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Past aquatic environments in the Levant inferred from stable isotope compositions of carbonate and phosphate in fish teeth


Here we explore the carbon and oxygen isotope compositions of the co-existing carbonate and phosphate fractions of fish tooth enameloid as a tool to reconstruct past aquatic fish environments and harvesting grounds. The enameloid oxygen isotope compositions of the phosphate fraction (δ18OPO4) vary by as much as ~4‰ for migratory marine fish such as gilthead seabream (Sparus aurata), predominantly reflecting the different saline habitats it occupies during its life cycle. The offset in enameloid Δ18OCO3-PO4 values of modern marine Sparidae and freshwater Cyprinidae from the Southeast Mediterranean region vary between 8.1 and 11.0‰, similar to values reported for modern sharks. The mean δ13C of modern adult S. aurata and Cyprinus carpio teeth of 0.1±0.4‰ and -6.1±0.7‰, respectively, mainly reflect the difference in δ13C of dissolved inorganic carbon (DIC) of the ambient water and dietary carbon sources. The enameloid Δ18OCO3-PO4 and δ13C values of ancient S. aurata (Holocene) and fossil Luciobarbus sp. (Cyprinidae; mid Pleistocene) teeth agree well with those of modern specimens, implying little diagenetic alteration of these tooth samples. Paired δ18OPO413C data from ancient S. aurata teeth indicate that hypersaline water bodies formed in the Levant region during the Late Holocene from typical Mediterranean coastal water with high evaporation rates and limited carbon input from terrestrial sources. Sparid tooth stable isotopes further suggest that coastal lagoons in the Eastern Mediterranean had already formed by the Early Holocene and were influenced by terrestrial carbon sources. Overall, combined enameloid oxygen and carbon isotope analysis of fish teeth is a powerful tool to infer the hydrologic evolution of aquatic environments and assess past fishing grounds of human populations in antiquity.


Fish live in nearly all aquatic environments worldwide, from mountain streams to the abyssal depths of the world’s oceans. Mineralised in isotopic equilibrium with body fluid, the phosphate oxygen isotope composition (δ18OPO4) of fish bioapatite is a function of both the temperature and the oxygen isotope composition (δ18OWater) of ambient water [14]. Bony as well as cartilaginous fish mineralize enameloid as crown-forming dental tissue. Enameloid consists of bioapatite that often has a high fluorine content, similar to that of fluorapatite[57], especially in sharks [8]. However, enameloid F-content in bony fish (teleosts) can vary between species and within species, depending to developmental stage [5, 9]. The teleost fish analysed in this study belong to species with high F-content (Sparus ~3%) as well as low F-content (Carpio ~0.3%) [9]. Fish tooth enameloid is highly resistant to diagenetic alterations following post-mortem burial [10]. Fish tooth δ18OPO4 values therefore serve as excellent proxies for ancient aquatic habitats (i.e. rivers, lakes, lagoons and oceans; e.g., [1120]) and enable us to infer the salinity of past fishing grounds [2122].

During biomineralisation of teeth some carbonate ions replace phosphate (Type-B) and OH-F sites (Type-A) in the crystal lattice of bioapatite (hydroxylapatite): i.e., Ca5(PO4, CO3)3(OH, F, CO3) [23]. The oxygen and carbon isotope composition (δ18OCO3 and δ13C values, respectively) of this structural carbonate can be used as a proxy to infer fish (palaeo-) environments (e.g.,[4, 1920, 2425]), by indicating different sources of carbon (marine vs. fresh water/terrestrial). However, compared to the phosphate fraction (δ18OPO4), the δ18O of the carbonate fraction of tooth enamel(oid) is more prone to diagenetic alteration [12, 2527]. Moreover, δ18OCO3 values exhibit high, unsystematic variability up to 5–6‰ in modern shark teeth, even within tooth samples from the same individual, suggesting that δ18OCO3 in fish teeth might not be an ideal palaeo-environmental proxy [24]. However, δ18OCO3—which systematically co-varies with δ18OPO4 in mammal teeth [28]—is frequently used in conjunction with δ18OPO4 as a tool to assess the degree of diagenetic alteration in fossil tooth samples (e.g.,[2930]).

Large, but systematic variations in δ13C values have also been observed between the dentine and the enameloid of modern shark teeth [24]. The positive δ13C values of enameloid (1.6 to 4.8‰) compared to significantly lower values in dentine (-6.4 to -2.3‰) suggest that enameloid approaches isotopic equilibrium with ambient dissolved inorganic carbon (DIC), likely through diffusional exchange with seawater DIC [24]. The δ13C values of structural carbonate of enameloid in fossil shark teeth agree well with those documented in modern species, suggesting that in vivo carbon isotope compositions are preserved [1920, 25]. Thus, in cases where the original stable isotope compositions of fish teeth are preserved, the δ13C values of the enameloid may serve as a proxy for changes in the isotopic composition of DIC of the ambient water [1920, 2425]. Yet, the range of enameloid δ13C values in modern bony fish, which is essential to assess fossil fish δ13C values, is presently limited.

