially record a gradual increase followed by a sudden
decrease, unlike the pattern illustrated by our Sr isotope
data. Holser (1977) suggested a scenario in which
rapid ocean mixing of an isolated brine pool could
explain the rapid rise of seawater S34S in the Early
Triassic; however, the Sr isotope minimum preceded
the S34Sminimum by several million years and the two
events are clearly separated by the P/Tr boundary.
6. Conclusions
A detailed record of seawater 87Sr/86Sr for the Permian indicates that a minimum value for Phanerozoic
time occurred in the Capitanian stage of the Late Permian. The isotopic composition of seawater for 10 Ma
prior to this point decreased at an average rate of
0.000062 Ma-‘, while values from the Capitanian into
the Triassic increased at an average rate of 0.000097
Ma- ‘. This increase is roughly equivalent to the steepest portions of the seawater 87Sr/86Srcurve for the past
40 Ma, and is approximately two and a half times
greater than the aver,age increase over that time period.
Model results indicate that these variations were primarily caused by fluctuations in the riverine flux of Sr
during the decreasing segment of the curve, and a combination of a change in the riverine flux and the isotopic
ratio of that flux to the oceans during the increasing
segment. Limited Nd isotope data support the theory
that the isotopic ratio of the flux became more “continental” during the Late Permian; however, it is difficult
to quantify the extent of this isotopic change. Continental weathering raites probably increased in the Late
Permian in conjunction with an increase in atmospheric
CO*. Thus, in contrast to the large-scale seawater “Sr/
*%r increase in the Cenozoic, which appears to be
controlled by weathering and climate changes associated with uplift of ,the Himalaya, the increase in the
Permian seems to have been influenced by changes in
weathering associatjed with other climate forcing factors, such as an increase in the concentration of atmospheric COZ, which resulted in global warming and
increased humidity. Erosion of mountains in the equatorial region, and their affect on the development of
monsoonal circulati,on and the transport of rain into the
continental interiors,, may also have played a role.
95
Acknowledgements
We particularly want to thank B.R. Wardlaw
(USGS, Reston, Virginia) for allowing us to dissolve
precious conodont samples from impoverished P/Tr
fauna. He generously provided all of the conodonts
analyzed for this study, as well as advice on P/Tr stratigraphy. D. Clark kindly supplied several conodont
specimens that were used for the leaching experiment.
We also want to thank C. MacIssac for assistance in
the laboratory, and G.W. Lugmair for analytical advice.
D.H. Erwin, W.T. Holser and G. Retallack read and
provided helpful comments on an early draft of this
paper, and L. Derry plus two anonymous reviewers
greatly improved the final version. This research was
supported by NSF (EAR9 1- 183 19).
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