Conjugated polymers based on metalla-aromatic building blocks

Conjugated polymers usually require strategies to expand the range of wavelengths absorbed and increase solubility. Developing effective strategies to enhance both properties remains challenging. Herein, we report syntheses of conjugated polymers based on a family of metalla-aromatic building blocks via a polymerization method involving consecutive carbyne shuttling processes. The involvement of metal d orbitals in aromatic systems efficiently reduces band gaps and enriches the electron transition pathways of the chromogenic repeat unit. These enable metalla-aromatic conjugated polymers to exhibit broad and strong ultraviolet–visible (UV–Vis) absorption bands. Bulky ligands on the metal suppress π–π stacking of polymer chains and thus increase solubility. These conjugated polymers show robust stability toward light, heat, water, and air. Kinetic studies using NMR experiments and UV–Vis spectroscopy, coupled with the isolation of well-defined model oligomers, revealed the polymerization mechanism.

Conjugated polymers are macromolecules usually featuring a backbone chain with alternating double and single bonds (1–3). These characteristics allow the overlapping p-orbitals to form a system with highly delocalized π-electrons, thereby giving rise to intriguing chemical and physical properties (4–6). They have exhibited many applications in organic light-emitting diodes, organic thin film transistors, organic photovoltaic cells, chemical sensors, bioimaging and therapies, photocatalysis, and other technologies (7–10). To facilitate the use of solar energy, tremendous efforts have been devoted in recent decades to developing previously unidentified conjugated polymers exhibiting broad and strong absorption bands (11–13). The common strategies for increasing absorption involve extending π-conjugation by incorporating conjugated cyclic moieties, especially fused rings; modulating the strength of intramolecular charge transfer between donor and acceptor units (D–A effect); increasing the coplanarity of π conjugation through weak intramolecular interactions (e.g., hydrogen bonds); and introducing heteroatoms or heavy atoms into the repeat units of conjugated polymers (11–16). Additionally, appropriate solubility is a prerequisite for processing and using polymers and is usually achieved with the aid of long alkyl or alkoxy side chains (12, 17).Aromatic rings are among the most important building blocks for conjugated polymers. In addition to aromatic hydrocarbons, a variety of aromatic heterocycles composed of main-group elements have been used as fundamental components. These heteroatom-containing conjugated polymers show unique optical and electronic properties (4–10). However, while metalla-aromatic systems bearing a transition metal have been known since 1979 due to the pioneering work by Thorn and Hoffmann (18), none of them have been used as building blocks for conjugated polymers. The HOMO–LUMO gaps (E_g) of metalla-aromatics are generally narrower (Fig. 1) than those of their organic counterparts (19–22). We reasoned that this feature should broaden the absorption window if polymers stemming from metalla-aromatics are achievable.Open in a separate windowFig. 1.Comparison of traditional organic skeletons with metalla-aromatic building blocks (the computed energies are in eV). (A) HOMO–LUMO gaps of classic aromatic skeletons. (B) Carbolong frameworks as potential building blocks for novel conjugated polymers with broad absorption bands and improved solubility.In recent years, we have reported a series of readily accessible metal-bridged bicyclic/polycyclic aromatics, namely carbolong complexes, which are stable in air and moisture (23–25). The addition of osmium carbynes (in carbolong complexes) and alkynes gave rise to an intriguing family of d_π–p_π conjugated systems, which function as excellent electron transport layer materials in organic solar cells (26, 27). These observations raised the following question: Can this efficient addition reaction be used to access metalla-aromatic conjugated polymers? It is noteworthy that incorporation of metalla-aromatic units into conjugated polymers is hitherto unknown. In this contribution, we disclose a polymerization reaction involving M≡C analogs of C≡C bonds, which involves a unique carbyne shuttling strategy (Fig. 2A). This led to examples of metalla-aromatic conjugated polymers (polycarbolongs) featuring metal carbyne units in the main chain. On the other hand, the development of polymerization reactions plays a crucial role in involving certain building blocks in conjugated polymers (28–32). These efficient, specific, and feasible polymerizations could open an avenue for the synthesis of conjugated polymers.Open in a separate windowFig. 2.Design of polymers and synthesis of monomers. (A) Schematic illustration of the polymerization strategy. (B) Preparation of carbolong monomers. Insert: X-ray molecular structure for the cations of complex 3. Ellipsoids are shown at the 50% probability level; phenyl groups in PPh₃ are omitted for clarity.     

From the Cover: Volcanic history of the Imbrium basin: A close-up view from the lunar rover Yutu

We report the surface exploration by the lunar rover Yutu that landed on the young lava flow in the northeastern part of the Mare Imbrium, which is the largest basin on the nearside of the Moon and is filled with several basalt units estimated to date from 3.5 to 2.0 Ga. The onboard lunar penetrating radar conducted a 114-m-long profile, which measured a thickness of ∼5 m of the lunar regolith layer and detected three underlying basalt units at depths of 195, 215, and 345 m. The radar measurements suggest underestimation of the global lunar regolith thickness by other methods and reveal a vast volume of the last volcano eruption. The in situ spectral reflectance and elemental analysis of the lunar soil at the landing site suggest that the young basalt could be derived from an ilmenite-rich mantle reservoir and then assimilated by 10–20% of the last residual melt of the lunar magma ocean.The surface of the Moon is covered by regolith, a mixed layer of fine-grained lunar soil and ejecta deposits, which is crucial to understanding the global composition of the Moon. The lunar regolith has also recorded the complex history of the surface processes, and it is the main reservoir of ³He and other solar wind gases. The thickness of the lunar regolith was estimated to be from 2 to 8 m in the maria and up to 8–16 m in the highland areas using various methods (1), including crater morphology (2, 3), seismology with low spatial resolution (4), radar wave scattering (5), and microwave brightness temperature (6). However, no in situ measurement of spectral reflectance, elemental compositions, lunar regolith thickness, or subsurface structures has been carried out.The surface of the Moon is dominated with numerous large basins. They were formed about 3.9 Ga (7, 8), probably by the late heavy bombardment, and then filled with dark lava flows derived from partial melting of the lunar mantle, within a period mainly during 3.8–3.1 Ga (7). The Imbrium basin is the largest and was formed on Procellarum KREEP [potassium (K), rare earth elements (REE), and phosphorus (P)] Terrane (9), a unique terrain highly enriched in U, Th, and K radionuclides and other incompatible trace elements referred to as KREEP (10) and considered as the last residual melt of the Lunar Magma Ocean (11). The presence of the KREEPy materials, indicated by high concentrations of radionuclides U, Th, and K (9), around the rims of the Imbrium basin suggests that they are likely the basin-forming ejecta deposits. At least three main lava flows, dated from 3.5 Ga to 2.0–2.3 Ga (7, 12), have been recognized in Mare Imbrium with distinct FeO and TiO₂ concentrations (13, 14), which brought up interior information of this KREEP-rich terrain. The old and low-Ti basalt unit has been sampled by the Apollo 15 mission that landed at the eastern rim of the Imbrium basin. Information of other lava flows in Mare Imbrium was obtained only by remote sensing from orbit. On December 14, 2013, Chang’e-3 successfully landed on the young and high-Ti lava flow in the northeastern Mare Imbrium, about 10 km south from the old low-Ti basalt unit (Fig. 1).Open in a separate windowFig. 1.The landing site of Chang’e-3 (red cross), on the high-Ti basalt (dark gray) near the boundary in contact with the low-Ti basalt (light gray). The background image was taken by Chang’e-1.The lunar rover Yutu (named for the jade rabbit on the Moon in a Chinese fairy tale) was equipped with an active particle-induced X-ray spectrometer (APXS), a visible to near-infrared (450–945 nm) imaging spectrometer and short-wave infrared (900–2,395 nm) spectrometer (VNIS), and a lunar penetrating radar (LPR), accompanied by a stereo camera and a navigating camera. Originally, the mission planned to have the lunar rover measure chemical and mineral compositions of the lunar soil and various types of ejecta rocks and to carry out a LPR profile of the lunar regolith and subsurface structures in the first 3 mo. The mission was scheduled to extend up to 1 y and to explore the old low-Ti lava flow ∼10 km north. Unfortunately, some of Yutu’s mechanical parts failed to move just before the rover prepared for sleeping at the end of the second month due to unknown faults probably in the control system. During the first 2 mo, Yutu successfully carried out two APXS and four VNIS analyses of the lunar soil and performed a 114-m-long LPR profile along the rover track in the landing area (Fig. 2). These in situ measurements provide insights into the volcanic history of Mare Imbrium and the ground-truth data for calibration of the orbital data.Open in a separate windowFig. 2.Chang’e-3 landing site and the rover Yutu’s track. Crater A is blocky, indicating penetration through the regolith. Crater B is the largest one without blocks in the landing area. The APXS (LS1–LS2) and VNIS (CD5–CD8) analysis positions and the rover navigation points are marked. The image was composed from the series images taken by the Chang’e-3 landing camera.     

Pictet–Spengler reaction-based biosynthetic machinery in fungi

The Pictet–Spengler (PS) reaction constructs plant alkaloids such as morphine and camptothecin, but it has not yet been noticed in the fungal kingdom. Here, a silent fungal Pictet–Spenglerase (FPS) gene of Chaetomium globosum 1C51 residing in Epinephelus drummondhayi guts is described and ascertained to be activable by 1-methyl-l-tryptophan (1-MT). The activated FPS expression enables the PS reaction between 1-MT and flavipin (fungal aldehyde) to form “unnatural” natural products with unprecedented skeletons, of which chaetoglines B and F are potently antibacterial with the latter inhibiting acetylcholinesterase. A gene-implied enzyme inhibition (GIEI) strategy has been introduced to address the key steps for PS product diversifications. In aggregation, the work designs and validates an innovative approach that can activate the PS reaction-based fungal biosynthetic machinery to produce unpredictable compounds of unusual and novel structure valuable for new biology and biomedicine.Microbes and plants produce a multitude of unpredictably structured organic molecules known as secondary metabolites (natural products), from which more than half of globally marketed drugs have been developed (1–3). Large-scale genomic mining has indicated that microbial secondary metabolites are substantially underestimated because many biosynthetic genes remain silent or less active in the laboratory cultivation conditions (4, 5). Accordingly, there has long been an urgent need to develop a new strategy that enables microorganisms to produce more unforeseeable bioactive compounds, which are important to the drug discovery efforts to combat life-threatening diseases (6, 7), and to the complexity-based driving force for synthetic and material chemistry (8–10).Characterized by forming a piperidine ring through a condensation of β-arylethylamine with an aldehyde, Pictet–Spengler (PS) reaction contributes greatly to the framework diversification of important alkaloidal phytochemicals such as morphine, camptothecin, and reserpine (Fig. 1A), with plant-derived Pictet–Spenglerase (called strictosidine synthetase, STR) mechanistically addressed (11). The PS mechanism has been presumed to involve in the tetrahydroisoquinoline antibiotic biosynthesis in the bacterium Streptomyces lavendulae (12, 13), and likely in the biosynthetic pathway of hyrtioreticulin F in the marine sponge Hyrtios reticulatus (14). However, surprisingly, nothing is known concerning the PS reaction in the fungal kingdom.Open in a separate windowFig. 1.Alkaloids derived from the PS reaction. (A) Representatives for PS reaction-based phytochemicals. (B) 1-MT has been found to be a potent up-regulator for the FPS gene expression, and a suitable FPS substrate for its PS condensation with flavipin to yield unnatural natural products (1−8) with unprecedented skeletons.Most if not all Chaetomium fungi in the Chaetomiaceae family produce l-tryptophan–derived alkaloids, but “refuse” to generate any PS reaction-based secondary metabolite (15–18). However, a comparative genomic analysis has clarified that C. globosum 1C51 does have an FPS gene (CHGG_06703, STR-like) (SI Appendix, Fig. S25), but remains silent or poorly activated in the laboratory cultivations because no PS-derived secondary metabolite has been detected in the fungal culture. Therefore, this C. globosum 1C51 strain was adopted here to test for the activation of its “unworking” PS reaction-based biosynthetic machinery. As a result, 1-methyl-l-tryptophan (1-MT) was demonstrated to be able to up-regulate the FPS expression and condense with the fungal aldehyde flavipin (3,4,5-trihydroxy-6-methyl phthalaldehyde) to form unexpectedly a family of skeletally unprecedented alkaloids, trivially named chaetoglines A−H (1−8) (Fig. 1B). A gene-implied enzyme inhibition (GIEI) strategy, derived from the hypothesis-based enzyme modulation described elsewhere (19, 20), was introduced to identify the key diversification steps for the PS reaction-derived compounds (Figs. 2–4). Chaetoglines B (2) and F (6) have been found to be more antibacterial than tinidazole (a coassayed drug prescribed in clinic for bacterial infections) against pathogenic anaerobes Veillonella parvula, Bacteroides vulgatus, Streptococcus sp., and Peptostreptococcus sp. Moreover, alkaloid 6 is potently inhibitory on acetylcholinesterase (AChE), an effective target enzyme exploited for the treatment of Alzheimer’s disease (21, 22).Open in a separate windowFig. 2.LC-MS profile-based comparisons for the chaetogline production in monooxygenase inhibitor exposed fungal cultures: (A) for 3–6; (B) for 1–2 and 7–8. The ESI-MS spectra (C) of 1−8 (①) displayed the corresponding protonated and Na⁺-liganded molecular ions. Samples were ethyl acetate extracts derived from 1-MT supplemented C. globosum 1C51 cultures without (②) and with the separate exposure to PB, PR, and MMI at 0.1 (③∼⑤) and 1.0 mM (⑥∼⑧), respectively.Open in a separate windowFig. 4.Proposed generation of the fungal PS-derived products. Catalyzed by FPS, 1-MT and flavipin (a fungal aldehyde) undergo PS reaction to form chaetoglines A−H (1−8) in concert with tailing reactions including oxidation, decarboxylation, and Aldol reaction. The Schiff base intermediate 9 tends to tautomerize via 10 and 11 to give chaetogline C (3) that can be methyl-esterified into chaetogline D (4). Intramolecular cyclization of 10 gives 12, which is oxidizable into chaetoglines A (1) and E (5), the latter yielding chaetogline F (6) after the oxidative and decarboxylative aromatization. Chaetogline E (5) can also be oxidized to intermediate 13, which gives 14 after condensing presumably via the decarboxylative Aldol reaction (34, 35) with 1-M-IAA derived from 1-MT by the fungi (SI Appendix, Fig. S24). Intermediate 14 undergoes intramolecular cyclization, monooxygenation, and isomerization to form chaetogline G (7), which after decarboxylation gives chaetogline H (8), a precursor of chaetogline B (2).     

