Palaeoworld 16 (2007) 16–30
Research paper
The Capitanian (Permian) Kamura cooling event:
The beginning of the Paleozoic–Mesozoic transition
Yukio Isozaki a,∗ , Hodaka Kawahata b , Kayo Minoshima c
b
a Department of Earth Science and Astronomy, The University of Tokyo, Komaba, Meguro, Tokyo 153-8902, Japan
Graduate School of Frontier Sciences and Ocean Research Institute, The University of Tokyo, Minamidai, Nakano, Tokyo 164-8639, Japan
c Geological Survey of Japan, AIST, Tsukuba 305-8567, Japan
Received 4 January 2007; received in revised form 12 May 2007; accepted 15 May 2007
Available online 25 May 2007
Abstract
The Capitanian (late Guadalupian) high positive plateau interval of carbonate carbon isotope ratio (␦13 Ccarb ) was recognized
lately in a mid-Panthalassan paleo-atoll limestone in Japan as the Kamura event. This unique episode in the late-middle Permian
indicates high productivity in the low-latitude superocean likely coupled with resultant global cooling. This event ended shortly
before the Guadalupian–Lopingian (middle-late Permian) boundary (ca. 260 Ma); however, its onset time has not been ascertained
previously. Through a further analysis of the Wordian (middle Guadalupian) to lower Capitanian interval in the same limestone at
Kamura in Kyushu, we have found that the ␦13 Ccarb values started to rise over +4.5‰ and reached the maximum of +7.0‰ within the
Yabeina (fusuline) Zone of the early-middle Capitanian. Thus the total duration of the Kamura event is estimated over 3–4 million
years, given the whole Capitanian ranging for 5.4 million years. This 3–4 million years long unique cooling event occurred clearly
after the Gondwana glaciation period (late Carboniferous to early Permian) in the middle of the long-term warming trend toward
the Mesozoic. This cooling may have been a direct cause of the end-Guadalupian extinction of low-latitude, warm-water adapted
fauna including the large fusulines (Verbeekinidae), gigantic bivalves (Alatoconchidae), and rugose corals (Waagenophyllidae). The
Kamura event marks the first sharp excursion of ␦13 Ccarb values in the volatile fluctuation interval that lasted for nearly 20 million
years from the late-Middle Permian until the early-Middle Triassic. This interval with high volatility in ␦13 Ccarb values represents
the transition of major climate mode from the late Paleozoic icehouse to the Mesozoic–Cenozoic greenhouse regime. The endPaleozoic double-phased extinction occurred within this interval and the Capitanian Kamura event is regarded as the prelude to this
transition.
© 2007 Nanjing Institute of Geology and Palaeontology, CAS. Published by Elsevier Ltd. All rights reserved.
Keywords: Guadalupian; C isotope; Panthalassa; Permo-Triassic boundary; Extinction; Productivity
1. Introduction
The terminal Paleozoic mass extinction represents the
greatest in magnitude throughout the Phanerozoic life
∗ Corresponding author. Tel.: +81 3 5454 6608;
fax: +81 3 3465 3925.
E-mail address: [email protected] (Y. Isozaki).
history (e.g., Erwin, 1993, 2006); however, it was not
long time ago when its double-phased nature became
widely recognized. Jin et al. (1994) and Stanley and
Yang (1994) first pointed out that the Permian biodiversity declined in two steps separated clearly from each
other; i.e., first at the Middle-Late Permian boundary
(=Guadalupian–Lopingian boundary; G–LB) and second at the Permo-Triassic boundary (P–TB) sensu stricto
(or Changhsingian–Induan boundary).
1871-174X/$ – see front matter © 2007 Nanjing Institute of Geology and Palaeontology, CAS. Published by Elsevier Ltd. All rights reserved.
doi:10.1016/j.palwor.2007.05.011
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
In contrast to the P–TB issue, not much attention
has been paid to the G–LB event; however, the significance of the G–LB event was re-emphasized from a
different aspect relevant to the superocean Panthalassa.
The timing of the end-Guadalupian extinction apparently
coincides with the onset of the superanoxia in Panthalassa, i.e., another global scale geologic phenomenon
across the P–TB (Isozaki, 1997a, 2007). In addition to
the faunal turnover in mid-oceanic plankton (radiolarians) detected in deep-sea chert, shallow marine sessile
benthos (fusulines) also sharply declined in diversity
across the G–LB in mid-Panthalassan paleo-atoll complex (Isozaki and Ota, 2001; Ota and Isozaki, 2006).
These positively suggest the global nature of the G–LB
extinction and causal environmental change.
The mid-oceanic paleo-atoll carbonates also recorded
secular change in stable carbon isotope composition.