In this study, we use carbon and oxygen isotopic compositions of co-existing carbonate and phosphate fractions of fish tooth enameloid to infer the aquatic environments in which (pre-)historic humans caught their fish. For this purpose, different fish species from the modern Mediterranean (Sparidae) and freshwater fish from Lake Kinneret (Cyprinidae) were analyzed for their δ18OCO3, δ18OPO4 and δ13CCO3 values. These results are then compared to values obtained from fish teeth (belonging to the same species) recovered from different archaeological sites and periods in the southern Levant. To examine the possible effect of diagenesis on enameloid carbon and oxygen isotopic compositions, we also analysed some fossil pharyngeal teeth (1st molariform tooth) recovered in deposits of lake ‘Ubediya (Lower palaeolithic, ca. 1.4 Ma; Jordan Rift Valley). Combined δ18OPO413C and δ18OCO313C data of fish tooth enameloid is explored as a tool to reconstruct the hydrological history of past fish habitats.

Materials and methods


Teeth of five extant fish species from three different aquatic habitats in the Levant were analysed for their enameloid carbon and oxygen isotope composition: Sparus aurata (n = 3, 9 teeth) and Pagrus caeruleostictus (n = 2, 6 teeth) (Sparidae, Perciformes, seabream) from the Mediterranean Sea (Haifa Bay, Israel), Scarus sp. (Scaridae; parrotfish, n = 1, 3 teeth) from the Red Sea, and the domesticated Cyprinus carpio (Cyprinidae; common carp, n = 2, 6 teeth) from the freshwater of Lake Kinneret in northern Israel (Table 1). Each of these species has a distinct diet: S. aurata feeds mainly on shellfish, mussels and oysters; P. caeruleostictus feeds on bivalves, crustaceans and fish; Scarus sp. feeds on corals and sponges; C. carpio feeds on plants, insects, crustaceans and benthic worms. Enameloid δ18OCO3, δ18OPO4 and δ13CCO3 values of these fish are used as reference data for comparison with those of ancient fish teeth of the same species.

Table 1. Stable isotope composition of enameloid of modern fish teeth (The Sparidae δ18OPO4 values were published in Sisma-Ventura et al. [17]).

The archaeological samples used in this study are the first molariform teeth of Sparus aurata (n = 45) of the family Sparidae (Perciformes, seabream). These fish teeth were excavated from 12 well-dated archaeological sites in Israel [22], covering key periods of the Holocene: (1) the Pre-Pottery Neolithic sub-periods (A, B and C) dated to the Early Holocene (11,700–8,500 yr BP); (2) the Chalcolithic, (7,500–5,700 yr BP); (3) the Late Bronze Age, (3,500–3,200 yr BP); (4) the Iron Age (3,200 and 2,700 yr BP); and the Byzantine Age (1,680–1,360 yr BP) (Table 2). Additionally, five fossil pharyngeal teeth (1st molariform) belonging to Luciobarbus sp. (Cyprinidae), recovered from 'Ubeidiya in the Jordan Rift Valley and dated to ca. 1.4-million-years ago [31] were analysed (Table 3).

Table 2. Stable isotope composition of enameloid from ancient Sparus aurata (Sparidae) first molariform teeth (The Sparidae δ18OPO4 values were published in Sisma-Ventura et al. [17]).

Table 3. Stable isotope composition of enameloid from fossil Cyprinidae teeth.

The fossil fish were identified to species level using reference collections housed at the University of Haifa (Israel), the National Natural History Collections at the Hebrew University and The Zooarchaeological collections at the Autonomous University of Madrid (Spain).

Oxygen and carbon isotope analyses of the carbonate fraction in bioapatite

The enameloid part (~0.2–0.4 mm layer) of each individual tooth crown was separated from the dentine using a diamond-head micro-dental drill, washed three times with distilled water and dried overnight at 50°C. Each sample was crushed and ground to powder using an agate mortar and pestle. Organic matter was removed from the sample powder by soaking in 30% H2O2 over night as described in Gehler et al. [32]. This pretreatment was tested on shark’s teeth (dentine + enameloid) and seemed to have minor effect on the δ13C values, which were comparable to those obtained using different pretreatments, such as NaOCl + acetic acid, as well as untreated samples. [24] The pretreatment effect on the δ18OCO3 values, observed in shark’s teeth [24] was minimised [26] by removing the dentine entirely from the enameloid. Due to low average CO32− content (2–3 wt %), approximately 2 mg of enameloid powder was reacted with 100% H3PO4 at 70°C in a Thermo Scientific KIEL IV automated carbonate device. The δ18OCO3 and δ13C values of the samples were then determined on the CO2 using a Finnigan Delta Plus gas source mass spectrometer in dual inlet mode at the stable isotope laboratory of the Geoscience Center at the University of Göttingen, Germany.