Studies on enmetazobactam clarify mechanisms of widely used β-lactamase inhibitors

β-Lactams are the most important class of antibacterials, but their use is increasingly compromised by resistance, most importantly via serine β-lactamase (SBL)-catalyzed hydrolysis. The scope of β-lactam antibacterial activity can be substantially extended by coadministration with a penicillin-derived SBL inhibitor (SBLi), i.e., the penam sulfones tazobactam and sulbactam, which are mechanism-based inhibitors working by acylation of the nucleophilic serine. The new SBLi enmetazobactam, an N-methylated tazobactam derivative, has recently completed clinical trials. Biophysical studies on the mechanism of SBL inhibition by enmetazobactam reveal that it inhibits representatives of all SBL classes without undergoing substantial scaffold fragmentation, a finding that contrasts with previous reports on SBL inhibition by tazobactam and sulbactam. We therefore reinvestigated the mechanisms of tazobactam and sulbactam using mass spectrometry under denaturing and nondenaturing conditions, X-ray crystallography, and NMR spectroscopy. The results imply that the reported extensive fragmentation of penam sulfone–derived acyl–enzyme complexes does not substantially contribute to SBL inhibition. In addition to observation of previously identified inhibitor-induced SBL modifications, the results reveal that prolonged reaction of penam sulfones with SBLs can induce dehydration of the nucleophilic serine to give a dehydroalanine residue that undergoes reaction to give a previously unobserved lysinoalanine cross-link. The results clarify the mechanisms of action of widely clinically used SBLi, reveal limitations on the interpretation of mass spectrometry studies concerning mechanisms of SBLi, and will inform the development of new SBLi working by reaction to form hydrolytically stable acyl–enzyme complexes.

β-Lactamases are a major mechanism of resistance to the clinically vital β-lactam antibiotics, with >2,000 different β-lactamases reported (1). β-Lactamases are grouped into classes A, C, and D, which employ a nucleophilic serine in catalysis (serine β-lactamases, SBLs), and class B, which employ metal ions in catalysis (2). Presently, SBLs are the most important β-lactamases from a clinical perspective. SBL inhibitors (SBLi) have been developed for use in combination with a β-lactam antibiotic, with tazobactam (3), sulbactam (4), and clavulanic acid (5) being the most widely used SBLi. These SBLi all contain a β-lactam ring which reacts with SBLs to produce an acyl–enzyme complex (AEC) intermediate, as is also the case for efficient SBL substrates (Fig. 1A). With efficient substrates the β-lactam–derived AEC is readily hydrolyzed. With SBLi the reaction bifurcates at the AEC stage; in addition to hydrolysis, reaction of the AEC via opening of the β-lactam fused five-membered ring occurs to give one or more relatively hydrolytically stable species (Figs. 1B and ​and2).2). The nature of these species is central to SBLi inhibition and has been studied by crystallography (6–11) and ultraviolet-visible (UV/Vis) (10, 12) and Raman (6, 7, 9, 12–15) spectroscopy, as well as different types of mass spectrometry (MS) (10, 16–22).Open in a separate windowFig. 1.Sulfone derivatives of penicillins are potent clinically used mechanism-based inhibitors of SBLs. (A) Outline mechanism for penicillin hydrolysis as catalyzed by SBLs; reaction proceeds via an AEC, which is efficiently hydrolyzed. (B) Sulfone derivatives of penicillins are SBLi that react to give one or more hydrolytically stable complex(es), the nature of which was the focus of our work.Open in a separate windowFig. 2.Pathways for reactions of penam sulfones with SBLs. Following initial acyl–enzyme 2 formation the main transient inactivation pathway occurs via thiazolidine ring opening to give species 3-5 which are relatively stable to hydrolysis. Fragmentation of 3-5 can occur in rare cases and is promoted by acid to give 6-8 or heat to give 11. In rare cases fragmentation of 2-5 can result in irreversible inactivation of the SBL to give 9 and 10. Efficient hydrolysis of the β-lactam occurs to give a β-amino acid product 12, which in solution fragments to give 13-16. Our results imply biologically relevant inhibition involves 3-5, or equivalent mass species.The structures of tazobactam and sulbactam are closely related to those of the penicillins; they differ by lack of a C-6 side chain, functionalization of the pro-S methyl group (in case of tazobactam), and by oxidation of the thiazolidine to a sulfone. These differences result in a loss of useful antibacterial activity but a gain of potent SBL inhibition. Although the presence of sulfur in drugs is common [e.g., sulfonamide antibiotics (23)] and there is growing interest in covalently acting drugs (24, 25), sulfones are rare in drugs and, as far as we are aware, sulbactam and tazobactam are the only clinically approved sulfone-containing drugs working by covalent reaction with their targets (26–28).Since the clinical introduction of the pioneering SBLi, β-lactamases have evolved and SBLi use is increasingly compromised by extended spectrum β-lactamases (ESBLs) and inhibitor-resistant SBLs (29). Efforts have been made to develop new SBLi, including those with and without a β-lactam. The latter include diazabicyclooctanes (30) and cyclic boronates (31, 32). However, β-lactam–containing SBLi remain of most clinical importance. Among SBLi in clinical development, enmetazobactam (formerly AAI-101; Fig. 1) is of particular interest because it is a “simple” N-methylated derivative of the triazole ring of tazobactam (33). In combination with cefepime, enmetazobactam is reported to manifest substantially better antimicrobial properties against class A ESBL-producing strains than the commonly used piperacillin/tazobactam combination (20, 33, 34).We report studies on the mechanism of SBL inhibition by enmetazobactam using denaturing and nondenaturing (native) MS methods, NMR spectroscopy, and crystallography. The results led us to reevaluate the mechanisms of SBL inhibition by the clinically important sulfone-containing SBLi, i.e., tazobactam and sulbactam, and reveal limitations on the interpretation of MS studies concerning SBL inhibition.     

From the Cover: Eleventh-century shift in timber procurement areas for the great houses of Chaco Canyon

An enduring mystery from the great houses of Chaco Canyon is the origin of more than 240,000 construction timbers. We evaluate probable timber procurement areas for seven great houses by applying tree-ring width-based sourcing to a set of 170 timbers. To our knowledge, this is the first use of tree rings to assess timber origins in the southwestern United States. We found that the Chuska and Zuni Mountains (>75 km distant) were the most likely sources, accounting for 70% of timbers. Most notably, procurement areas changed through time. Before 1020 Common Era (CE) nearly all timbers originated from the Zunis (a previously unrecognized source), but by 1060 CE the Chuskas eclipsed the Zuni area in total wood imports. This shift occurred at the onset of Chaco florescence in the 11th century, a time with substantial expansion of existing great houses and the addition of seven new great houses in the Chaco Core area. It also coincides with the proliferation of Chuskan stone tools and pottery in the archaeological record of Chaco Canyon, further underscoring the link between land use and occupation in the Chuska area and the peak of great house construction. Our findings, based on the most temporally specific and replicated evidence of Chacoan resource procurement obtained to date, corroborate the long-standing but recently challenged interpretation that large numbers of timbers were harvested and transported from distant mountain ranges to build the great houses at Chaco Canyon.The high desert landscape of Chaco Canyon, New Mexico was the locale of a remarkable cultural development of Ancestral Puebloan peoples, including the construction of some of the largest pre-Columbian buildings in North America (1) (Fig. 1). The monumental “great houses” of Chaco Canyon reflect an elaborate socioecological system that spanned much of the 12,000-km² San Juan Basin from 850 to 1140 Common Era (CE) (2). These massive stone masonry structures required a wealth of resources to erect, including an estimated 240,000 trees incorporated as roof beams, door and window lintels, and other building elements (3). The incongruity of the great houses located in a nearly treeless landscape has led archaeologists and paleoecologists to investigate the origins of timbers used in construction (4–9). Beyond the simple curiosity driving this question, the answer has important implications for understanding the complexities of human–environmental interactions, the sociopolitical organization, and the economic structure of Chacoan society (10–12).Fig. 1.Aerial view of Pueblo Bonito, the largest of the Chaco Canyon great houses. Image courtesy of Adriel Heisey.The first excavators of the great houses in the early 20th century speculated that construction timbers were harvested locally, perhaps resulting in deforestation of the surrounding landscape (13). Paleoecological studies conducted during the late 1970s and early 1980s, however, showed that ponderosa pine (Pinus ponderosa), the primary tree species used in construction, was not abundant enough at the relatively low elevations (1,800–2,000 m above sea level) of Chaco Canyon and nearby mesas to support timber demand (14–16). Spruce (Picea spp.) and fir (Abies spp.), which account for tens of thousands of construction beams, have been absent from Chaco Canyon for at least 12,000 y and could have only been logged from distant, higher-elevation sites (2,500–3,450 m above sea level) (4). An inadequate supply of timbers in Chaco Canyon and its immediate surroundings during Puebloan occupation strongly suggests long-distance procurement from surrounding mountain ranges, where all three conifers now grow in abundance. This inference was corroborated by strontium isotope (⁸⁷Sr/⁸⁶Sr)-based sourcing. Through a comparison of ⁸⁷Sr/⁸⁶Sr values from great-house timbers to ₈₇Sr/₈₆Sr values from conifer stands growing today in mountains surrounding the San Juan Basin, two studies concluded that the Chuska Mountains (75 km west) and Mount Taylor (85 km southeast) were the most likely sources for spruce, fir, and ponderosa pine trees (6, 7). Recently, the explanation of long-distance timber transport and the related interpretations of ⁸⁷Sr/⁸⁶Sr evidence have been challenged and an alternative has been proposed that most great-house timbers (particularly ponderosa pine) were just as likely to have originated from nearby and low-elevation sites within, east, and south of Chaco Canyon (8, 9).We assessed probable timber origins independently from previous efforts by applying tree-ring width-based sourcing techniques to a set of 170 beams from our archives at the University of Arizona. These beams comprise six tree species from seven great-house structures (17) (Fig. S1). Each site chronology, as the average of 40–100 trees, represents tree-ring growth patterns peculiar to an individual landscape. This method of identifying the probable origin of timbers has been applied widely in Europe in the study of archaeological and nautical timbers and artifacts, musical instruments, and paintings on oak panels (17–20). These techniques are underused in North America, but recent efforts in the northeastern United States have revealed distant, inland sources for 18th- and 19th-century nautical timbers (21, 22).Table S1.Number of sourced beams by species and great-house structureFig. S1.An example of sourcing a great-house beam via tree-ring-width methods. (A) An individual beam (black line), the ponderosa pine JPB-88 from Pueblo del Arroyo, and the Chuskas chronology (red line). (B) Bivariate plot comparing the ring-width indices of ...Tree-ring sourcing can only be applied where tree growth patterns are distinguishable between the potential locations of origin. In the southwestern United States tree growth primarily responds to regionally coherent winter precipitation (23, 24), and as a consequence trees across the region tend to share roughly half of their interannual variability (25). Differences between site chronologies are predominantly attributed to variations in topography and subregional-scale climate conditions (26).We compared great-house beams to the site chronologies of eight potential harvesting areas surrounding the San Juan Basin. Chaco Canyon was not included as one of our sites because it lacked enough remnant wood from the Chaco era to build a local site chronology. To assess the efficacy and accuracy of the tree-ring sourcing method within the San Juan Basin, we tested whether tree-ring growth patterns could be distinguished between the various mountain ranges surrounding Chaco Canyon by applying sourcing methods to living trees of known origin (SI Text and Fig. S2).Table S2.Tree-ring sites used to evaluate tree-ring-based sourcing in the San Juan BasinFig. S2.Evaluation of dendroprovenance in the San Juan Basin. (A–F) Each tile provides a different test for a set of modern trees (green triangles). Circle sizes are proportional to the number of trees (labeled in the circle) sourcing to a given location. ...     