Musashi et al. (2001, 2007) and Isozaki et al. (2007)
first documented the secular change in carbonate carbon
isotopic ratio (␦13 Ccarb ) of mid-Panthalassa across the
P–TB and the G–LB, respectively. Besides the boundary negative shifts both at P–TB and G–LB properly
predicted from previous studies (e.g., Baud et al., 1989;
Holser et al., 1989; Wang et al., 2004), a unique high
productivity interval in the Capitanian (late Guadalupian) was newly detected on the basis of the appreciable
length of high positive ␦13 Ccarb (between +5 and +6‰)
interval (Isozaki et al., 2007; Fig. 1). As such high positive values over +5.0‰ are quite rare in the Phanerozoic
record except for several unique events in the Paleozoic (e.g., Veizer et al., 1999; Saltzman, 2005), they
named this Capitanian episode the “Kamura event”,
emphasizing its significance of global cooling and relevant extinction of large fusulines and gigantic bivalves
in low-latitude Panthalassa (Isozaki et al., 2007). In
the fusuline-tuned section, the waning history of the
Kamura event was clearly documented in high resolution, whereas the earlier history including the onset
timing was not yet revealed, owing to the absence of continuous exposure in the previously studied section. This
left a big chasm in our understanding of the major environmental change in the late Guadalupian, in particular
the cause and processes of the Kamura cooling event.
This study aimed to clarify the earlier stage of the
Kamura event, particularly focusing on the onset timing, and to bracket the total duration of the event. In
the same Kamura area in Kyushu, Japan, we analyzed
␦13 Ccarb chemostratigraphy of two other sections that
expose much lower parts of the Guadalupian (Wordian
to lower Capitanian) mid-oceanic paleo-atoll carbonates. This article reports the ␦13 Ccarb measurements and
discusses their implications to the Capitanian environ-
17
mental change and relevant extinction event. A particular
emphasis is given to the Kamura event in the context
of a long-term change in environmental regime during
the nearly 20 million years of the Paleozoic–Mesozoic
transition.
2. Geologic setting
The Permian and Triassic limestone at Kamura
(Takachiho town, Miyazaki prefecture; Fig. 2) in Kyushu
forms a part of an ancient mid-oceanic atoll complex
primarily developed on a mid-oceanic paleo-seamount
(Sano and Nakashima, 1997; Isozaki and Ota, 2001; Ota
and Isozaki, 2006). This limestone, like many other Permian limestones in Japan, occurs as an allochthonous
block incorporated in the Middle-Upper Jurassic disorganized mudstone/sandstone of the Jurassic accretionary
complex in the Chichibu belt (the tectonic outlier of the
Mino-Tanba belt; Isozaki, 1997b). The limestone blocks
in the Kamura area retain parts of the primary midoceanic stratigraphy that ranges in age from the Wordian
(middle Guadalupian) to Norian (Late Triassic) with several sedimentary breaks in the Triassic part (Kambe,
1963; Kanmera and Nakazawa, 1973; Watanabe et al.,
1979; Koike, 1996; Ota and Isozaki, 2006).
The Permian part consists of bioclastic limestone with
a typical Tethyan shallow marine fauna that includes
various fusulines, smaller foraminifera, large-shelled
bivalves, gastropods, brachiopods, rugose corals, and
calcareous algae. The Permian rocks are stratigraphically divided into the Guadalupian Iwato Formation (ca.
70 m thick) and the overlying Lopingian Mitai Formation
(ca. 30 m thick). Fusulines are the most abundant, and
they provide a basis for subdividing the Iwato Formation
into four biostratigraphic units; i.e., the Neoschwagerina
Zone, Yabeina Zone, Lepidolina Zone, and a barren interval, in ascending order (Ota and Isozaki, 2006; Isozaki
and Igo, in preparation). The overlying Lopingian Mitai
Formation is subdivided into two fusuline zones, i.e.,
the Codonofusiella-Reichelina Zone and Palaeofusulina
Zone (Kanmera and Nakazawa, 1973; Ota and Isozaki,
2006). All these fusuline assemblages and associated
fossils (rugose corals and large-shelled bivalves of Family Alatoconchidae; Isozaki, 2006) indicate that the
seamount was located in a low-latitude warm-water
domain in the superocean Panthalassa under a tropical
climate.
The Neoschwagerina Zone is correlated with the
Wordian (middle Guadalupian) of Texas and with the
Murgabian in Transcaucasia (Leven, 1996; Wilde et al.,
1999), while the Yabeina Zone, Lepidolina Zone, and
most of the barren interval are correlated with the Capi-
18
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
Fig. 1. Schematic diagram showing the late Guadalupian Kamura event documented by high positive ␦13 Ccarb values at Kamura in Japan (modified
from Isozaki et al., 2007) (A), and the composite Permian secular curve of ␦13 Ccarb values modified from Korte et al. (2005) (B). Road: Roadian,
Wor: Wordian. Note that the Guadalupian large fusuline and bivalve fauna became extinct in the middle of the Kamura cooling event, whereas
the post-extinction radiation of the Lopingian small fusulines started during the subsequent warming period. In contrast to the waning history of
the Kamura event, its onset timing and processes were unknown previously. In (B), two possible paths (broken lines) for the Guadalupian secular
change of ␦13 Ccarb values were shown by Korte et al. (2005); the lower for the Tethyan domain, the upper for the Delaware basin in Texas. The
Capitanian Kamura event recorded much higher positive ␦13 Ccarb values between +5.0 and +7.0‰ in Kamura, suggesting the positive excursion of
global context in the late Guadalupian. See text and Figs. 4 and 5 for details.