Measured isotope ratios were normalised to an in-house carbonate standard (Solnhofen) which has been calibrated against NBS-19. The analytical precision of this method is better than ±0.1‰ for both δ18OCO3 and δ13C. Measured isotope ratios are reported in δ-notation, i.e., as the deviation in per mil (‰) from the international measurement standards Vienna Standard Mean Ocean Water (VSMOW; δ18OCO3) and Vienna Pee Dee Belemnite (VPDB; δ13C): where R represents the 18O/16O or 13C/12C ratio.

The analytical precision of the in-house carbonate standard was ±0.08‰ (1σ) in δ18OCO3 and ±0.05‰ (1σ) in δ13C (n = 16). For untreated samples of the NBS 120c Florida phosphate rock standard, we obtained a δ18OCO3 value of 30.1±0.12‰ and a δ13C value of -6.3±0.05‰ (n = 16). The reproducibility of the in-house bioapatite standard AG LOX (modern African elephant enamel; [27]) for untreated samples was 29.9±0.14‰ for δ18OCO3 and -11.75±0.04‰ for δ13C (n = 16). These mean values compare well with the values of 30.30±0.17‰ for δ18OCO3 and -11.78±0.12‰ (n = 23) for δ13C reported for untreated AG LOX by Wacker et al. [33].

Oxygen isotope analyses of the phosphate fraction in bioapatite

The phosphate fraction of the samples was separated using a method modified after Dettmann et al. [34] and described in detail by Tütken et al. [35] and Sisma-Ventura et al. [22]. Five milligrams of pretreated enameloid powder were placed on vibrating table and digested in 0.8 ml HF (2 M) for 12 h. The samples were centrifuged, and the remaining supernatant solution with the dissolved phosphate was separated from the CaF2 precipitates. The HF solution transferred into a new vial and was neutralised with 25% NH4OH using bromothymol blue as pH indicator, and Ag3PO4 precipitation was completed by adding 0.8 ml of 2 M AgNO3. The settled Ag3PO4 crystals were centrifuged, and the supernatant solution containing excess AgNO3 was removed. The Ag3PO4 precipitate was then rinsed five times with Milli-Q water and dried overnight in an oven at 50°C.

Ag3PO4 aliquots of 0.5 mg were placed into silver capsules and analysed in triplicate by means of high temperature reduction using a Finnigan TC-EA coupled via a Conflo III to a Micromass 100 GC-IRMS at the University of Mainz, or by a Finnigan Delta Plus XL GC-IRMS at the Universities of Tübingen and Lausanne, according to the method of Vennemann et al. [36]. The raw δ18OPO4 values were normalised to an Ag3PO4 standard with a certified value of 21.7‰ (silver phosphate P/N IVA33802207, batch no. 180097, distributed by IVA Analysentechnik, Germany). The analytical precision for this standard was better than ±0.3‰ (1σ). For untreated NIST SRM 120c Florida phosphate rock standard reference material, we obtained a δ18OPO4 value of 21.9±0.3‰ (n = 15). This value compares well with the value of around 21.7‰ initially proposed by Lécuyer et al. [12] and measured in most other laboratories as compiled in Chenery et al. [37].


Oxygen and carbon isotope compositions of carbonate and phosphate in modern and ancient fish tooth enameloid

Individual δ18OCO3, δ18OPO4 and δ13C data from the analyses of bulk enameloid powder from five modern Sparidae and two modern Cyprinidae are summarised in Table 1. The δ18OCO3 values of six teeth from two adult S. aurata (> 0.65–1.0 kg) are very similar and range between 32.3 and 34.2‰, while those of two adult P. caeruleostictus (> 0.9–2.5 kg, n = 6) range between 31.4 and 33.5‰. A similar δ18OCO3 range between 31.1 and 32.5‰, was obtained for three teeth of a single juvenile sample of S. aurata (<0.5 kg). The ancient Sparidae teeth yield variable δ18OCO3 values (n = 45), spanning a broad range between 29.5 and 36.6‰ (Table 2).

The modern samples of Sparidae from the southeast Mediterranean littoral yield a relatively narrow δ18OPO4 range of about 1.0‰ between 22.5 and 23.4‰ (n = 15). In contrast, δ18OPO4 values between 21.1 and 26.4‰ were measured for the ancient S. aurata teeth (n = 45).

A narrow range between -0.5 and 0.6‰ is also observed for the δ13C values of modern adult S. aurata specimens (Table 1). A lower δ13C value, averaging -2.3±0.6‰ was found for teeth of a single juvenile sample of S. aurata (<0.5 kg). For adult P. caeruleostictus the δ13C values range between -0.7 and 0.6‰. No significant differences are present between the average δ13C values of the adult S. aurata and P. caeruleostictus samples. The ancient Sparidae teeth (Table 2) yield variable δ13C values, covering a large range from -5.8 to 6.1‰.

