Late Cenozoic C4 Expansion in the Central Inner Mongolia and Paleoenvironmental Evolution of the Qaidam Basin, China

Employing stable carbon and oxygen isotope analyses (including d13C and d18O of structural carbonate in mammalian tooth enamel and d18Op of fish bone samples) to reconstruct paleoecology and paleoenvironments, this dissertation strives to further our understanding of two important yet contentious issues: the late Miocene and Pliocene C4 expansion, and the late Cenozoic uplift history of the Tibetan Plateau. We collected and analyzed a diverse group of mammalian tooth enamel samples including horses, elephants, rhinos, deer, and giraffes from the central Inner Mongolia area (late Oligocene to modern) and Qaidam Basin (late Miocene to modern), and fish bone samples from the Qaidam Basin (late Miocene to modern). The major results are as follows: (1) In the central Inner Mongolia area, the d13C values of 91 tooth enamel samples of early late-Miocene age or older, with the exception of two 13 Ma rhino samples (-7.8 and -7.6‰) and one 8.5 Ma suspected rhino sample (-7.6‰), were all less than -8.0‰ (VPDB), indicating that there were no C4 grasses present in their diets and thus probably few or no C4 grasses in the ecosystems of the central Inner Mongolia prior to ~8 Ma. However, 12 out of 26 tooth enamel samples of younger ages (~7.5 Ma to ~3.9 Ma) have d13C values higher than -8.0‰ (up to -2.4‰), indicating that herbivores in the area had variable diets ranging from pure C3 to mixed C3-C4 vegetation during that time interval. The presence of C4 grasses in herbivores' diets (up to ~76% C4) suggests that C4 grasses were a significant component of the local ecosystems in the latest Miocene and early Pliocene, consistent with the hypothesis of a global factor as the driving mechanism of the late Miocene C4 expansion. Today, C3 grasses dominate grasslands in the central Inner Mongolia area. (2) In the Qaidam Basin, the d13C values of mammalian tooth enamel samples generally show only small variations and are mostly less than 8‰ for modern samples and less than -7‰ for fossils, except a rhino tooth CD0722 from Shengou (late Miocene, with a current best estimated age of ~8–10 Ma) that yielded d13C values up to -4.1‰ - an unambiguous indication of a significant intake of C4 plants in its diet (up to ~56%). If the Qaidam Basin was as arid as today during the late Miocene and early Pliocene, this would indicate that the animals had pure or nearly pure C3 diets and the local ecosystems were likely composed of pure or nearly pure C3 vegetation and that the lone rhino (CD0722) was more likely a migrant primarily lived in places where C4 plants were present and migrated to the Shengou area. If, however, the Qaidam Basin was warmer and more humid during the late Miocene and early Pliocene than today as suggested by geological evidence, then at least a few more samples from Shengou (late Miocene) and Huaitoutala (early Pliocene) in addition to CD0722 indicate significant dietary intakes of C4 plants based on their tooth enamel d13C values. This suggests that the Qaidam Basin very likely had more C4 plants in the local ecosystems during the late Miocene and early Pliocene than today. Moreover, the Qaidam Basin probably also had much denser vegetation at that time to support additional large mammals such as rhinos and elephants. Nonetheless, the C4 plants seemed to have not been consistently utilized because C3 plants, which were more nutritious and easier to digest than C4 plants, were readily available. Today, C3 plants dominated the sparse vegetation in the Qaidam Basin. The d18O values of these samples did not increase monotonously with time. However, the range of variation seems to have increased considerably since the early Pliocene, indicating increased aridification in the basin. The mean d18O values of large mammals and the reconstructed d18O compositions of local meteoric water display a significant negative shift from the Tuosu Nor-Quanshuiliang interval (~11-11.2 Ma) to the Shengou-Naoge interval (~9-9.5 Ma), which is consistent with the marine d18O record showing a cooling trend in the same period as suggested by a positive excursion in the d18O values of benthic forams (Zachos et al., 2001). (3) The oxygen isotope compositions of phosphate (d18Op) from fish bone samples from the Qaidam Basin showed statistically significant enrichment in d18Op from the Tuxi-Shengou-Naoge interval (late Miocene) to the Yahu interval (early Pliocene) and from the Yahu interval to the present day. This is most likely reflecting increases in the d18O of lake water over time. Estimated water temperatures for the modern Qinghai Lake from fish bone d18Op and measured lake water d18Ow using the Longinelli and Nuti (1973) equation range from 19.3 to 23.1 (± 0.3) °C, about 4 to 11°C higher than the reported average temperature of surface water during the summer (~12-15°C). This indicates that the Qinghai Lake fish was not living in the saline Qinghai Lake itself exclusively, but spent at least part of its life in estuary or in the fresh water of an inflow river(s). Temperatures calculated from the average fish bone d18Op values and the average d18Ow derived from structural carbonate d18O of large mammal tooth enamel samples were much too low to be reasonable. Temperatures estimated from d18Op of fish bones and d18Ow estimated from d18O of co-ocurring large mammal tooth enamel samples (including using the highest d18Ow to represent the dry season during which the large mammals presumably had to drink from the lake/lakes) were all lower than the average temperature of the modern Qinghai Lake surface water during the summer, and some were too low to be reasonable – even without considering the observed cooling trend since the late Miocene (which means that lake water probably had higher temperatures during the late Miocene and early Pliocene). These indicate that the fish and the large mammals were not in equilibrium with the same water, which is expected as the d18O values of enamel from large mammals have been shown to generally track the d18O values of meteoric water, whereas the d18O value of lake water as recorded in fish bones often deviates significantly from the d18O of meteoric (source) water due to environmental/hydrological factors such as evaporation. Using the relationship between salinity and d18Ow observed for the modern Qinghai Lake and its surrounding lakes and ponds and assuming that this relationship was applicable in the geological past, we calculated the paleosalinities of lake waters to be from 0.13-0.26 g/l (4°C) to 0.54-1.03 g/l (20°C) for the Tuxi-Shengou-Naoge interval, and from 0.93-2.03 g/l (4°C) to 3.74-8.13 g/l (20°C) for the Yahu interval. The estimated salinity for Yahu is most likely too low, as the extraordinarily thickened skeleton that left little room for muscles would probably require much higher salinity (as well as Ca concentration). In summary, these data, in combination with data from other localities in North China and around the world, provide support for a global factor for driving the late Miocene C4 expansion as suggested by Cerling et al. (1997). The central Inner Mongolia and the Qaidam Basin were warmer and more humid during the late Miocene and early Pliocene, and had significant C4 components in the local ecosystems, but they had different moisture sources: the former from the western Pacific Ocean, while the latter from the Indian Ocean and Bay of Bengal. Further significant uplift of the Tibetan Plateau (at least portions of it) drew some moisture away from the central Inner Mongolia area, diverting it to the Chinese Loess Plateau (CLP) and later the Linxia Basin. The uplift also blocked most of the moisture from the Indian Ocean and Bay of Bengal as well as that from the western Pacific Ocean from reaching the Qaidam Basin. The result is the retreat or significant reduction of C4 grasses in the local ecosystems of the central Inner Mongolia area and the Qaidam Basin after early Pliocene, and the further expansion of C4 plants in the CLP and later Linxia Basin.