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
19
Fig. 2. Index map and stratigraphic columns of the three studied sections in the Kamura area, Kyushu. Not to scale. The present chemostratigraphic
research focused on the Wordian and lower Capitanian parts of the Iwato Formation exposed in Sections 1 and 3. Refer to Ota and Isozaki (2006)
and Isozaki et al. (2007) for more details of the area and Section 2. Loping.: Lopingian; Wuch.: Wuchiapingian; C-R: Codonofusiella-Reichelina.
20
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
tanian (upper Guadalupian) of Texas and with Midian
in Transcaucasia (Ota and Isozaki, 2006). The stratigraphic relationship between the Yabeina Zone and
the Lepidolina Zone has long been controversial (e.g.,
Toriyama, 1967; Ishii, 1990), however, our recent study
clarified that the former stratigraphically underlies the
latter within the Iwato Formation (Isozaki and Igo, in
preparation). The Codonofusiella-Reichelina Zone corresponds to the Wuchiapingian (Lower Lopingian) in
South China. For details of fusuline biostratigraphy and
age assignment, see Ota and Isozaki (2006) and Isozaki
(2006).
The Iwato Formation is exposed in three sections in
Kamura; i.e., Sections 1–3 from the east to the west
(Fig. 2). Section 2 (32◦ 44 58 N, 131◦ 20 02 E; Fig. 2)
at south of Shioinouso displays a continuous outcrop
of the upper Iwato Formation and the lower Mitai Formation that spans across the G–LB (Ota and Isozaki,
2006). In this section, a unique high positive plateau
in the Lepidolina Zone/barren interval and the following sharp negative shift in ␦13 Ccarb were documented
(Isozaki et al., 2007).
In the present study, we analyzed two additional sections in the Kamura area that expose the lower part of
the Iwato Formation; i.e., Sections 1 and 3 (Kambe,
1963; Murata et al., 2003; Isozaki, 2006; Fig. 2). Section 1 (32◦ 45 12 N, 131◦ 20 55 E) to the southeast of
Saraito village is composed of 57 m-thick limestone that
belongs to the Neoschwagerina Zone and Yabeina Zone,
whereas Section 3 (32◦ 45 05 N, 131◦ 19 52 E) to the
northeast of Hijirikawa is 8 m thick and entirely belongs
to the Yabeina Zone (Fig. 2). Detailed biostratigraphy of
these two sections is under scrutiny and results will be
published elsewhere (Isozaki and Igo, in preparation).
Among the three sections in Kamura, Section 1 represents the stratigraphical lowest, whereas Section 2 the
highest (Fig. 2). A slight stratigraphic gap may exist
between Sections 2 and 3; however, the similarity in
lithofacies suggests that the possible gaps are considerably small, if at all. The same fauna and lithofacies in
Sections 1 and 3 likewise indicate that a possible gap is
much smaller or even absent.
3. Samples and analytical methods
We collected dark gray to black limestone specimens
of the Guadalupian Iwato Formation for stable carbon
and oxygen isotope measurements at 34 horizons; i.e.,
21 from Section 1 and 13 from Section 3. Rocks of the
two sections are unmetamorphosed and mostly fresh,
and those with strong weathering and with many calcite
veins were screened out in the field and in the labora-
tory under the microscope. The black limestone of the
Iwato Formation has TOC around 0.1 wt% (Isozaki et
al., 2007).
The micritic part of wackestone from each horizon
was milled by microdrill after examining under the
microscope. Approximately 100 ␮m of the aliquot
samples were reacted with 100% H3 PO4 at 90 ◦ C
in an automated carbonate device (Multiprep) coupled with a Micromass Optima mass spectrometer
at the Geological Survey of Japan, AIST. Here,
␦13 C = [((13 C/12 Csample )/(13 C/12 Cstandard )) − 1] × 1000,
and ␦18 O = [((18 O/16 Osample )/(18 O/16 Ostandard )) − 1] ×
1000. All isotopic data are reported as per mil (‰)
relative to Vienna Pee Dee belemnite (V-PDB) standard.
The internal precision was 0.03‰ and 0.04‰ (1␴)
for ␦13 C and ␦18 O, respectively, based on replicate
measurements of 23 consecutive samples of the NBS-19
calcite standard (Suzuki et al., 2000).
