Shallow-water temperature seasonality in the middle Cretaceous mid-latitude northwestern Pacific

Temperature seasonality during the middle Cretaceous provides vital information about climate dynamics and ecological traits of organisms under the conditions of the “supergreenhouse” Earth. However, sub-annual scale paleotemperature records in the mid-latitude region remain limited. In this study, sclerochronological and stable oxygen isotope (δ¹⁸O) analyses of bivalve fossils from the northwestern Pacific (paleolatitude: 44°N) were used to estimate their life history and sub-annual scale temperature patterns of the middle Cretaceous. The materials studied included Cucullaea (Idonearca) delicatostriata and Aphrodina pseudoplana recovered from middle Turonian (middle Cretaceous) shallow marine deposits in Hokkaido, northern Japan. Growth increment width and shell δ¹⁸O of C. (I.) delicatostriata revealed that the growth rate was temporally maximized and then minimized, which can be interpreted as representing spring and winter growth, respectively. Approximately 25 fortnightly growth increments occurred within that cycle, suggesting that shell formation proceeded continuously throughout the year. Based on shell δ¹⁸O values, shallow-water temperatures from 28°C to 35°C with 7°C seasonality were estimated, under the assumption that seawater δ¹⁸O values were annually invariant at −1‰ relative to VSMOW. This temperature seasonality in the middle Cretaceous is more than 5°C smaller than the seasonality of modern shallow-water environments at the same latitudes. These findings, taken together with previous studies of other oceanic regions, suggest that the Northern Hemisphere had low seasonal shallow-water temperature variation of up to 10°C in the middle Cretaceous.

1 Introduction

The middle Cretaceous “supergreenhouse” period (Aptian–Santonian; 125–83 Ma; Huber et al., 2018; Petrizzo et al., 2022) is characterized by a warm global climate (Vandermark et al., 2007; Hay, 2008); high pCO₂ levels reaching approximately 1,000 ppm (Foster et al., 2017), nearly approximating the IPCC 6^th Assessment Report SSP3-7.0-lowNTCF scenario (Meinshausen et al., 2020); an absence of polar ice sheets (Moriya et al., 2007; MacLeod et al., 2013); and a much flatter equator-to-pole temperature gradient than today (Barron, 1983; Huber et al., 1995; Pucéat et al., 2007). This warm middle Cretaceous period is also associated with drastic faunal and floral transitions (Vermeij, 1977; Lidgard and Crane, 1988). Based on these characteristics, the middle Cretaceous period provides insights into a high-pCO₂ Earth, such as a possible future with accelerated anthropogenic CO₂ emissions. To properly characterize climate in the middle Cretaceous supergreenhouse period, seasonal-scale paleoenvironmental reconstruction is crucial because aspects of climate dynamics, ecology, and evolution of organisms have been controlled by seasonality throughout Earth’s history (Denton et al., 2005; Marshall and Burgess, 2015).

Fossil bivalve shells are commonly used to reconstruct sub-annual scale climatic fluctuations over geological timescales (Ivany, 2012; Walliser et al., 2015; Posenato et al., 2022). Stable oxygen isotope ratios (δ¹⁸O values) of marine bivalve shell carbonates (mainly aragonite and calcite) can be used for seawater temperature reconstruction because bivalve shells are precipitated in oxygen isotope equilibrium with ambient seawater (Epstein et al., 1953; Grossman and Ku, 1986). Bivalves form disconformities and translucent bands within their shells (called growth lines) when their shell growth stops or slows due to cyclic or irregular environmental or physiological changes (Kennish and Olsson, 1975; Ohno, 1989; Schöne et al., 2003). Furthermore, annual shell growth rates of bivalves ontogenetically decrease due to reductions in daily growth rates (Jones and Quitmyer, 1996), which are mainly observed in regions with small seasonality (Schöne et al., 2005; Reza Mirzaei et al., 2017) and increases in growth cessation interval, which can be observed worldwide (Hall et al., 1974; Tanabe, 1988; Jones et al., 1989; Tanabe et al., 2017, 2020). The timing of shell formation, temporal shell growth rate, and age of bivalves can be estimated by identifying the timing of growth line formation on shells and measuring the widths and numbers of these lines (Jones, 1983; Sato, 1997; Kubota et al., 2017).

Fossil bivalves with primary low-Mg calcite shells have often been used for climate reconstruction in deep time (Beard et al., 2015; de Winter et al., 2017, 2021) because low-Mg calcite is less susceptible to diagenetic alteration than aragonite. For example, δ¹⁸O values of rudists with large, thick, and low-Mg calcite shells have been preferentially used for seasonal-scale paleotemperature reconstruction in the middle Cretaceous paleoenvironmental studies (Steuber et al., 2005; Walliser and Schöne, 2020; Schmitt et al., 2022). However, because rudists were reef/bioherm builders that primarily inhabited low-latitude shallow oceans, seasonal-resolution sea surface temperature (SST) records for the middle Cretaceous have mostly been reported from the low-latitude Tethys Ocean (Steuber et al., 2005; Walliser and Schöne, 2020; Huck and Heimhofer, 2021). Because seasonality varies with latitude, SST seasonality must be investigated at low latitudes, as well as middle and high latitudes, to elucidate its global seasonal climatic trends. Furthermore, the mid-latitude region is the best area for identifying maximal seasonal temperature fluctuations (Ivany, 2012; Ivany and Judd, 2022). Despite its importance, however, few middle Cretaceous studies have included reconstruction of paleo-seawater temperature and its seasonality from carbonate fossil shells in mid-latitude regions. On the other hand, aragonitic shells produced by mollusks with sufficient thickness for sclerochronological and geochemical analyses from shell cross sections (e.g., Aphrodina, Cucullaea, and Glycymeris) were common at low latitudes, as well as middle and high latitudes, during the middle Cretaceous (Ando and Kodama, 1998; Squires, 2010; Hoşgör and Yılmaz, 2019). The sub-annual scale paleoenvironment, including shallow-water temperature, can be reconstructed from aragonitic fossil shells in deep time if evaluation of fossil preservation shows that shell aragonite is well-preserved (Walliser et al., 2015; Hall et al., 2018; Meyer et al., 2018; Posenato et al., 2022).

The Cretaceous Aptian–Maastrichtian Yezo Group in Hokkaido (northern Japan), deposited in the middle latitudes of the northwestern Pacific, contains a variety of well-preserved marine fossils, including ammonites, bivalves, and gastropods (Suzuki et al., 1998; Hikida et al., 2003; Yabumoto et al., 2012). The Yezo Group was mainly deposited in the deep ocean. However, some parts, such as the Mikasa Formation (latest Albian–Turonian), were deposited in a shallow ocean with abundant shallow-water bivalves (Tashiro, 1995; Ando and Kodama, 1998). This study presents detailed sclerochronological and oxygen isotope records of bivalve fossils from the Mikasa Formation to reveal seasonal-scale shallow-water temperature records and their life history.

2 Geological setting

Aptian–Maastrichtian forearc (Yezo Basin) sediments of the paleo-Asian continental eastern margin (Takashima et al., 2004), called the Yezo Group, are widely distributed from Soya to Urakawa areas, central Hokkaido, northern Japan (Figure 1A). Paleomagnetic studies showed that the paleolatitude of the Yezo Basin was approximately 44°N in the Late Cretaceous (Tamaki and Itoh, 2008; Kitagawa et al., 2016). An exception is the Urakawa area, which was at approximately 16.7°N during the Campanian (Late Cretaceous) before being transported to its present position (Tamaki et al., 2008). The Cenomanian–Campanian Yezo Basin was affected by a southward shallow-water current system comparable to the modern Oyashio Current based on neodymium isotope (¹⁴³Nd/¹⁴⁴Nd ratio) analysis of fish bone and tooth fossils from the Yezo Group (Moiroud et al., 2013). The mean annual SST in the middle–late Turonian Yezo Basin has been estimated as 26°C–29°C from planktonic foraminiferal δ¹⁸O data (Moriya, 2011).

Figure 1

Figure 1 Geological map of Hokkaido modified from Takashima et al. (2004) (A) and the Mikasa area modified from Kikuchi (2018) (B). (C) Map of the Ponbetsu River section modified from Ando and Kodama (1998), and details of the sampling locality. Me, Twa, Twb, Twc, and Twd in (C) represent sedimentary units of the Mikasa Formation in the Ponbetsu River section from Ando (1987). (D) Photograph of the sampled shellbed.