Modern carp teeth enameloid yield relatively narrow ranges between 21.0 and 21.7‰ for δ18OPO4, between 29.4 and 30.8‰ for δ18OCO3, and between -5.2 and -6.6‰ for δ13C (n = 4; Table 1). Fossil carp enameloid values range between 16.8 and 20.7‰ for δ18OPO4, 25.5 and 29.2‰ for δ18OCO3, and -6.4 and -3.1‰ for δ13C (n = 5; Table 3).

The theoretical δ18OPO4 range of Sparidae bioapatite

The calculation of the equilibrium δ18OPO4 range of bioapatite forming in the southeast Mediterranean littoral and in hypersaline lagoons, evolving from typical seawater of this region, is based on the temperature-dependent relation for the oxygen isotope fractionation during biomineralisation of apatite by Lécuyer et al. [4]: where δ18OPO4 and δ18OSeaWater correspond to the isotope compositions of bioapatite and seawater relative to VSMOW, respectively. This relationship is valid for the temperature range of 8°C < T < 32°C. This equation from Lécuyer et al. [4] was established using bioapatite from both modern lingulids and sharks, yielding consistent temperatures from bioapatites equivalent to those from co-existing carbonates, therefore, providing the most reliable estimates of aquatic palaeotemperatures for bioapatites. Hence, we use the Lécuyer et al. [4] equation for calculating the equilibrium δ18OPO4 range of bioapatite. Likewise, the equation by Pucèat et al. [3] for the temperature-dependent phosphate oxygen isotope fractionation is widely used to calculate the equilibrium δ18OPO4 range of bioapatite. Surface water temperatures of the East Mediterranean range from 15°C in late winter (February-March) to 30°C in summer (July-August). The range of δ18OSeaWater recorded from the East Mediterranean is relatively small [3839], varying between 1.4‰ (February-March) and 1.8‰ (July-August). The calculated range of expected bioapatite δ18OPO4 forming in the southeast Mediterranean is slightly less than 3.0‰ and varies between 21.1 and 23.7‰. The δ18OPO4 values of modern Sparidae which were caught in the southeast Mediterranean littoral vary between 21.5 and 23.4‰, therefore agreeing well with the calculated range [2122]. Significantly higher δ18OPO4 values, between 23.6 and 25.4‰ [2], were measured in teeth of fish from the hypersaline Bardawil lagoon of Northern Sinai, Egypt. The Bardawil lagoon (Fig 1) is a large (30 km long, 14 km max. width), shallow (0.3–3 m deep) hypersaline coastal lagoon, separated from the Mediterranean Sea by a narrow sandbar. The Bardawil is connected to the sea via two small natural inlets (Boughaz Zaranik). Water exchange in the lagoon is controlled by Mediterranean Sea tides with a mean height of 50 cm. As a result, it has an elevated salinity level and δ18OWater values around 3.7‰ (range: 1.8‰ near the Mediterranean inlet, reflecting inflowing seawater, up to 7.2‰ due to high evaporation rates [2, 22]). Here we use the theoretical range of δ18OPO4 values expected for both ecosystems (southeast Mediterranean and the Bardawil lagoon) as endmembers to infer the palaeo-environments in which the fish were caught.

Fig 1. Study site map with archaeological sites from which the ancient Sparus aurata teeth were excavated with according archaeological periods: PPNA = Pre-pottery Neolithic, LBA = Late Bronze Age, Iron Age, Byzantine period, as well as fossil carp teeth from the Early Pleistocene palaeolake Ubeidiya.

See text for exact age ranges of the periods.


Enameloid δ18OPO4 and δ18OCO3 of Sparidae and Cyprinidae teeth

The difference between the δ18OCO3 and the δ18OPO4 in bioapatite of about +9‰ is in agreement with the theoretical predictions of equilibrium precipitation of the phosphate and the carbonate fractions simultaneously from the same body fluids according to the phosphate-water and carbonate-water temperature equations [1, 4, 30]. However, the difference between δ18OCO3 and δ18OPO4 (i.e., Δ18OCO3-PO4) of tooth enamel can vary between species and sometimes also within a single tooth ([30] and refs. therein). For example, Δ18OCO3-PO4 in tooth enamel of modern mammals ranges between 7.2 and 10.4‰ (e.g., [30, 40]) and can even reach a variation of up to 2.2‰ within the same tooth ([30], and refs. therein). Similarly high variability was also observed for Δ18OCO3-PO4 in modern shark teeth (enameloid and dentine), which were found to vary between 6.9 and 11.8‰ [24]. The intra jaw inter-tooth Δ18OCO3-PO4 variation in a single modern shark can also vary by more than 2‰ [24].