4. Results
Table 1 lists all the measurements of ␦13 Ccarb and
carb of 47 samples from 34 horizons from Sections
1 and 3 in Kamura. Figs. 3 and 4 show secular changes
in ␦13 Ccarb values plotted on the stratigraphic columns
of Sections 1 and 3, respectively. All ␦13 Ccarb values
showed a wide range from +3.55 to +6.97‰, whereas
␦18 Ocarb values fluctuated between −7.58 to −12.36‰,
which might be partly due to a slight diagenetic alteration, however, the correlation between ␦13 Ccarb and
␦18 Ocarb indicates that they behaved independently. Thus
we consider that the ␦13 Ccarb values were not likely
affected by secondary alteration but reflect the primary
isotopic composition of the inorganic carbon reservoir in ancient seawater, in which the carbonates were
deposited.
At Section 1, the ␦13 Ccarb values range between +3.5
and +5.2‰. This section is divided into two parts in terms
of ␦13 Ccarb values; i.e., segment Sr-1 (Neoschwagerina
Zone; 32 m) and the overlying segment Sr-2 (Yabeina
Zone; 3.5 m) (Fig. 3). In the segment Sr-1, the ␦13 Ccarb
values gradually and steadily increased from +3.5 to
+4.3‰. On the other hand in the segment Sr-2, the
␦13 Ccarb values fluctuated between +4.4 and +5.2‰.
Although the boundary between the segments Sr-1 and
Sr-2 is covered, a general trend of increasing ␦13 Ccarb
values can be recognized in Section 1. The sample SrB35
in the segment Sr-2 marked the lowest horizon of high
positive ␦13 Ccarb values over +5.0‰ in the Iwato Formation.
At Section 3, the ␦13 Ccarb values range between
+5.0 and +7.0‰. This section is chemostratigraphi␦18 O
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
Table 1
Analytical results of ␦13 Ccarb and ␦18 Ocarb normalized to the Vienna
Pee Dee belemnite of the Guadalupian Iwato Formation in the Kamura
area, Kyushu
␦13 Ccarb (‰)
␦18 Ocarb (‰)
Section 3 (Hijirikawa)
Yabeina Zone (13 horizons)
Hj-0.5
7.50
Hj0
6.75
Hj1
6.15
Hj1.5
5.90
Hj2-1
5.45
Hj2-2
5.45
Hj3
5.15
Hj4
4.55
Hj5
4.15
Hj5.5
3.75
Hj6.1
3.35
Hj7.1-1
2.40
Hj7.1-2
2.40
Hj10
0.70
Hj11
0.20
5.332
5.817
5.425
5.936
6.872
6.970
5.791
5.681
5.191
5.502
5.005
5.046
5.008
4.776
4.857
−10.001
−8.480
−8.366
−11.370
−10.052
−8.559
−7.578
−9.013
−9.941
−8.697
−8.616
−10.414
−10.085
−11.163
−10.913
Section 1 (Saraito)
Yabeina Zone (8 horizons)
B41
56.8
B40
56.4
B39-1
56.1
B39-2
56.1
B38
55.3
B38Y1
55.3
B38Y2
55.3
B37
55.0
B37Y2
55.0
B37Y3
55.0
B36
54.7
B35
54.1
B34
53.8
4.575
4.478
5.051
5.235
5.013
5.056
4.941
4.932
4.774
4.917
4.704
5.050
4.430
−10.049
−8.744
−10.426
−10.824
−10.893
−10.789
−11.677
−11.356
−11.978
−11.891
−8.933
−9.350
−8.081
Neoschwagerina Zone (13 horizons)
B12-1
33.9
4.213
B12-2
33.9
4.384
B9
31.0
4.314
B6-1
28.0
4.042
B6-2
28.0
4.267
B5-1
27.7
4.128
B5-2
27.7
4.173
B4-1
27.0
4.123
B4-2
27.0
4.177
B3
26.0
4.201
51
25.5
3.935
B2
24.0
4.072
B1-1
21.0
3.550
B1-2
21.0
3.648
X-1
15.0
3.714
X-2
15.0
3.969
07-05
4.5
3.959
07-03
3.4
3.546
Z4
0.0
3.536
−10.435
−10.715
−11.773
−10.775
−9.899
−11.003
−11.245
−11.414
−12.362
−7.970
−10.139
−8.329
−10.890
−10.570
−10.602
−7.753
−9.192
−8.841
−9.747
Sample
Horizon (m)
21
cally divided into two parts; i.e., segment Hi-1 (3.5 m+)
and the overlying segment Hi-2 (2.2 m+) (Fig. 4). The
segment Hi-1 is characterized by a gradual increase
in ␦13 Ccarb values, whereas the Hi-2 by a reversed
decrease. The sample Hj-2 with +7.0‰ marked the highest ␦13 Ccarb value in the Iwato Formation.
In summary, the current C isotope analysis clarified
the following two facts: (1) the ␦13 Ccarb values keep
increasing from the Neoschwagerina Zone (segment Sr1; Wordian) to the Yabeina Zone (segments Sr-2 and
Hi-1; lower Capitanian) except for the upper part of the
Yabeina Zone (segment Hi-2); (2) all the ␦13 Ccarb values
of the Capitanian Iwato Formation range above +4.4‰
up to the highest value of +7.0‰ in the upper Yabeina
Zone.