The Albian–Turonian Mikasa Formation, the Yezo Group is characterized by sandstone with hummocky cross-stratification (HCS) (Ando, 1987). The Mikasa Formation is distributed across the Sorachi–Ikushunbetsu Anticline (Figure 1B) (Ando, 1990b). The Mikasa Formation has various sedimentological facies that were deposited from the lagoon to the outer shelf, mainly under shallow-water conditions (Ando, 1990a). The Mikasa Formation is almost continuously exposed along the Ponbetsu River in Ponbetsu-cho, Mikasa City, central Hokkaido (Ponbetsu River section; Ando, 1987) (Figures 1C, 2) and is divided into five units based on sedimentary facies (Matsuno et al., 1963), designated Me (main part of the Middle Yezo Group, unit e), Twa, Twb, Twc, and Twd (Tw denotes the Trigonia sandstone in the western wing of the Sorachi–Ikushunbetsu Anticline) (Figures 1C, 2; Yabe, 1926; Matsumoto, 1951; Shimizu et al., 1955; Takashima et al., 2004). The Twc unit is approximately 32 m-thick and consist of fine- to very fine-grained sandstone with hummocky cross-stratification (HCS). This unit is considered to have been deposited from the lower shoreface to the shallow inner shelf, in waters tens of meters deep (Ando and Kodama, 1998). The Twc unit has been correlated with the middle Turonian based on the occurrence of an age-indicative ammonoid Subprionocyclus neptuni (Matsumoto, 1965) and inoceramid Inoceramus hobetsensis (Ando, 1987). The Twd unit is approximately 100 m-thick and mainly composed of medium- to coarse-grained sandstone with HCS (Ando and Kodama, 1998). The occurrence of I. hobetsensis in the lowest part and I. teshioensis in the uppermost part of the Twd unit indicates that the Twd unit was deposited in the middle to late Turonian (Ando, 1987).

Figure 2

Figure 2 Columnar section with sedimentary setting and deposition age of the Mikasa Formation at the Ponbetsu River section modified from Ando (1987, 1990b) and Ando and Kodama (1998). Black bars with photographs of bivalve fossils represent the stratigraphic occurrences of identified bivalves in the studied shellbed according to Ando and Kodama (1998). Gray zones indicate the suggested stratigraphic ranges of our studied materials.

3 Materials and methods

3.1 Field sampling

A boulder of calcite-cemented fine-grained sandstone containing abundant mollusk fossils (coquina) was collected alongside a forest road near the Ponbetsu River section (Figures 1C, D). In this rock, disarticulated bivalves (Aphrodina pseudoplana (Figures 2, 3D), Apiotrigonia (A.) mikasaensis (Figure 2), Cucullaea (Idonearca) delicatostriata (Figures 3A–C), and possibly Leptosolen sp.) and gastropods (possibly naticids) were observed. Most fossils from the coquina were identified as C. (I.) delicatostriata, and other species were observed as only one or a few individuals.

Figure 3

Figure 3 Photographs of the Cucullaea (Idonearca) delicatostriata (A: PU-1, B: PU-2, C: PU-5) and Aphrodina pseudoplana (D) shells used in this study. Scale bar = 10 mm.

Although this coquina was found as an isolated unrooted block, the shellbed was determined to be derived from the Twc unit or the lowermost part of the Twd unit because large boulders of fine-grained sandstone rich in C. (I.) delicatostriata were exclusively found near the outcrop of the Twc unit and lowermost part of the Twd unit in the Ponbetsu River section (Figure 1C); moreover, Apiotrigonia (A.) mikasaensis, a trigoniid bivalve restricted to the Turonian age (Tashiro, 1979), was observed in the collected coquina (Figure 2). In the studied succession, only Twc and Twd units are correlated with the Turonian, and the occurrences of A. (A.) mikasaensis in the Ponbetsu River section were limited to the upper Twc and lower Twd units (Ando and Kodama, 1998). These findings indicated that the studied shellbed originated from Twc or lower Twd units (Figure 2).

3.2 Sample preparation

The surfaces of the bivalve shells were carefully cleaned and exposed from the sandstone matrix. For this study, three shells of C. (I.) delicatostriata (Figures 3A–C; labeled PU-1, PU-2, and PU-5) and one shell of A. pseudoplana (Figure 3D) were selected for subsequent growth line observation and geochemical analysis, considering that these species have sufficiently thick shells for observation and geochemical sampling from cross-sections. Additionally, the abundant occurrence of C. (I.) delicatostriata allows intraspecific comparisons, thus improving data reliability. Shell surfaces were coated with epoxy resin to avoid mechanical damage during preparation. Shells were diagonally cut along their maximum growth direction using a diamond saw (Figures 4A, B). One side was polished with silicon carbide up to #3,000 (Figures 4A, B). Regarding C. (I.) delicatostriata, the other side of the cross-section was used to make a thin section (0.1 mm thickness) for growth line observation. Small pieces of shells were collected from the studied shells to evaluate the extent of diagenetic alteration and determine the diagenetic level according to preservation index (PI) (Cochran et al., 2010; Knoll et al., 2016) via scanning electron microscopy (SEM) (Figures 4A, B). The polished shell surface was etched with 5% acetic acid for 15 min prior to microstructure observation with SEM.

Figure 4

Figure 4 Sample preparation of Cucullaea (Idonearca) delicatostriata (A) and Aphrodina pseudoplana (B) shells. The valves were cut perpendicular to the maximum growth direction. One side of the cutting slab was polished to make a polished slab and the other side of each C. (I.) delicatostriata shell was thin-sectioned. Small fragments of shell were sampled by cracking each shell and then coating the cracked surface with platinum. The coated shell was used for preservation state estimation based on the PI scale (Cochran et al., 2010; Knoll et al., 2016). The surfaces of the thin section and polished shell were wetted and covered with cover glass during observation (C). Small holes in the polished sections are the remains of samples used for stable isotope analysis. OSL, outer shell layer; ISL, inner shell layer; HP, hinge plate.

3.3 Evaluation of the extent of diagenetic alteration

Evaluation of diagenetic effects on shell microstructure and mineralogy is essential for reliable paleoenvironmental reconstruction because diagenesis alters the original chemical composition of the shell (Cochran et al., 2010; Pederson et al., 2019). Shell microstructure and mineralogy of the studied shells were analyzed to determine the extent of diagenetic alteration. Thin sections and polished sections of the shells were carefully observed with a binocular microscope and a petrographic microscope to identify macro-scale diagenetic products, such as mosaic calcite crystals (Maliva et al., 2000; Pederson et al., 2019).

3.3.1 Raman spectroscopic analysis

Raman spectroscopic analysis was conducted to identify the mineral composition of the studied fossils using a confocal Raman micro-spectrometer (Nicolet Almega XR; Thermo Fisher Scientific, Bremen, Germany) in the Petrology and Mineralogy Laboratory of Nagoya University, Nagoya, Japan. A 532 nm Nd-YAG laser at a power of approximately 10 mW was irradiated to the thin sections and the polished sections of studied shells through a confocal microscope (BX51; Olympus, Tokyo, Japan) with a 100× objective (Olympus Mplan-BD 100X; numerical aperture, 0.90). The scattered light was collected using backscatter geometry from a 25 µm pinhole and a holographic notch filter, dispersed using 2,400 lines/mm, and analyzed using a Peltier-cooled charge-coupled device detector composed of 256 × 1024 pixels (Andor Technology, Belfast, Ireland). Raman spectra were collected in six periods of 10 s each for point analysis. Spectra between 100 and 1,290 cm⁻¹ were used to compare the spectra of calcium carbonate minerals (Parker et al., 2010; Wehrmeister et al., 2010).

3.3.2 SEM observation

The extent of diagenesis was determined using the PI, which was proposed for the evaluation of diagenetic alteration based on shell microstructure (Cochran et al., 2010; Knoll et al., 2016). First, the shell microstructure of a surface-etched polished slab coated with platinum was observed with SEM (SU6600; Hitachi High-Technologies, Tokyo, Japan) in the Laboratory of Geobiology at Nagoya University. Then, cracked shell surfaces coated with platinum were observed with SEM, and the preservation status was scored according to the PI (Cochran et al., 2010; Knoll et al., 2016).