Drought,agricultural adaptation,and sociopolitical collapse in the Maya Lowlands

Paleoclimate records indicate a series of severe droughts was associated with societal collapse of the Classic Maya during the Terminal Classic period (∼800–950 C.E.). Evidence for drought largely derives from the drier, less populated northern Maya Lowlands but does not explain more pronounced and earlier societal disruption in the relatively humid southern Maya Lowlands. Here we apply hydrogen and carbon isotope compositions of plant wax lipids in two lake sediment cores to assess changes in water availability and land use in both the northern and southern Maya lowlands. We show that relatively more intense drying occurred in the southern lowlands than in the northern lowlands during the Terminal Classic period, consistent with earlier and more persistent societal decline in the south. Our results also indicate a period of substantial drying in the southern Maya Lowlands from ∼200 C.E. to 500 C.E., during the Terminal Preclassic and Early Classic periods. Plant wax carbon isotope records indicate a decline in C₄ plants in both lake catchments during the Early Classic period, interpreted to reflect a shift from extensive agriculture to intensive, water-conservative maize cultivation that was motivated by a drying climate. Our results imply that agricultural adaptations developed in response to earlier droughts were initially successful, but failed under the more severe droughts of the Terminal Classic period.The decline of the lowland Classic Maya during the Terminal Classic period (800–900/1000 C.E.) is a preeminent example of societal collapse (1), but its causes have been vigorously debated (2–5). Paleoclimate inferences from lake sediment and cave deposits (6–11) indicate that the Terminal Classic was marked by a series of major droughts, suggesting that climate change destabilized lowland Maya society. Most evidence for drought during the Terminal Classic comes from the northern Maya Lowlands (Fig. 1) (6–8, 10), where societal disruption was less severe than in the southern Maya Lowlands (12, 13). There are fewer paleoclimate records from the southern Maya Lowlands, and they are equivocal with respect to the relative magnitude of drought impacts during the Terminal Classic (9, 11, 14). Further, the supposition that hydrological impacts were a primary cause for societal change is often challenged by archaeologists, who stress spatial variability in societal disruption across the region and the complexity of human responses to environmental change (2, 3, 12). The available paleoclimate data, however, do not constrain possible spatial variability in drought impacts (6–11). Arguments for drought as a principal cause for societal collapse have also not considered the potential resilience of the ancient Maya during earlier intervals of climate change (15).Open in a separate windowFig. 1.Map of the Maya Lowlands indicating the distribution of annual precipitation (64) and the location of paleoclimate archives discussed in the text. The locations of modern lake sediment and soil samples (Fig. 2) are indicated by diamonds.For this study, we analyzed coupled proxy records of climate change and ancient land use derived from stable hydrogen and carbon isotope analyses of higher-plant leaf wax lipids (long-chain n-alkanoic acids) in sediment cores from Lakes Chichancanab and Salpeten, in the northern and southern Maya Lowlands, respectively (Fig. 1). Hydrogen isotope compositions of n-alkanoic acids (δD_wax) are primarily influenced by the isotopic composition of precipitation and isotopic fractionation associated with evapotranspiration (16). In the modern Maya Lowlands, δD_wax is well correlated with precipitation amount and varies by 60‰ across an annual precipitation gradient of 2,500 mm (Fig. 2). This modern variability in δD_wax is strongly influenced by soil water evaporation (17), and it is possible that changes in potential evapotranspiration could also impact paleo records. Accordingly, we interpret δD_wax values as qualitative records of water availability influenced by both precipitation amount and potential evapotranspiration. These two effects are complementary, since less rainfall and increased evapotranspiration would lead to both increased δD_wax values and reduced water resources, and vice versa.Open in a separate windowFig. 2.Scatter plot showing the negative relationship between annual precipitation and δD_wax-corr measured in modern lake sediment and soil samples (Fig. 1). Results from Lake Chichancanab (CH) and Salpeten (SP) are indicated. The black line indicates a linear regression fit to these data, with regression statistics reported at the bottom of the plot. Large squares indicate mean values for each sampling region, with error bars indicating SEM in both δD_wax-corr and annual precipitation. The black error bar indicates the 1σ error for δD_wax-corr values (SI Text). Original δD_wax data from ref. 17. VSMOW, Vienna Standard Mean Ocean Water.Plant wax carbon isotope signatures (δ¹³C_wax) in sediments from low-elevation tropical environments, including the Maya lowlands, are primarily controlled by the relative abundance of C₃ and C₄ plants (18–20). Ancient Maya land use was the dominant influence on the relative abundance of C₃ and C₄ plants during the late Holocene, because Maya farmers cleared C₃ plant-dominated forests and promoted C₄ grasses, in particular, maize (21–24). Thus, we apply δ¹³C_wax records as an indicator of the relative abundance of C₄ and C₃ plants that reflects past land use change (SI Text). Physiological differences between plant groups also result in differing δD_wax values between C₃ trees and shrubs and C₄ grasses (16), and we use δ¹³C_wax records to correct for the influence of vegetation change on δD_wax values (25) (δD_wax-corr, SI Text and Fig. S1).Plant waxes have been shown to have long residence times in soils in the Maya Lowlands (26). Therefore, age−depth models for our plant wax isotope records are based on compound-specific radiocarbon ages (Fig. 3), which align our δD_wax records temporally with nearby hydroclimate records derived from other methodologies (26) (SI Text and Fig. S2). The mean 95% confidence range for the compound-specific age−depth models is 230 y at Lake Chichancanab and 250 y at Lake Salpeten. Given these age uncertainties, we focus our interpretation on centennial-scale variability (26). The temporal resolution of our plant wax isotope records is lower than speleothem-derived climate records (8, 9), but combining plant wax records from multiple sites allows comparisons of climate change and land use in the northern and southern Maya Lowlands, which would otherwise not be possible. In addition, plant wax isotope records extend to the Early Preclassic/Late Archaic period (1500–2000 B.C.E.), providing a longer perspective on climate change in the Maya Lowlands than most other regional records (6, 8–11).Open in a separate windowFig. 3.Plant wax (green; left) and terrigenous macrofossil (red; right) age−depth models for (A) Lake Chichancanab and (B) Lake Salpeten. The age probability density of individual radiocarbon analyses is shown. The black lines indicate the best age model based on the weighted mean of 1,000 age model iterations (62). Colored envelopes indicate 95% confidence intervals. Cal, calendar.     

Mandibular and dental characteristics of Late Triassic mammaliaform Haramiyavia and their ramifications for basal mammal evolution

As one of the earliest-known mammaliaforms, Haramiyavia clemmenseni from the Rhaetic (Late Triassic) of East Greenland has held an important place in understanding the timing of the earliest radiation of the group. Reanalysis of the type specimen using high-resolution computed tomography (CT) has revealed new details, such as the presence of the dentary condyle of the mammalian jaw hinge and the postdentary trough for mandibular attachment of the middle ear—a transitional condition of the predecessors to crown Mammalia. Our tests of competing phylogenetic hypotheses with these new data show that Late Triassic haramiyids are a separate clade from multituberculate mammals and are excluded from the Mammalia. Consequently, hypotheses of a Late Triassic diversification of the Mammalia that depend on multituberculate affinities of haramiyidans are rejected. Scanning electron microscopy study of tooth-wear facets and kinematic functional simulation of occlusion with virtual 3D models from CT scans confirm that Haramiyavia had a major orthal occlusion with the tallest lingual cusp of the lower molars occluding into the lingual embrasure of the upper molars, followed by a short palinal movement along the cusp rows alternating between upper and lower molars. This movement differs from the minimal orthal but extensive palinal occlusal movement of multituberculate mammals, which previously were regarded as relatives of haramiyidans. The disparity of tooth morphology and the diversity of dental functions of haramiyids and their contemporary mammaliaforms suggest that dietary diversification is a major factor in the earliest mammaliaform evolution.Haramiyidans are among the first mammaliaforms to appear during the Late Triassic in the evolutionary transition from premammalian cynodonts. Their fossils have a cosmopolitan distribution during the Late Triassic to the Jurassic (1–8), tentatively with the youngest record in the Late Cretaceous of India (9). Most of these occurrences are of isolated teeth. For this reason, Haramiyavia clemmenseni (1) holds a special place in mammaliaform phylogeny: It is the best-preserved Late Triassic haramiyid with intact molars, nearly complete mandibles, and also postcranial skeletal elements (Figs. 1 and ​and22 and SI Appendix, Figs. S1–S4) (1). By its stratigraphic provenance from the Tait Bjerg Beds of the Fleming Fjord Formation, East Greenland (Norian-Rhaetic Age) (7), Haramiyavia is also the oldest known haramiyid (5, 7). Haramiyids, morganucodonts, and kuehneotheriids are the three earliest mammaliaform groups that are distinctive from each other in dental morphology and masticatory functions (10–12).Open in a separate windowFig. 1.(A and B) Composite reconstruction of Haramiyavia clemmenseni right mandible in lateral (A) and medial (B) views. Dark red: original bone with intact periosteal surface; brown: broken surface of preserved bone or remnant of bone; light blue: morphologies preserved in mold outlines or clear impression. (C) Morganucodon mandible in medial view.Open in a separate windowFig. 2.Molar features of Haramiyavia. (A) Right M1–M3 in occlusal view (medio-lateral orientation by the zygomatic root and the palate). (B) Occlusal facets of upper molars. (C) Lingual view of M1–M3. (D) Root structures of upper molars (M1 and M2 show three partially divided anterior roots connected by dentine and two posterior roots connected by dentine; M3 has two anterior and two posterior roots connected respectively by dentine). These roots have separate root canals. (E) Buccal view of M1–M3. All roots are bent posteriorly, suggesting that crowns shifted mesially, relative to the roots, during the tooth eruption, also known as mesial drift of teeth (arrowhead), typical of successive eruption of multirooted postcanines. (F) SEM photograph of lower m3 in a posterior occlusal view. (G) Approximate extent of wear facets by orthal occlusion (a1 cusp in embrasure of upper molars) (blue) and palinal movement of b2–b4 cusps sliding across the median furrow of upper molars (green). (H) Lingual view of m3. There are no wear facets on lingual side of cusps a1–a4. (I) Buccal view of m3 showing wear facets on the buccal sides of cusps b1–b4 and on apices.Haramiyids are characterized by their complex molars with longitudinal rows of multiple cusps. The cusp rows occlude alternately between the upper and lower molars. Primarily because of similarities in molar morphology, haramiyids are considered to be related to poorly known theroteinids of the Late Triassic (5, 13) and eleutherodontids of the Middle to Late Jurassic (14–17). Collectively haramiyids and eleutherodontids are referred to as “haramiyidans” (10, 14, 15, 18, 19). Recent discoveries of diverse eleutherodontids or eleutherodontid-related forms with skeletons from the Tiaojishan Formation (Middle to Late Jurassic) of China (18–20) have greatly augmented the fossil record of haramiyidans, ranking them among the most diverse mammaliaform clades of the Late Triassic and Jurassic.Historically, it has been a contentious issue whether haramiyidans (later expanded to include theroteinids and eleutherodontids) are closely related to the more derived multituberculates from the Middle Jurassic to Eocene (13) or represent a stem clade of mammaliaforms excluded from crown mammals (21, 22). The conflicting placement of haramiyidans was attributable in part to the uncertainties in interpreting the isolated teeth of most Late Triassic haramiyids (21, 22). More recent phylogenetic disagreements have resulted from different interpretations of mandibular characters in Haramiyavia (17–20, 23–25), which has not been fully described (figure 2 in ref. 1).Here we present a detailed study of the mandibles and teeth of Haramiyavia from the exhaustive documentation during initial fossil preparation (Fig. 1 and SI Appendix, Figs. S1–S4), from scanning electron microscopy (SEM) images, and from computed tomography (CT) scans and 3D image analyses of the two fossil slabs with mandibles (MCZ7/95A and B), plus a referred specimen of upper molars in a maxilla (MCZ10/G95) (Figs. 2 and ​and3,3, SI Appendix, Figs. S5–S8 and Tables S2 and S3, and Movie S1). These new data are informative for testing alternative mammaliaform phylogenies (Fig. 4 and SI Appendix) and are useful for reconstructing evolutionary patterns of feeding function in the earliest mammaliaforms.Open in a separate windowFig. 3.Molar occlusion of haramiyids. (A) In Haramiyavia the upper and lower molars form an en echelon pattern, a series of parallel and step-like occlusal surfaces in lingual and buccal views (based on 3D scaled models from CT scans of MCZ7/G95 and MCZ10/G95). (B–E) In Haramiyavia are shown the occlusal paths of cusps a1–a4 of the lingual row (B), cusps b2–b4 of the buccal row (shown with the lingual half of the tooth cut away) (C), and tooth orientation and the cut-away plane (D). During the orthal occlusion phase, the tallest lingual cusp a1 occludes into the embrasure of the preceding and the opposing upper molars (B and E), and the tallest buccal cusp b2 occludes into the upper furrow and behind the A1–B1 saddle of the upper molar (C). During the palinal occlusion phase, cusps b1–b4 of the buccal row slide posteriorly in the upper furrow, and in the upper row cusps B5–B1 slide in the lower furrow (lower molars with blue and green shading, superpositioned by flipped upper molar in clear outlines) (E). (F) Extent of wear on molars during the orthal phase (blue) and the palinal movement (green) produced by OFA simulation (Movie S1). (G–I) In Thomasia, reconstruction of upper and lower molar series on the basis of wear surfaces and tooth crown morphology (revised from refs. 2, 3, and 10). (G) The en echelon occlusal surfaces in lingual view. (H) Orthal occlusion (blue) is followed by palinal occlusal movement (green). (I) Occlusal wear facets of molars. Facets worn by orthal occlusion are shown in blue, and facets worn by palinal occlusion are shown in gray hatching. Cusp and facet designations are after refs 3 and 6.Open in a separate windowFig. 4.Hypotheses concerning the phylogenetic relationship of Haramiyavia and timing estimates of the basal diversification of crown mammals. (A) Haramiyavia is a close relative of multituberculates, both nested in the crown Mammalia. This hypothesis (haramiyidan node position 1) was based on a misinterpretation of a previous illustration of a fragment of the mandible (17, 18). (B) Haramiyavia is a stem mammaliaform, as determined by incorporating the features preserved on both mandibles into phylogenetic estimates (haramiyidan node position 2). (C) Placement of Haramiyavia and other haramiyidans among mammaliaforms according to this study. Many mandibular features were treated as unknown by studies favoring a Late Triassic diversification of mammals (18, 23). A more complete sampling of informative features revealed by this study now has overturned the previous placement. Clades: crown Mammalia (node a); Mammaliaformes (node b); haramiyidans (node 1 or 2, alternative positions); Eleutherodontida (node c). The rescored datasets and analyses are presented in SI Appendix.     