5. Discussion
This study confirms the development of the Kamura
event in the late Guadalupian, and suggests that the
interval of this unique event has ranged stratigraphically
further downward. We will discuss here the geological implications of the new dataset, focusing on the
onset timing and total duration of the Kamura event with
respect to the end-Guadalupian environmental changes
and mass extinction, and particularly to the transition of
climatic regime from the late Paleozoic icehouse (Gondwana glaciation) to Mesozoic greenhouse.
5.1. Onset of the Kamura event
The present study has clarified that the lower part
of the Iwato Formation (Wordian Neoschwagerina Zone
and lower Capitanian Yabeina Zone) is thoroughly characterized by positive values of ␦13 Ccarb over +3.5‰
(Figs. 3 and 4). In particular, all the ␦13 Ccarb values of the Yabeina Zone both in Sections 1 and 3
exceed +4.4‰, and they range mostly in a high positive domain between +5.0 to +6.0‰. The Yabeina Zone
of the Iwato Formation in general has more or less
the same isotopic signature as the overlying Lepidolina
Zone and barren interval in which the Kamura event
was originally recognized (Isozaki et al., 2007). Thus
the interval of the Kamura event with ␦13 Ccarb values
over +5.0‰ ranges stratigraphically downward to the
Yabeina Zone.
On the other hand, the Wordian Neoschwagerina
Zone recorded relatively lower ␦13 Ccarb values between
+3.5 and +4.2‰, thus the Kamura event had not
yet started in the Wordian. However, the ␦13 Ccarb
record of the Neoschwagerina Zone clearly demonstrates
a steadily upward-increasing pattern, suggesting that
22
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
Fig. 3. Chemostratigraphy of stable carbon isotope of carbonates of Section 1 near Saraito in Kamura. Legends for columnar section are the same
as those for Fig. 2. This section is divided into two chemostratigraphic segments: Sr-1 and Sr-2.
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
23
Fig. 4. Chemostratigraphy of stable carbon isotope of carbonates of Section 3 near Hijirikawa in Kamura. Legends for columnar section are the
same as those for Fig. 2. This section is divided into two chemostratigraphic segments: Hi-1 and Hi-2.
the oceanographic condition started to shift gradually
already in the Wordian toward the extreme state of the
Capitanian with unusual enrichment of 13 C in seawater.
The sample SrB35 in the Yabeina Zone in Section 1 marks the lowest horizon with ␦13 Ccarb values
over +5.0‰, suggesting the lower limit of the interval
of the Kamura event. Unfortunately much lower horizon of the Yabeina Zone is covered and the boundary
between the Neoschwagerina Zone and Yabeina Zone
has not been observed in Kamura. Nonetheless, the
general secular trend of the ␦13 Ccarb record positively
indicates that the Kamura event has first emerged around
the Wordian/Capitanian boundary (265.8 Ma according
to the latest geological timescale by Gradstein et al.,
2004; Fig. 5). It is noteworthy that a strange condition
has appeared in the middle of the superocean around
the Wordian/Capitanian boundary because the Kamura
event may mark the first episode of large isotopic excursion in the Permian (Fig. 1B). Although the trigger for
this oceanographic change is unknown at present, the
24
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
Fig. 5. Schematic summary of the ␦13 Ccarb chemostratigraphy of the Guadalupian Iwato Formation and Early Lopingian Mitai Formation in Kamura,
showing the total range of the Kamura event. Not to scale. Note the main extinction occurred in the middle of the high positive plateau interval of
␦13 Ccarb values. C-R: Codonofusiella-Reichelina.
onset timing of the Kamura event should be checked
further carefully in continuous sections elsewhere in
order to examine whether or not this event started synchronously throughout the world.
5.2. Duration of the Kamura event
As discussed above, the Kamura event apparently
ranged through three successive fusuline zones of the
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
Capitanian; i.e., the Yabeina Zone, Lepidolina Zone,
and barren interval in ascending order (Fig. 5). The
basal part of the Yabeina Zone in Kamura is missing, whereas the uppermost part of the barren interval
is free from high positive ␦13 Ccarb values. Thus the
Kamura event likely spanned throughout almost the
entire Capitanian, except for the uppermost and possibly the lowest parts. This is supported by the data
from the GSSP of the G–LB at Penglaitan and Tieqiao,
South China, as there is no high positive plateau recognized in the uppermost Capitanian immediately below
the conodont-defined G–LB (Wang et al., 2004; Jin et
al., 2006).
Although detailed chronology of the three fusuline
zones of the Capitanian has not yet been established,
given the whole Capitanian ranging for 5.4 million years
from 265.8 Ma to 260.4 Ma (according to the timescale
by Gradstein et al., 2004), the total duration of the
Kamura event is estimated to be more than a half of the
Capitanian, probably 3–4 million years. Such a remarkable period characterized by an unusual positive ␦13 Ccarb
excursion has never been recognized in the Permian.