3.4 Growth line analysis

Growth lines are observable as black-colored disconformities and translucent bands in accretionary-formed biogenic carbonates, which are caused by growth reduction or cessation. In bivalves, growth lines can be observed as concentric lines on the interior and exterior of the shells of various taxa; they were formed under environmental deterioration (e.g., tidal emersion and temperature extremes) or physical deterioration (Schöne et al., 2004; Schöne and Giere, 2005; de Winter et al., 2020). Growth lines tend to become darker with longer growth cessation. Growth lines act as records of the shell growth rate over various timescales (e.g., daily, fortnightly, and annually) (Kennish and Olsson, 1975; Ohno, 1989), indicating the timing of biological and environmental events (Kennish and Olsson, 1975; Sato, 1995). Therefore, growth lines can be used to elucidate organisms’ growth strategies (Tanabe, 1988; Sugiura et al., 2014; Posenato et al., 2022).

Growth lines were observed on thin sections of each shell under crossed Nicol polarization and polished slabs of each shell under stereoscopic microscopy. To reduce diffuse reflections and obtain clear images, the surfaces of thin sections and polished slabs were immersed in water and then covered with a cover glass (Figure 4C). Enlarged images of thin sections and polished surfaces taken using microscopes were stitched into single TIFF images with Adobe Photoshop software (Adobe, San Jose, CA, USA). Throughout the study, growth lines in the outer shell layer were used for investigation because these growth lines can be observed more clearly than growth lines in the inner shell layer and hinge plate. To accurately distinguish the shell growth lines, the greyscale value of the shell cross-section was measured along the direction of shell growth, in accordance with the methods of Carré et al. (2005) and Otter et al. (2019). In fossil shells, black stains may appear in areas other than growth lines due to shell coloration or crack formation during fossilization. To remove background colors, the image of the polished section was first subjected to black and white conversion followed by a rolling ball algorithm (Sternberg, 1983) for background subtraction (10-pixel diameter). The width of growth increments (space between neighboring growth lines) was measured along the growth direction. All image processing operations were performed using ImageJ (https://imagej.nih.gov/ij/).

3.5 Stable isotope analysis

Carbonate powders (approximately 100 µg) were collected from the outer shell layers of polished shells along their growth direction using a drill bit 0.4 mm in diameter (Minitor Co., Ltd., Osaka, Japan). Stable oxygen and carbon isotope compositions of bivalve shells (δ¹⁸O_shell and δ¹³C_shell, respectively) were measured with a Thermo Fisher Delta V Advantage isotope ratio mass spectrometer (Bremen, Germany) coupled to a ThermoQuest Kiel-III automated carbonate device (Bremen, Germany) at the Institute of Geology and Paleontology, Tohoku University, Sendai, Japan. Carbonate samples were reacted with 100% phosphoric acid at approximately 72°C. Both isotope ratios were expressed in conventional notation (δ, ‰) and calibrated to the NBS-19 international standard relative to Vienna Pee Dee Belemnite. The external precision (1σ) of the measured isotopic values was calculated based on repeated measurements of a laboratory reference sample (JCt-1) (Okai et al., 2004). The external precision (1σ) values of δ¹⁸O_shell and δ¹³C_shell were 0.05‰ and 0.03‰, respectively. Therefore, the error in seawater temperatures reconstructed based on the δ¹⁸O_shell values measured in this study was approximately 0.9°C at the 95% significance level.

3.6 Paleotemperature estimation

Mollusks precipitate aragonite their shells at or near oxygen isotope equilibrium with ambient seawater (Epstein et al., 1953; Grossman and Ku, 1986; Watanabe and Oba, 1999). Seawater paleotemperature was estimated using the equation proposed by Grossman and Ku (1986) with a correlation adjustment of −0.27‰ (Dettman et al., 1999):

$\begin{array}{ll}\text{T}(°\text{C})=20.60-4.34({\text{δ}}^{18}{\text{O}}_{\text{shell}}-({\text{δ}}^{18}{\text{O}}_{\text{seawater}}-0.27))& (1)\end{array}$

where T represents seawater temperature, and δ¹⁸O_shell and δ¹⁸O_seawater represent oxygen isotope values of shell carbonates (relative to VPDB) and seawater where the bivalves were grown (relative to VSMOW), respectively. In this study, −1‰ was used for δ¹⁸O_seawater according to the average δ¹⁸O_seawater value of ocean water during the Late Cretaceous ice-free world (Shackleton and Kennett, 1975) and numerically simulated δ¹⁸O_seawater values for the middle Cretaceous northwestern Pacific coastal region (Zhou et al., 2008). The validity of assuming a constant oxygen isotope ratio in the surrounding seawater is discussed in Section 5.3.

We reconstructed the monthly averaged paleotemperature of our study area from δ¹⁸O_shell values using the algorithm developed by Judd et al. (2018). Input parameter values in our study are shown in Supplementary Table 1, but the following conditions were assumed when setting the parameters:

(1) The numbers of solar days within a fortnight cycle and one year were 14.7 and 365, respectively.

(2) The days of the year with the coldest and warmest temperatures (T_pha) within one year were regarded as January 1 and July 1, respectively, in all cases.

(3) The day of the year with maximum growth (G_pha) was regarded as the day in the middle of the fortnight cycle with the widest fortnightly growth increment.

We used younger age records of C. (I.) delicatostriata shells for the calculations because younger age records have paleotemperature and seasonal growth rate with sufficient time resolution.

4 Results

4.1 State of preservation

The studied shells showed no macro-scale diagenetic alterations in thin sections and polished sections (Figures 5A, B, E), aside from the presence of dark green needle-like chlorite in the outer part of the outer shell layer of A. pseudoplana (Figure 5E). These needle-like crystals extend perpendicular to the shell surface, and none straddle the growth lines.

Figure 5

Figure 5 Polished section, thin section, and scanning electron microscopy (SEM) images of the studied mollusk materials. (A) Polished surface of C. (I.) delicatostriata (PU-1). (B) Thin section of C. (I.) delicatostriata (PU-2). Growth lines can be observed in both (A, B). “Major” growth lines are indicated by light green arrows, and “minor” growth lines can be observed between each pair of “major” growth lines. (C, D) SEM images of C. (I.) delicatostriata. (C) Crossed-lamellar fabric in the outer shell layer. (D) Complex crossed-lamellar fabric in the inner shell layer. (E) Polished surface of A. pseudoplana showing clear growth lines. White arrows indicate chlorite-replaced parts. (F–H) SEM images of A. pseudoplana. (F) Irregular simple prismatic structures in the growth line (sandwiched between the light green lines); (G) crossed-acicular fabric in the OSL; and (H) fine complex crossed-lamellar fabric in the ISL.

The crossed-lamellar structure in the outer shell layer (Figure 5C) and complex crossed-lamellar structure in the inner shell layer (Figure 5D) were clearly observed in C. (I.) delicatostriata shells. In the A. pseudoplana shell, a crossed-acicular structure in the outer shell layer (Figure 5G), fine complex crossed-lamellar structures in the inner layer (Figure 5H), and irregular simple prismatic structures along the growth lines (Figure 5F) were preserved.

Raman spectroscopic analysis showed clear peaks at 153, 181, 206, 702, 706, and 1,085 cm⁻¹ in all studied shells (Figure 6). The peaks at 153, 181, 206, 702, and 706 cm⁻¹ together indicate the presence of aragonite (Parker et al., 2010); the peak at 1,085 cm⁻¹ is the intrinsic Raman band of carbonate mineral (Wehrmeister et al., 2010). Considering that modern bivalves in the order Arcoida and the family Veneridae, which include C. (I.) delicatostriata and A. pseudoplana, respectively, produce aragonite shells (Kitamura, 2018), the two species also likely formed primarily aragonite shells. Therefore, the studied shells showed preservation of primary microstructure and mineralogy (aragonite) as indicated by SEM images and Raman spectroscopy (Figure 6).