The Δ18OCO3-PO4 values of modern Sparidae and Cyprinidae enameloid samples vary between 8.1 and 11.0‰, yielding an average of 9.5±0.7‰, (n = 8 modern fish; Table 1). These values agree well with the average Δ18OCO3-PO4 of 9.1±1.5‰ (n = 12) reported for modern shark teeth [24]. Moreover, similar to sharks, high unsystematic variability (~3‰) was observed in the δ18OCO3 values of modern Sparidae teeth, while δ18OPO4 values vary less by <1‰ in the same teeth (Fig 2A). The high correlation between the δ18OCO3 and the δ18OPO4 in modern teeth (r2 = 0.73), indicates that both fractions precipitate simultaneously from the same body water pool [25]. However, the more variable δ18OCO3 values in modern bony fish and shark teeth in comparison to that of the δ18OPO4 (Fig 2A) suggests that the two fractions could be subjected to different exchange rates, thus integrating different ranges of ambient conditions. Further support for this distribution can be seen in the data obtained from ancient S. aurata teeth, particularly from the LB-Ir age. The larger scatter of the data obtained from the carbonate fraction of the ancient S. aurata teeth may, therefore, still reflect the in vivo distribution of δ18OCO3 values [1920, 24], rather than the effect of diagenetic alteration.

Fig 2. The difference between δ18O of co-existing carbonate and phosphate fractions in enameloid of modern and ancient Sparidae and Cyprinidae teeth.

(a) modern sharks, Sparidae and Cyprinidae (b) modern and ancient Sparidae as well as Cyprinidae. The black bold trend line: represents the δ18OPO418OCO3 relationship of modern mammals after Pellegrini et al. [25]. Note, the Δ18OCO3-PO4 prediction interval (PI) uncertainty of ±1.5‰ (stippled line) reported by Pellegrini et al. [25] agrees well with the range of modern shark teeth [19] and all bony fish data of Sparidae and Cyprinidae fall in this range too. The modern δ18OPO418OCO3 relationship of Sparidae and Cyprinidae, although yielding high correlation, covers a narrow δ18OPO418OCO3 range to allow comparison with the δ18OPO418OCO3 relationship of mammals. More data is needed to establish the δ18OPO418OCO3 relationship in modern bony fish.

The Δ18OCO3-PO4 values of ancient S. aurata teeth (ranging between 6.8 and 11.7‰, Fig 2B) and fossil Cyprinidae teeth (ranging between 8.2 and 10.0‰) are, in most cases, in good agreement with expected values for Δ18OCO3-PO4 in modern sharks [24] (Fig 2A). Furthermore, they are similar to Δ18OCO3-PO4 of mammalian bioapatite, for which diagenetic alteration of oxygen isotope composition is better quantified [2830]. The 1.4-million-year-old fossil teeth of Luciobarbus sp. (Cyprinidae) from the palaeolake of 'Ubeidiya are also well preserved as they fall in the expected δ18OPO418OCO3 range for modern bioapatite (Table 3). Thus, the Δ18OCO3-PO4 offset is similar for enamel (biogenic hydroxylapatite) and enameloid (biogenic fluorapatite).

We calculated the offset between the predicted and measured Δ18OCO3-PO4 values using the δ18OCO3 to δ18OPO4 relationship by Pellegrini et al. [30] which is based on all available modern mammal data. The calculated offset for ancient S. aurata (1.0±0.7‰) and fossil Luciobarbus sp. (Cyprinidae) (0.6±0.36‰) from the Δ18OCO3-PO4 regression falls within the predicted interval of uncertainty for the δ18OCO318OPO4 regression for mammals (Fig 2B). This suggests that the majority of enameloid δ18OCO3 (and by implication δ13C) values of ancient S. aurata teeth and fossil Cyprinidae teeth are well preserved and likely were not affected by any significant diagenetic alteration.

Enameloid δ13C of Sparidae teeth

δ13C values in animals vary as a function of environment and biochemical pathways. The largest difference in δ13C occurs between organic (e.g., lipids, carbohydrates, proteins, and collagen) and inorganic (e.g., shell carbonate or carbonate within bioapatite) carbon components (e.g., [4143]). The δ13C values of structurally bound carbonate in the bioapatite of the bones and teeth of herbivorous mammals are ~6–15‰ higher than the dietary sources of these animals (e.g., [4345]). With similar carbon isotope fractionation between tooth enamel bioapatite, breath CO2, and diet, the inter-species differences in the δ13C values of structurally bound carbonate in the bioapatite of mammals result primarily from differences in digestive physiology [4445]. While the δ13C of dentine in shark teeth derives predominantly from a dietary carbon source, the enameloid of the same teeth is typically characterised by 6–8‰ higher δ13C values [24]. Secondary, in vivo diffusional processes and isotopic exchange with ambient water DIC during tooth formation have been suggested as possible explanations for the elevated δ13C values of shark tooth enameloid [24].