Early guenon from the late Miocene Baynunah Formation,Abu Dhabi,with implications for cercopithecoid biogeography and evolution

A newly discovered fossil monkey (AUH 1321) from the Baynunah Formation, Emirate of Abu Dhabi, United Arab Emirates, is important in a number of distinct ways. At ∼6.5–8.0 Ma, it represents the earliest known member of the primate subfamily Cercopithecinae found outside of Africa, and it may also be the earliest cercopithecine in the fossil record. In addition, the fossil appears to represent the earliest member of the cercopithecine tribe Cercopithecini (guenons) to be found anywhere, adding between 2 and 3.5 million y (∼50–70%) to the previous first-appearance datum of the crown guenon clade. It is the only guenon—fossil or extant—known outside the continent of Africa, and it is only the second fossil monkey specimen so far found in the whole of Arabia. This discovery suggests that identifiable crown guenons extend back into the Miocene epoch, thereby refuting hypotheses that they are a recent radiation first appearing in the Pliocene or Pleistocene. Finally, the new monkey is a member of a unique fauna that had dispersed from Africa and southern Asia into Arabia by this time, suggesting that the Arabian Peninsula was a potential filter for cross-continental faunal exchange. Thus, the presence of early cercopithecines on the Arabian Peninsula during the late Miocene reinforces the probability of a cercopithecoid dispersal route out of Africa through southwest Asia before Messinian dispersal routes over the Mediterranean Basin or Straits of Gibraltar.Cercopithecine monkeys (Order Primates, Superfamily Cercopithecoidea, Family Cercopithecidae, Subfamily Cercopithecinae), also known as cheek-pouch monkeys, are the most speciose and widely distributed group of living Old World primates. Recent molecular estimates date the divergence of Cercopithecinae from Colobinae (leaf-eating monkeys) to between 17.6 Ma (range 21.5–13.9 Ma) and 14.5 Ma (range 16.2–12.8 Ma) and the origin of crown Cercopithecinae to around 11.5 Ma (range 13.9–9.2 Ma) (1, 2). However, the earliest known fossil cercopithecines only appear much later, around 7.4 Ma in the Turkana Basin of East Africa (3, 4).Cercopithecine monkeys are divided into two tribes: Cercopithecini, including African guenons (Allenopithecus, Miopithecus, Chlorocebus, Erythrocebus, Allochrocebus, Cercopithecus), and Papionini, which includes African and Eurasian macaques (Macaca) as well as African papionins (Papio, Lophocebus, Rungwecebus, Theropithecus, Mandrillus, Cercocebus). Of the living cercopithecines, only two genera are known outside of the African continent, both of them papionins: Papio (found on the Arabian Peninsula) and Macaca (found throughout Southern and Southeast Asia, and introduced in Gibraltar). The earliest fossil cercopithecines known outside of Africa are attributed to the genus Macaca and appear to be latest Miocene or early Pliocene in age (∼6.0–5.0 Ma) (Fig. 1) (5–9). Until now, no guenons, extant or extinct, have ever been known outside of the African continent.Open in a separate windowFig. 1.Hypothesized cercopithecoid dispersal routes out of Africa in relation to the known late Miocene fossil record. The oldest cercopithecine, Parapapio lothagamensis (light blue circles), is known from ∼7.4–6.1 Ma in the Turkana Basin and Tugen Hills, Kenya (3, 4, 41). An unnamed fossil papionin (purple circle) is known from the late Miocene of Ongoliba, Congo (5, 57). Macaca spp. (dark blue circles) are located throughout North Africa at sites ranging in age from ∼6.5–5.5 Ma (5, 8, 58, 59), and Macaca spp. first appear in Europe ∼6.0–5.3 Ma and in China in the early Pliocene (5–9). The oldest colobine outside of Africa, Mesopithecus (green circles), is known from a number of late Miocene sites securely dated between ∼8.5 and 5.3 Ma in Greece, Macedonia, Italy, Ukraine, Iran, Afghanistan, possibly Pakistan, and China (46–48). Three dispersal routes for cercopithecoids can be hypothesized: route 1 imagines a dispersal event over the Straits of Gibraltar or Mediterranean Basin into Europe and Asia; route 2 postulates a dispersal event through the Arabian Sinai Peninsula; and route 3 suggests a migration over the Arabian Straits of Bab el Mandeb. The discovery of AUH 1321 and AUH 35 in Abu Dhabi at >6.5–8 Ma (red circle), contemporaneous with the first appearance of Mesopithecus sp. in Eurasia and ∼1–2 million y earlier than the appearance of Macaca spp. in Eurasia, suggests scenarios 2 and 3 were possible before scenario 1. None of these scenarios is mutually exclusive and may have occurred in combination or succession.Three possible routes can be reasonably hypothesized for cercopithecine (and cercopithecoid) dispersal out of Africa and into Europe and Asia during the late Miocene: (i) over the Mediterranean Basin or Straits of Gibraltar to the north/northwest, (ii) across the Arabian Sinai Peninsula to the northeast, or (iii) across the Arabian Straits of Bab el Mandeb to the east (Fig. 1). Fossil Macaca specimens from the terminal Miocene of Spain and Italy have been suggested to provide evidence for the use of a route across the Mediterranean Basin or the Straits of Gibraltar via an ephemeral land bridge either immediately before—or perhaps associated with—the drop in Mediterranean sea levels during the Messinian (∼6.0–5.3 Ma) (8–19). Paleontological evidence for an Arabian route has been lacking, but paleogeographic and paleoenvironmental work on circum-Arabia suggests that the region did not present a persistent ecological barrier to some amount of intercontinental exchange during the late Miocene (20). In fact, an established land connection through Sinai was probably present during this time period, and oceanic spreading is not estimated to have begun in the southern Red Sea until around 5 Ma, with progressive development of open marine conditions throughout the Pliocene. Thus, before 6.5 Ma, a southern route in the region of the Straits of Bab el Mandeb was also possible (Fig. 1) (21).Although Arabia is a large area of the earth, fossil monkeys have so far been represented by only a single specimen, an isolated male lower canine (AUH 35), discovered in 1989 by A.H. and Peter Whybrow in the late Miocene Baynunah Formation, Abu Dhabi, United Arab Emirates (22–25). The specimen came from Jebel Dhanna, site JDH-3 (JD-3 in refs. 24 and 26) (Fig. 2), a locality now lost to industrial development. Because male cercopithecid lower canines are not metrically identifiable beyond the Family level of classification (23), AUH 35 was described as a cercopithecid with indeterminate affinities. Here we report the discovery of a second monkey specimen from the Baynunah Formation in Abu Dhabi (AUH 1321), found almost 20 y after the first. AUH 1321 clearly represents a cercopithecine and, because it is dated to between 6.5 and 8.0 Ma, it is the oldest cercopithecine yet known outside of Africa and possibly the oldest cercopithecine in the fossil record. Thus, the discovery of AUH 1321 provides the earliest paleontological evidence of cercopithecine dispersal out of continental Africa and possibly hints at an Arabian cercopithecoid dispersal route into Eurasia during the Late Miocene (Fig. 1). Furthermore, we believe AUH 1321 can be attributed to the Cercopithecini (guenons) and, therefore, it represents the only record of this tribe, living or fossil, yet known outside of Africa.Open in a separate windowFig. 2.Map illustrating the location of the two fossil sites in the Baynunah Formation that have produced fossil monkeys. Top Right Inset shows the location of the SHU 2–2 excavation (kite aerial photography by Nathan Craig).     

Entropic factors provide unusual reactivity and selectivity in epoxide-opening reactions promoted by water

Despite the myriad of selective enzymatic reactions that occur in water, chemists have rarely capitalized on the unique properties of this medium to govern selectivity in reactions. Here we report detailed mechanistic investigations of a water-promoted reaction that displays high selectivity for what is generally a disfavored product. A combination of structural and kinetic data indicates not only that synergy between substrate and water suppresses undesired pathways but also that water promotes the desired pathway by stabilizing charge in the transition state, facilitating proton transfer, doubly activating the substrate for reaction, and perhaps most remarkably, reorganizing the substrate into a reactive conformation that leads to the observed product. This approach serves as an outline for a general strategy of exploiting solvent-solute interactions to achieve unusual reactivity in chemical reactions. These findings may also have implications in the biosynthesis of the ladder polyether natural products, such as the brevetoxins and ciguatoxins.Given its simple structure and low molecular weight, water is a remarkably complex substance. Several landmark investigations (1) have revealed the ability of water to self-assemble to form sophisticated dynamic hydrogen bond networks, which accounts for its unusual properties, such as its high boiling point and high surface tension (2). Despite the unique properties that it offers and that nature has exploited, water is generally eschewed in synthetic chemistry largely because of chemical incompatibility with many commonly used reagents and the low aqueous solubility of many organic molecules coupled with the attendant assumption that homogeneity is required for reactivity. Although several important examples of remarkable reactivity have been documented for reactions carried out in water (3–5), on the surface of water (6, 7), or in micelles suspended in water (8, 9), utilization of the “biological solvent” in organic reactions remains uncommon.We became acutely aware of the remarkable properties of water when we discovered that neutral aqueous solutions of epoxy alcohol 1a underwent a spontaneous and selective “endo” cyclization reaction to form 6,6-fused bicyclic product 2a (Fig. 1) (10–12). Exceptional reactivity and selectivity were observed only when two criteria were satisfied: The substrate contained a six-membered tetrahydropyran ring, and the solvent used was water at pH 7.0. These surprising results were counter to a set of general empirical rules put forth by Baldwin for cyclization reactions, which state that the smaller ring product resulting from what is commonly called an exo cyclization is favored for similar reactions (13). The tetrahydropyranol ring, therefore, appeared to “template” the endo cyclization pathway to form the larger ring product with water playing a critical yet heretofore unknown role.Open in a separate windowFig. 1.Endo selective cyclization of epoxy alcohols templated by a tetrahydropyran ring(s) and promoted by water.Moreover, 2a is a substructure found in a large family of natural products commonly referred to as the ladder polyethers (e.g., brevetoxin B, S) (14–16). These constituents of harmful algal blooms (a.k.a., red tide) have attracted significant attention (17) because of the remarkable structural regularities that temper their apparent complexity. Several biogenetic pathways for the construction of rings of these natural products have been proposed and generally involve a cascade of cyclizations by way of epoxide-opening reactions (Fig. 2) (18–23). However, this hypothesis requires that all of the epoxide ring-opening events proceed with atypical endo regioselectivity.Open in a separate windowFig. 2.Biosynthetic proposal for the formation of the ladder polyether natural product brevetoxin B, a potent neurotoxin.Our discovery that reactions of 1a in water preferentially lead to the endo product 2a rather than exo product 3a provided a possible means to overcome this obstacle. Further support was provided when we demonstrated that cascade reactions akin to those proposed for the biosynthesis of the natural products was possible with the selective endo cyclization of di- and triepoxide analogs, such as 4 and 6, respectively (Fig. 1). Once again selective reactions were only observed when reactions were carried out in neutral water.These striking results and their possible relevance to the biogenesis of the ladder polyether natural products prompted us to investigate the mechanism of the cyclization reaction. Here we report the culmination of these efforts. They reveal an intimate connection between water and the tetrahydropyran ring (template) of 1a, provide support for the feasibility of the biosynthetic proposal for the ladder polyether natural products, and demonstrate the importance of solvent–substrate interactions for promoting selectivity. We believe that this last concept may have broad implications and a different means for chemists to attain unusual selectivity for a variety of chemical reactions.     