Figure 6

Figure 6 Raman spectra acquired at 100–1,290 cm⁻¹ from the outer shell layer of all studied shells along with aragonite Raman spectra (aragonite R150021 with a 532 nm laser; red) and calcite Raman spectra (calcite R150075 with a 532 nm laser; gray) from RRUFF (rruff.info). Black dashed lines represent the positions of 153, 181, 206, 702, 706, and 1,085 cm⁻¹. Peaks at 153, 181, 206, 702, and 706 cm⁻¹ indicate the presence of aragonite and the peak at 1,085 cm⁻¹ indicates carbonate mineral.

In high-magnification SEM images, irregular thickness and fusions of crystals were observed in C. (I.) delicatostriata (Figures 7A–C). The PI of C. (I.) delicatostriata crossed-lamellar structures was 3 (good) for PU-1 (Figure 7A), 1 (poor) for PU-2 (Figure 7B), and 3 for PU-5 (Figure 7C) (Knoll et al., 2016). The PI of the A. pseudoplana shell could not be determined because the PI of the crossed-acicular structure has not been established (Figure 7D).

Figure 7

Figure 7 High-resolution images of the studied Cucullaea (Idonearca) delicatostriata (crossed-lamellar structure; PU-1, A; PU-2, B; PU-5, C) and Aphrodina pseudoplana (crossed-acicular structure; D) shells. Dissolution marks and irregular thickness of crystals can be observed in (A–C). Preservation index values of 3 for PU-1, 1 for PU-2, and 3 for PU-5 were assigned. Indistinct structures can be observed in (D), but the PI value could not be calculated because a PI for crossed-acicular structures has not been established.

4.2 Growth line observation

Growth line visibility in the C. (I.) delicatostriata shell was compared between thin sections and polished sections of the same individuals (Figure 8). In the younger portion of the shells, growth lines of varying distinctness were clearly observed in both thin sections (Figure 8B) and polished sections (Figure 8D). Observation of polished sections under reflected light was more suitable for detecting faint growth lines than thin sections due to the brightness of the views. Near the ventral margin, prominent growth lines were observed in the thin sections (Figure 8C), while these lines were not clearly visible in the polished sections (Figure 8E). These results indicated that thin sections are suitable for observing growth lines throughout all ontogenetic stages, while polished sections are suitable for observing growth lines in younger portions of the shells. To monitor the shell growth pattern of C. (I.) delicatostriata throughout its lifetime, the most effective approach is to use thin sections and measure the width of the prominent growth increment along the growth direction.

Figure 8

Figure 8 Comparison of growth line visibility in the outer shell layer between thin sections (A–C) and polished sections (D–F) of a C. (I.) delicatostriata shell (PU-5). The growth lines are clearly visible in the younger portions of both thin sections (B) and polished sections (D). In the shell near the ventral margin, growth lines are clearly visible in the thin sections (C) but are indistinct in the polished sections (E). Sampling holes in the polished sections (D–F) are the remains of samples taken for stable isotope analysis.

The growth lines observed in thin sections and polished sections can be divided into two categories, prominent (“major”) and faint (“minor”) growth lines (Figures 5A, B, 9), according to their darkness. Major growth lines were darker than minor growth lines according to processed greyscale levels (Figure 9). The processed greyscale values more clearly reflected the blackness of growth lines than the raw greyscale values because the color inversion process and rolling ball algorithm subtracted background color heterogeneity (Figure 9). Notably, approximately 15 minor growth lines with low processed greyscale values were observed between neighboring major growth lines with high processed greyscale values in the polished sections (Figure 9).

Figure 9

Figure 9 Polished section of the younger portion of a C. (I.) delicatostriata (PU-1) shell and greyscale values (raw and processed) along its growth direction in polished sections (pink line). Processed greyscale values were obtained by inverting the raw greyscale data and then applying the rolling ball algorithm. Raw greyscale data are shown as a black bold line and processed greyscale data are represented by red lines. Note that a high raw greyscale value is indicated by a light color, while a high processed greyscale value is represented by a dark color. The vertical black lines on the graph indicate the positions where growth lines were visually observed on the analyzed greyscale line. Solid lines indicate “major” growth lines and dashed lines indicate “minor” growth lines. The gray zones on the greyscale graph indicate the positions of dark staining. Numbers on the black arrows between lines represent the numbers of growth lines observed in the polished section. The estimated tidal condition (timing of spring tide and neap tide) is noted below the greyscale graphs.

4.3 Growth increment width

“Major” growth increment widths along the growth direction in thin sections of three C. (I.) delicatostriata shells are shown in Figure 10. The growth increment width varied among individuals due to intraspecific differences in growth rate and differences in the angles between the maximum growth direction and the cutting plane of each shell during thin section preparation.

Figure 10

Figure 10 Fortnightly growth increment width of C. (I.) delicatostriata, δ¹⁸O_shell and δ¹³C_shell values of all studied shells with temperatures reconstructed based on Grossman and Ku (1986), and photographs of thin sections used for growth increment width measurement and polished slabs used for stable isotope analysis. Numbers above arrows indicate the number of growth lines between each pair of arrows. Approximately 25 lines are observed between the arrows, indicating a period of around 1 year. Bold black arrows on the graph of fortnightly growth increment width indicate periods of ontogenetic growth reduction. Squares on the graph of δ¹⁸O_shell represent the positions of the annual growth increments used for monthly averaged paleotemperature calculations (Figure 13; Supplementary Table 2). For paleotemperature reconstruction from shell δ¹⁸O, a value of −1‰ was used for δ¹⁸O_seawater, obtained from Shackleton and Kennett (1975) and Zhou et al. (2008).

4.4 δ¹⁸O_shell and δ¹³C_shell values

The δ¹⁸O_shell profiles of the three C. (I.) delicatostriata shells (PU-1, 2, 5) and one A. pseudoplana shell showed clear cyclic variations of various amplitudes (Figure 10). The ranges of δ¹⁸O_shell values were from −4.40‰ to −3.06‰ (average = −3.65‰, σ [standard deviation] = 0.40‰) for PU-1, from −5.12‰ to −3.31‰ (average = −4.38‰, σ = 0.43‰) for PU-2, from −5.33‰ to −2.93‰ (average = −3.70‰, σ = 0.39‰) for PU-5, and from −4.49‰ to −3.21‰ (average = −3.85‰, σ = 0.30‰) for A. pseudoplana (Table 1).

Table 1

Table 1 Stable carbon and oxygen isotope values, paleotemperature estimated by Grossman and Ku (1986), and preservation status, as indicated by the preservation index (PI) according to Cochran et al. (2010) and Knoll et al. (2016), of the studied shells.

In contrast, the δ¹³C_shell profiles of all analyzed shells did not show significant cyclicity compared with δ¹⁸O_shell (Figure 10). The ranges of δ¹³C_shell values were from −4.02‰ to 0.31‰ (average = −2.06‰, σ = 0.94‰) for PU-1, from −6.62‰ to −1.85‰ (average = −3.71‰, σ = 1.13‰) for PU-2, from −3.41‰ to −0.28‰ (average = −1.78‰, σ = 0.71‰) for PU-5, and from −2.59‰ to 0.48‰ (average = −0.78‰, σ = 0.54‰) for A. pseudoplana (Table 1).

5 Discussion

5.1 Evaluation of fossil preservation

5.1.1 State of preservation and shell microstructure

Shells composed primarily of aragonite are susceptible to diagenetic alteration and thus rarely retain their primary mineral composition and microstructure (Walliser et al., 2015; Posenato et al., 2022). With regard to the shell microstructure of C. (I.) delicatostriata, the crossed-lamellar structure in the outer shell layer (Figure 5C) and complex crossed-lamellar structure in the inner shell layer (Figure 5D) are same as those of C. labiata, an extant Cucullaea species (Taylor et al., 1969). The outermost part of the outer shell layer was absent although the shell was buried in a sandstone matrix (Figure 8C), indicating that the fossil shells were physically abraded or chemically dissolved before or during the fossilization process.

The A. pseudoplana shell showed a crossed-acicular structure in the outer shell layer (Figure 5G) and fine complex crossed-lamellar structure in the inner layer (Figure 5H). The needle-like crystals of chlorite present in the shell were considered pseudomorphs of biogenic aragonite produced by diagenetic alteration based on their morphology, distribution, and orientation. Silicification of biogenic carbonate with preserved microstructures has been observed in other areas (Suzuki et al., 1998; Zakhera et al., 2001; Foster et al., 2022). The arrangement of growth lines and crystals in the outermost part of the outer shell layer of A. pseudoplana is similar to some venerid bivalves (Chioninae and Mercenaria stimpsoni: Shimamoto, 1986); its shell microstructure is similar to the arcticid bivalve Arctica islandica (Dunca et al., 2009; Höche et al., 2021).