It is also noteworthy that the highest ␦13 Ccarb value
7.0‰ was detected in the sample Hj-2 in the upper
part of the Yabeina Zone, as no-such high positive value
has ever been reported from the Permian rocks (e.g.,
Grossman, 1994; Scholle, 1995; Korte et al., 2005). Thus
the maximum ␦13 Ccarb value in the Yabeina Zone suggests that the Kamura event may have culminated in the
early-middle Capitanian, held the similar condition for
a while, and finally collapsed quickly in the late Capitanian.
In general, extremely high positive ␦13 Ccarb values
indicate extraordinarily high productivity in the ocean.
As oceanic productivity is strongly controlled by nutrient availability, constant supply particularly of limiting
elements such as P and N is necessary to maintain longlasting high productivity. Saltzman (2005) compiled all
available ␦13 Ccarb measurements from the Paleozoic
(middle Cambrian to Carboniferous) rocks in the Great
Basin of western USA, and demonstrated a composite Paleozoic secular curve that is punctuated by nine
unique events of remarkable positive excursions by over
+3.0‰ lasted for more than a few million years. He
concluded that these nine events, detected also in different continents, represent intermediate cool climate
intervals between typical greenhouse and icehouse periods. By lowering the sea surface temperature, oceanic
circulation can be accelerated to bring sufficient nutrients from the deep ocean to the surface and this will
result in high primary productivity by blooming of
phytoplankton/cyanobacteria. As to the nutrient condi-
25
tion, however, the high productivity will not hold for a
long time because of negative feedback mechanism, if
world oceans are nitrogen-limited. In contrast, under a
phosphorous-limited condition, high primary productivity coupled with preferential organic carbon burial will
continue to keep seawater ␦13 C in high positive values for certain duration until effective recycling of P
stops (Saltzman, 2005). Although the Silurian (Ireviken)
event remains still controversial (Cramer and Saltzman,
2007), other eight Paleozoic cases with prominent positive ␦13 Ccarb excursion all suggest the appearance of cool
climate.
Accordingly, the late Guadalupian Kamura event was
nominated as the 10th case in the Paleozoic characterized by a remarkable positive ␦13 Ccarb excursion, and
the Kamura event likewise represents a transient cool
interval that appeared in the late Guadalupian (Isozaki
et al., 2007). The unique lithofacies of the Iwato Formation dominated by black to dark gray, organic-rich
(TOC ∼0.1 wt%) wackestone probably reflects the high
productivity in surface waters, as most of the Permian
paleo-atoll limestone in Japan has much lower TOC
less than 0.01 wt%. It is worth noting that this event
marks the first cooling episode, after the Gondwana
glaciation ended in the Artinskian (late Cisuralian; Jones
and Fielding, 2004), in the middle of the long-term
warming trend toward the generally warm Mesozoic
era.
In good accordance with the above interpretation,
the lately compiled Permian sea-level fluctuation curve
demonstrates that the Permian lowest-stand occurred
around the G–LB (Hallam and Wignall, 1999; Tong et
al., 1999). A major hiatus on the top of the Guadalupian Maokou Formation has been recognized extensively
in South China, and the top of the well-known Permian Reef complex in west Texas is unconformably
covered by the Lopingian evaporites (e.g., Mei and
Wardlaw, 1996). The “Permian chert event” in high
latitudes (Beauchamp and Baud, 2002) likely supports
the appearance of a cool period in the Guadalupian,
too.
5.3. Critical cooling
The end-Guadalupian is regarded as a timing of one of
the two major extinction events of the terminal Paleozoic
era (Jin et al., 1994; Stanley and Yang, 1994). Isozaki
(1997a, 2007) emphasized the geological significance
of the G–LB event from a viewpoint of the timing coincidence between two global geological phenomena; i.e.,
the biotic extinction and the onset of the P–TB superanoxia in the superocean. As to the cause of the G–LB
26
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
event, a global environmental change triggered by the
large-scale volcanism of the Emeishan Trap in South
China is currently favored by many workers (e.g., Chung
et al., 1998; Ali et al., 2002; Wignall, 2001); however,
details including possible direct kill mechanisms are not
yet fully clarified.
In this regard, the extinction of the Guadalupian
fauna in the middle of the Capitanian Kamura event
appears critical. The clear extinction pattern of large
fusulines (Verbeekinidae) and bivalves (Alatoconchidae) in the Capitanian part of the Iwato Formation
(Isozaki, 2006; Ota and Isozaki, 2006) suggests that
the claimed cooling of the Kamura event might have
played the key role in the kill scenario, in particular for
the creatures well adapted to a warm tropical climate
in low-latitude areas (Isozaki et al., 2007). The occurrence/distribution of the middle Permian Verbeekinidae,
Alatoconchidae, and Waagenophyllidae (rugose coral)
families was restricted in low-latitude shallow seas in
Tethys and Panthalassa, and the gigantism in fusulines and Alatoconchidae bivalves was probably due
to the symbiosis with photosynthetic algae/bacteria in
such oligotrophic environment in mid-ocean in general
(Isozaki, 2006).