Late Miocene episodic lakes in the arid Tarim Basin,western China

The Tibetan Plateau uplift and Cenozoic global cooling are thought to induce enhanced aridification in the Asian interior. Although the onset of Asian desertification is proposed to have started in the earliest Miocene, prevailing desert environment in the Tarim Basin, currently providing much of the Asian eolian dust sources, is only a geologically recent phenomenon. Here we report episodic occurrences of lacustrine environments during the Late Miocene and investigate how the episodic lakes vanished in the basin. Our oxygen isotopic (δ¹⁸O) record demonstrates that before the prevailing desert environment, episodic changes frequently alternating between lacustrine and fluvial-eolian environments can be linked to orbital variations. Wetter lacustrine phases generally corresponded to periods of high eccentricity and possibly high obliquity, and vice versa, suggesting a temperature control on the regional moisture level on orbital timescales. Boron isotopic (δ¹¹B) and δ¹⁸O records, together with other geochemical indicators, consistently show that the episodic lakes finally dried up at ∼4.9 million years ago (Ma), permanently and irreversibly. Although the episodic occurrences of lakes appear to be linked to orbitally induced global climatic changes, the plateau (Tibetan, Pamir, and Tianshan) uplift was primarily responsible for the final vanishing of the episodic lakes in the Tarim Basin, occurring at a relatively warm, stable climate period.Once part of the Neo-Tethys Sea, indicated by the Paleogene littoral-neritic deposits with intercalated marine strata (1), the Tarim Basin in western China (Fig. 1) has experienced dramatic environmental and depositional changes during the Cenozoic. The eventual separation of the basin from the remnant sea probably occurred during the middle to late Eocene (1, 2), and since then, terrestrial sedimentation and environment have prevailed. Today, the basin has relatively flat topography with elevation ranging between 800 and 1,300 m above sea level (asl), surrounded by high mountain ranges with average elevation exceeding 4,000 m. Hyperarid climate prevails with mean annual precipitation <50 mm and evaporation ∼3,000 mm. Active sand dunes occupy 80% of the basin, forming the Taklimakan Desert, the largest desert in China and the second largest in the world (3). Thick desert deposits in the basin also provide a major source for dust storms occurring in East Asia (4).Open in a separate windowFig. 1.Digital elevation model of the Tarim Basin and surrounding mountain ranges. The studied 1,050-m sediment core is retrieved from Lop Nor (1), at a relatively low elevation (∼800 m asl) in the eastern basin. Also indicated are locations of previous studied sections, near Sanju (2), Kuqa (3), and Korla (4) in the basin.The Tibetan Plateau uplift, long-term global cooling, and the associated retreat of the remnant sea during the Cenozoic, through their complex interplay, may have all contributed to the enhanced aridification in the Asian interior and eventually the desert formation (2, 5–11). Although the onset of Asian desertification is proposed to have started in the earliest Miocene (12, 13), prevailing desert environment in the Tarim Basin, currently providing much of the Asian eolian dust sources (4), is only a geologically recent phenomenon (9). Previous lithological and pollen studies (9, 14, 15) suggest that the currently prevailing desert environment started probably at ∼5 Ma. However, what kind of environmental conditions prevailed before that and how the dramatic changes occurred largely remain elusive.To better decipher the aridification history, we used a 1,050-m-long, continuous sediment core retrieved from Lop Nor (39°46''0''''N, 88°23''19''''E) in the eastern basin (Fig. S1). The core site has a relatively low elevation (∼800 m asl) (Fig. 1) and thus effectively records basin environment. Previous study sites came from elevated basin margins (9, 14). The core mainly consists of lacustrine sediments with associated fluvial-eolian sands (Fig. 2). The core chronology was established previously based on paleomagnetic polarity (15). The total 706 remanence measurements on the continuous sediment profile allow straightforward correlation with the CK95 geomagnetic polarity timescale (16) and identification of 14 normal (N1–N14) and 13 reversal (R1–R13) polarity zones over the last 7.1 Ma (Fig. S1). The CK95 timescale is largely consistent (within a few ky) with that inferred from marine archives (17) over the last 5.23 Ma, and before 5.23 Ma a practical measure of chronological uncertainty was estimated to be within ∼100 ky (18), assuring that the terrestrial records can be directly compared with marine records and orbital changes at the timescale of >100 ky. The derived chronology yields an average sedimentation rate of ∼200 m/Ma at lower sections (>1.77 Ma), allowing high-resolution studies of the desertification history at the critical interval.Open in a separate windowFig. 2.Records of δ¹¹B, δ¹⁸O, TOC, and CaCO₃ changes in the Lop Nor profile. Representative photos show lithological changes mainly from lacustrine bluish gray argillaceous limestone to fluvial-eolian brown/red clayey silt, as shown in the lithological column with visual colors indicated (15). High levels and large fluctuations of proxy values occurred only before ∼4.9 Ma, and since then, those remain low and within a small range, largely similar to modern conditions. The transition from episodic occurrences of lacustrine phases to prevailing desert environments thus appears to be permanent and irreversible.Boron and oxygen isotopes from carbonates are important environmental indicators (19–25) and used here to infer the environmental evolution in the Tarim Basin. We also present total organic carbon (TOC), calcium carbonate (CaCO₃), ostracod, and grain size records to substantiate the isotopic evidence (SI Materials and Methods). Our δ¹¹B profile shows substantial, stepwise changes over the last ∼7.1 Ma (Fig. 2). Carbonate δ¹¹B values were −5.0 ± 1.6‰ (n = 3) before 6.0 Ma, increased rapidly to ∼11‰ at 4.9–6.0 Ma, and then stayed at roughly the same level (10.7 ± 2.2‰, n = 25) for the remaining 4.9 Ma. Higher-resolution δ¹⁸O, TOC, and CaCO₃ profiles generally confirm the pattern observed in the low-resolution δ¹¹B one (Fig. 2). The δ¹⁸O values remained low, ranging from −10‰ to −4‰ over the last 4.9 Ma. However, δ¹⁸O values frequently oscillated between −10‰ and 5‰ before that. Similarly, the TOC profile shows consistently low organic carbon content (0–0.2%) after 4.9 Ma and large fluctuations (0–1.0%) before then. The CaCO₃ profile also indicates consistently low values (0–25%) after 4.9 Ma and large fluctuations (0–50%) earlier (Fig. 2).The multiple proxy records strongly suggest that critical environmental changes must have occurred at ∼4.9 Ma. δ¹¹B values of carbonates from marine sources differ substantially from those of nonmarine carbonates (19–21). δ¹¹B values after 4.9 Ma are close to those from marine carbonates, but values before 6 Ma fall into the range of lacustrine carbonates (22). Positive δ¹⁸O values before 4.9 Ma also indicate lacustrine environments at that time. Carbonates from modern lakes in arid and semiarid regions of northwestern China show similar positive δ¹⁸O values (23), due to strong evaporation processes. High TOC and CaCO₃ contents (Fig. 2) further support that lacustrine environments existed in the basin before ∼4.9 Ma. δ¹⁸O values after 4.9 Ma are comparable to those in Cenozoic soil carbonates (24) and ancient marine carbonates in the Tarim Basin (25). However, the accompanying carbonate δ¹³C values throughout the record, ranging from −4‰ to 1‰ (Dataset S1), are significantly higher than those from Cenozoic soil carbonates reported (26), essentially ruling out the possibility of soil carbonate source. Using modern prevailing desert environment in the basin as an analog, the combined δ¹¹B and δ¹⁸O evidence thus suggests that the sediment deposits in the basin after 4.9 Ma must be eolian-fluvial in origin and their sources, at least carbonate grains, came from weathered ancient marine carbonates in nearby regions.Sedimentological and stratigraphic patterns in other exposed sections from different parts of the basin (9, 14) share great similarity with the Lop Nor core profile (Fig. S1). Episodic lacustrine mudstones and/or siltstones during the Late Miocene were present in all sections and were replaced by fluvial-eolian deposits later. Studies of ostracod assemblages (27) also suggest a shallow paleolake with brackish water environments in the northern basin during the Late Miocene. Changes in the depositional environment from our Lop Nor profile alone could be plausibly explained by a shift in basin center due to tectonic compressions, as evidenced from the slightly uplifted central basin (Fig. 1). However, similar temporal changes occurring basin-wide at ∼4.9 Ma argue against it. Instead, our results, together with previous studies (5, 14, 15, 27), suggest that paleolakes were widely present in the low lands of the basin during the Late Miocene, much different from currently prevailing desert environments with a few scattered small lakes. The existing evidence, although still limited (Fig. 1), would point to the occurrence of a possible megalake in the Tarim Basin during the Late Miocene.Three high-resolution records, δ¹⁸O, TOC, and CaCO₃, further suggest that lacustrine environments before ∼4.9 Ma were not permanent (Fig. 2). These large fluctuations indicate frequent switches between lacustrine and fluvial-eolian environments in the basin. High proxy values, δ¹⁸O in particular, appear to indicate lacustrine environments, whereas low values, similar to ones after 4.9 Ma, correspond to fluvial-eolian deposits. This is consistent with lithological features at this interval, showing argillaceous limestone intercalated with clayey layers (15), the occurrence of ostracod assemblages (Fig. 3) from lacustrine sediments, grain size changes (Fig. S2), and detrital carbonate grains identified in photomicrographs of fluvial-eolian deposits (Fig. S3).Open in a separate windowFig. 3.δ¹⁸O fluctuations linked to eccentricity and obliquity orbital variations at 4.5–7.1 Ma. Lacustrine phases (high δ¹⁸O) generally correspond to periods of high eccentricity and obliquity. Fluvial-eolian environment (low δ¹⁸O, highlighted with gray bars), developed more around 6.5, 6.1, 5.7, and 5.2 Ma, at a ∼400-ky eccentricity beat. The red bar indicates the last occurrence of lacustrine environment at ∼4.9 Ma. Occurrences of ostracod assemblages with total number >100 are also indicated.To further investigate such episodic changes, we performed spectral analysis on the δ¹⁸O record over the interval 4.5–7.1 Ma. Strong spectral power at orbital frequencies were identified, with periods of ∼400 ky throughout the interval, ∼41 ky particularly at 6–6.5 Ma, and ∼100 ky at 5–6 Ma and 6.5–7.1 Ma (Fig. S4). Precessional ∼20-ky power might also have existed but was relatively weak and discontinuous. Lacustrine phase as indicated by high δ¹⁸O values and ostracod assemblages generally occurred at periods of high eccentricity and obliquity (28) (Fig. 3). At 6–6.5 Ma, δ¹⁸O shows clear correspondence to orbital obliquity variation (Fig. 3A). Additionally, the number of low δ¹⁸O values (fluvial-eolian environment) occurred more around 6.5, 6.1, 5.6, and 5.2 Ma, at a ∼400-ky beat following orbital eccentricity variation (Fig. 3B). The cluster of high δ¹⁸O values (>0‰) at 4.9–5.0 Ma signals the last occurrence of lacustrine environments in the basin. Our orbital association thus allows us to precisely determine the timing of desert formation at ∼4.9 Ma, ∼400 ky (an eccentricity cycle) later than the age inferred from basin margins (9, 14) and yet all occurring at eccentricity minima (Fig. 3B). As high eccentricity and obliquity generally correspond to warm conditions at orbital timescales, lacustrine (wet) phase could be associated with warm periods, consistent with the notion that cooler conditions would reduce moisture in the atmosphere and enhance continental drying (10, 11). We recognize that the chronological uncertainty from the geomagnetic polarity timescale before 5.23 Ma, within ∼100 ky (18), could confound our association of wet phase with high obliquity, although it is unlikely affected at the 400-ky eccentricity beat. However, the opposite association, wet phase with low obliquity, and the combination with high eccentricity, would require a different, yet unknown mechanism that is inconsistent with the orbital theory of Pleistocene ice ages.Superimposed on the orbitally episodic changes, the δ¹⁸O record also shows a long-term trend of deteriorating lacustrine conditions at 4.9–7.1 Ma. As δ¹⁸O values indicate two depositional environments, lacustrine and fluvial-eolian, the range of δ¹⁸O changes (between −10‰ and 5‰) does not vary much over this period (Fig. 3). Rather, the duration of high δ¹⁸O vs. low values would reflect the long-term trend. The mean δ¹⁸O values over 40-ky and 400-ky intervals both show a decreasing trend, with dominant lacustrine phase before 6.1 Ma, more developed fluvial-eolian environment at 5.7–6.1 Ma, a return to slightly better lacustrine environment at 5.3–5.7 Ma, and lacustrine phase permanently vanished around 4.9 Ma (Fig. 4).Open in a separate windowFig. 4.Long-term δ¹⁸O changes compared with global climatic conditions and regional tectonic activities. Global benthic δ¹⁸O (29) and northwestern Pacific SST (30) records show minimal long-term changes at 4–7 Ma, whereas occurrences of detrital apatite fission track ages (35) in northern western Kunlun peaked, and the Tarim episodic lakes gradually vanished. The dark thick lines are their 400-ky running means, and the light green line on δ¹⁸O represents the 40-ky running mean.The gradual disappearance of the Tarim episodic lakes could be potentially explained by the two driving forces, plateau uplift (the Tibetan Plateau, Pamir Plateau, and Tianshan) and long-term global cooling. However, the long-term global climate was relatively warm and stable during this period (Fig. 4). The global benthic δ¹⁸O record (29) shows that much of the Miocene cooling occurred between 15 and 11 Ma, and the cooling between 8 and 5 Ma was minimal. Supporting this view, sea surface temperature records from the northwestern Pacific (30) show that almost no cooling occurred between 6 and 4 Ma (perhaps further to 3 Ma) (Fig. 4). Particularly, the mean global climate (29, 30) was even warmer at 4.1–4.5 Ma than at 5.7–6.1 Ma, whereas lacustrine phase permanently disappeared after ∼4.9 Ma (Fig. 4), indicating decoupling of the lake evolution from global climate. Therefore, long-term global cooling might have played a subordinate role in the lake disappearance.Instead, the growth of surrounding mountain ranges (Tibetan, Pamir, and Tianshan) may have blocked moisture from the west and south, changed air circulations, and eventually led to the permanent lake disappearance within the basin. Today, the Tarim Basin receives limited moisture from westerlies (31) through the Pamir and Tianshan Ranges (and perhaps from the Indian Ocean in summertime as well). Although the Indo–Eurasian convergence since the Late Eocene resulted in high elevations of the Tibetan Plateau and, to a lesser degree, surrounding mountains including Pamir and Tianshan by the mid-Miocene time (8), tectonic activities in broad areas around the Tarim Basin appear to be rejuvenated since the Late Miocene. Tectonic deformations during the Late Miocene–Early Pliocene inferred from growth strata, sedimentary facies changes, and low-temperature thermochronologic studies occurred in Tianshan to the north of the Tarim Basin, in the Kunlun Mountains to its south and the Pamir to its west (32–35). Syntectonic growth strata from the foreland basins of the Kunlun and Tianshan Ranges (32) show that strong crust shortening and potential mountain uplift initiated ∼6.5–5 Ma and lasted to the Early Pleistocene (Fig. S5), similarly reported in northern Pamir (33). Cenozoic sequences in the Pamir–Tianshan convergence zone, changing from an arid continental plain to an intermountain basin by ∼5 Ma, support surface uplift of the west margin of the Tarim Basin (34). Occurrence of detrital apatite fission track ages from West Kunlun Ranges (35) also peaked at ∼4.5 Ma (Fig. 4). Thus, the reactivated uplift of Pamir and West Kunlun Ranges and northward movement of Pamir during the Late Miocene–Pliocene would progressively block moisture into the basin and enhance regional aridity to a certain threshold to terminate lacustrine environments even during warm periods with favorable orbital configuration. Although inconsistencies indeed exist in linking regional climatic and environmental changes to tectonic events (8), the final vanishing of the Tarim episodic lakes is better explained by tectonic factors.Therefore, our multiple-proxy results consistently show that the Taklimakan sand sea began to form at ∼4.9 Ma and that the transition into prevailing desert environment was permanent and irreversible. The episodic occurrences of lacustrine environments at favorable climatic conditions (warm periods) during the Late Miocene suggest that uplifted mountain ranges then were not high enough to effectively block moisture from being transported into the basin. The transition from episodic lacustrine environments to prevailing desert deposits was gradual, from 7.1 Ma (or earlier beyond our record) to 4.9 Ma, for which we suggest that although long-term global cooling enhanced the overall aridification in the Asian interior during the Late Cenozoic (10, 11), plateau uplift played a more important role in finally drying up the episodic lakes within a relatively warm, stable climate period, thus decoupling regional climate temporally from a global trend. Our high-resolution records thus demonstrate that regional climate in the Tarim Basin reached a critical state in the Late Miocene, with the dual effects from global climate conditions and regional tectonic settings then. With global climate remaining relatively stable and warm entering the Pliocene, by ∼3–4 Ma (Fig. 4), drying up of episodic lakes at ∼4.9 Ma could thus be largely attributed to rejuvenated tectonic activities.     