Dissolution traces and irregular crystal thickness were observed in high-magnification SEM images (Figure 7). A PI value of 3 was assigned for PU-1 and PU-5, whereas a value of 1 was assigned for PU-2; all of these values were lower than the best PI of 5 (classified as “excellent”; Cochran et al., 2010; Knoll et al., 2016), suggesting that the studied bivalves were affected by low-level diagenetic alteration. The partial chloritization of the surface of the A. pseudoplana shell is consistent with diagenetic alteration. However, chloritization of shell aragonite did not reach the parts used for stable isotope analysis in this study (Figure 5E).

The studied shells were buried in fine-grained sandstone. Biogenic carbonates in such coarse host rocks are considered highly susceptible to diagenetic alteration due to their high permeability and porosity. Nevertheless, the studied shells retained the initial shell microstructure and mineralogy although they had been slightly dissolved. The good state of preservation may be attributable to the well-cemented sandstone, which effectively limited pore water reaching the fossil sites. Well-preserved biogenic carbonate fossils buried in sediments with reduced influence of pore water are consistent with the finding that fossils of planktonic foraminifera in clay-rich sediments showed better preservation than such fossils in coarser-grained sediments (Pearson et al., 2001).

Yamamoto et al. (2017) reported a nonlinear relationship between the preservation of shell microstructure and geochemical composition, based on analyses of cathodoluminescence and SEM images, as well as measurements of the trace element concentrations and carbon and oxygen isotope compositions of Middle Pleistocene (approximately 0.8 Ma) brachiopod shells. They showed that the original isotopic composition was almost retained in many brachiopod shells that were altered due to meteoric diagenesis. Because there remains no robust criterion to assess diagenetic alteration, it is necessary to apply and cross-check multiple criteria. We suggest criteria that include assessment of whether the shells possess expected carbon and oxygen isotope compositions. Consequently, we concluded that the studied shells with PI of 3 retained their original carbon and oxygen isotope compositions.

5.1.2 Diagenetic effects on oxygen and carbon isotopes of the fossil shells

Previous studies suggested that geochemical and isotopic compositions of biogenic carbonate fossils with PI values > 3 are suitable as paleoenvironmental proxies (Cochran et al., 2010; Ryan et al., 2021; Tajika et al., 2023). The δ¹⁸O_shell and δ¹³C_shell values of PU-1 and PU-5, which have PI values of 3, were distributed in a similar area of the cross-plots (Figure 11). This similarity indicated that these shells have a similar preservation state and have not been affected by significant diagenetic alteration. Meanwhile, the δ¹⁸O_shell and δ¹³C_shell values of PU-2, with a low PI value of 1, were lower than the other shells (by approximately 1‰ and 2‰, respectively) (Figure 11). These results suggest that PU-2 was more diagenetically altered and affected to a greater extent by isotopic exchange of C and O between its shell aragonite and pore water with low δ¹⁸O and δ¹³C values. Pore water could supply lower δ¹⁸O and δ¹³C than contemporaneous seawater due to dissolution of decomposed organic matter (Cochran et al., 2010). Therefore, the isotopic records of PU-2 are excluded from further paleoenvironmental discussion due to diagenetic alteration. There is evidence from other areas that carbonate fossils strongly affected by diagenetic alteration have outlier δ¹⁸O_shell and δ¹³C_shell values compared with well-preserved fossils from the same locality (Cochran et al., 2010; Jones et al., 2022).

Figure 11

Figure 11 Cross-plots of δ¹⁸O and δ¹³C values of the studied C. (I.) delicatostriata (PU-1, 2, 5) and A. pseudoplana shells. Small translucent plots represent the original isotope values and large opaque plots show the average isotope ratio of each shell. Bars indicate the ranges of both isotope ratios for each shell.

Although the isotopic composition of PU-2 was somewhat modified, its δ¹⁸O values showed cyclicity similar to other studied shells (PU-1 and PU-5) with a better state of preservation, which may reflect seasonal environmental changes (Figure 10). Although Cochran et al. (2010) showed that stable isotope composition was modified in samples, they did not analyze isotopic composition along the growth axis or present any isotopic profiles. Given pristine shell microstructures, it is possible that variations in stable isotope composition, including seasonal variations, are preserved in shell carbonate even at low PI. Original seasonal δ¹⁸O patterns can be preserved despite geochemical alteration. For example, heating can shift shell δ¹⁸O values by 1.5‰, equivalent to 6°C–7°C in paleotemperature reconstruction, while the primary seasonal pattern is preserved (Moon et al., 2021). However, the stratigraphic horizon from which the studied shells were collected does not show any indications of exposure to high temperatures due to burial diagenesis or hydrothermal activity. Therefore, it is unlikely that the oxygen isotope profiles of the studied shells were shifted toward a smaller value while maintaining their initial seasonality.

5.2 Life history of C. (I.) delicatostriata

5.2.1 Timing of growth line formation in the shells of C. (I.) delicatostriata

The growth of carbonate shells (including bivalves) and skeletons is controlled by both tidal (Ohno, 1989; Schöne et al., 2003; Schöne and Giere, 2005) and daily rhythms (de Winter et al., 2023). In bivalves inhabiting tidally influenced areas, shell precipitation occurs during high tide; their growth is paused due to the stresses caused by tidal emersion and ebb currents during low tide, leading to the formation of growth lines (Evans, 1972; Ohno, 1989; Schöne et al., 2003; Reza Mirzaei et al., 2014). As tidal activity occurs twice per day, two growth lines are formed daily in the shells of bivalves inhabiting intertidal areas (Richardson, 1988; Ohno, 1989; Schöne et al., 2003). Bivalves in constantly immersed areas, such as subtidal zones, experience weaker tidal stresses and produce one growth line per day (Ohno, 1989; Reza Mirzaei et al., 2014). During spring tides, which occur on a fortnightly cycle (14.7 days), the stress at low tide is greater, which leads to more prominent growth line formation in both intertidal and subtidal bivalves (Azzoug et al., 2012; Reza Mirzaei et al., 2014). Irradiance (i.e., light intensity) is regarded as a major factor affecting bivalve shell formation controlled by daily rhythms (Ohno, 1989; Clark, 2005; Sano et al., 2012).

The most important consideration for establishing bivalve shell chronology from shell growth lines in geologic time is whether the environmental rhythms that drive modern shell growth line formation (i.e., tidal and daily rhythms) in the study periods were similar to today (Wierzbicki et al., 2023). The daily rhythm interval has lengthened at a rate of 1 s/5000 years on average (i.e., decreasing number of days per year) throughout geological time due to the decline in Earth’s spin velocity (Darwin, 1880; Lantink et al., 2022). According to sclerochronological and geochemical analyses on rudist shells, day length was 23.5 h (372 days per year) in the Late Cretaceous (Campanian) (de Winter et al., 2020). The tidal rhythm in the middle Cretaceous was approximately similar to the present (14.8 days for semilunar month) because lunar-related rhythms, which are strongly related to tidal rhythm, were presumably more stable than solar rhythms throughout Earth’s history (Williams, 2000; Bhattacharya and Jha, 2014; Wierzbicki et al., 2023). Calculations based on the above data indicated that the numbers of solar days, lunar days, and fortnightly growth lines per year during the middle Cretaceous were 372, 352.5, and 25.1, respectively. The numbers of growth lines according to solar days and lunar days per single fortnight cycle in the middle Cretaceous were 15.6 and 14.8, respectively.

Two different types of growth lines were recognized in the studied C. (I.) delicatostriata shells; specifically, distinct “major” and relatively indistinct “minor” growth lines were present, with an alternating pattern of a single “major” growth line followed by 15–16 “minor” growth lines (see Section 4.2) (Figure 9). The repeated appearance of prominent growth lines alternating with 15–16 faint growth lines can be explained as the “minor” growth lines corresponding to daily growth lines and “major” growth lines corresponding to fortnightly growth lines in modern bivalves (Carré et al., 2005; Azzoug et al., 2012). This character is consistent with the expected Cretaceous daily and fortnightly rhythm relations outlined above. Both “major” and “minor” growth lines presumably formed in response to tidal cycles; therefore, the growth lines of C. (I.) delicatostriata can be used for reliable inner shell chronology and as records of temporal shell growth.