The carbon intake of adult Sparidae is mainly derived from their diet; food items include the soft tissues of bivalves, gastropods and crustaceans [46]. Changing dietary carbon sources at different ontogenetic stages of the fish life cycle may explain the somewhat lower mean δ13C value of -2.3‰ found in a single juvenile fish (< 0.5 kg). Indeed, wild S. aurata have been observed to forage for shellfish and zooplankton [47] in a manner that varies with the size of the fish [48]. For instance, soft-bodied animals such as polychaetes and small crustaceans predominantly occur in the stomach contents of small S. aurata (5–9 cm in length), while larger, hard-shelled prey such as barnacles, bivalves and other teleost fish are consumed by larger S. aurata (10 to > 25 cm in length) [48].

Typical fractionation between carbonate and dissolved carbon in the form of bicarbonate ions is 2.5‰ at 20°C [49]. Hence, for typical Mediterranean seawater DIC δ13C values of 1.0±0.3‰ [3839] carbonate δ13C values of up to 3.5‰ would approach isotopic equilibrium with DIC. Such positive δ13C values between 1.6 and 4.8‰ have been recorded in the carbonate fraction of enameloid bioapatite of modern and ancient shark teeth [1920, 24]. A similar value was also measured in this study in a single sample of Scarus sp. (Scaridea, δ13C = 3.8‰) from the Red Sea (raised in a seawater aquarium-open system). Structurally bound carbonate δ13C values in enameloid bioapatite of modern Sparidae samples did not reach such high values. Thus, their body water bicarbonate pool seems still to be influenced to some degree by their dietary carbon sources, while approaching the isotopic composition of DIC in their habitat.

We estimate the percentage of metabolic carbon (ffood) in the carbonate δ13C values of modern Sparidae and Cyprinidae teeth (Fig 3) by using a simple mixing calculation: where δ13C of possible food sources for Sparidae range between -8.5 and -23‰ [5051] and for Cyprinidae between -15 and -34‰ [52]. The δ13C of Mediterranean DIC is 1.0±0.3‰ [3839] and between -4.0 and -7.0‰ for Lake Kinneret [53]. The fDIC and the fFood are the percent contribution of DIC and food, respectively, to the enameloid δ13C value (δenameloid), and α is the fractionation factor between the dissolved carbon sources and the δenameloid value (assuming a 2.5‰ fractionation). The ffood in the δ13C of juvenile Sparidae range between 20% and 50%, while for adult fish (both Sparidae and Cyprinidae) it ranges between 10% and 30% (Fig 3). Assuming that the structural carbonate in the enameloid of bony fish has a similar carbon isotope fractionation as calcite (assuming 1.0‰ fractionation factor), and not aragonite (as assumed for sharks [24]), then the metabolic effect of both Sparidae and Cyprinidae would be much lower than estimated and their δ13C values would further approach isotopic equilibrium with the water DIC. The Scaridea δ13C value seems to approach isotopic equilibrium with the Red Sea DIC, but may also relate to its durophageous diet feeding on corals and sponges.

Fig 3. Estimated fraction of metabolic carbon from the diet (ffood) in the carbonate δ13C values of modern Sparidae and Cyprinidae tooth enameloid from mixing calculation.

For the δ13Cdiet values of the food sources see text for details. Theoretical equilibrium δ13CCO3 of 3.5‰ and -2‰ for carbonate in fish enameloid for the South East Mediterranean (SEM) and Lake Kinneret, respectively, were calculated for an equilibrium fractionation of 2.5‰ (20°C), based on the δ13CDIC of 1‰ and -5‰, respectively, for the two different water bodies [3334, 48].

The δ13C of the freshwater Cyprinidae enameloid samples range between -6.6 and -5.2‰ while the lake water δ13CDIC values vary between -7.0 and -4.0‰ (mean δ13CDIC of around -5.5‰), with the Jordan River input ranging from -7.3 to -6.5‰ [53]. Thus, carbonate δ13C values of up to -2.0‰ would appear to approach isotopic equilibrium with the lake water DIC. Similar to Sparidae, Cyprinidae samples did not reach such high values, supporting the hypothesis that bioapatite δ13C values represent a mixture of ambient water δ13CDIC and δ13Cdiet.

The main control on the δ13CDIC in coastal lagoons is the mixing of fresh and marine waters entering the lagoon; meaning it is also a useful palaeosalinity indicator [54]. This is reflected in the co-variation of δ18OPO413C and δ18OCO313C values in tooth enameloid of S. aurata, which migrate between water bodies of different salinities [4648]. Thus, variations in the δ18OH2O and δ13CDIC values of the lagoonal water are expected to be recorded in the bioapatite of the S. aurata teeth.