Little evidence for enhanced phenotypic evolution in early teleosts relative to their living fossil sister group

Since Darwin, biologists have been struck by the extraordinary diversity of teleost fishes, particularly in contrast to their closest “living fossil” holostean relatives. Hypothesized drivers of teleost success include innovations in jaw mechanics, reproductive biology and, particularly at present, genomic architecture, yet all scenarios presuppose enhanced phenotypic diversification in teleosts. We test this key assumption by quantifying evolutionary rate and capacity for innovation in size and shape for the first 160 million y (Permian–Early Cretaceous) of evolution in neopterygian fishes (the more extensive clade containing teleosts and holosteans). We find that early teleosts do not show enhanced phenotypic evolution relative to holosteans. Instead, holostean rates and innovation often match or can even exceed those of stem-, crown-, and total-group teleosts, belying the living fossil reputation of their extant representatives. In addition, we find some evidence for heterogeneity within the teleost lineage. Although stem teleosts excel at discovering new body shapes, early crown-group taxa commonly display higher rates of shape evolution. However, the latter reflects low rates of shape evolution in stem teleosts relative to all other neopterygian taxa, rather than an exceptional feature of early crown teleosts. These results complement those emerging from studies of both extant teleosts as a whole and their sublineages, which generally fail to detect an association between genome duplication and significant shifts in rates of lineage diversification.Numbering ∼29,000 species, teleost fishes account for half of modern vertebrate richness. In contrast, their holostean sister group, consisting of gars and the bowfin, represents a mere eight species restricted to the freshwaters of eastern North America (1). This stark contrast between teleosts and Darwin''s original “living fossils” (2) provides the basis for assertions of teleost evolutionary superiority that are central to textbook scenarios (3, 4). Classic explanations for teleost success include key innovations in feeding (3, 5) (e.g., protrusible jaws and pharyngeal jaws) and reproduction (6, 7). More recent work implicates the duplicate genomes of teleosts (8–10) as the driver of their prolific phenotypic diversification (8, 11–13), concordant with the more general hypothesis that increased morphological complexity and innovation is an expected consequence of genome duplication (14, 15).Most arguments for enhanced phenotypic evolution in teleosts have been asserted rather than demonstrated (8, 11, 12, 15, 16; but see ref. 17), and draw heavily on the snapshot of taxonomic and phenotypic imbalance apparent between living holosteans and teleosts. The fossil record challenges this neontological narrative by revealing the remarkable taxonomic richness and morphological diversity of extinct holosteans (Fig. 1) (18, 19) and highlights geological intervals when holostean taxonomic richness exceeded that of teleosts (20). This paleontological view has an extensive pedigree. Darwin (2) invoked a long interval of cryptic teleost evolution preceding the late Mesozoic diversification of the modern radiation, a view subsequently supported by the implicit (18) or explicit (19) association of Triassic–Jurassic species previously recognized as “holostean ganoids” with the base of teleost phylogeny. This perspective became enshrined in mid-20th century treatments of actinopterygian evolution, which recognized an early-mid Mesozoic phase dominated by holosteans sensu lato and a later interval, extending to the modern day, dominated by teleosts (4, 20, 21). Contemporary paleontological accounts echo the classic interpretation of modest teleost origins (22–24), despite a systematic framework that substantially revises the classifications upon which older scenarios were based (22–25). Identification of explosive lineage diversification in nested teleost subclades like otophysans and percomorphs, rather than across the group as a whole, provides some circumstantial neontological support for this narrative (26).Open in a separate windowFig. 1.Phenotypic variation in early crown neopterygians. (A) Total-group holosteans. (B) Stem-group teleosts. (C) Crown-group teleosts. Taxa illustrated to scale.In contrast to quantified taxonomic patterns (20, 23, 24, 27), phenotypic evolution in early neopterygians has only been discussed in qualitative terms. The implicit paleontological model of morphological conservatism among early teleosts contrasts with the observation that clades aligned with the teleost stem lineage include some of the most divergent early neopterygians in terms of both size and shape (Fig. 1) (see, for example, refs. 28 and 29). These discrepancies point to considerable ambiguity in initial patterns of phenotypic diversification that lead to a striking contrast in the vertebrate tree of life, and underpins one of the most successful radiations of backboned animals.Here we tackle this uncertainty by quantifying rates of phenotypic evolution and capacity for evolutionary innovation for the first 160 million y of the crown neopterygian radiation. This late Permian (Wuchiapingian, ca. 260 Ma) to Cretaceous (Albian, ca. 100 Ma) sampling interval permits incorporation of diverse fossil holosteans and stem teleosts alongside early diverging crown teleost taxa (Figs. 1 and ​and2A2A and Figs. S1 and ​andS2),S2), resulting in a dataset of 483 nominal species-level lineages roughly divided between the holostean and teleost total groups (Fig. 2B and Fig. S2). Although genera are widely used as the currency in paleobiological studies of fossil fishes (30; but see ref. 31), we sampled at the species level to circumvent problems associated with representing geological age and morphology for multiple congeneric lineages. We gathered size [both log-transformed standard length (SL) and centroid size (CS); results from both are highly comparable (Figs. S3 and ​andS4);S4); SL results are reported in the main text] and shape data (the first three morphospace axes arising from a geometric morphometric analysis) (Fig. 2A and Figs. S1) from species where possible. To place these data within a phylogenetic context, we assembled a supertree based on published hypotheses of relationships. We assigned branch durations to a collection of trees under two scenarios for the timescale of neopterygian diversification based on molecular clock and paleontological estimates. Together, these scenarios bracket a range of plausible evolutionary timelines for this radiation (Fig. 2B). We used the samples of trees in conjunction with our morphological datasets to test for contrasts in rates of, and capacity for, phenotypic change between different partitions of the neopterygian Tree of Life (crown-, total-, and stem-group teleosts, total-group holosteans, and neopterygians minus crown-group teleosts), and the sensitivity of these conclusions to uncertainty in both relationships and evolutionary timescale. Critically, these include comparisons of phenotypic evolution in early crown-group teleosts—those species that are known with certainty to possess duplicate genomes—with rates in taxa characterized largely (neopterygians minus crown teleosts) or exclusively (holosteans) by unduplicated genomes. By restricting our scope to early diverging crown teleost lineages, we avoid potentially confounding signals from highly nested radiations that substantially postdate both genome duplication and the origin of crown teleosts (26, 32). This approach provides a test of widely held assumptions about the nature of morphological evolution in teleosts and their holostean sister lineage.Open in a separate windowFig. 2.(A) Morphospace of Permian–Early Cretaceous crown Neopterygii. (B) One supertree subjected to our paleontological (Upper) and molecular (Lower) timescaling procedures to illustrate contrasts in the range of evolutionary timescales considered. Colors of points (A) and branches (B) indicate membership in major partitions of neopterygian phylogeny. Topologies are given in Datasets S4 and S5. See Dataset S6 for source trees.Open in a separate windowFig. S1.Morphospace of 398 Permian–Early Cretaceous Neopterygii. Three major axes of shape variation are presented. Silhouettes and accompanying arrows illustrate the main anatomical correlates of these principal axes, as described in Open in a separate windowFig. S2.Morphospace of 398 Permian–Early Cretaceous Neopterygii, illustrating the major clades of (A) teleosts and (B) holosteans.Open in a separate windowFig. S3.Comparisons of size rates between (A) holosteans and teleosts, (B) crown teleosts and all other neopterygians, (C) crown teleosts and stem teleosts, (D) crown teleosts and holosteans, and (E) stem teleosts and holosteans. Comparisons were made using the full-size SL dataset, a CS dataset, and a smaller SL dataset pruned to exactly match the taxon sampling of the CS dataset. Identical taxon sampling leads the CS and pruned SL datasets to yield near identical results. Although the larger SL dataset results often differ slightly, the overall conclusion from each pairwise comparison (i.e., which outcome is the most likely in an overall majority of trees) is identical in all but one comparison (E, under molecular timescales).Open in a separate windowFig. S4.Comparisons of size innovation between (A) holosteans and teleosts, (B) crown teleosts and all other neopterygians, (C) crown teleosts and stem teleosts, (D) crown teleosts and holosteans, and (E) stem teleosts and holosteans. Comparisons were made using the full-size SL dataset, a CS dataset, and a smaller SL dataset pruned to exactly match the taxon sampling of the CS dataset. Comparisons of size innovation are presented for K value distributions of the three datasets resemble each other closely.     

Cleavage and polyadenylation specificity factor 30: An RNA-binding zinc-finger protein with an unexpected 2Fe–2S cluster

Cleavage and polyadenylation specificity factor 30 (CPSF30) is a key protein involved in pre-mRNA processing. CPSF30 contains five Cys₃His domains (annotated as “zinc-finger” domains). Using inductively coupled plasma mass spectrometry, X-ray absorption spectroscopy, and UV-visible spectroscopy, we report that CPSF30 is isolated with iron, in addition to zinc. Iron is present in CPSF30 as a 2Fe–2S cluster and uses one of the Cys₃His domains; 2Fe–2S clusters with a Cys₃His ligand set are rare and notably have also been identified in MitoNEET, a protein that was also annotated as a zinc finger. These findings support a role for iron in some zinc-finger proteins. Using electrophoretic mobility shift assays and fluorescence anisotropy, we report that CPSF30 selectively recognizes the AU-rich hexamer (AAUAAA) sequence present in pre-mRNA, providing the first molecular-based evidence to our knowledge for CPSF30/RNA binding. Removal of zinc, or both zinc and iron, abrogates binding, whereas removal of just iron significantly lessens binding. From these data we propose a model for RNA recognition that involves a metal-dependent cooperative binding mechanism.Zinc-finger proteins (ZFs) are a large class of proteins that use zinc as structural cofactors (1–4). ZFs perform a variety of functions ranging from the modulation of gene expression through specific interactions with DNA or RNA to the control of signaling pathways via protein–protein interactions. The general feature that defines a ZF protein is the presence of one or more domains that contain a combination of four cysteine and/or histidine residues that serve as ligands for zinc. When zinc binds to these ligands, the domain adopts the structure necessary for function (1–4).ZFs are typically identified by the presence of cysteine and histidine residues in regular repeats and are categorized into classes based upon the number of cysteine and histidine residues, and the spacing between the residues (1, 2). At least 14 distinct classes of ZFs have been identified to date. ZFs are highly abundant, with more than 3% of the proteins in the human genome annotated as ZFs, based upon their sequences (1, 5–9). In some cases, there are considerable in vitro and in vivo data that support the annotation of proteins as ZFs, whereas in other cases the only evidence that a protein is a ZF comes from its amino acid sequence. The best-studied class of ZFs comprises the “classical” ZFs. This class is composed of ZFs that contain a Cys₂His₂ domain (CysX_2–5CysX_12–13HisX_3–5His). Classical ZFs adopt an alpha-helical/antiparallel beta-sheet structure when zinc is coordinated and bind DNA in a sequence-specific manner (2, 4). The remaining classes of ZFs are collectively called “nonclassical” ZFs (1). One class of nonclassical ZFs is the Cys₃His class (CysX_7–9CysX_4–6CysX₃His). The first protein of this class to be identified was tristetraprolin, which contains two Cys₃His domains and regulates cytokine mRNAs via a specific ZF domain/RNA binding interaction (1). With the publication of genome sequences this domain has been found in a myriad of proteins. The National Center for Biotechnology Information (NCBI) conserved domain architecture tool identifies 404 distinct proteins (both hypothetical and experimentally validated) that contain this domain, and humans contain at least 60 (Fig. 1). As a class, these proteins are predicted to be involved in RNA regulation; however, the function(s) of most of these proteins have not yet been established (1, 2, 10, 11).Open in a separate windowFig. 1.Survey of the CCCH domain containing proteins in H. sapiens.One important Cys₃His ZF protein is cleavage and polyadenylation specificity factor 30 (CPSF30) (2, 12). CPSF30 contains five Cys₃His domains. CPSF30 is part of a complex of proteins, collectively called CPSF, that are involved in the polyadenylation step of pre-mRNA processing (16). The other members of CPSF are CPSF160, CPSF73, CPSF100, Fip1, and Wdr33 (16). Polyadenylation is a 3′ end maturation step that all eukaryotic mRNAs (except histones) undergo (Fig. 2) (12). It involves endonucleolytic cleavage of the pre-mRNA followed by addition of a polyadenosine tail. Polyadenylation occurs at a specific region of the pre-mRNA called the polyadenylation cleavage site (PAS). The PAS consists of an upstream element with the conserved sequence AAUAAA (also called the AU-hexamer), a stretch of bases where cleavage occurs, after which a conserved GU-rich or U-rich sequence is present (usually between 40–60 nt after the cleavage site) (12, 13). CPSF73 is the endonuclease that cleaves the RNA; the roles of the other CPSF proteins are less clear (12, 13). Initially, CPSF160 was identified as the protein within the CPSF complex that recognizes the AU-hexamer (14–16); however, two recent studies using cell-based methods found that CPSF160 does not play this role (17, 18). Instead, CPSF30 and Wdr33 were identified as the proteins involved in AU-hexamer recognition (17, 18). These findings are intriguing in light of evidence that the H1N1 human influenza virus protein NS1A targets CPSF30 to obstruct cellular mRNA processing (19–21), suggesting that the link between NS1A and cellular mRNA processing is RNA recognition by CPSF30.Open in a separate windowFig. 2.Cartoon of pre-mRNA processing, with possible roles of CPSF30 highlighted in blue.Given these cell-based results that CPSF30 is involved in recognition of the AU-hexamer of pre-mRNA along with our emerging understanding that CCCH-type ZFs are a general ZF motif involved in AU-rich RNA sequence recognition, we sought to determine whether CPFS30 directly recognizes the pre-mRNA AU-hexamer sequence via its CCCH domains by isolating CPSF30 and examining its RNA binding properties at the molecular level. CPSF30 contains five CCCH domains, and our hypothesis was that CPSF30 would bind five zinc ions at these CCCH domains and selectively recognize the AU-rich hexamer of pre-mRNA. To our surprise, CPSF30 was a reddish-colored protein upon isolation and purification, which suggested the presence of an iron cofactor. Here, we report that CPSF30 contains a 2Fe–2S site, with a CCCH ligand set, in addition to zinc. We also report that CPSF30 selectively recognizes the polyadenylation hexamer (AAUAAA) of pre-mRNA in a cooperative and metal-dependent manner. These findings are discussed in the context of CCCH “zinc” domains, iron, and recognition of AU-rich RNA sequences.     