5.2.2 Shell growth of C. (I.) delicatostriata

The δ¹⁸O_shell values and growth lines of the studied shells can be used to delineate the life history of C. (I.) delicatostriata. The studied shells (PU-1, PU-2, and PU-5) exhibited distinct cyclicity of δ¹⁸O_shell values, with an amplitude of approximately 1.5‰ over approximately 25 “major” (i.e., fortnightly) growth lines (Figure 10). Because the tidal cycle during the Cretaceous period was approximately similar to the present (Section 5.2.1 and Bhattacharya and Jha, 2014), the presence of approximately 25 fortnightly growth lines per year suggested that the bivalves continued to form shell carbonates throughout the year (Berry and Barker, 1968; Azzoug et al., 2012). The lowest and highest values of δ¹⁸O_shell within these yearly growth line intervals represent summer and winter values, respectively; the main factor controlling the δ¹⁸O values is seawater temperature at the sites in which the shell grow. Therefore, the observed δ¹⁸O_shell curves along the 25 fortnightly growth lines reflect seasonal seawater temperature variations. Bivalves that form their shell throughout the year typically inhabit the tropics (Thébault et al., 2007; Yamanashi et al., 2016) and submarine caves (Kitamura et al., 2012). Therefore, C. (I.) delicatostriata may have also inhabited an environment suitable for shell formation throughout the year (i.e., seasonal variation in seawater temperature).

The growth increment width and δ¹⁸O_shell values of C. (I.) delicatostriata (top panels in Figure 10) represent temporal growth rate records. In younger portions of the shells, the fortnightly growth increment width shows an increase up to the maximum of 0.4 mm with a subsequent decrease (minimum = 0.05 mm) in 1 year, corresponding to 25 growth lines. The largest growth increment width was formed during the period spanning the δ¹⁸O_shell-based temperature minimum (winter) to its maximum (summer) (Figure 10), corresponding to the spring season. Conversely, the narrowest growth increments coincide with the coldest paleotemperatures and are regarded as representing winter (Figure 10). These observations suggest that the growth rate of C. (I.) delicatostriata reached its maximum in spring and minimum in winter. This growth pattern of bivalve shells is consistent with the patterns observed in extant bivalves inhabiting shallow-ocean areas and has been attributed to high spring phytoplankton abundance (Sato, 1997).

The “major” growth increment width and its variation significantly decreased toward the ventral shell margin (< 0.1 mm) (Figure 10), representing reductions in the shell growth rate and its seasonality during ontogeny. The presence of 25 “major” (i.e., fortnightly) growth increments within each annual growth cycle suggests that the shell was continuously precipitated throughout the year without any period of growth cessation. Although ontogenetic growth decline is common in modern bivalves (Jones and Quitmyer, 1996; Tanabe et al., 2017, 2020), ontogenetic growth characteristics in C. (I.) delicatostriata, which was continuous throughout the year associated with fortnightly (and daily) cycles, are similar to the characteristics of the extant tropical shallow-water bivalve Tegillarca granosa (Reza Mirzaei et al., 2017).

5.3 Shallow-water temperature and its seasonality in the mid-latitude northwestern Pacific region during the Turonian

Based on the δ¹⁸O_shell values of the two bivalve species and the relationships between temperature and aragonite δ¹⁸O established by Grossman and Ku (1986; Equation 1), the annual shallow-water temperature in the middle Turonian Yezo Basin ranged from 27.8°C to 34.9°C (Figure 12; Table 1). Analysis of sedimentary facies suggested that C. (I.) delicatostriata and A. pseudoplana inhabited shallow depths less than some tens of meters (Ando and Kodama, 1998; Komatsu, 2013) (Figure 2). Because water temperature at these depths would not substantially differ from the temperature at the sea surface (Tanabe et al., 2017), the temperatures calculated in this study should reflect shallow-water temperature and its seasonality in the middle Turonian northwestern Pacific Ocean.

Figure 12

Figure 12 Comparison of shallow-water temperature and its seasonality calculated among various equations of temperature and aragonite δ¹⁸O. Aragonite δ¹⁸O was derived from δ¹⁸O_shell values of studied bivalve shells. TSI and TSO in Kodama et al. (2021), KTd-1, KTd-4, and KTd-5 in Yamanashi et al. (2016), and NTM, NTGs, IHH, and NHH in Duprey et al. (2014) represent the names of shells. Seasonal temperature maxima and minima were based on the maximum and minimum δ¹⁸O_shell values of the studied shells. Mean annual temperature was calculated from monthly averaged δ¹⁸O_shell values (Supplementary Table 2). The gray zone on the temperature plot represents the mean annual sea surface temperature (SST) of the middle–late Turonian Yezo Basin based on planktonic foraminiferal δ¹⁸O (Moriya, 2011). The δ¹⁸O_seawater value used for temperature estimation was −1.00‰.

Seawater temperatures estimated from the shell δ¹⁸O values and their seasonality can vary depending on assumed δ¹⁸O_seawater values and their seasonal fluctuation at the site of shell growth. For example, for an actual δ¹⁸O_seawater value 1‰ lower than the assumed value, the seawater temperature calculated based on the Equation 1 would increase by approximately 4°C. The δ¹⁸O_seawater value varies depending on salinity because it increases with massive evaporation, whereas it decreases with strong precipitation and freshwater input due to ¹⁸O depletion in the vapor and meteoric water (Craig, 1961). Based on the modern relationship of δ¹⁸O_seawater and salinity, a salinity change of 1 causes a change of approximately 0.3‰ in δ¹⁸O_seawater (Fairbanks et al., 1997; Abe et al., 2009; Takayanagi et al., 2015). Recent clumped isotope analysis of modern bivalve shells suggested that seasonal changes in δ¹⁸O_seawater values make a significant contribution to those changes in δ¹⁸O_shell values, thus affecting δ¹⁸O_shell-based seawater temperature estimates (Caldarescu et al., 2021). Furthermore, clumped isotope analyses of Cretaceous rudists and oysters revealed that δ¹⁸O_seawater values at that time were not necessarily −1‰, which has been traditionally used for paleotemperature estimation from δ¹⁸O values of biogenic carbonates during the Cretaceous in shallow oceans (de Winter et al., 2021; Jones et al., 2022). de Winter et al. (2021) also reported that large seasonal δ¹⁸O_seawater fluctuations (approximately 1‰), which showed the highest values near the hottest season and lowest values near the coldest season, were present in the Campanian shallow ocean, resulting in underestimation of seawater temperature seasonality. However, it is difficult to assume that δ¹⁸O_seawater changes over seasonal cycles.

The seasonal salinity ranges in the modern subtidal zone, especially at depths of more than 10 m, are small (salinity< 2) (Wanamaker et al., 2011; Kubota et al., 2017; Tanabe et al., 2017). This change corresponds to the δ¹⁸O_seawater change of approximately 0.6‰, when converting salinity into δ¹⁸O_seawater; it may cause an error of< 3°C in reconstructing seawater temperatures (Grossman and Ku, 1986; Chauvaud et al., 2005). Therefore, some studies have assumed that seasonal δ¹⁸O_seawater fluctuations were annually constant when reconstructing water temperature from δ¹⁸O_shell values of bivalves inhabiting water with almost constant salinity (Wanamaker et al., 2011; Kubota et al., 2017; Tanabe et al., 2017). The studied species, C. (I.) delicatostriata and A. pseudoplana, presumably inhabited normal marine environments with limited freshwater influence because they are found in lower shoreface to inner shelf deposits, which likely formed at depths of 10 m to some tens of meters, and not in freshwater and estuary deposits such as oyster beds (Ando and Kodama, 1998; Komatsu, 2013) (Figure 2). The absence of neighboring prominent and faint growth lines in C. (I.) delicatostriata (Section 5.2.1 and Figure 9) also indicated that they did not live in the intertidal zone, which is commonly affected by freshwater input (Ohno, 1989; Tojo and Masuda, 1999; Miyaji et al., 2007; Reza Mirzaei et al., 2014). Massive freshwater inflow events, such as typhoons or heavy rains, may have occurred in the depositional area of the middle Turonian Mikasa Formation because of the humid climate conditions in the northwestern Pacific region, suggested by terrestrial organic δ¹³C and numerical climate simulation, and the presence of HCS deposits (likely formed by storm events) in this formation (Ando, 1987; Hasegawa, 2003; Hasegawa et al., 2012; Higuchi et al., 2021). However, such short-term precipitation events are unlikely to be recorded in the δ¹⁸O profiles of the studied shells due to their temporal resolution (more than one week), which is not sufficiently high for detection. Therefore, our seasonal δ¹⁸O_shell variations likely reflect shallow-water temperature seasonality at the time the shells inhabited.