Identifying the environmental history of fish ecosystems from paired δ18OPO4-CO3 and δ13C values

Fig 4A and 4B show paired δ18OPO413C and δ18OCO313C values from ancient S. aurata and fossil Luciobarbus sp. (Cyprinidae) teeth from the Levant, compared with the values of modern fish from the southeast Mediterranean coastal water and from Lake Kinneret, respectively. For comparison with pre-industrial values, modern fish tooth δ13C values were adjusted to the east Mediterranean Suess effect of ~1.2‰ on seawater δ13CDIC (the depletion trend in oceanic δ13CDIC, due to the emission of CO2 from burning fossil fuels) obtained from the vermetid reef carbonate δ13C composite record [38, 55]. The ancient S. aurata teeth that bear the δ18OPO4 signature typical for bioapatite formed in Mediterranean coastal water [2122] also fit the pre-industrial δ13C range for modern fish from this habitat. These values are mainly observed in the Early Holocene specimens, suggesting that the fish were caught in typical coastal water [22]. However, teeth with typical hypersaline δ18OPO4 values (> 23.7‰, [2122] similar to those of extant fish from the Bardawil lagoon [2], show typical marine δ13C values (Fig 4B).

Fig 4.

Covariance of (a) δ18OCO313C and (b) δ18OPO413C values of ancient and modern Sparidae as well as Cyprinidae tooth enameloid. The modern fish δ13C data is corrected for the Suess effect of -1.2‰ [33, 50] to pre-industrial values to enable comparison with the ancient Sparidae and Cyprenidae of the Southeast-Mediterranean region. Note that increasing δ18O values often also coincide with higher δ13C values recording evaporative changes in salinity and carbon source of the lagoonal water bodies these fish were caught from. A freshwater habitat (Lake Kinneret) of modern Cyprinidae is reflected primarily by the low δ13C, but also by δ18O values lower than those of marine Sparidae. Note that already in PPN some lagoons were mostly brackish-marine as indicated by low δ13C and δ18O values of the Sparus teeth but a few high values also imply the existence of hypersaline lagoons.

The highest δ13C values in the ancient S. aurata teeth are observed during the Mid-Late Holocene (Bronze Age and Iron Age) and are consistent with high δ18OPO4 values typical for tooth formation in hypersaline lagoons [2, 2122]. The increase in the δ13CDIC of closed, evaporative and warm water bodies, like the Bardawil lagoon, may be related to degassing of CO2 and seasonal biological effects, resulting in the high correlation between δ18OPO4 and δ13C in modern fish teeth (Fig 4B). Both, photosynthesis and degassing remove 12C-rich CO2 from the DIC pool of seawater leading to higher δ13CDIC values ([56] and refs. therein). These effects are amplified when they occur in a limited isotopic reservoir (i.e. a restricted and shallow water bodies), such as a long-term lakes or lagoons [5758]. Thus, similar to the coupling of δ18OCO3 and δ13C in carbonates of closed water bodies [57], the positive relationship between δ18OPO413C (r2 = 0.42; p<0.001) and δ18OCO313C (r2 = 0.41; p<0.001) values in the bioapatite of ancient and modern Sparidae teeth can be used as proxies for the formation of coastal lagoons.

The δ18OPO4 and δ13C values of the ancient fish teeth indicate that the hypersaline habitats of S. aurata in the Levant region formed during the Late Holocene from typical southeast Mediterranean coastal water under high evaporation rates and a limited carbon reservoir, similar to the present day Bardawil lagoon in Northern Sinai. This lagoon represents the main hypersaline habitat of Sparidae in the Levantine basin today [59] as well as in the past millennia [2122]. These evaporatively-influenced lagoonal habitats evolved during the Mid-to-Late Holocene transition (~3.5 ka ago), when sea level stabilised close to its present-day level [60] approximately 3,620±160 years BP [61]. Sea-level stabilisation resulted in the formation of the perennial shallow hypersaline Bardawil lagoon along the northern Sinai coast, due to the establishment of long-shore currents that transported Nile sands which built up blocking sandbars [58, 62].