From the Cover: 1.36 million years of Mediterranean forest refugium dynamics in response to glacial–interglacial cycle strength

The sediment record from Lake Ohrid (Southwestern Balkans) represents the longest continuous lake archive in Europe, extending back to 1.36 Ma. We reconstruct the vegetation history based on pollen analysis of the DEEP core to reveal changes in vegetation cover and forest diversity during glacial–interglacial (G–IG) cycles and early basin development. The earliest lake phase saw a significantly different composition rich in relict tree taxa and few herbs. Subsequent establishment of a permanent steppic herb association around 1.2 Ma implies a threshold response to changes in moisture availability and temperature and gradual adjustment of the basin morphology. A change in the character of G–IG cycles during the Early–Middle Pleistocene Transition is reflected in the record by reorganization of the vegetation from obliquity- to eccentricity-paced cycles. Based on a quantitative analysis of tree taxa richness, the first large-scale decline in tree diversity occurred around 0.94 Ma. Subsequent variations in tree richness were largely driven by the amplitude and duration of G–IG cycles. Significant tree richness declines occurred in periods with abundant dry herb associations, pointing to aridity affecting tree population survival. Assessment of long-term legacy effects between global climate and regional vegetation change reveals a significant influence of cool interglacial conditions on subsequent glacial vegetation composition and diversity. This effect is contrary to observations at high latitudes, where glacial intensity is known to control subsequent interglacial vegetation, and the evidence demonstrates that the Lake Ohrid catchment functioned as a refugium for both thermophilous and temperate tree species.

Identification and protection of past forest refugia, supporting a relict population, has gained interest in light of projected forest responses to anthropogenic climate change (1–4). Understanding the past and present composition of Mediterranean forest refugia is central to the study of long-term survival of tree taxa and the systematic relation between forest dynamics and climate (5). The Quaternary vegetation history of Europe, studied for over a century, is characterized by successive loss of tree species (6–8). Species loss was originally explained by the repeated migration across east–west oriented mountain chains during glacial–interglacial (G–IG) cycles (9). Later views gave more importance to the survival of tree populations during warm and arid stages in southern refugia (10, 11). Tree survival likely depends on persistence of suitable climate and tolerable levels of climate variability, as well as niche differentiation and population size at the refugium (12), although the precise relation between regional extinctions, climate variability, and local edaphic factors is not well known (13). Mediterranean mountain regions are considered to serve as forest refugia over multiple glacial cycles and frequently coincide with present-day biodiversity hotspots (14). Across the Mediterranean, increases in aridity and fire occurrence have impacted past vegetation communities (15–18). Comprehensive review of available Quaternary Mediterranean records indicates that Early (2.58 to 0.77 Ma) and Middle Pleistocene (0.77 to 0.129 Ma) tree diversity was higher compared to the present (13, 19–21). Particularly drought intolerant, thermophilic taxa were more abundant and diverse (8) but with strong spatial and temporal variations in tree diversity across the region. Long-term relationships between refugia function and environmental change over multiple G–IG cycles are hard to quantify due to the rarity of long, uninterrupted records.The Early–Middle Pleistocene Transition (EMPT), between 1.4 and 0.4 Ma (22), is of particular importance for understanding the relation between past climate change, vegetation dynamics, and biodiversity in the Mediterranean region. The EMPT is characterized by a gradual transition of G–IG cycle duration from obliquity (41 thousand years; kyr) to eccentricity (100 kyr) scale with increasing amplitude of each G–IG cycle (e.g., refs. 23, 24). The EMPT was accompanied by long-term cooling of the deep and surface ocean and was likely caused by atmospheric CO₂ decline and ice-sheet feedbacks (25–30). In Europe, the EMPT is associated with pronounced vegetation changes and local extinction and isolation of small tree populations (31).Here, we document vegetation history of the last 1.36 Ma in the Lake Ohrid (LO) catchment, located at the Albanian/North Macedonian border at 693 m above sea level (m asl, Fig. 1), the longest continuous sedimentary lake record in Europe (32, 33). The chronology of the DEEP core (International Continental Scientific Drilling Program site 5045-1; 41°02’57’’ N, 20°42’54’’ E, Fig. 1) is based on tuning of biogeochemical proxy data to orbital parameters with independent tephrostratigraphic and paleomagnetic age control (32, 33). The Balkan Peninsula has long been considered an important glacial forest refugium for presently widespread taxa such as Abies, Picea, Carpinus, Corylus, Fagus Ostrya, Quercus, Tilia, and Ulmus (7, 34–36). More than 60% of the Balkans is currently located >1,000 m asl (36), providing steep latitudinal and elevational gradients to support refugia under both cold and warm conditions. Today, the LO catchment is dominated by (semi) deciduous oak (Quercus) and hornbeam (Carpinus/Ostrya) forests. Above 1,250 m elevation, mixed mesophyllous forest with montane elements occurs (Fagus and at higher elevations Abies), which above 1,800 m elevation develops into subalpine grassland with Juniperus shrubs (see ref. 37 for site details). Isolated populations of Pinus peuce and Pinus nigra currently grow in the area (37–40).Open in a separate windowFig. 1.(A) Location of LO and TP on the Balkan Peninsula. (B) Local setting around LO, bathymetry (81), and DEEP coring site (adapted from ref. 32).Previous analysis of pollen composition of the last 500 kyr at the DEEP site revealed that the LO has been an important refugium. Arboreal pollen (AP) is deposited continuously and changes in abundance on multimillennial timescales in association with G–IG cycles, whereas millennial-scale variability is tightly coupled to Mediterranean sea-surface temperature variations (37, 41–45). Subsequent studies confirm the refugial character of the site recording Early Pleistocene (1.365 to 1.165 Ma) high relict tree diversity and abundance—and significant hydrological changes, including an increase in lake size and depth (38). Here, we present a continuous palynological record from LO with millennial resolution (∼2 kyr) back to 1.36 Ma to assess the systematic relationships between tree pollen abundance, forest diversity, and G–IG climate variability.Our objective is as follows: 1) infer the impact of past climate variability on local vegetation across the EMPT, 2) estimate tree species diversity in the catchment, and 3) examine how the amplitude and duration of preceding G–IG intervals affected the vegetation development and plant species diversity in this refugial area.     

Tooth wear and dentoalveolar remodeling are key factors of morphological variation in the Dmanisi mandibles

The Plio-Pleistocene hominin sample from Dmanisi (Georgia), dated to 1.77 million years ago, is unique in offering detailed insights into patterns of morphological variation within a paleodeme of early Homo. Cranial and dentoalveolar morphologies exhibit a high degree of diversity, but the causes of variation are still relatively unexplored. Here we show that wear-related dentoalveolar remodeling is one of the principal mechanisms causing mandibular shape variation in fossil Homo and in modern human hunter–gatherer populations. We identify a consistent pattern of mandibular morphological alteration, suggesting that dental wear and compensatory remodeling mechanisms remained fairly constant throughout the evolution of the genus Homo. With increasing occlusal and interproximal tooth wear, the teeth continue to erupt, the posterior dentition tends to drift in a mesial direction, and the front teeth become more upright. The resulting changes in dentognathic size and shape are substantial and need to be taken into account in comparative taxonomic analyses of isolated hominin mandibles. Our data further show that excessive tooth wear eventually leads to a breakdown of the normal remodeling mechanisms, resulting in dentognathic pathologies, tooth loss, and loss of masticatory function. Complete breakdown of dentognathic homeostasis, however, is unlikely to have limited the life span of early Homo because this effect was likely mediated by the preparation of soft foods.Although patterns of dental micro- and macrowear and wear-related pathologies are amply documented in the hominin fossil record (1–5), processes of in vivo dentoalveolar remodeling (3, 6, 7) and their potential influence on dentognathic morphology are only beginning to be studied in fossil hominins (8). In modern human hunter–gatherer populations, remodeling of dentoalveolar hard tissue is triggered mainly by dental wear, aging, pathologies, and trauma. Wear-related remodeling can be understood as a mechanism of in vivo modification that maintains masticatory function (3, 7, 9–14). Three main processes are typically identified (Fig. 1):

• i)Wear-induced reduction of dental crown height leads to alterations in masticatory biomechanics. This triggers alveolar bone remodeling, yielding dislocation of dental structures and eventual continuous eruption of all teeth. As an effect, occlusal contact between upper and lower teeth is maintained (15–17), and the position and orientation of the occlusal plane relative to the temporomandibular joints (TMJs) is held approximately constant, thus preventing the “wear-out” (18) of the TMJs.
• ii)The reduction of mesiodistal crown dimensions through interproximal dental wear triggers alveolar bone remodeling in the mesiodistal direction. This leads to mesial drift of the postcanine dentition and shortening of the dental arcade (19).
• iii)In the anterior dentition, remodeling induced by interproximal wear results in increased lingual tipping; that is, teeth become more upright relative to the alveolar plane (7).

As an effect of mesial drift, interproximal contacts between adjacent postcanine teeth are preserved. Similarly, as an effect of lingual tipping, interproximal contacts between incisors are preserved. During an individual’s lifetime, the combined, accumulated effects of continuous eruption, mesial drift, and lingual tipping tend to result in an edge-to-edge bite in the front dentition (7) and in significant changes in mandibular morphology (13, 20).Open in a separate windowFig. 1.Mechanisms of in vivo dentoalveolar remodeling. Continuous eruption (CE) is tracked by the distance between the mandibular canal and root apices. Mesial drift (MD) is tracked by the length of the posterior dental arcade from M2 to P3. Lingual tipping (LT) is tracked by the angle of inclination of the anterior teeth (incisors and canines) relative to the mandibular canal.The site of Dmanisi, Georgia, has yielded a crucial sample of early Pleistocene hominin fossils along with a rich vertebrate fauna and mode I (Oldowan) lithic implements (21, 22). Site occupation began shortly after 1.85 Ma (million years ago) (23); the fossils are dated to 1.77 Ma (21, 23–26). The Dmanisi hominin sample comprises five individuals that document dentognathic development from adolescence to old age, thus providing a unique opportunity to study wear-related dentognathic variation in a sample that comes from a single point in space and geological time (26–29). Various hypotheses have been proposed to explain the high degree of variation seen in the Dmanisi mandibles, ranging from intrataxon sexual dimorphism to intertaxon variation (25, 27–31). Here we focus on in vivo dentognathic remodeling as a potential mechanism contributing to the remarkable dentognathic variation in Dmanisi and differentiate normal remodeling processes from pathologic alterations. Using direct observations and data derived from computed tomography (CT) and scanning electron microscopy (SEM), we quantify dental wear (DW), continuous eruption (CE), mesial drift (MD), and lingual tipping (LT) in the following groups: the Dmanisi mandibles [specimens D2735, D211, D2600, D3900 (Fig. 2 and Fig. S1) and KNM-WT 15000 (grouped as early Pleistocene Homo)]; the middle Pleistocene mandibles from Tighenif (n = 3) and Atapuerca Sima de los Huesos (SH) (n = 4); and mandibles of modern hunter–gatherer populations from Australia (n = 26) and Greenland (n = 15) (Table S1), which all show substantial intragroup variation in dental wear. Measurement protocols are specified in Materials and Methods and in Table S2; abbreviations are listed in Table S3. It is well known that dental wear rates (amount of wear per year of life) vary substantially between populations as an effect of differences in food properties (e.g., abrasiveness) and paramasticatory tooth use (32, 33). To permit comparisons between different populations, remodeling rates are thus calculated per wear stage.Open in a separate windowFig. 2.Dmanisi mandibles. Right lateral and occlusal views are photographs taken from original specimens; left lateral views are CT-based 3D reconstructions highlighting internal structures. The canines and incisors of D2735 were found in isolation and digitally reinserted into their sockets. Arrow indicates toothpick lesion.     

Groundwater sapping as the cause of irreversible desertification of Hunshandake Sandy Lands,Inner Mongolia,northern China

In the middle-to-late Holocene, Earth’s monsoonal regions experienced catastrophic precipitation decreases that produced green to desert state shifts. Resulting hydrologic regime change negatively impacted water availability and Neolithic cultures. Whereas mid-Holocene drying is commonly attributed to slow insolation reduction and subsequent nonlinear vegetation–atmosphere feedbacks that produce threshold conditions, evidence of trigger events initiating state switching has remained elusive. Here we document a threshold event ca. 4,200 years ago in the Hunshandake Sandy Lands of Inner Mongolia, northern China, associated with groundwater capture by the Xilamulun River. This process initiated a sudden and irreversible region-wide hydrologic event that exacerbated the desertification of the Hunshandake, resulting in post-Humid Period mass migration of northern China’s Neolithic cultures. The Hunshandake remains arid and is unlikely, even with massive rehabilitation efforts, to revert back to green conditions.Earth’s climate is subject to abrupt, severe, and widespread change, with nonlinear vegetation–atmosphere feedbacks that produced extensive and catastrophic ecosystem shifts and subsequent cultural disruption and dispersion during the Holocene (1–7). In the early and middle Holocene, northern China’s eastern deserts, including much of the currently sparsely vegetated and semistabilized Hunshandake (Figs. 1 and ​and2),2), were covered by forests (8), reflecting significantly wetter climate associated with intensification of monsoon precipitation by up to 50% (6).Open in a separate windowFig. 1.Geographical location of the Hunshandake Sandy Lands (A) and its area (encircled by red line in B). The black rectangle in B marks the location of the enlarged maps C and D on the Right, and the green rectangle shows the location of Fig. 2. Map C shows the localities of water samples, and map D shows the localities of sections with stratigraphy presented in Fig. 3. The sand–paleosol section P (Fig. 3) is on the southern margin, and the site Bayanchagan marks the coring site of ref. 8. Rivers with headwaters in the Hunshandake likely formed by groundwater sapping are marked in blue. Drainages to the southwest and west are currently undergoing groundwater sapping, with substantial spring-driven flow found at the current river base level.Open in a separate windowFig. 2.(Left) Holocene lakes and channels in the Hunshandake and lake extent at selected epochs. Upper, middle, and lower lakes are indicated by points A, B, and C, respectively. Xilamulun River (point D) drains to the east. Groundwater-sapping headcuts at the upper reaches of incised canyons (point E) suggest a mid-Holocene interval of easterly surface flow, followed by groundwater drainage beginning at the ca. 4.2 ka event. Northern and central channels at point E are currently abandoned, and groundwater sapping has migrated to the southerly of the three channels shown. (Right) Cross-sections of the predrainage shift, northerly drainage into Dali Lake (Localities shown on the Left), showing the increase in widths of channels downstream (Vertical exaggeration ∼30:1).Monsoonal weakening, in response to middle-to-late Holocene insolation decrease, reduced precipitation, leading to a green/sandy shift and desertification across Inner Mongolia between ca. 5,000 and 3,000 y (years) ago (6). However, variations in the timing of this transition (9, 10) suggest local/regional thresholds or possibly environmental tipping by stochastic fluctuations. The impacts of this wet-to-dry shift in the Hunshandake, expressed as variations in surface and subsurface hydrology coincident with the termination of the formation of thick and spatially extensive paleosols, and the impacts of a ca. 4.2 ka (1 ka = 1,000 years) mid-Holocene desiccation of the Hunshandake on the development of early Chinese culture remain poorly understood and controversial (6, 11). Here we report for the first time to our knowledge on variability in a large early-to-middle Holocene freshwater lake system in China’s Hunshandake Sandy Lands and associated vegetation change, which demonstrates a model of abrupt green/desert switching. We document a possible hydrologic trigger event for this switching and discuss associated vegetation and hydrologic disruptions that significantly impacted human activities in the region.     