The reconstructed paleotemperature in our study was consistent with shallow-water paleotemperature records (28°C–34°C) from the late Cenomanian mid-latitude (36°–45°N paleolatitude) North American Interior Seaway, derived from clumped isotope analysis on oyster fossil species (Pycnodonte newberryi shells from the Mancos Shale Formation at 36°N paleolatitude, P. newberryi shells from the Tropic Shale Formation at 37°N paleolatitude, P. kellumi shells from the Frontier Formation at 43°–45°N paleolatitude, and P. kellumi shells from the Greenhorn Formation at 44°N paleolatitude) using the temperature calculation formula of Petersen et al. (2019) (Jones et al., 2022). In contrast, planktonic foraminiferal δ¹⁸O from the western extension of the Mikasa Formation indicated that the mean annual sea surface water temperature of the middle to late Turonian Yezo Basin ranged from 26°C to 29°C (Moriya, 2011). These temperatures are lower than temperatures determined from the studied shells (Figure 12). Such differences may have been due to variations in seawater temperatures during the Turonian and differences in growth site environments (e.g., depths) between benthic bivalves and planktonic foraminifera. Furthermore, the planktonic foraminiferal temperatures reported by Moriya (2011) do not necessarily reflect sea surface temperatures because planktonic foraminifera inhabit areas close to the sea surface and areas within the mixed layer. However, considering that Moriya (2011) did not provide detailed geological or paleontological data on the studied samples or information about the oxygen isotope thermometer used, it is not possible to identify the factors responsible for these discrepancies.

Differences in reconstructed temperatures between bivalves and foraminifera could have been partly due to differences in the equations representing the relationships between (δ¹⁸O_shell − δ¹⁸O_seawater) and seawater temperature (i.e., oxygen isotope thermometers). The equations proposed in previous studies differ among taxa and environmental settings (Nishida, 2020; Kodama et al., 2021). We calculated and compared shallow-water temperatures and their seasonality in the Turonian Yezo basin using 24 equations established for aragonite δ¹⁸O in various modern marine calcifiers (Figure 12). These equations were derived from the arcid bivalve Scapharca broughtonii (Horibe and Oba, 1972; Nishida et al., 2015), venerid bivalve Mesodesma donacium (Carré et al., 2005), giant clam Tridacna squamosa (Arias-Ruiz et al., 2017; Kodama et al., 2021; Wang et al., 2022), T. derasa (Yamanashi et al., 2016), T. gigas (Elliot et al., 2009; Welsh et al., 2011; Yan et al., 2013; Ayling et al., 2015; Ma et al., 2020), T. maxima (Duprey et al., 2014; Komagoe et al., 2018), Hippopus hippopus (Watanabe and Oba, 1999; Aubert et al., 2009), a combination of giant clam species (Aharon, 1983; Duprey et al., 2014), the freshwater gastropod Peregriana peregra (White et al., 1999), various benthic foraminifera (Dunbar and Wefer, 1984), and various biogenic aragonites (Böhm et al., 2000).

Although the reconstructed seawater temperatures differed according to the equations used, they fell within a range from 22°C to 39°C, with seasonality of <8°C (Figure 12). In some cases (e.g., Dunbar and Wefer, 1984), the reconstructed temperatures were consistent with temperatures estimated from planktonic foraminifera for the same time period (Moriya, 2011), but more commonly, the reconstructed temperatures were higher than temperatures from planktonic foraminifera (e.g., Grossman and Ku, 1986 and Yan et al., 2013). The seawater temperatures reconstructed using equations derived from bivalves taxonomically close to C. (I.) delicatostriata (Scapharca broughtonii: Horibe and Oba, 1972) and A. pseudoplana (Mesodesma donacium: Carré et al., 2005) were slightly lower but roughly similar to temperatures derived from the studied bivalves, suggesting that the two fossil species precipitated their shells under physiological conditions similar to the conditions experienced by closely related extant species. The reconstructed seawater temperature of the middle Turonian Yezo Basin was significantly higher than the temperatures of the modern mid-latitude ocean (Figure 13). For example, the modern eastern Pacific Ishikari area of Hokkaido, Japan (44°N), has a surface seawater temperature of 4.8°C–22.1°C (monthly averaged data from Japan Meteorological Agency, 2024a). These findings suggest that extremely warm shallow water was present in the mid-latitude region during the middle Cretaceous “supergreenhouse” Earth.

Figure 13

Figure 13 Comparison of monthly averaged shallow-ocean temperature and its seasonality in the western to northwestern Pacific regions among the middle Turonian, middle Cretaceous derived from this study (PU-1, PU-5, Averaged), and recent ocean derived from Japan Meteorological Agency (2024a) (Hokkaido) and Japan Meteorological Agency (2024b) (Okinawa). PU-1 and PU-5 refer to Cucullaea (Idonearca) delicatostriata shells examined in this study and averaged indicates averaged temperatures of PU-1 and PU-5. Mean annual temperature was calculated from monthly averaged temperature values (Supplementary Table 2).

Monthly averaged sub-annual scale δ¹⁸O_shell data along the growth direction of two C. (I.) delicatostriata specimens (PU-1 and PU-5) (Figure 10) indicated that seasonal shallow-water temperature variation was approximately 6°C in the middle Turonian mid-latitude northwestern Pacific Ocean (Figure 13; Supplementary Table 2). This value is substantially smaller than temperature ranges in modern mid-latitude regions, such as Hokkaido, where the monthly averaged sea surface temperature changes by approximately 17°C throughout the year (Japan Meteorological Agency, 2024a) (Figure 13). Low seasonal variability of SST is typically observed in modern low-latitude tropical ocean areas, such as 22.7°C–29.3°C (6.6°C seasonality) derived from eastern Pacific data collected around southern Okinawa Island, the Ryukyu Islands, Japan (26°N; Japan Meteorological Agency, 2024b) (Figure 13). This finding is consistent with previous studies arguing that paleo-Hokkaido was located in a subtropical climate during the middle Cretaceous (e.g., Hay and Floegel, 2012).

5.4 Shallow-water temperatures during the Turonian greenhouse Earth

Shallow-water temperatures and their seasonal variations during the Turonian were estimated based on δ¹⁸O_shell values (Figure 14) of rudist shells from the low-latitude Tethys Ocean (2–10 m water depth,< 30°N paleolatitude) in some previous studies. Distefanella sp. from the middle Turonian sediments in Zakynthos Island, Greece (91.1 Ma, 16.5°N paleolatitude) indicated one year seawater temperatures of 28.5°C–34.4°C with a seasonal variation of 5.9°C (Steuber et al., 2005). Hippurites resectus from the Gosau Group (the upper Turonian to the lower Coniacian, 89–91 Ma, 27.4°–27.5°N paleolatitude) and Vaccinites inequicostatus from the same group and Theresienstein, Austria, indicate seawater temperatures of 24.1°C–29.7°C (five years) and 25.9°C–35.0°C (1.5–2 years), respectively (Steuber et al., 2005; Walliser and Schöne, 2020). These temperatures were estimated using the equation of Anderson and Arthur (1983), assuming a constant δ¹⁸O_seawater value of −1‰ throughout the year. The shallow-water temperature estimated in this study roughly coincided with the previous estimations; the common range was 28°C–35°C (Figure 14). One exception to this trend was the minimum temperature recorded in the Gosau Group (24.1°C), which was lower than the temperatures in the Mikasa Formation, although the Mikasa Formation was deposited in a higher latitude setting than the Gosau Group.