During evaporation, the lighter 16O isotope is preferentially fractionated into the gas phase, this results in the gas phase having ~9‰ (at 25°C, the mean annual surface temperature of water bodies in the study area) lower δ18O values then the water mass from which it evaporated [63]. This, for the most part, is an inreversable reaction. As such, the residual water mass becomes more fractionated in a way that could be described through Rayleigh distillation equations [64]. As such, if we know the initial δ18O value of the water mass, and how fractionated it became, we can calculate the loss of water to the vapor phase. As the salt does not evaporate, the residual water mass will retain the original amount of salt. Dividing the initial amount ot salt by the final mass of water will allow us to calculate the salinity. Taking the observed 4 to 5‰ difference in δ18O values between the modern (representing the non-evaporated water mass) and Late Holocene S. aurata tooth assemblages, it is thus possible to calculate the fraction of water lost (Fig 5A). Within a reasonable ambient temperature range for the SE Mediterranean region (15 to 30°C), this translates to a 33% to 46% loss of the initial water mass–indicating significant evaporation and restricted water exchange between lagoons and the open Mediterranean Sea during the Late Holocene. Assuming an initial seawater salinity of ~37‰, this loss would place the salinity in these basins to between 54‰ and 68‰ (Fig 5B), representing the upper range of habitable conditions for S. aurata [2, 59]. This is a maximum estimate, given that about 2.5‰ (of the 5‰ range) in Sparus teeth δ18OPO4 of the Bardawil lagoon can potentially be attributed to seasonal changes in water temperatures [22].

Fig 5.

Calculated change in δ18Owater for different water temperatures (15–40°C) relative to initial conditions at 37‰ (PSU) for Mediterranean seawater as a function of (a) fraction of the initial water mass lost to vapor and (b) as a function of salinity. Horizontal bar illustrates the maximum range of δ18OPO4 difference observed in the Sparidae dataset, assuming this reflects only δ18OH2O (i.e. salinity) differences in the ambient water; vertical grey bars represent the water loss and salinity, respectively, inferred from this maximum Sparus δ18OPO4 range for different water temperatures.

Our stable isotope results of S. aurata tooth enameloid indicate that an early phase of lagoon formation in the Levant was already taking place during the Pre-Pottery Neolithic period (Fig 4) between 11,000 and 6,000 yr BC (Early-mid Holocene, Greenlandian to mid Northgrippian), despite the lack of sedimentological evidence for such lagoons. However, δ13C values suggest that these lagoons were likely also fed by terrestrial sources. In fact, Sparus can tolerate a wide salinity range, including brackish water [46]. While most of the δ18OPO4 values of the PPN (Early Holocene) fish fall within the expected range of bioapatite formed in isotopic equilibrium with Mediterranean temperature and seawater δ18O values, some of the PPN samples have relatively low δ13C and δ18OPO4 values. These low δ13C and δ18OPO4 values fall along the regression between modern fish from the Mediterranean and Lake Kinneret (Fig 4B). This implies a freshwater influx (with low δ13CDIC values) into past brackish habitats of Sparus during the PPN, Early Holocene. This provides evidence that S. aurata were caught in different lagoonal settings between the Early and Late Holocene.

Overall, the pairing of δ18OPO413C and δ18OCO313C in fish tooth enameloid thus represent a potential tool for investigating the hydrological history of marine to freshwater lagoonal settings and assessing the environments of past fish habitats. This isotopic approach can provide valuable information about past human fish exploitation patterns and enable us to distinguish marine from lagoonal fishing grounds.


The environmental history of fish ecosystems in the southeast Mediterranean during the Early to Late Holocene was studied by investigating the stable isotope composition of Sparidae tooth enameloid. The provenance of Sparidae from marine versus hypersaline lagoonal water was recorded in the δ18OPO4, δ18OCO3 and δ13C values of modern and ancient fish teeth, which reflects the wide range of saline habitats encountered during their migratory life cycle. The enameloid δ18OCO3 as well as the Δδ18OCO3-PO4 offset and δ13C values of Holocene Sparidae teeth from the Levant are consistent with those of modern fish, indicating only minor diagenetic alteration. The δ13C of Sparidae tooth enameloid reflects the water DIC of the fish habitat and its diet. Thus, carbon cycling in past fish environments can be characterised using the δ13C of fish teeth.

Paired δ18OPO413C and δ18OCO313C values of ancient Sparidae tooth enameloid were used to infer the environmental history of fish from coastal and lagoonal settings in the Levant. Our results indicate that Late Holocene hypersaline lagoons evolved from typical southeast Mediterranean coastal water under high rates of evaporation and limited carbon input from terrestrial sources. This lagoon formation occurred in association with the postglacial stabilisation of the sea level about 3,500 years ago, however, first lagoons had already formed during the Early Holocene and were likely influenced by freshwater input. Overall, the pairing of δ18OPO413C and δ18OCO313C of fish tooth enameloid is therefore, a potential tool to infer past hydrological balance and environmental shifts of fish ecosystems. This enables us to gain new insights into past fishing grounds that were exploited by (pre-)historic humans.


We thank Omri Lernau for providing the archaeological fish samples. We thank Torsten Vennemann (University of Lausanne) and Michael Maus (University of Mainz) for performing the phosphate oxygen isotope analysis of the silver phosphate samples. Jennifer Leichliter is acknowledged for proof reading of the English. We thank Ofer Bar Yosef for providing us the access to 'Ubeidiya fish remains.

We thank the editor Emmanuelle Puceat and ananonymous reviewers for their kind assistance in evaluating this paper.


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