From the Cover: Carbon isotopes characterize rapid changes in atmospheric carbon dioxide during the last deglaciation

An understanding of the mechanisms that control CO₂ change during glacial–interglacial cycles remains elusive. Here we help to constrain changing sources with a high-precision, high-resolution deglacial record of the stable isotopic composition of carbon in CO₂ (δ¹³C-CO₂) in air extracted from ice samples from Taylor Glacier, Antarctica. During the initial rise in atmospheric CO₂ from 17.6 to 15.5 ka, these data demarcate a decrease in δ¹³C-CO₂, likely due to a weakened oceanic biological pump. From 15.5 to 11.5 ka, the continued atmospheric CO₂ rise of 40 ppm is associated with small changes in δ¹³C-CO₂, consistent with a nearly equal contribution from a further weakening of the biological pump and rising ocean temperature. These two trends, related to marine sources, are punctuated at 16.3 and 12.9 ka with abrupt, century-scale perturbations in δ¹³C-CO₂ that suggest rapid oxidation of organic land carbon or enhanced air–sea gas exchange in the Southern Ocean. Additional century-scale increases in atmospheric CO₂ coincident with increases in atmospheric CH₄ and Northern Hemisphere temperature at the onset of the Bølling (14.6–14.3 ka) and Holocene (11.6–11.4 ka) intervals are associated with small changes in δ¹³C-CO₂, suggesting a combination of sources that included rising surface ocean temperature.Over thirty years ago ice cores provided the first clear evidence that atmospheric CO₂ increased by about 75 ppm as Earth transitioned from a glacial to an interglacial state (1, 2). After decades of research, the underlying mechanisms that drive glacial–interglacial CO₂ cycles are still unclear. A tentative consensus has formed that the deglaciation is characterized by a net transfer of carbon from the ocean to the atmosphere and terrestrial biosphere, through a combination of changes in ocean temperature, nutrient utilization, circulation, and alkalinity. Partitioning these changes in terms of magnitude and timing is challenging. Estimates of the glacial–interglacial carbon cycle budget are highly uncertain, ranging from 20–30 ppm for the effect of rising ocean temperature, 5–55 ppm for ocean circulation changes, and 5–30 ppm for decreasing iron fertilization (3, 4), with feedbacks from CaCO₃ compensation accounting for up to 30 ppm (5, 6).A precise history of the stable isotopic composition of atmospheric carbon dioxide (δ¹³C-CO₂) can constrain key processes controlling atmospheric CO₂ (7, 8). A low-resolution record from the Taylor Dome ice core (9) identified a decrease in δ¹³C-CO₂ at the onset of the deglacial CO₂ rise that was followed by increases in both CO₂ and δ¹³C-CO₂ (Fig. 1). A higher-resolution record from the European Project for Ice Coring in Antarctica Dome C (EDC) ice core (10) provided additional support for the rapid δ¹³C-CO₂ decrease associated with the initial CO₂ rise, and box modeling indicated that this decrease was consistent with changes in marine productivity. The record also included other rapid changes in δ¹³C-CO₂, albeit at low precision, supporting large variations of organic carbon fluxes, notably a sharp increase in δ¹³C-CO₂ during the Bølling–Allerød (BA) interval attributed to carbon uptake by the terrestrial biosphere. A combined record including higher-precision EDC and Talos Dome data (11) documented a δ¹³C-CO₂ decrease beginning near 17.5 ka. This shift in δ¹³C-CO₂ was interpreted to indicate that some process in the Southern Ocean (SO), possibly changes in upwelling, drove the initial CO₂ rise. This previous work did not resolve high-frequency variability in the δ¹³C-CO₂ records that may be essential for discerning mechanisms of change.Open in a separate windowFig. 1.Carbon isotope records during the last deglaciation. Taylor Glacier δ¹³C-CO₂ data from this study (red). Previous work from Taylor Dome (gray open circles) (9), Grenoble EDC data (open green squares) (10), Bern EDC data (orange circles) (11, 45), sublimation measurements from EDC (blue triangles), and Talos Dome (purple squares) with an estimate of the 1-sigma uncertainty from a compilation of previous ice core δ¹³C-CO₂ data (11).Here we use an analytical method (12) that employs dual-inlet isotope ratio mass-spectrometry to obtain precision approaching that of modern atmospheric measurements [∼0.02‰ 1-sigma pooled SD based on replicate analysis compared with ∼0.05–0.11‰ for previous studies (9–11)]. We extracted atmospheric gases from large (400–500 g) samples taken from surface outcrops of ancient ice at Taylor Glacier, Antarctica, at an average temporal resolution of 165 y between 20 and 10 ka, and subcentury resolution during rapid change events. This resolution allows us to delineate isotopic fingerprints of rapid shifts in CO₂ that were previously impossible to resolve. Our study complements recent precise observations of CO₂ concentration variations during the last deglaciation, which revealed abrupt centennial-scale changes (13) (Fig. 2).Open in a separate windowFig. 2.Carbon cycle changes of the last deglaciation. WAIS Divide continuous CH₄ (green) (14) and discrete CO₂ (blue) (13) concentration data plotted with Taylor Glacier CO₂ and δ¹³C-CO₂ data (this study) (red markers, black line is a smoothing spline), the five-point running Keeling intercept with shading indicating the R² for each time interval. Blue bars indicate intervals of rapid CO₂ rise identified in the WAIS Divide ice core (13).During the initial 35-ppm CO₂ rise from 17.6 to 15.5 ka, we find a 0.3‰ decrease in δ¹³C-CO₂ that is interrupted by a sharp minimum coincident with rapid increases in CO₂ and CH₄ around 16.3 ka (13, 14) (Fig. 2). The 16.3-ka feature in the CO₂ and CH₄ concentration records, which corresponds to a 0.1‰ negative excursion in δ¹³C-CO₂, has been plausibly tied to the timing of Heinrich event 1 (13, 14) and signals a mode switch in the deglacial CO₂ rise. The subsequent slower rise in CO₂ from 15.5 to 14.8 ka is not accompanied by large changes in δ¹³C-CO₂. Across the Oldest Dryas to Bølling transition (14.6–14.3 ka) and coincident with a 10-ppm CO₂ increase and large CH₄ increase, we resolve a 0.08‰ increase in δ¹³C-CO₂ (Fig. 2). Rapid increases in CO₂ and CH₄ at the Younger Dryas (YD) to Preboreal transition (11.6–11.4 ka) are associated with minor variability δ¹³C-CO₂. On the other hand, the onset of the YD (12.8–12.5 ka) is characterized by a small rise in CO₂ associated with a 0.15‰ decrease in δ¹³C-CO₂ that appears tightly coupled to the timing of the large CH₄ decrease. The recovery from this excursion is characterized by increasing CO₂ and δ¹³C-CO₂. Broadly, our data confirm the results of Schmitt et al. (11) (Fig. 1). However, some of the large swings in δ¹³C-CO₂ indicated by the earlier EDC record (10), may be inaccurate and require reexamination.     

Evolution of a genetic polymorphism with climate change in a Mediterranean landscape

Many species show changes in distribution and phenotypic trait variation in response to climatic warming. Evidence of genetically based trait responses to climate change is, however, less common. Here, we detected evolutionary variation in the landscape-scale distribution of a genetically based chemical polymorphism in Mediterranean wild thyme (Thymus vulgaris) in association with modified extreme winter freezing events. By comparing current data on morph distribution with that observed in the early 1970s, we detected a significant increase in the proportion of morphs that are sensitive to winter freezing. This increase in frequency was observed in 17 of the 24 populations in which, since the 1970s, annual extreme winter freezing temperatures have risen above the thresholds that cause mortality of freezing-sensitive morphs. Our results provide an original example of rapid ongoing evolutionary change associated with relaxed selection (less extreme freezing events) on a local landscape scale. In species whose distribution and genetic variability are shaped by strong selection gradients, there may be little time lag associated with their ecological and evolutionary response to long-term environmental change.Ongoing changes in regional climates are pushing many species to shift their distribution toward higher latitudes and altitudes (1–7). Such changes in species distribution, with an expansion in previously hostile areas and contraction in areas becoming less favorable, can occur rapidly both in plants and animals (2, 3, 5, 6). As a result, major changes in community composition due to differential migration rates may occur (8). Indeed, habitat fragmentation may prevent many species from showing such a distributional response to climate change. As a result, only those species that can respond by phenotypic plasticity or genetically based local adaption will persist (9). In animal and plant species, phenotypic plasticity of phenological traits can allow individuals to adjust to climate change (1, 10, 11). In addition to changes in distribution and plasticity, an evolutionary response to climate change may occur if species evolve a genetically based adaptation to climate change (12, 13). It is important to distinguish this genetic response from a plastic response of individuals if we are to fully understand the evolutionary potential of species to evolve with climate change (14). Adaptive trait variation in relation to climate change has been shown in the classic study of Drosophila (15, 16) and in experimental and natural populations of a small number of plant species (17–20). However, in some cases, the evolutionary response to climate change may be slow due to genetic constraints (21) causing a time lag between the environmental change and an observed evolutionary response. Understanding how species track climate change by genetically based adaptive trait variation and which traits facilitate the evolution of such adaption is important; such issues determine which species may persist locally and which may shift their distribution (22, 23).In this study, we test for an evolutionary response of a genetic polymorphism in essential oil composition in Mediterranean wild thyme, Thymus vulgaris, to reduced selection associated with a warming of extreme winter freezing events over a sharp spatial climatic gradient. Our study was done in and around the Saint Martin-de-Londres basin (43°48′N, 03°46′E), which covers an area of ∼80 km² and whose southern limits are ∼20 km north of Montpellier in the Mediterranean climate region of southern France. The center of the basin (lowest altitude, 145 m) is surrounded by calcareous hills, ranging from 300 to 658 m. The study area has a Mediterranean climate with summer drought but also severe winter freezing temperatures within the basin as a result of a dramatic temperature inversion (Fig. 1). In this area, there are six different chemotypes that are the expression of a genetically controlled polymorphism in T. vulgaris (24). Two phenolic chemotypes (carvacrol and thymol) are largely dominant on the slopes outside of the basin on stony soils above 250-m elevation and four nonphenolic chemotypes (linalool, thuyanol-4, α-terpineol, and geraniol) occur within the basin below 200-m elevation on deeper, more humid soils (25–27), where they experience the winter temperature inversion. There is thus a sharp cline in chemotype frequency over only 3–5 km that goes from 100% of either phenolic or nonphenolic chemotypes to 100% of the other form, with a narrow transitional zone (Fig. 2). In short, nonphenolic chemotypes show marked adaptation to habitats, which in the past have frequently experienced extreme freezing temperatures in early winter, whereas phenolic chemotypes are sensitive to intense early-winter freezing and occur in habitats where extreme summer drought can exclude nonphenolic chemotypes (28, 29).Open in a separate windowFig. 1.Coldest annual temperature from 1955 to 2010 at the weather station of Saint Martin-de-Londres (filled squares), which occurs in the zone dominated by freezing-tolerant nonphenolic chemotypes, and from 1970 to 2010 at the Centre d''Ecologie Fonctionnelle et Evolutive–Centre National de la Recherche Scientifique experimental gardens on the northern periphery of Montpellier (open circles), where natural thyme populations are dominated by freezing-sensitive phenolic chemotypes.Open in a separate windowFig. 2.Spatial distribution and chemical composition of sampled thyme populations in the early 1970s and the 36 populations that were resampled in the present study along six transects. The chemical composition of resampled populations indicated on the map is that which was observed in the initial study. Black circles, phenolic populations; open circles, nonphenolic populations; gray circles, mixed populations. On each transect, the six populations are connected by a dashed line and are represented by a slightly larger circle than populations that were not resampled in the present study.In the part of the study area where nonphenolic chemotypes dominated populations in the early 1970s, extreme winter temperatures have become less severe (Fig. 1) with a significant increase in temperature of extreme freezing events (r = 0.36, n = 56, P < 0.01). In the 20 y before and during the initial study, winter temperatures fell below levels (−15 °C in December) that would exclude phenolic chemotypes from sites dominated by nonphenolic chemotypes (29) on five occasions. In the 16 y following the initial study, three such events were recorded. In the last 20 y, no such extreme events have been recorded in the zone dominated by nonphenolic chemotypes. The last extreme freezing event that could impact the composition of thyme populations occurred 25 y ago.Here, we test the hypothesis that phenolic chemotypes (thymol and carvacrol) now occur in sites where they were previously absent or have increased their frequency in transitional sites due to a relaxation of selection normally associated with extreme early-winter freezing temperatures. To do so, we compared the chemotype composition of populations observed in the early 1970s (26) to that in 2009–2010 for 36 populations sampled along six transects. Each transect is <10 km long, each containing six populations, with two “phenolic,” “mixed,” and “nonphenolic” populations (Fig. 2). To provide an indication of whether population-level changes are due to within-population adaptation or migration among populations, we also examine whether any increases in the abundance of phenolic chemotypes are primarily in nonphenolic or mixed populations that are spatially the closest to preexisting phenolic populations.