In the late Capitanian, large fusulines were screened
out by size (Wilde, 2002; Yang et al., 2004; Ota
and Isozaki, 2006), aberrant Alatoconchidae bivalves
became totally extinct (Isozaki, 2006), and the diversity of rugose corals declined remarkably (Wang and
Sugiyama, 2000). The possible drop in sea surface
temperature in low latitudes may have caused a total
malfunction of photosymbiosis factory shared by the
above-mentioned “tropical trio” that were too much
adapted to warm-water environments to survive the
change. At Tieqiao, the last occurrence of large fusuline Metadoliolina (Verbeekinidae) was confirmed in the
Jinogondolella xuanhanensis Zone with ␦13 Ccarb values of +3 to +4‰ (Jin et al., 2006), suggesting that the
extinction of large fusulines slightly delayed in South
China probably owing to the local variability in water
temperature.
27
In addition, some gastropods and brachiopods
behaved similarly as the trio. For example, the
occurrence of extraordinarily large gastropods, such
as Bellerophon (13 cm in diameter), Pleurotomaria
(18 cm × 16 cm), and Murchisonia (40 cm in height),
were reported from the Capitanian limestone in Akasaka
(Hayasaka and Hayasaka, 1953), whereas all the gastropods from the overlying Lopingian are no more
than 1 cm in diameter. It is also noteworthy that some
early-middle Permian brachiopods originated in middlehigh paleolatitude domains (Attenuatella, Waagenites,
Strophalosiina, Comuqia) migrated to low-latitudes and
made their first appearance in the paleoequatorial zone at
the end of the Capitanian (Shen and Shi, 2002). Although
the Permian gastropods and brachiopods as a whole did
not experienced a remarkable diversity loss at the G–LB
(e.g., Pan and Erwin, 1994; Shen and Shi, 2002), these
observations likewise support the appearance of a cool
interval in the Capitanian Tethys and Panthalassa.
The G–LB is placed not at the extinction level of
the Guadalupian fauna but at the first appearance datum
(FAD) of the Wuchiapingian index conodont Clarkina
postbitteri postbitteri as defined at the stratotype section
(GSSP) at Penglaitan in South China (Jin et al., 1998;
Henderson et al., 2002). Owing to the absence of conodonts, the G–LB in Kamura is set at the horizon ca.
11 m above the main extinction level in the upper part of
the barren interval on the basis of the first appearance of
the Lopingian fusulines and ␦13 Ccarb chemostratigraphical correlation (Ota and Isozaki, 2006; Isozaki et al.,
2007). The main extinction occurred not at the G–LB
per se but in a much lower horizon in the midst of the
positive ␦13 Ccarb excursion interval. Thus an appreciable
time has elapsed between the end-Guadalupian extinction and the following radiation of the Lopingian fauna
in shallow mid-Panthalassa (Fig. 1A).
The present study demonstrated that the oceanic carbon cycle started to change in mid-Panthalassa around
265 Ma, at least by 4–5 million years earlier than the
G–LB (ca. 260 Ma). Should the claimed cooling have
been responsible for the extinction of the Guadalupian
Fig. 6. Secular change of ␦13 Ccarb in the Paleozoic and early Mesozoic, compiled from Saltzman (2005) for the Cambrian to Carboniferous, from
Korte et al. (2005) and this study for the Permian, from Payne et al. (2004) and Gradstein et al. (2004) for the Triassic, and from Palfy et al. (2001)
and Katz et al. (2005) for the Jurassic. Note the four distinct intervals of volatility with high positive ␦13 Ccarb excursion; i.e., Late Cambrian, Late
Ordovician–Silurian, Late Devonian–Early Carboniferous, and the Middle Permian–Middle Triassic (=Paleozoic–Mesozoic transitional interval;
PMT-interval). The PMT-interval from the Capitanian (Late Middle Permian) to Anisian (Early Middle Triassic) ranged for ca. 20–25 million
years, representing the transition from the late Paleozoic icehouse, centered by the Pennsylvanian to Early Permian Gondwana glaciation, to
the Mesozoic/Cenozoic greenhouse. The PMT-interval recorded a period of a transient cool climate with a P-limited oceanographic condition
between icehouse and greenhouse modes, and the Kamura event marks the beginning of the mode change from the Paleozoic icehouse to the
Mesozoic greenhouse. It is noteworthy that the two major mass extinctions (at the G–LB and P–TB) occurred during the PMT-interval, and that the
PMT-interval chronologically overlaps the P–TB superanoxic period (Isozaki, 1997a).
28
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
fauna, the ultimate cause of the environmental change
must have appeared by the early Capitanian time.
The Emeishan Trap volacanism in South China was
recently dated 256–259 Ma (Zhou et al., 2002). This
early Lopingian age is obviously too young for the trap
to be responsible for the environmental change in the
early Capitanian (ca. 265 Ma). In general, large-scale
basaltic volcanism likely drives the opposite consequence; i.e., global warming, rather than 3–4 million
year-long cooling. Thus the superficial correlation
between the end-Guadalupian extinction and the trap
volcanism needs re-consideration. Refer to Isozaki and
Ota (2007) for more details of relative timing between
the G–LB extinction and the Emeishan Trap volcanism.