Figure 14

Figure 14 Shallow-water temperature records including seasonality during the Turonian compiled from this and previous studies ([a] Steuber et al. (2005); [b] Walliser and Schöne (2020); [c] this study), along with a paleogeographic map of the late Turonian (90 Ma) Northern Hemisphere modified from Nakajima et al. (2017) and Herman et al. (2019) with shallow-water current systems from Moiroud et al. (2013) and Mansour and Wagreich (2022). Shallow-water temperature data for individuals from the same region are grouped together as local shallow-water temperatures, with the Gosau Group and Theresienstein combined due to their close localities (Walliser and Schöne, 2020). The paleogeographic position of each locality is based on previous studies (Steuber et al., 2005; Walliser and Schöne, 2020) and paleomagnetic records for this study (Tamaki and Itoh, 2008). The δ¹⁸O_seawater value for temperature estimation from shell carbonate δ¹⁸O was −1.00‰ for all data according to the original article (Steuber et al., 2005; Walliser and Schöne, 2020).

There are multiple possible explanations for the temperature differences outlined above, such as differences in δ¹⁸O_seawater values between Tethys and Pacific Oceans. According to a numerical climate model of the global δ¹⁸O_seawater value distributions, δ¹⁸O_seawater values in the middle Cretaceous Tethys Ocean were approximately +0.5‰. This value is approximately 1.5‰ higher than the northwestern Pacific region at the same time (Zhou et al., 2008), as noted by Walliser and Schöne (2020). Consistent with the results of Zhou et al. (2008), the estimated shallow-water temperatures of the Gosau Group and Theresienstein depositional area were 31°C–43°C, higher than results for the Pacific derived in the present study. Warmer temperatures in the Tethys Ocean than the Pacific Ocean are consistent with planktonic foraminiferal biogeography of the middle Turonian Yezo Group with loss of frequent Tethyan type (tropical–subtropical) foraminifera and transitions to Transition type (warm–temperate) foraminifera bioprovinces (Nishi et al., 2003).

The effects of shallow-ocean currents can also explain the lower minimum seawater temperatures observed in the low-latitude Tethys Ocean (Walliser and Schöne, 2020) compared with the mid-latitude northwestern Pacific (this study). Cold currents flowed from the Arctic to the western Tethys Ocean, including the depositional area of the Gosau Group (Walliser and Schöne, 2020). In contrast, during the Turonian, the Pacific Ocean was separated from the Arctic Ocean by the Bering Land Bridge (Herman and Spicer, 1996, 2010; Herman et al., 2019). In this ocean–land configuration, the cold Arctic water could not flow into the Pacific Ocean, leading to warm shallow-water temperatures in the mid-latitude Pacific Ocean. The seasonal variations of shallow-water temperatures were commonly smaller than 7°C in the low-latitude Tethys (including the Gosau Group and Theresienstein area) and the mid-latitude Pacific (Figure 12).

Ivany (2012) noted that the seasonality of SST in the modern Northern Hemisphere is greatest at 40°–45°N. Our study site (44°N) exhibited good alignment with this maximal SST (or shallow-water temperature) seasonality zone of the Turonian Northern Hemisphere, indicating that maximum SST (or shallow-water temperature) seasonality in the mid-Cretaceous was similar to seasonality in the modern tropics. Additionally, the meridional temperature gradients and seasonal SST (or shallow-water temperature) ranges in the Northern Hemisphere were flatter and smaller in the mid-Cretaceous than at present. Low-temperature seasonality in the Northern Hemisphere during the Turonian is consistent with paleontological evidence from high-latitude plant and vertebrate fossils. Herman et al. (2019) reported conifer- and angiosperm-dominant floras from the Derevyannye Gory Formation (Turonian–Coniacian) of the New Siberia Island, Russia (70°N paleolatitude) and they concluded that low-temperature seasonality was present in the Turonian–Coniacian Arctic. Vandermark et al. (2007) reported that the extinct diapsid reptile Champsosaurus fossils from the Turonian–Coniacian sediments in Axel Heiberg Island, northern Canada (79°N, paleolatitude), was dominated by subadult-stage individuals, suggesting that temperatures suitable for reproduction and growth of reptiles were present during the Turonian–Coniacian at high latitudes due to a warm mean annual temperature with small temperature seasonality. The absence of ice sheets was presumably responsible for the reduced seasonal temperature variations on a global scale in the mid-Cretaceous (Steuber et al., 2005; Ivany, 2012). Under ice-free Earth conditions, the positive feedback of ice albedo becomes absent or is greatly weakened compared with today. In this case, the ice-free climate maintains warm shallow-water temperatures even in winter, resulting in slight temperature seasonality. Conversely, during the Eocene when the Antarctic ice sheet was ice-free, as in the Cretaceous, the annual SST range was greater than during the icehouse period in Antarctica (Buick and Ivany, 2004). There is a need for further data acquisition and numerical modeling of the latitudinal gradients and annual ranges of SST and air temperature during the middle Cretaceous and other greenhouse periods.

6 Conclusions

Growth increment and stable isotope analyses were performed on newly obtained bivalve fossils with preserved biogenic aragonite and shell microstructures from the middle Turonian Mikasa Formation, Yezo Group, in Hokkaido, Japan, to reveal the life history characteristics of the bivalves and sub-annual scale paleotemperature variations of mid-latitude shallow-water areas in the northwestern Pacific region during the Turonian. In Cucullaea (Idonearca) delicatostriata, approximately 25 fortnightly growth lines were observed per year, indicated by δ¹⁸O_shell, demonstrating that the annual cycles obtained from growth line analysis are consistent with those obtained from δ¹⁸O_shell ratios. Growth increment width and δ¹⁸O_shell values of C. (I.) delicatostriata revealed that they continuously produced their shells throughout the year, with the seasonal growth rate maximized in spring and minimized in winter. Reconstructed temperatures from bivalve δ¹⁸O_shell values ranged from 28°C to 35°C and showed significantly smaller seasonality compared with temperatures in modern Hokkaido. The δ¹⁸O_seawater value was assumed to be −1‰ without seasonal variations due to the limited influence of freshwater in the studied section. A comparison of temperature seasonality records with the same period from the Tethys Ocean showed similar values, indicating that tropical ocean conditions with low seasonality in mid-latitude regions were not limited to the Tethys Ocean but also occurred in the northwestern Pacific Ocean. Low shallow-water temperature seasonality in the mid-latitudes suggests that seasonal temperature fluctuations in the Northern Hemisphere were small in the middle Turonian “supergreenhouse” period.

Data availability statement

The original contributions presented in the study are included in the article/Supplementary Material. Further inquiries can be directed to the corresponding author.

Author contributions

SI: Writing – original draft, Conceptualization, Investigation, Methodology, Visualization, Data curation. HT: Investigation, Writing – review & editing, Data curation, Funding acquisition. YI: Writing – review & editing, Data curation, Funding acquisition. ST: Supervision, Writing – review & editing, Conceptualization, Funding acquisition. TO: Supervision, Writing – review & editing, Conceptualization, Funding acquisition, Investigation.

Funding

The author(s) declare financial support was received for the research, authorship, and/or publication of this article. This research was supported by Grant-in-Aid for Scientific Research (B: No. 19H02014 to TO, 23H01287to ST, 21K18642 and 23H01263 to YI).

Acknowledgments

We express our appreciation to the editor and three reviewers for their time and valuable comments. We would like to thank Y. Koketsu (Nagoya University) for assistant of Raman spectroscopic analysis. We also thank T. Miyaji (National Ainu Museum) for useful advice about growth line analysis. We gratefully acknowledge the Sorachi Forest Management Office for permitting us to conduct field survey. We thank S. Hayashi (Nagoya University) and Y. Ando (Mizunami Fossil Museum) for helping SEM observation. We also thank H. Oda (Tohoku University) for helping stable isotope analysis. We like to thank D. Nakai (Nagoya University) for advice about improving growth line observation. This work was partly supported by World Premier International Research Center Initiative (WPI), MEXT, Japan.

Conflict of interest

The authors declare that the research was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

Publisher’s note

All claims expressed in this article are solely those of the authors and do not necessarily represent those of their affiliated organizations, or those of the publisher, the editors and the reviewers. Any product that may be evaluated in this article, or claim that may be made by its manufacturer, is not guaranteed or endorsed by the publisher.

Supplementary material

The Supplementary Material for this article can be found online at: https://www.frontiersin.org/articles/10.3389/fmars.2024.1324436/full#supplementary-material

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