5.4. δ13 Ccarb volatility in the Paleozoic–Mesozoic
transition interval
The Guadalupian major environmental change of
global context has appeared around 265 Ma (early
Capitanian) several million years earlier than the endGuadalupian mass extinction at the latest Capitanian.
In a long-term viewpoint, the Capitanian Kamura event
is particularly significant because it marks not only the
onset of wild ␦13 Ccarb fluctuations across the P–TB into
the early Anisian, middle Triassic, but also the first
remarkable positive ␦13 Ccarb excursion after a nearly 85
million years of relative quiescence (Fig. 6). In addition to the well-known sharp negative shift across the
P–TB (Baud et al., 1989; Holser et al., 1989; Musashi
et al., 2001), more positive and negative excursions of
greater magnitude occurred particularly in the early Triassic (Payne et al., 2004). As to the Lopingian, similar
␦13 Ccarb fluctuations are likely expected from the preliminary results (Baud et al., 1996; Shao et al., 2000;
Korte et al., 2005); however, the dataset is too immature
to document detailed secular change particularly for the
Wuchiapingian (lower Lopingian).
In sharp contrast to the Paleozoic–Mesozoic transition (PMT)-interval, such a wild fluctuation varying from
+8 to −3‰ has not been recognized in the Cisuralian to
early Guadalupian nor in the middle Triassic and later
part of the Mesozoic (Fig. 6). In fact, ␦13 Ccarb value
over +5.0‰ has never been recorded throughout the
Mesozoic and Cenozoic (e.g., Veizer et al., 1999; Katz
et al., 2005). Thus a volatile change in global carbon
cycle relevant to oceanography is restricted solely to the
∼20 million year-long PMT-interval from the Capitanian
(ca. 265 Ma) to early Anisian (ca. 245 Ma).
As pointed out by Saltzman (2005), there are three
other intervals in the Paleozoic that are characterized
by volatile fluctuations of ␦13 Ccarb values; i.e., the
Late Cambrian, Late Ordovician to Silurian, and Late
Devonian to Early Carboniferous (Fig. 6). The first
two intervals was described as a transient cool interval
between two greenhouse periods when the globe was
almost running into an icehouse but did not. Unlike these,
the rest two correspond to bona fide transient cool periods
between an icehouse and a greenhouse period. The late
Paleozoic (Carboniferous to Early Permian) was dominated by the icehouse climate centered by the Gondwana
glaciation, while the Mesozoic in total was governed by
warm greenhouse climate (e.g., Frakes et al., 1992).
It is noteworthy that this PMT-interval with high
␦13 Ccarb volatility approximately overlaps the superanoxic period in the superocean (Isozaki, 1997a).
Regardless of climatic modes, the deep-sea cherts both
of the Pennsylvanian–Guadalupian (icehouse interval)
and the Middle Triassic to Jurassic (greenhouse interval)
were well oxygenated. This indicates that the growth and
retreat of superanoxia have been controlled not solely
by the climate-dependent, global oceanic circulation but
also by other factors.
At any rate, a major re-organization of global
oceanography, including the global carbon cycle,
occurred during the PMT-interval, and this clearly separated the ancient regime of the Paleozoic and the new
one of the Mesozoic. The causes and processes of the
two major mass extinction events, at the G–LB and at
P–TB, should better be explained in the scope of such
long-term geological context.
After all, the late Guadalupian Kamura event preludes all these drastic change in the PMT-interval from
the Paleozoic to post-Paleozoic world, and the ultimate
trigger(s) of this major mode change in environment
can be found neither in the strict G–LB nor P–TB
intervals but likely in the upper Guadalupian rock
records.
6. Summary
The present study on the mid-Panthalassan paleo-atoll
complex clarified the following new aspects of the late
Guadalupian environmental change relevant to the mass
extinction:
(1) The Kamura event with high positive ␦13 Ccarb values
ranged for nearly 3–4 million years in the Capitanian
(ca. 265–260 Ma), late Guadalupian.
(2) The end-Guadalupian extinction occurred in the
middle of the Kamura cooling event.
(3) The Kamura event marks the beginning of the major
mode change of global climate and oceanography
from the Paleozoic to post-Paleozoic regime.
Y. Isozaki et al. / Palaeoworld 16 (2007) 16–30
Acknowledgements
This article is dedicated to late Lao Jin (Prof. Jin
Yugan) for honoring his great contributions to the Permian study. We thank Shen Shuzhong, Chen Siwei and
an anonymous reviewer for their constructive reviews,
Hisayoshi Igo for identification of fusulines, Teruhisa
Kasuya for drafting in part, plus Susumu Nohda and
Tomomi Kani for their help in fieldwork. This research
was supported by the Grant-in-Aid of Japan Society
of Promoting Science (no. 16204040 to YI and no.
17253006 to HK).
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